Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform - Special Publication 340 (Geological Society Special Publication)

Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

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

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

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: SOSSON , M., KAYMAKCI , N., STEPHENSON , R. A., BERGERAT , F. & STAROSTENKO , V. (eds) 2010. Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340. YEGOROVA , T., BARANOVA , E. & OMELCHENKO , V. 2010. The crustal structure of the Black Sea from the reinterpretation of deep seismic sounding data acquired in the 1960s. In: SOSSON , M., KAYMAKCI , N., STEPHENSON , R. A., BERGERAT , F. & STAROSTENKO , V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 43–56.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 340

Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform

EDITED BY

M. SOSSON University of Nice-Sophia Antipolis, France

N. KAYMAKCI Middle East Technical University, Turkey

R. A. STEPHENSON University of Aberdeen, UK

F. BERGERAT CNRS, France

and V. STAROSTENKO National Academy of Sciences of Ukraine, Ukraine

2010 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2010. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/10/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-308-0 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]

Foreword and acknowledgements This volume belongs to a series of four regionallythemed Geological Society of London Special Publications presenting some of the results of the Middle East Basin Evolution Programme (MEBE). The other Special Publications in the series are focused on the South Caspian –Central Iran area (Brunet et al. 2009), the Zagros –East Arabian margin (Leturmy & Robin 2010) and the Levant (Homberg & Bachmann 2010). MEBE was a four-year research consortium (2003–2006), comprising a multidisciplinary study of the post-Palaeozoic tectonic evolution of the Middle East, spanning the Arabian –Peri-Arabian and Caucasian –Caspian areas and funded by several major oil companies and French research organizations (see below). It was led by E. Barrier (CNRS-Universite´ Pierre et Marie Curie, Paris, France) and M. Gaetani (University of Milan, Italy). More details can be found in the preface of Brunet et al. (2009). Workshops were a key element of MEBE for integrating and synthesizing the programme’s diverse, multidisciplinary results. This volume evolved from two such workshops: in Kiev in February 2006 by MEBE’s Black Sea Working Group, led by R. Stephenson (VU University Amsterdam, Netherlands; now University of Aberdeen, UK) and in Ankara in September 2006 by the Caucasus Working Group, led by M.

Sosson (CNRS-Universite´ Nice-Sophia Antipolis, Valbonne, France), with the workshops themselves respectively organized by V. Starostenko, Institute of Geophysics, National Academy of Sciences of Ukraine, Kiev) and N. Kaymakci (Middle East Technical University, Ankara, Turkey). The editors would like to thank and acknowledge the efforts of many colleagues who chose to publish their results in this volume and also to those who submitted papers that do not appear in the final volume. As always, such a book would not be possible without peer-review and the efforts of the reviewers who so willingly contributed their time and expertise are very much appreciated. These include (with apologies if anyone has been inadvertently missed): S. Adamia, Ph. Agard, A. Aldanmaz, M. Allen, J. Angelier, K. Atakan, E. Barrier, N. Bragin, M.-F. Brunet, M. Burianyk, R. Carbonell, J. Cartwright, M. Chiari, B. Davis, R. Ebert, S. Egan, C. Homberg, B. H. Jacobsen, R. Klemd, A. Okay, B. Ma¨chtle, M. Bielik, M. Grad, L. Matenco, T. McCann, J. Mosar, A. Nikishin, G. Nolet, L. Popescu, A. Robertson, Y. Rolland, R. Sachsenhofer, A. Saintot, A. Tremblay, S. Vandycke, J. Woodside, E. Yig˘itbas¸ and A. Zanchi. The following companies and organizations are thanked for funding the MEBE Programme.

References BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) 2009. South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312. HOMBERG , C. & BACHMANN , M. (eds) 2010. Evolution of the Levant Margin and Western Arabia Platform

since the Mesozoic. Geological Society, London, Special Publications, 341. LETURMY , P. & ROBIN , C. (eds) 2010. Tectonic and Stratigraphic Evolution of Zagros and Makran during the Mesozoic –Cenozoic. Geological Society, London, Special Publications, 330.

Contents Foreword and acknowledgements SOSSON , M., KAYMAKCI , N., STEPHENSON , R., BERGERAT , F. & STAROSTENKO , V. Sedimentary basin tectonics from the Black Sea and Caucasus to the Arabian Platform: introduction

vii 1

STEPHENSON , R. & SCHELLART , W. P. The Black Sea back-arc basin: insights to its origin from geodynamic models of modern analogues

11

YEGOROVA , T. & GOBARENKO , V. Structure of the Earth’s crust and upper mantle of the West- and East-Black Sea Basins revealed from geophysical data and its tectonic implications

23

YEGOROVA , T., BARANOVA , E. & OMELCHENKO , V. The crustal structure of the Black Sea from the reinterpretation of deep seismic sounding data acquired in the 1960s

43

STAROSTENKO , V. I., RUSAKOV , O. M., SHNYUKOV , E. F., KOBOLEV , V. P. & KUTAS , R. I. Methane in the northern Black Sea: characterization of its geomorphological and geological environments

57

BERGERAT , F., VANGELOV , D. & DIMOV , D. Brittle deformation, palaeostress field reconstruction and tectonic evolution of the Eastern Balkanides (Bulgaria) during Mesozoic and Cenozoic times

77

HIPPOLYTE , J.-C., MU¨ LLER , C., KAYMAKCI , N. & SANGU , E. Dating of the Black Sea Basin: new nannoplankton ages from its inverted margin in the Central Pontides (Turkey)

113

KHRIACHTCHEVSKAIA , O., STOVBA , S. & STEPHENSON , R. Cretaceous –Neogene tectonic evolution of the northern margin of the Black Sea from seismic reflection data and tectonic subsidence analysis

137

KONERDING , C., DINU , C. & WONG , H. K. Seismic sequence stratigraphy, structure and subsidence history of the Romanian Black Sea shelf

159

MC C ANN , T., CHALOT -PRAT , F. & SAINTOT , A. The Early Mesozoic evolution of the Western Greater Caucasus (Russia): Triassic –Jurassic sedimentary and magmatic history

181

ADAMIA , SH ., ALANIA , V., CHABUKIANI , A., CHICHUA , G., ENUKIDZE , O. & SADRADZE , N. Evolution of the Late Cenozoic basins of Georgia (SW Caucasus): a review

239

MOSAR , J., KANGARLI , T., BOCHUD , M., GLASMACHER , U. A., RAST , A., BRUNET , M.-F. & SOSSON , M. Cenozoic –Recent tectonics and uplift in the Greater Caucasus: a perspective from Azerbaijan

261

USTAO¨ MER , T. & ROBERTSON , A. H. F. Late Palaeozoic–Early Cenozoic tectonic development of the Eastern Pontides (Artvin area), Turkey: stages of closure of Tethys along the southern margin of Eurasia

281

SOSSON , M., ROLLAND , Y., MU¨ LLER , C., DANELIAN , T., MELKONYAN , R., KEKELIA , S., ADAMIA , S., BABAZADEH , V., KANGARLI , T., AVAGYAN , A., GALOYAN , G. & MOSAR , J. Subductions, obduction and collision in the Lesser Caucasus (Armenia, Azerbaijan, Georgia), new insights

329

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CONTENTS

ROLLAND , Y., GALOYAN , G., SOSSON , M., MELKONYAN , R. & AVAGYAN , A. The Armenian Ophiolite: insights for Jurassic back-arc formation, Lower Cretaceous hot spot magmatism and Upper Cretaceous obduction over the South Armenian Block

353

DANELIAN , T., ASATRYAN , G., SAHAKYAN , L., GALOYAN , GH ., SOSSON , M. & AVAGYAN , A. New and revised radiolarian biochronology for the sedimentary cover of ophiolites in the Lesser Caucasus (Armenia)

383

AVAGYAN , A., SOSSON , M., KARAKHANIAN , A., PHILIP , H., REBAI , S., ROLLAND , Y., MELKONYAN , R. & DAVTYAN , V. Recent tectonic stress evolution in the Lesser Caucasus and adjacent regions

393

KAYMAKCI , N., INCEO¨ Z , M., ERTEPINAR , P. & KO c¸, A. Late Cretaceous to Recent kinematics of SE Anatolia (Turkey)

409

KUSCU , I., KUSCU , G. G., TOSDAL , R. M., ULRICH , T. D. & FRIEDMAN , R. Magmatism in the southeastern Anatolian orogenic belt: transition from arc to post-collisional setting in an evolving orogen

437

OBERHA¨ NSLI , R., CANDAN , O., BOUSQUET , R., RIMMELE , G., OKAY , A. & GOFF , J. Alpine high pressure evolution of the eastern Bitlis complex, SE Turkey

461

¨ ZACAR , A. A., ZANDT , G., GILBERT , H. & BECK , S. L. Seismic images of crustal variations O beneath the East Anatolian Plateau (Turkey) from teleseismic receiver functions

485

Index

497

Sedimentary basin tectonics from the Black Sea and Caucasus to the Arabian Platform: introduction MARC SOSSON1*, NURETDIN KAYMAKCI2, RANDELL STEPHENSON3, FRANC¸OISE BERGERAT4 & VITALY STAROSTENKO5 1

University of Nice-Sophia Antipolis, CNRS, OCA, UMR Ge´oazur, 250 rue A. Einstein, 06560 Valbonne, France 2

Middle East Technical University, Department of Geological Engineering, 06531-Ankara, Turkey

3

University of Aberdeen, School of Geosciences, Meston Building, King’s College, Aberdeen AB24 3UE, UK

4

CNRS, UMR 7193, ISTeP, Case 117, 4 place Jussieu, F-75252 Paris cedex 05, France 5

National Academy of Sciences of Ukraine, Institute of Geophysics, Kiev, Ukraine *Corresponding author (e-mail: [email protected]) Abstract: The Palaeozoic to recent evolution of the Tethys system gave way to the largest mountain chain of the world extending from the Atlantic to Pacific oceans – the Alpine– Himalayan Mountain chain, which is still developing as a result of collision and northwards convergence of continental blocks including Apulia in the west, the Afro-Arabian Plate in the middle and the Indian Plate in the east. This Special Publication addresses the main problems of the middle part of this system incorporating the Balkans, Black Sea and Greater Caucasus in the north and the Afro-Arabian Plate in the south. Since the Early Mesozoic a number of small to large scale oceanic basins opened and closed as the intervening continental fragments drifted northwards and diachronously collided with and accreted to the southern margin of the Eurasian Plate. Despite the remarkable consequences of this, in terms of subduction, obduction and orogenic processes, little is known about the timing and palaeogeographic evolution of the region. This includes the amounts of shortening and interplay between synconvergent extension and compression, development of magmatic arc and arc-related basins and the timing and mechanism of their deformation. The chapters presented in this Special Publication present new information that help to fill some of the gaps of the puzzle.

The tectonic history of the sedimentary basins of the Black Sea–Caucasus and surrounding areas and the geodynamic processes governing the formation and deformation of these basins, as well as the development of related mountain belts, corresponded to the goals of two thematic working groups of the Middle East Basin Evolution (MEBE) Programme (Black Sea and Caucasus working groups led by R. Stephenson and M. Sosson respectively). The scientific results of these working groups, which form the contents of this Special Publication, along with some complementary studies that took place outside MEBE (e.g. Adamia et al. 2010; ¨ zacar et al. 2010; Ustao¨mer & Kuscu et al. 2010; O Robertson 2010), are briefly summarized and collectively discussed in terms of the key objectives of the MEBE programme in the Black Sea– Caucasus area.

Regional tectonic issues in the Black Sea – Caucasus – Arabian Platform corridor The Alpine–Himalayan belt evolved through a series of tectonic events related to the opening and closing of the Tethys Ocean, producing the largest mountain belt of the world, running from the Atlantic to the Pacific oceans. The basins associated with this belt contain invaluable information related to mountain-building processes and are also the locus of rich hydrocarbon accumulations; however, knowledge about the geological evolution of the region is limited compared to what is ultimately available. This has been mainly due to the difficulty and inaccessibility of cross-country studies. This Special Publication is dedicated in part to redressing this situation in the segment of the

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 1–10. DOI: 10.1144/SP340.1 0305-8719/10/$15.00 # The Geological Society of London 2010.

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M. SOSSON ET AL.

Alpine– Himalayan belt running from Bulgaria to Armenia and from Ukraine to the Arabian Platform (Figs 1 & 2). Considering the degree of ambiguity and number of conflicting models proposed for the region, many fundamental issues pertaining to the evolution of sedimentary basins and mountain belts remain to be (re)solved. It was the aim of the MEBE programme to elucidate some of these issues using modern methods and ideas. At the outset of the MEBE programme, project teams working in the Black Sea– Caucasus region identified four main themes for attention and elucidation, involving four broad periods of distinct tectonic evolution of this area. These, from oldest to youngest, were as follows. (1)

How pre-Cimmerian tectonics may have or may not have controlled or influenced the subsequent deformational history of the Black Sea –Caucasus area. There is considerable uncertainty regarding the geometry and evolution of the southern margin of Europe generally during the Palaeozoic and Early Mesozoic. It appears as though there was no Late Palaeozoic accretionary event (i.e. ‘Scythian Orogeny’), adding to the European continent, but that widespread extensional tectonics of older crustal basement prevailed throughout much of the Black Sea –Caucasus area during this time. Allied to this are issues such as the crustal affinity of the mid– Black Sea ridge and the nature of the basement of

(2)

(3)

the Crimean Orogeny and the relationship of these, for example, with the crust of the Transcaucasus area. The significance of Cimmerian ‘Orogeny’ tectonics throughout the Black Sea–Caucasus area, as expressed by the reportedly widespread occurrence of various Late Triassic –Jurassic unconformities. The ‘midCimmerian’ unconformity was thought to be the most profound. A key question that was asked was whether Cimmerian tectonic events were related to continental collision or rather a response to (minor?) plate geometry reorganizations and/or to subductionaccretion ‘anomalies’ broadly affecting the kinematics of convergence. Cretaceous extensional tectonics, being the main phase of development of the Black Sea basin, and its relationship with earlier extensional events in the Jurassic (and earlier) such as the Greater Caucasus (and, further afield, South Caspian) basin. The Black Sea is generally interpreted as a back-arc basin but, in fact, fundamental issues such as the presence or absence of a related magmatic arc and the orientation of the related subduction zone remain vague at best. There are obvious implications for the geometry of extension and rifting within the western and eastern Black Sea basins and the role of broader plate configuration and kinematics in controlling this.

Fig. 1. Shaded relief image obtained from SRTM data (www2.jpl.nasa.gov/srtm) and approximate locations of the papers presented in this volume.

INTRODUCTION

3

Fig. 2. Tectonic outline of eastern Mediterranean region (modified from Kaymakci et al. 2010).

(4)

Mid-Eocene and younger shortening that is widespread, affecting almost all of the Balkanides –Black Sea –Caucasus– eastern Anatolian area. Was the onset of this deformation – and discrete pulses of compression thereafter – almost synchronous throughout the area or is there clear evidence of diachroneity? Are observations that suggest essentially coeval active extension in the eastern Black Sea basin and its margins at this time evidence for a penultimate regime of extension in what was otherwise a broad zone of convergence throughout the Mesozoic and Cenozoic. Accordingly, it may provide important inferences on the geometry of indentation of convergence during the Palaeogene.

Results This Special Publication includes 20 multidisciplinary studies covering topics in structural geology/ tectonics, geophysics, geochemistry, palaeontology, petrography, sedimentology, stratigraphy, as well as subsidence and lithospheric modelling, reporting results obtained during MEBE and related projects in the circum Black Sea and peri-Arabian regions.

All are aimed at addressing the general issues highlighted above. These papers are presented in five sections. The first two sections deal with the crustal and lithospheric structure of the Black Sea basin and its margins, using new potential field and seismic reflection and refraction data and other geophysical observations, including the reinterpretation of ‘old’ seismic refraction data, linked with information obtained from old and new boreholes. Onshore geological studies are also represented. Many of the papers in these sections discuss their results at the scale of the lithospheric processes responsible for back-arc basin formation and inversion and how these help to understand the observed tectonic styles and timing of key events in the Black Sea region, involved in the origin and evolution of the eastern and western Black Sea basins. The important role of ‘Cimmerian’ tectonics in the region prior to the formation of the Black Sea itself, from the Late Triassic until the Early Cretaceous, also receives attention. The third section of the volume focuses on documenting the geological characteristics and geodynamic evolution of the sedimentary basins related to the Greater Caucasus, located in the internal part of the Alpine –Himalayan belt within the Eurasian plate. The fourth and fifth sections

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comprise newly obtained data acquired from fieldwork and modern analytical studies on collected field materials, including geochronological dating, petrology and geochemistry, as well as structural mapping and palaeostress determinations. A comprehensive compilation of the information from the pre-existing literature, as related to the tectono-stratigraphic and geodynamic evolution of the external part of the Caucasus area (Eastern Pontides, Lesser Caucasus and South–East Anatolia), is also found in this section.

Black Sea: crustal and lithosphere structure The Black Sea is generally considered to be a Cretaceous –Palaeogene back-arc basin and, in the first paper of the collection, Stephenson & Schellart point out that this means that its formation, behind a subduction zone, must be viewed in the context of subduction-scale geodynamic processes. They argue that observations and models of modern (recent and currently active) subduction zones formed as a result of asymmetric subduction rollback are compatible with the geometry and history of the Black Sea as a whole. Although such a model predicts a slightly earlier onset of development of extensional structures in the western compared to the eastern Black Sea sub-basin, with the youngest such activity being in the east and NE, it implies that the development of these would be essentially contemporaneous. Comparison of the regional structure of the Black Sea and its sub-basins to model predictions also implies that the proto-basement of the Black Sea probably comprised strong, cratonic, lithosphere. Yegorova & Gobarenko, in the second paper of the volume, take a look at this very issue of basement lithosphere structure and affinity from the point of view of a comprehensive analysis of available geological and geophysical information on the Black Sea region. Their key result is that there are observable differences between the lithosphere underlying the western and eastern sub-basins of the Black Sea. This is based in part on new local seismic tomography results that suggest that the lithosphere underlying the latter has a lower average P-wave velocity (and higher temperature?) than the former. Yegorova & Gobarenko consider the implications of this for the distribution of seismicity around the Black Sea and conclude that this may partly explain why earthquakes in the east mainly occur on the northern margin of the Black Sea while those to the west mainly occur on the southern margin. One of the key geophysical observables used in the analysis of Yegorova & Gobarenko, consisted of new crustal-upper mantle velocity models presented by Yegorova et al. based on old Soviet-era

Deep Seismic Sounding (DSS) data. The old DSS data were re-analysed on two north– south running profiles, one from the northern margin (Odessa Shelf) of the western Black Sea into its deep basin and one from the Azov Sea south to the central part of the eastern Black Sea sub-basin. It is the first and only time that any of the old Black Sea DSS data have been studied using modern computer-based ray-tracing methods. The results provide more robust estimates of the thickness of sedimentary and crustal layers and, importantly, the seismic velocities that characterize the latter, than almost all of those previously available. Within the limits of the accuracy of the data and models, the crustal structure and thickness of sediments in the eastern Black Sea sub-basin is the same as that of the western Black Sea sub-basin. Starostenko et al. give an overview of the observed occurrences of methane gas seeps in the northern Black Sea, particularly in shelf-slope transition zone, and how these are related to their geomorphological and geological environments, including submarine canyons, mud volcanoes, and gas hydrates. It is argued that the gas hydrate stability zone in the Black Sea lies at a depth of 600–650 m below the sea floor. The authors consider that the methane in the Black Sea is mostly of abiogenic origin and that deep-seated faults, which are also thought to control the locations of submarine canyons, play a role in transporting gas to the sea floor. It is assumed that the faults observed there are predominantly deep-seated faults, which by definition cut down to crystalline basement levels, related to important crustal-scale structures, although no data are presented to support this assertion.

Margins of the Black Sea: crustal affinity, basin formation and inversion and neotectonics; Balkanides, Romanian and Bulgarian shelves, Odessa Shelf and Azov Sea; central Pontides of Turkey Bergerat et al. present new brittle tectonics analysis and stress tensor reconstructions that add considerable knowledge about the Mesozoic and Cenozoic geodynamic evolution of the superimposed basin systems of the eastern Balkanides. Evidence of Jurassic–Early Cretaceous tectonic events are presented, especially in the Strandja Zone, in which NE– SW extension followed by ENE –WSW compression is recorded, possibly corresponding to Early Alpine phases of rifting and shortening, respectively. Late-Alpine tectonic phases correspond to the Late Cretaceous –Middle Eocene period dominated by: 1) the back-arc basin/ island-arc system of the East Balkan/Srednogorie zones, with the main direction of extension,

INTRODUCTION

Turonian–Early Campanian, being north –south to NNE –SSW; 2) the inversion of extensional structures within this zone after the Maastrichtian (Laramian phase); and, finally, 3) a major shortening phase that began in the latest Middle Eocene (Illyrian phase) with the development of large shearing and thrusting structures. Compression was directed during these two last phases NE –SW and north–south, respectively. The paper by Hippolyte et al. deals with the Western Black Sea Basin. These authors provide new, high precision palaeontological data related to the timing of the opening of the western Black Sea and the Eocene inversion of the Central Pontides. New nanoplankton ages reveal that subsidence and rifting started in the Late Barremian and accelerated during the Aptian. Rifting in the western Black Sea Basin lasted about 40 Ma (from the Late Barremian until the Coniacian). The syn-rift sequence terminates with shallow marine sands in the inner, inverted, Black Sea margin. The authors propose that the uppermost Albian to Turonian corresponds to a period of erosion or non-deposition and that the Coniacian –Santonian succession represents the results of rapid thermal post-rift subsidence on the basin margin. They infer that the collision of the Central Pontides and the Tauride – Anatolian Block began during the Early Eocene and that this led to compressional deformation and sedimentation in piggyback basins in the Central Pontides. Khriachtchevskaia et al. interpret recently acquired industry seismic reflection and newly correlated well data from the Odessa Shelf and the Azov Sea, on the northern margin of the Black Sea, to elucidate the timing and style of tectonic events shaping the architecture of the Black Sea. Structural and subsidence analyses demonstrate that active extension (rifting) began by Aptian –Albian times and ended in the Santonian in the Late Cretaceous, essentially consistent, given correlation uncertainties, with the observations of Hippolyte et al. derived from the opposite margin of the Black Sea. Discrete inversion events are documented to have begun in the late Middle Eocene, similar to what Bergerat et al. found in Bulgarian eastern Balkanides, and to have ended by the Middle Miocene. Konerding et al. focus on the Miocene and younger sequence stratigraphy of the Romanian sector of the Black Sea shelf. Older (Black Sea rifting) extensional structures inverted during Middle Eocene –Oligocene times, as documented by Bergerat et al. to the (south)west and Khriachtchevskaia et al. to the (north)east in part define the architecture of the Mio-Pleistocene successions with sediment input and, to a lesser extent sea-level changes, also playing a role according to the

5

authors. There is evidence of extensional tectonic activity from the latest Miocene (Pontian), after the cessation of inversion tectonics documented by Khriachtchevskaia et al.

Greater Caucasus: basin and tectonic evolution in Russia, Georgia and Azerbaijan McCann et al. analyse structural, sedimentological, petrological and geochemical data from numerous field localities in the Fore-Caucasus and Central and Western Greater Caucasus mountain belt. This is a high Alpine fold-and-thrust belt on the southern margin of the East European Platform and this study is aimed at elucidating the earlier Triassic –Jurassic Cimmerian tectonic and palaeogeographic evolution of the area. The authors document that a broad asymmetric basin, with associated emergent volcanic islands, dominated the area in the Jurassic and interpret this as the result of incipient back-arc rifting. Similar rifting episodes in the Pontides and South Caspian Sea areas suggest the redevelopment of a common Tethys subduction zone at this time to the south of a Late Triassic –Early Jurassic Eo-Cimmerian accretionary belt. The associated volcanic arc subsequently widened, taken to indicate a temporary shallowing of the northward subducting slab until the Late Jurassic, when sediment accommodation space was rapidly created to form the Crimea-Greater Caucasus sedimentary basin. An overview of the lithological and structural characteristics of basins associated with a later phase of development of the Greater Caucasus, specifically the Late Cenozoic collisional basins in front of the Greater Caucasus in Georgia, as well as the coeval magmatic events, is given by Adamia et al. on the basis of numerous studies carried out during past decades. With the help of field and well data, complemented by some interpreted seismic reflection lines, these authors present the detailed lithostratigraphy of each of the Upper Cenozoic formations of these basins, including the Rioni, Kartli, Kura and others. The foldand-thrust belts of the Greater and Lesser Caucasus ranges, separated by the Transcaucasian intermontane depression, formed during syn-collisional (Oligocene– Middle Miocene) and post-collisional (the Late Miocene– Quaternary) phases where, previously, back-arc basins had developed. Adamia et al. also describe the post-collisional, Late Miocene to Pleistocene magmatic activity of the central part of the area. The paper by Mosar et al. describes the orogenic processes that led to the present mountain belt of the Greater Caucasus, starting in early Cenozoic times, accelerating during the Plio-Pleistocene, and being still active as demonstrated by GPS and earthquake

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studies. The chain is a doubly verging fold-andthrust belt, with pro- and retro-wedges actively propagating into the foreland sedimentary basins, the Kura to the south and the Terek to the north, respectively. The mountain range can be subdivided into several zones with different uplift magnitudes and rates with significantly heterogeneous strain partitioning. The authors show that clear links exist between that geomorphology, seismicity and tectonics that can be observed in the Greater Caucasus.

Lesser Caucasus and Eastern Pontides: obduction and collision stages in Armenia, Azerbaijan, Georgia and Turkey Ustao¨mer & Robertson is the first of five papers that deal with the Eastern Pontides, in Turkey, and the Lesser Caucasus area, in Armenia, Azerbaijan and Georgia. They document new data from the Artvin region in the Eastern Pontides located near the SE margin of the Black Sea basin. They propose that a large continental block rifted away from the southern margin of Eurasia in a suprasubduction zone environment during the Early– Middle Jurassic and the intervening basin is filled with a thick sequence of terrigenous clastics, debris flow deposits and turbidites, and deep-sea deposits associated with radiolarites and volcanic rocks and mafic intrusions. The Middle Jurassic is dominated by subduction influenced basalts and volcaniclastics. They elegantly document that this basin was inverted during the Late –Middle Jurassic NeoCimmerian deformation that resulted in local uplift and erosion, this unconformity being later covered by an Upper Jurassic continental to shallow marine succession. They also provide information about the timing of onset of Eastern Pontides magmatic arc activity, resulting in a huge accumulation of magmatic and volcanogenic rocks, during the Late Cretaceous. They propose that the region was telescoped during the Middle Eocene continental collision and claim that the geological evolution of the Artvin area correlates with the Pontides further west and with the southern and northern Transcaucasus to the east. They also argue that there is no evidence of ‘Palaeotethyan’ ophiolites in the Eastern Pontides region. Sosson et al. present new data on the structure of the Lesser Caucasus and its geodynamic evolution. From SW to NE they recognize three main domains: 1) the autochthonous South Armenian Block (SAB), a Gondwana-derived terrane, 2) the ophiolitic Sevan–Akera (Se –Ak) suture zone and 3) the Eurasian plate. According to field data and geochemical analyses, they find evidence for two Middle Jurassic to Late Cretaceous subductions

zones, one related to Neotethys consumption beneath the Eurasian margin and an intra-oceanic one (in a supra-subduction zone context) responsible for the opening of a back-arc basin correlatable with the ophiolites of the Lesser Caucasus. New dating of nannofossils allow a Late Coniacian to Santonian age for obduction, responsible for the widely exposed ophiolitic nappe in front of the Se –Ak suture zone, to be specified. The collision of the SAB with Eurasia started during the Paleocene. The completion of oceanic lithosphere subduction beneath Eurasia and the onset of collision between the Arabian plate with the SAB to the south took place from the Middle Eocene to the Miocene, during which time there southward propagation of shortening featured by folding and thrusting all along the belt. Rolland et al. describe new geological, petrological, geochemical data and radiometric ages on the Se –Ak, Stepanavan and Vedi ophiolitic massifs in Armenia. The authors conclude that a single large ophiolite unit has been obducted onto the SAB. Ophiolite rock assemblages suggest a slow-spreading rate in Early–Middle Jurassic times. The age of the ophiolite is constrained by 40 Ar/39Ar dating that has provided a magmatic crystallization age of 178.7 + 2.6 Ma. The top-tothe-south obduction was likely initiated along the margin of the back-arc domain and was transported as a whole onto the SAB in Coniacian –Santonian times, as reported by Sosson et al. Final closure of the basin is Late Cretaceous in age as dated by metamorphic rocks. Danelian et al. provide new information relevant to the palaeogeographic and geodynamic evolution of the Tethyan realms preserved in the Lesser Caucasus on the basis of new and revised radiolarian biochronology for the sedimentary cover of Armenian ophiolites, including those associated with the Se –Ak suture zone discussed by Sosson et al. and Rolland et al. Danelian et al. also provide a review of previously existing data for these sedimentary rocks. The oldest available ages come from the Se–Ak and suggest that ocean crust formation had occurred by the Late Triassic, somewhat older than Ar dates reported by Rolland et al. Data from both the Se– Ak and Vedi ophiolites provide new evidence for Middle Jurassic (Bajocian) oceanic environment continuing until at least the Late Jurassic (Mid–Late Oxfordian to Late Kimmeridgian/Early Tithonian). The paper by Avagyan et al. deals with the recent tectonic stress evolution in the area. The present stress field, derived from the kinematics of active faults, corresponds to a strike–slip regime with both transtensional and transpressional characteristics. However, both the micro fault kinematic data and the distribution of Neogene to Quaternary

INTRODUCTION

volcanic clusters demonstrate changes of stress field orientation through time since the beginning of the Arabia –Eurasian collision. A NW–SE orientation of compression was dominant between the Palaeogene and the late Early Miocene, a NE –SW one prevailed between the Late Miocene and the Quaternary. The stress field has accordingly changed from a general north– south compressional regime (expressed as thrusting and reverse faulting) to a transtensional-transpressional one (expressed by strike –slip faulting with various vertical components).

East and SE Anatolia: stratigraphic and tectonic evolution of the Peri-Arabian suture zone Kaymakci et al. present new data and a comprehensive review of kinematic evolution of the SE Anatolian Orogen based on a palaeostress inversion study for the Late Cretaceous to Recent. They nicely illustrate that the region experienced five different phases of deformation. Among these, the first and the third deformation phases are characterized by extension while the other three phases are related to compression. They propose that the first three phase are related to subduction related processes that include slab roll-back, subduction of young oceanic crust and slab-detachment/slab tear processes. They relate the last two phases to the northwards convergence of the Arabian Plate into Eurasia since the end of the Middle Miocene. The paper by Kuscu et al. presents new geochemical and geochronological results about the Late Cretaceous to Middle Eocene igneous rocks in the southeastern-eastern Anatolian orogenic belt. These authors have studied the major intrusive and extrusive rocks in this belt using high-precision geochronology (U –Pb and 40Ar/39Ar) and provide a complete set of new geochemical data. The integration of these geochronological and geochemical data from an evolving orogen provides the basis to reconstruct the temporal and spatial transition from subduction-related to post-collision and to late-orogenic magmatism in the region. Kuscu et al. suggest that the subduction-related magmatism is rooted to the closure of the Neotethys ocean during Late Cretaceous to Palaeogene times, whereas postcollision and late orogenic-within plate-related magmatism is driven by the collision of a northern promontory of the SE Anatolian orogenic belt with northerly derived ophiolitic rocks during the Palaeogene. Oberha¨nsli et al. present the alpine evolution of the eastern Bitlis complex in southeastern Anatolia according to new field studies, including the

7

discovery of HP–LT rocks. These authors demonstrate that the Bitlis complex, now accreted to the Tauride –Anatolide –South Armenian (TASA) block, was derived from north Arabia. Metamorphic studies in its cover sequences have allowed constraints to be placed on the thermal evolution of the massif: the regionally distributed HP –LT metamorphic evolution is documented by glaucophane, relics of carpholite in chloritoid-bearing schists and pseudomorphs after aragonite in marbles. A new metamorphic age of these assemblages is determined from 40Ar/39Ar dating as 74 + 2 Ma. This indicates that the Bitlis complex is a terrane detached from an indenter belonging to the Arabian plate that was subducted and stacked to form a nappe complex during the closure of the southern branch of Neotethys and that during its Late Cretaceous to Cenozoic evolution it never was heated to temperatures over 450 8C. The final paper of this volume is related to ¨ zacar et al. crustal structure in Eastern Anatolia. O use teleseismic P-wave receiver functions to map out Moho depths in the region and illustrate that the crust is thinnest under the Arabian plate and gradually becomes thicker towards the TASA block in the NE. These authors also provide new data from the western part of Eastern Anatolia showing that the crust is thicker near the Bitlis suture and towards the interiors of Anatolian plate. In addition, the Moho displays pronounced topography and, apparently, multiple fragmentations with anomalously high Vp/Vs along the North Anatolian Fault and near the youngest volcanic units inferred to indicate the presence of partial melting. A prominent low velocity zone in the lower crust indicative of a decoupling zone between the crust and the upper mantle in Eastern Anatolia is also inferred.

Synthesis and summary Pre-Cimmerian and Cimmerian history Palaeotethys is a triangular remnant oceanic basin with its tip located around present-day Iberia and widening eastwards. The polarity of subduction of the oceanic crust of Palaeotethys is a matter of debate. Nevertheless, a number of continental blocks rifted away from the northern margin of Gondwana, on its southern boundary, and these blocks drifted northwards during the Mesozoic giving rise to the opening of the Neotethys (also known as Mesozoic Tethys) ocean behind them. The evolution and closure of the Neotethys has dominated the subsequent tectonics of the Black Sea –Caucasus region and, as such, is the main concern of this book. However, the contentious issue of the polarity and timing of subduction of

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Palaeotethys (‘Cimmerian’ tectonic events) during the Mesozoic and initiation of Neotethyan subduction (‘Alpine’ tectonic events) are closely related. The general assumption among different researchers is that early to mid-Mesozoic events on the southeastern margin of the European continent are related to the closure of the Palaeotethys ocean although no conclusive evidence related to the age of oceanic remnants and its epiophiolitic cover have been reported from the region. This issue has also been addressed in GSL Special Publication No. 312 (Brunet et al. 2009), also based on results of the MEBE programme. The new results in this issue from the eastern Pontides and Greater Caucasus (McCann et al. 2010; Ustao¨mer & Roberston 2010) support models in which the subduction of Palaeotethys terminated at the end of Triassic to earliest Jurassic times. The evidence includes the presence of widespread volcanic rocks in Crimea and Caucasus that imply that the subduction of Neotethys had already commenced in the Early –Middle Jurassic. The location of crustal accretion associated with Palaeotethys closure, determined here to be the eastern Pontides and Greater Caucasus, has implications for the pre-Black Sea position of the margin of the European continent. The geological results show that the early Mesozoic (Palaeotethys subduction) accretionary margin of Europe was indeed to the south of the Scythian Platform (Fig. 2). This is consistent with the inferences of Stephenson & Schellart, who concluded that in general the basement lithosphere of the Black Sea, in which it formed, is strong (craton-like) and also the geophysical results of Yegorova & Gobarenko who found that the crust is colder and stronger beneath the western Black Sea sub-basin than the eastern one. Cimmerian deformations are also well imprinted in the Eastern Pontides (Ustao¨mer & Robertson) and the Crimean and Greater Caucasus region (McCann et al. 2010). Rifting in the region was initiated during the early Middle Jurassic with deformed units unconformably overlain by the Upper Jurassic deposits. These results are comparable to those of the Transcaucasus region to the SE where it seems that a similar tectonic evolution occurred during this period (Sosson et al. 2010). Metamorphosed rock units related to Cimmerian events are also found in the Strandja unit of Eastern Balkanides (Bergerat et al. 2010).

Back-arc basins: extension and magmatism The Mid-Jurassic –Cretaceous period witnessed two subduction processes and two remnant suture zones in Anatolia (Fig. 1). The Izmir-AnkaraErzincan-Sevan Suture Zone demarcates the

former position of the northern subduction. The ophiolites associated with this suture zone are considered to be derived from the Northern Neotethys Ocean. The Bitlis– Zagros Suture Zone demarcates the former position of the Southern Neotethys Ocean. The subduction zones associated with these suture zones were characterized by slab roll-back, slab detachment and slab-tear processes that possibly gave way to development of suprasubduction zone ophiolites, subduction related magmatism mainly in the Pontides (Ustao¨mer & Robertson), and opening of the western and eastern Black Sea sub-basins (Hippolyte et al. 2010; Stephenson & Shellart 2010). The newly obtained 40Ar/39Ar ages and geochemical data (Kuscu et al. 2010; Oberha¨nsli et al. 2010) indicate that similar processes were operating in the southern Neotethys ocean, which is also characterized by the Supra-Subduction Zone (SSZ) ophiolite generation and widespread magmatic activity (starting from 98.9 Ma). The SSZ ophiolites of the northern Neotethys Branch were obducted onto the Tauride–Anatolide –South Armenian (TASA, Fig. 2) microplate during Late Coniacian – Santonian as evidenced by stratigraphic and geochronological data (Rolland et al. 2010; Sosson et al. 2010) and by the ages of the plutons intruding both these ophiolitic nappes and their footwall (Kuscu et al. 2010).

Inversion: compression and magmatism The extensional events, including the formation of the Black Sea basin, that took place mainly during middle-Late Cretaceous times (Bergerat et al. 2010; Kaymakci et al. 2010; Khriachtchevskaia et al. 2010; Konerding et al. 2010) were replaced as subduction terminated in the northern Neotethys and collision of the intervening continental terranes began, leading to inversion tectonics throughout the circum-Black Sea area. The onset of inversion tectonics is fairly precisely dated by structural relationships seen in seismic data on the northern margin of the Black Sea as late Middle Eocene (Khriachtchevskaia et al. 2010), following a period of relative quiescence. However, extension following Late Cretaceous obduction processes on the SE margin of the TASA block in SE Turkey and on the Arabian Plate continued until the Middle Eocene possibly due to roll-back of the slab related to the southern Neotethys Oceanic (Kaymakci et al. 2010). The collision in the north gave way to deposition of thick Palaeogene flysch to molasse on the Pontides and TASA Block (Hippolyte et al. 2010; Sosson et al. 2010; Ustao¨mer & Robertson 2010). Compressional deformation and inversion along the Bitlis –Zagros suture zone also started by the Late Eocene.

INTRODUCTION

From the Late Oligocene onwards slab roll-back, slab-detachment and finally collision and further convergence of the Arabian Plate gave way to the present-day tectonics that affects a very large area extending from the Arabian Plate in the south to Greater Caucasus and the Black Sea Basin in the north. Inversion tectonics in the immediate Black Sea area ended by the Middle to Late Miocene (Khriachtchevskaia et al.; Konerding et al. 2010), as indicated by evidence of renewed extensional tectonic activity in the western Black Sea area (Konerding et al. 2010).

Present configuration and active tectonics The collision between Arabia and Eurasia led to north–south shortening, thickening and uplift of the crust of the Taurides Anatolides and the Armenian highland as well as Lesser and Greater Caucasus regions, since the Neogene. This also gave way to the development of active thrust and strike– slip faults recently expressed by devastating earthquakes. Tomographical studies indicate that the depth to Moho beneath East Anatolia and the Bitlis region is deeper than the Arabian plate, possibly indicating the effect of convergence and the consequent shortening of Eurasia in the region ¨ zacar et al. 2010). As a result of this compression, (O the Greater Caucasus is uplifted and thrust over the flanking foreland basins northwards and southwards (Mosar et al. 2010), as a large, intraplate, crustalscale inversion structure. The Quaternary magmatic activity in the Lesser Caucasus and adjacent regions is aligned to major fault zones (Avagyan et al. 2010) while it is distributed in large area in eastern Turkey (Kuscu et al. 2010). The main active strike – slip faults in the Lesser Caucasus correspond to inherited structures related to the collision of the TASA block with Eurasia. On the other hand, some of the active normal faults within the TASA basement are most likely related to the early Mesozoic Neotethys rifting and/or synconvergence extensional deformation related to Late Cretaceous to Eocene back-arc extension (Avagyan et al. 2010; Kaymakci et al. 2010; Kuscu et al. 2010).

Conclusion The 20 papers presented in this Special Publication address various aspects of evolution of Tethys system in the area extending from the Balkans, Black Sea and Greater Caucasus to the Arabian Plate. They cover some of the main events related to closure of the Palaeotethys and opening of Neotethys oceans, Early to Late Jurassic Cimmerian events, the Cretaceous opening of Black Sea basins, formation of arcs and back arc basins along with the

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subduction of Neotethys, Late Cretaceous to Early Cenozoic convergence and collisional processes, as well as uplift and present day crustal and lithospheric structure in the region. The results of these papers indicate that the Palaeotethys Ocean completely closed at the end of Triassic –Early Jurassic times, while two branches of the Neotethys ocean were opening. The northern branch lay between the Pontides and the Tauride –Anatolide –South Armenian (TASA) Block and the southern branch lay between the TASA block and the Afro-Arabian Plate. There is convincing evidence that the subduction of Neotethys commenced possibly during the Jurassic, therefore much earlier than previously thought. Formation of arcs and back-arc basins took place during the Late Jurassic and Cretaceous. The Late Cretaceous and, especially, the early Cenozoic is dominated by compressional deformation, basin inversion, and widespread subduction to collision related magmatism. Uplift and full development of the mountain chains in the region began at this time and culminated by the Oligo-Miocene in response to the collision and further northwards convergence of the Arabian Plate along the Bitlis Zagros Suture zone. This ultimately gave way to the development of active thrust and strike–slip fault systems in the region.

References Adamia, Sh., Alania, V., Chabukiani, A., Chichua, G., Enukidze, O. & Sadradze, N. 2010. Evolution of the Late Cenozoic basins of Georgia (SW Caucasus): a review. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 239–259. Avagyan, A., Sosson, M. et al. 2010. Recent tectonic stress evolution in the Lesser Caucasus and adjacent regions. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 393–408. Bergerat, F., Vangelov, D. & Dimov, D. 2010. Brittle deformation, paleostress field reconstruction and tectonic evolution of the Eastern Balkanides (Bulgaria) during Mesozoic and Cenozoic times. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 77– 111. Brunet, M.-F., Wilmsen, M. & Granath, J. W. (eds) 2009. South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312. Hippolyte, J.-C., Mu¨ller, C., Kaymakci, N. & Sangu, E. 2010. Dating of the Black Sea Basin: new

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nannoplankton ages from its inverted margin in the Central Pontides (Turkey). In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 113– 136. Kaymakci, N., Inceo¨z, M., Ertepinar, P. & Koc¸, A. 2010. Late Cretaceous to Recent kinematics of SE Anatolia (Turkey). In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 409–435. Khriachtchevskaia, O., Stovba, S. & Stephenson, R. 2010. Cretaceous–Neogene tectonic evolution of the northern margin of the Black Sea from seismic reflection data and tectonic subsidence analysis. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 137–157. Konerding, C., Dinu, C. & Wong, H. K. 2010. Seismic sequence stratigraphy, structure and subsidence history of the Romanian Black Sea shelf. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 159–180. Kuscu, I., Kuscu, G. G., Tosdal, R. M., Ulrich, T. D. & Friedman, R. 2010. Magmatism in the southeastern Anatolian orogenic belt: transition from arc to postcollisional setting in an evolving orogen. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 437–460. McCann, T., Chalot-Prat, F. & Saintot, A. 2010. The Early Mesozoic evolution of the Western Greater Caucasus (Russia): Triassic –Jurassic sedimentary and magmatic history. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 181–238. Mosar, J., Kangarli, T. et al. 2010. Cenozoic– Recent tectonics and uplift in the Greater Caucasus: a perspective from Azerbaijan. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black

Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 261– 279. Oberha¨nsli, R., Candan, O. et al. 2010. Alpine high pressure evolution of the eastern Bitlis complex, SE Turkey. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 461– 483. ¨ zacar, A. A., Zandt, G., Gilbert, H. & Beck, S. L. O 2010. Seismic images of crustal variations beneath the East Anatolian Plateau (Turkey) from teleseismic receiver functions. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 485–496. Rolland, Y., Galoyan, G., Sosson, M., Melkonyan, R. & Avagyan, A. 2010. The Armenian Ophiolite: insights for Jurassic back-arc formation, Lower Cretaceous hot spot magmatism and Upper Cretaceous obduction over the South Armenian Block. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 353– 382. Sosson, M., Rolland, Y. et al. 2010. Subductions, obduction and collision in the Lesser Caucasus (Armenia, Azerbaijan, Georgia), new insights. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 329– 352. Stephenson, R. & Schellart, W. P. 2010. The Black Sea back-arc basin: insights to its origin from geodynamic models of modern analogues. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 11–21. Ustao¨mer, T. & Robertson, A. H. F. 2010. Late Palaeozoic–Early Cenozoic tectonic development of the Eastern Pontides (Artvin area), Turkey: stages of closure of Tethys along the southern margin of Eurasia. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 281– 327.

The Black Sea back-arc basin: insights to its origin from geodynamic models of modern analogues RANDELL STEPHENSON1* & WOUTER P. SCHELLART2 1

School of Geosciences, Meston Building, King’s College, University of Aberdeen, Aberdeen AB24 3UE, UK 2

School of Geosciences, Monash University, Clayton, VIC 3800, Australia *Corresponding author (e-mail: [email protected])

Abstract: The Black Sea is generally thought to be a back-arc basin with active extension (rifting) beginning in late Early Cretaceous times – although some fundamental issues such as the presence or absence of a related magmatic arc and the orientation of the related, driving, subduction zone remain vaguely defined at best. However, as shown here, the regional structure of the Black Sea is consistent with that predicted by geodynamic models of modern back-arc basin formation, in which extension is driven by slab roll-back. This includes an asymmetric distribution of horst and graben structures in the back-arc basin, the distribution and spacing of which is related to the strength of the underlying lithosphere, which forms the hanging wall of the subduction zone. By analogy, the intrabasinal structure of the Black Sea as a whole is explicable as the consequence of a single phase of asymmetric back-arc basin formation, not two separate phases independently responsible for its western and eastern segments, and its underlying lithosphere is rheologically strong, as predicted by recent models of Precambrian Europe and present-day tomography.

The Black Sea is generally considered to be a Cretaceous –Palaeogene back-arc basin related to the subduction of an ocean since closed in presentday Turkey (e.g. Zonenshain & Le Pichon 1986; Dercourt et al. 1993; Okay et al. 1994; Robinson & Kerusov 1997; Nikishin et al. 2003 among others). It formed within a continental lithosphere that is made up of a mosaic of fragments accreted during several tectonic events from the Precambrian to the early Mesozoic (cf. Pease et al. 2008). The Black Sea basin consists of two main deep depressions, referred to as its western and eastern sub-basins or segments; both are generally thought to have oceanic or sub-oceanic affinity (Starostenko et al. 2004; Shillington et al. 2009); these are separated by the Mid–Black Sea Rise (Fig. 1), which is composed of thinned continental crust (Starostenko et al. 2004; Stephenson et al. 2004; Yegorova & Gobarenko 2010). Extensional structures in both segments of the Black Sea have been inverted in a compressional stress regime established in Eocene –Oligocene times (cf. Khriachtchevskaia et al. 2010) that is regionally related to convergence processes and collision on the southern margin of Eurasia. There is an ongoing debate about the relative ages of the western and eastern segments of the Black Sea basin, with the former usually considered to be somewhat older than the latter (Cretaceous v. Palaeogene; cf. Nikishin et al. 2003). Although there is abundant seismic information indicating

that a thick succession of sediments lies in both segments, there are no robust data constraining the ages of the oldest of these (e.g. Starostenko et al. 2004; Stephenson et al. 2004; Yegorova & Gobarenko 2010). The orientation of the main rift structures defining each segment also appears to be somewhat oblique. The main rift of the eastern Black Sea sub-basin appears to run NW–SE, (sub)parallel to the trend of the Mid–Black Sea Rise to its west and with structures along its eastern margin with the Caucasus area. The main rift structures of the western Black Sea sub-basin are more poorly defined, because they are covered with a significant thickness of flat-lying post-rift sediments (Yegorova & Gobarenko 2010) but they are normally judged to be running SW– NE to west– east in accordance with the orientation of the Odessa Shelf margin to the north and the Bulgarian-Turkish margin to the south. What these discordances in geometry and timing could mean for the mechanism(s) of back-arc extension and mechanics of rifting in terms of the broader plate configuration and kinematics is poorly defined. The inferred obliquity between the two rift axes in the western and eastern Black Sea sub-basins, and that the latter might be younger than the former, has led some authors to incorporate two co-existing subduction zone segments running sub-parallel to the rift axes in their reconstructions with a cusp in between (e.g. Nikishin et al. 2001, 2003) or other significant complications in

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 11– 21. DOI: 10.1144/SP340.2 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Regional tectonic map of the Black Sea and surrounding regions, modified from Okay & Tu¨ysu¨z (1999).

plate kinematic models (e.g. Stampfli & Kozur 2006). Incorporation of these kinds of complexities clearly has significant implications for Tethyan geodynamics, slab geometry and palaeo-tectonic reconstructions. Here, we try to relate what is generally believed about Black Sea geology and evolution, including the relative timing of rifting and orientation of rift structures in its western and eastern segments, to geodynamic models and observations of modern back-arc basin systems. We hope to elucidate the processes responsible for its formation and possibly help define a template in which to advance interpretation of the generally under constrained observational dataset for the Black Sea.

The Black Sea: basement and regional tectonic setting The basement crust that was rifted to form the Black Sea basin in the Cretaceous is poorly understood. There is considerable doubt about the basement affinity of segments of the margins of the basin even where they remain above water. At best, it is possible only to speculate about the affinity of submarine continental crust, such as that lying under the Odessa Shelf and the Mid–Black Sea Rise.

Much of the crust to the north, west and south of the western Black Sea sub-basin and north of the eastern Black Sea sub-basin is of Precambrian age. This corresponds to tectonic domains referred to as the Scythian and Moesian platforms and the Istanbul Zone of the western Pontides, respectively (Fig. 1). Okay et al. (2008) consider the basement of the Istanbul Zone to have Panafrican-Cadomian affinity (Neoproterozoic). The (thinned) continental crustal basement of the Mid –Black Sea Rise is thought to be correlatable with Precambrian units found in the Transcaucasus area to the east (Saintot et al. 2006). The geology exposed on the Crimean Peninsula can be considered as the basement geology of the northern shelf of the Black Sea, uplifted during Cenozoic (Eocene and younger) compression as the relatively buoyant crust of the Mid–Black Sea Rise was thrust under it (Stephenson et al. 2004). The northern part of the Crimean Peninsula and areas to the east, beneath the Azov Sea (northern margin of the eastern Black Sea sub-basin) and north of the Greater Caucasus are underlain by Precambrian basement of the Scythian Platform. (Note that this differs significantly from a number of papers in which the crystalline basement of the Scythian Platform is referred to as being Late Palaeozoic in age [e.g. Nikishin et al. 2003]; however, there is

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no evidence – among, indeed, very little unambiguous data in general – to suggest that this is the case [cf. Saintot et al. 2006; Stephenson et al. 2004, 2006].) South of the Black Sea, east of the Istanbul Zone, the central and eastern Pontides can be considered as being accreted during the closure of a PalaeoTethyan ocean in latest Palaeozoic through Late Triassic (to earliest Jurassic) times. The basement of the Crimean mountains on the southernmost part of the Crimean Peninsula has sometimes also been regarded as a Pontide equivalent, hence representing crust newly formed in a Palaeo-Tethyan subduction-accretion complex. It might also comprise deformed sedimentary units deposited atop crust similar to the Scythian Platform crust to the north. The Crimean mountains display some deformation of Early Jurassic and Early Cretaceous age but there is no firm evidence of ocean closure and most of the deformation is Cenozoic. They can be considered as the western prolongation of the Greater Caucasus orogen, where the latter has involved inversion of structures on the northern margin of the Black Sea and on the Odessa Shelf (Khriachtchevskaia et al. 2010). There are also inversion structures on the southern margin of the Black Sea (e.g. Stephenson et al. 2004) and on the margins of the Mid–Black Sea Rise (S. M. Stovba, pers. comm. 2009). Pease et al. (2008) concluded that the late Precambrian border of the European craton (Baltica) lay somewhere in the vicinity of the present axis of the Black Sea. This would later represent the locus of a Neoproterozoic-Early Palaeozoic suture along which Cadomian terrains to the south, including units such as those Winchester et al. (2006) refer to as ‘Avalonian eastern extremities’, were accreted. It is beyond the scope of this chapter to solve the puzzle of the mosaic of crustal fragments that lay in the area of the Black Sea prior to its development but suffice to say that the Black Sea formed in continental crust characterized by pre-existing structures associated with various accretionary events from late Precambrian through Late Palaeozoic–Early Jurassic times. These were likely probably further accentuated prior to Black Sea rifting by widespread extensional and transtensional tectonics (and basin formation) affecting the southern margin of Eurasia from the Late Devonian through to the Triassic (Saintot et al. 2006; Stephenson et al. 2006) including, for example, the Ku¨re Basin of Stampfli & Kozur (2006). Figure 2 shows an early Cretaceous (middle Aptian) reconstruction of the Black Sea area modified from the MEBE (Middle East Basins Evolution) Atlas of palaeotectonic– sedimentary –palinspastic maps (Barrier & Vrielynck 2008). Subduction of the Neo-Tethys Ocean to the north under Eurasia

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began around the beginning of the Triassic, according to most palaeo-tectonic reconstructions (e.g. Stampfli & Kozur 2006). Thereafter, extension and basin formation was widespread within the Eurasian continental back-arc behind the inferred subduction zone; some reconstructions consider that new (true) ocean basins formed and closed in Triassic –Jurassic times (e.g. Stampfli & Kozur 2006) while others (e.g. Barrier & Vrielynck 2008) imply only that deep sedimentary basins, not necessarily floored by oceanic crust, formed and none of these occurred in the immediate vicinity of the eventual Black Sea. Back-arc rifting in the Greater Caucasus and south Caspian basins occurred during the Jurassic and these basins were undergoing post-rift thermal subsidence by the Aptian (Fig. 2). To the west, prior to the Aptian, according to the reconstructions of Barrier & Vrielynck (2008), the subduction system south of the present-day Black Sea was linked by a transform plate boundary to a mid-ocean rift system forming the northeastern Ligurian Ocean (south of the Bohemian Massif), equivalent to the (eastern) Alpine Tethys of Stampfli & Kozur (2006).

Geodynamic models of back-arc extension There is little argument that the Black Sea formed as a result of extension in a back-arc setting and that this occurred behind a retreating subduction zone (slab roll-back) that was 1500–2000 km in trench-parallel extent (c. east –west). Although Mesozoic plate reconstructions in the vicinity of the Black Sea vary in detail and complexity according to their scale, all of them incorporate in a fundamental way a long-lived convergent northern margin involving the consumption of the NeoTethyan Ocean along a subduction system dipping northwards under the Eurasian continent (e.g. Fig. 2). Accordingly, any hypothesis about Black Sea evolution – including the distribution and relative age of formation of the extensional structures that it comprises – should be compatible with what is known about the formation of recent and modern back-arc basin analogues. As such, we briefly review the state-of-the-art vis-a`-vis extension behind subduction zones, as primarily determined from observations and models of subduction in the Pacific Ocean.

Symmetric v. asymmetric back-arc basin opening Recent geodynamic models of subduction (Kincaid & Griffiths 2003; Schellart 2004; Funiciello et al. 2006; Stegman et al. 2006; Schellart et al. 2007; Di

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Fig. 2. Simplified palaeotectonic reconstruction of the Black Sea area in the early Cretaceous, based mainly on the MEBE Atlas middle Aptian map (Barrier & Vrielynck 2008; Map 5). Short, thick dashed blue lines represent the main rift axes developing at this time, eventually to coalesce as the Black Sea basin. V indicates arc volcanism, associated with active subduction (red line with teeth). Previously, subduction was active further to the west (dashed red line) related to rifting and sea-floor spreading in the NE Ligurian Ocean (LO) and back-arc rifting was active in Greater Caucasus (GCB) and South Caspian (SCB) basins (Barrier & Vrielynck 2008; Map 4 – Tithonian), which by the early Cretaceous have coalesced (thin, longer dashed blue line) into a thermally subsiding basin in a post-rift setting. The dotted curved line crossing corresponds to the 308N palaeo-latitude. Other abbreviations: EBB, axis of the eastern Black Sea Basin; KT, Karkinit Trough; MBR, Mid–Black Sea Rise; MP, Moesian Platform; WBB, axis of the western Black Sea Basin; UkS, Ukrainian Shield. Inset: The relative position of the subduction zone trench south of the Black Sea at this time (middle Aptian; 118 Ma) compared to the early Tithonian (156 Ma), red and yellow lines, respectively, according to the MEBE Atlas (Barrier & Vrielynck 2008). The dotted line is the Aptian 308N latitude and the Tithonian trench position has been drawn taking into account a southward drift of Eurasia of about 200 km in the intervening c. 30 Ma.

Giuseppe et al. 2008) have stressed the importance of the third, trench-parallel, dimension in such models. Such models, as well as kinematic investigations, indicate that maximum trench retreat velocities are found close to (c. ,1500 km) lateral slab edges (Schellart et al. 2007) because of an efficient torroidal-type return flow of mantle material around the lateral slab edge (e.g. Schellart 2004; Funiciello et al. 2006; Stegman et al. 2006; Schellart et al. 2007). The presence of a lateral subduction zone edge thus enhances variations in trench velocity along the trench axis, promoting asymmetric trench retreat and slab roll-back, and progressive growth of the slab tear. Govers & Wortel (2005) referred to such subvertical tear faults as subduction transform edge propagator (STEP) faults.

Figure 3 shows the differences in back-arc deformation structure attained from symmetric (Fig. 3a) and asymmetric (Fig. 3b) slab roll-back. In the former, the retreating boundary attains an arcshaped geometry that is convex towards the direction of retreat (see also Hatzfeld et al. 1997; Gautier et al. 1999; Schellart & Lister 2005). In the latter, the retreating boundary is analogous to an opening door with a hinge-point on one side (see also Schellart et al. 2002a, 2003). For the symmetric case, deformation close to the retreating boundary is characterized by normal faults striking (sub)parallel to the arc-shaped retreating boundary. For the asymmetric case, the normal faults strike at an angle with the straight retreating boundary and the direction of extension is orientated more

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Fig. 3. Experiments simulating back-arc deformation in overriding plate lithosphere as a consequence of an extensional boundary condition. (a) Experiment 7 from Schellart & Lister (2005) with symmetrical boundary conditions. (b) Experiment 3 from Schellart et al. (2003) with asymmetrical boundary conditions. For both experiments the model consists of three layers with 5.5 cm of high-density ( r ¼ 1.42  103 kg/m3) low-viscosity (dynamic viscosity h  102 Pa . s) glucose syrup at the bottom (cf. asthenosphere), overlain by 1.2 cm of high viscosity (h ¼ 2  104 Pa . s) silicone putty filled with iron powder ( r ¼ 1.22  103 kg/m3) (cf. viscous lower lithosphere), and on top 0.6 cm of frictional-plastic glass microspheres mixed with hollow glass microspheres ( r ¼ 1.22  103 kg/m3) (cf. brittle upper lithosphere). Spacing between individual grid lines is 3 cm.

towards the far end of the retreating boundary. This angle progressively increases going from the far end towards the hinge-point. The difference in deformation pattern in the overriding plate is obvious not only close to the retreating boundary, but also further away in the region located north of the gate. For the symmetric case, deformation is dominated by normal faulting and rift structures on either side of the (north–south striking) symmetry plane. For the asymmetric case, normal faults are present near the hinge-point and the region north of the gate, while strike –slip deformation is found near the far end of the retreating boundary.

Strength of the back-arc lithosphere Figure 4 shows the effect of overriding plate rheology on the development of back-arc basin fault patterns from two experiments by Schellart et al. (2002a). The model lithosphere is relatively strong in Figure 4a, b and relatively weak in Figure 4c, d (see caption for details). Two stages of evolution are shown: Figure 4a, c are early stages and Figure 4b, d are later stages. In the strong lithosphere case (Fig. 4a, b), the fault style is dominated by normal faulting, with deformation localized into a number of discrete rift structures with relatively undeformed lithosphere in between. Individual fault segments are long (longer than the thickness of the model lithosphere). In the weak lithosphere case (Fig. 4c, d), the fault style is dominated by both normal faulting and strike–slip faulting, with deformation

distributed over a large area. Normal faults predominantly occur close to the retreating boundary, while transtensional strike–slip faults occur more inboard. Individual normal fault segments are much shorter than in the strong lithosphere case. In both cases (strong and weak), deformation initiates close to the far end of the retreating boundary (Fig. 4a, c), and migrates towards the hinge-point and away from the subduction zone (Fig. 4b, d). For the weaker lithosphere, the area that eventually accommodates deformation is much larger than for the strong lithosphere.

Discussion Asymmetric trench retreat Asymmetric trench retreat has been suggested to explain the wedge-shaped geometry and complicated spreading patterns in the North Fiji Basin (Schellart et al. 2002b). It has also been proposed to explain the oblique normal fault patterns in the Okinawa Trough (Schellart et al. 2002a), as well as the oblique fault patterns in the Sea of Okhotsk and the wedge shaped geometry of the Kuril Basin (Schellart et al. 2003). Asymmetric trench retreat is rather common for present-day subduction zones, such as the Tonga subduction zone (Bevis et al. 1995; Schellart et al. 2002a), New Hebrides subduction zone (Schellart et al. 2002b), New Britain subduction zone (Wallace et al. 2005) and Ryukyu subduction zone (Schellart et al. 2002a; Nishimura et al. 2004). Of

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Fig. 4. Experiments simulating back-arc deformation in overriding plate lithosphere during asymmetric back-arc deformation for different integrated strengths of the lithosphere. Panels on the left-hand side show the photographs, while right-hand side panels show the line drawings. (a–b) Experiment 4 from Schellart et al. (2002a) with a strong

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Fig. 5. Schematic sketches of progressive asymmetric trench retreat, slab roll-back and back-arc opening in top view (top) and 3D perspective view (bottom) (modified from Schellart et al. 2002a). (a) Initial state with rectilinear subduction zone; (b) initiation of roll-back between two points or cusps where subducting slab is resisting to roll back; one end evolves into transform plate boundary due to formation of sub-vertical tear in subducting slab, while the other becomes a hinge-point around which the arc rotates; (c) progressive growth of tear and continued rotation around hinge-point. Black arrows indicate direction of trench retreat. Thick arrow in (c) indicates direction of upper mantle return flow to accommodate slab roll-back.

these, the Ryukyu subduction zone involves backarc rifting of continental lithosphere such as for the Black Sea. Further, the apparent duration of rifting in the Black Sea (c. 50 Ma?) is comparable to that of modern back-arc basins (few Ma to tens of Ma) (Schellart & Lister 2005), though indeed towards the upper limit.

Application to the Black Sea Asymmetric trench retreat as illustrated in Figures 3 and 4 induces an asymmetric kinematic boundary condition on the overriding plate, which has profound implications for deformation of the overriding plate. Such asymmetric trench retreat promotes the formation of an asymmetric back-arc basin with rift structures that strike obliquely to the

strike of the trench. Figure 5 is a schematic diagram showing how this might apply to the formation of the Black Sea. Roll-back initiates between points or cusps where it is being resisted for some reason. In the case of the Black Sea this is, in the east, the terminus of the Neo-Tethys subduction zone segment rolled-back previously to produce the south Caspian and (a bit later) the Greater Caucasus back-arc basins (Fig. 2). Why there is resistance to continued roll-back in this segment is unspecified but that it exists is consistent with the simple fact that back-arc rifting in these basins has ceased by the time rifting in the Black Sea segment began. The western cusp of the Black Sea segment roll-back coincides with the pre-existing zone of strike –slip kinematics (linking the NeoTethyan convergent boundary with a spreading

Fig. 4. (Continued) lithosphere dominated by brittle deformation; (a) early stage and (b) late stage. Lithosphere consists of 1.0 cm of high-viscosity silicone putty at the bottom filled with iron powder (dynamic viscosity h ¼ 2  104 Pa . s; density r ¼ 1.22  103 kg/m3), a 2 mm thick viscoplastic layer of silicone mixed with glass spheres and 0.6 cm of frictional plastic glass microspheres on top mixed with hollow glass microspheres (r ¼ 1.22  103 kg/m3). This gives a relatively strong brittle top layer and a relatively weak viscous bottom layer and, thus, a relatively high brittle to viscous strength ratio. (c – d) Experiment 2 from Schellart et al. (2002a, 2003) with a weak lithosphere dominated by viscous deformation; (c) early stage and (d) late stage. Lithosphere consists of 1.2 cm of high-viscosity silicone putty at the bottom filled with iron powder (h ¼ 2  104 Pa . s; r ¼ 1.22  103 kg/m3) and 0.4 cm of frictional plastic glass microspheres on top mixed with hollow glass microspheres ( r ¼ 1.22  103 kg/m3). This gives a relatively weak brittle top layer and relatively strong viscous bottom layer, so a lower brittle to viscous strength ratio than for the strong lithosphere experiment. In both experiments the model lithosphere is underlain by 5.5 cm of high-density ( r ¼ 1.42  103 kg/m3) low-viscosity (h  102 Pa . s) glucose syrup (cf. asthenosphere).

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system in the Alpine Tethys) and may represent the continued propagation of this boundary as a subvertical tear or STEP fault. This end of the Black Sea subduction segment would have evolved into (or would have continued to evolve as) a transform plate boundary – kinematically similar to the West Black Sea Fault postulated by Okay & Tu¨ysu¨z (1999) in their original version of Figure 1 – while the other became a hinge-point around which the arc rotated leading to the development of an asymmetric back-arc rift basin. The orientation of extensional structures, shown schematically in Figure 5 and based directly on the results of analogue models such as those illustrated in Figures 3 and 4, is strongly compatible with the orientation of rift structures, to the extent they are confidently known, in the Black Sea (cf. Fig. 1). (Note that the trend of the Karkinit Trough – KT in Fig. 2 – has been affected by the subsequent – Cenozoic – development of a Greater Caucasus foredeep.) Furthermore, the models show that these structures tend to be older at the non-hinge end of the asymmetric roll-back system (Fig. 4), which is consistent with suggestions that rifting in the western Black Sea pre-dates rifting in the eastern Black Sea. Though the actual observational evidence constraining the relative ages of structures in the western and eastern sub-basins of the Black Sea is ambiguous, the proposed asymmetric roll-back model (Fig. 5) can easily explain this commonly made inference (cf. Nikishin et al. 2003). The inset in Figure 2 gives an indication of the amount of roll-back the Black Sea segment of the Neo-Tethys subduction zone might have experienced according to the reconstructions of Barrier & Vrielynck (2008) in order to accommodate rift structures developed by the Aptian. Active backarc rifting in the Black Sea clearly continued after the Aptian, until at least the Late Cretaceous and perhaps later (Nikishin et al. 2003; Hippolyte et al. 2010; Khriachtchevskaia et al. 2010), after which post-rift thermal subsidence occurred, and the position of the subduction trench in the reconstructions is only schematic. Nevertheless, the amount of extension in the Black Sea, based on twodimensional subsidence analyses (Spadini et al. 1996, 1997; Meredith & Egan 2002; Cloetingh et al. 2003; Shillington et al. 2008) and on inferred crustal structure (Starostenko et al. 2004; Yegorova & Gobarenko 2010) is in the order of 100 –150 km, compatible with the amount of slab roll-back implied by our model. The model predicts that the amount of the extension in the western Black Sea would be greater than in the eastern Black Sea (cf. Fig. 4). Though the general physiography of the Black Sea and its two sub-basins (see Yegorova & Gobarenko 2010) appears compatible

with this prediction, it is difficult to reconcile with the various estimates based on subsidence modelling, which do not clearly show any such tendency. (Note that available data on the age of sediments in the deeper parts of the Black Sea do not allow the subsidence and crustal structure models to make any differentiation between Cretaceous rifting and extension and basin formation that may have occurred at an earlier time, such as in the Permo-Triassic.) The effect of basement rheology on the style of back-arc extension illustrated by the models (Fig. 4) clearly suggests that the Black Sea formed atop a relatively strong lithosphere. The relative widths and lengths of the main structures forming the Black Sea basin – the eastern Black Sea rift, the Mid –Black Sea Rise, the (inferred) constituent rifts of the western Black Sea, and the Karkinit Trough and associated grabens on the northern margin of the Black Sea (Fig. 2) mimic those predicted by the strong rheology modelling results (Fig. 4a, b). This is in keeping with the supposition that the Black Sea formed within continental lithosphere mainly of Precambrian age. The presentday thermal state of the lithosphere underlying the Black Sea is relatively cold, much more like the European craton to the north than to Anatolia to the south (e.g. Tesauro et al. 2009), and this is perhaps also compatible with a strong lithosphere in the Cretaceous. Spadini et al. (1996), from two-dimensional subsidence modelling, concluded that the lithosphere of the eastern sub-basin of the Black Sea was significantly weaker than that of the western sub-basin. We think that this might be more a consequence of Precambrian basement geology heterogeneities, as inferred by Yegorova & Gobarenko (2010) from various kinds of geophysical data, and that the eastern sub-basin is much more affected by Cenozoic (Greater Caucasus) shortening superimposed upon the pre-existing extensional geometries.

Summary and conclusions Subduction is, by definition, a large-scale geodynamic process and, in turn, the processes responsible for back-arc extension cannot be overly localized. This does not preclude changes in convergence rates along subduction zones, including diachroneity or the ‘catastrophic’ modification to subduction zones such as the development of STEP faults (e.g. Govers & Wortel 2005). It also does not preclude complexities developed in back-arc extensional settings resulting from pre-existing structural or compositional heterogeneity. However, such heterogeneity and any dependent complexity must also respond and be displayed at a scale appropriate

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to the driving (subduction) mechanism, which is at least at a lithosphere to upper mantle scale. Modern (recent and currently active) subduction zones do not readily display separate back-arc spreading events (in age and orientation) at a scale as small as the Black Sea. Conversely, much of what is confidently known about the Black Sea basin, at the basin scale, is highly compatible with what is predicted from models of modern back-arc basins forming as a result of asymmetric slab rollback. The asymmetric back-arc extension model explains the general shape of the Black Sea basin as well as the geometry of its regional intrabasinal structure, including its western and eastern subbasins and intervening Mid–Black Sea Rise. Accordingly, our main conclusions are as follows: (1) The Black Sea formed as a single entity as an asymmetric back-arc basin bordering a subducting slab displaying asymmetric counterclockwise slab roll-back. (2) As such, the development of the western and eastern sub-basins of the Black Sea are essentially contemporaneous, although the asymmetric roll-back model is completely compatible with an earlier onset of development of extensional structures in the western compared to the eastern sub-basin with the youngest such activity being in the east and NE. (3) The Black Sea back-arc basin formed within rheologically-strong continental lithosphere, in the context of analogue models of asymmetric slab roll-back, demonstrated by the regional geometry of its main constituent extensional (horst and graben-like) structures, including the Mid–Black Sea Rise. This is in keeping with tectonic reconstructions in which the underlying lithosphere rifted to produce the Black Sea basin was mostly assembled prior to late Precambrian–Early Palaeozoic times. (4) The asymmetric roll-back of the Black Sea requires the presence of a transform-like or ‘STEP’-fault plate boundary near its western limit during the Cretaceous, which has implications for tectonic reconstructions in this area, including the link between the Alpine Tethys and Neo-Tethys oceans. (5) The hinge of asymmetric slab roll-back forming the Black Sea basin was, at a regional scale, near its south-eastern corner. This implies that slab roll-back to the east of the Black Sea, leading to the early stage of development of the south Caspian Basin and the Greater Caucasus Basin in the Jurassic, encountered increasing resistance by the

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Cretaceous, when roll-back in the Black Sea segment of the Neo-Tethyan subduction system began.

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Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 137–157. Meredith, D. J. & Egan, S. S. 2002. The geological and geodynamic evolution of the eastern Black Sea basin: insights from 2-D and 3-D tectonic modelling. Tectonophysics, 350, 157– 179. Nikishin, A. M., Korotaev, M. V., Ershov, A. V. & Brunet, M.-F. 2003. The Black Sea basin: tectonic history and Neogene– Quaternary rapid subsidence modeling. Sedimentary Geology, 156, 149–168. Nikishin, A. M., Ziegler, P. A. et al. 2001. Mesozoic and Cainozoic evolution of the Scythian PlatformBlack Sea-Caucasus domain. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & CrasquinSoleau, S. (eds) Peri-Tethys Memoir 6, Peri-Tethyan Rift/Wrench Basins and Passive Margins. Memoires du Museum National d’Histoire Naturelle, 186, 295– 346. Nishimura, S., Hashimoto, M. & Ando, M. 2004. A rigid block rotation model for the GPS derived velocity field along the Ryukyu arc. Physics of the Earth and Planetary Interiors, 142, 185– 203. Okay, A. I. & Tu¨ysu¨z, O. 1999. Tethyan sutures of northern Turkey. In: Durand, B., Jolivet, L., Horva´th, F. & Se´ranne, M. (eds) The Mediterranean Basins, Tertiary Extension within the Alpine Orogen. Geological Society, London, Special Publications, 156, 475– 515. Okay, A. I., Bozkurt, E., Satır, M., Yig˘itbas¸, E., Crowley, Q. G. & Shang, C. K. 2008. Defining the southern margin of Avalonia in the Pontides: geochronological data from the Late Proterozoic and Ordovician granitoids from NW Turkey. Tectonophysics, 461, 252– 264. Okay, A. I., S¸engo¨r, A. M. C. & Go¨ru¨r, N. 1994. Kinematic history of the opening of the Black Sea and its effect on the surrounding regions. Geology, 22, 267–270. Pease, P., Daly, J. S. et al. 2008. Baltica in the Cryogenian, 850–630 Ma. In: Bogdanova, S. V., Li, X.-X., Moores, E. M. & Pisarevsky, S. A. (eds) Testing the Rodinia Hypothesis: Records in its Building Blocks. Precambrian Research, 160, 46–65. Robinson, A. G. & Kerusov, E. 1997. Stratigraphic and structural development of the Gulf of Odessa, Ukrainian Black Sea, implications for petroleum exploration. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Areas. American Association of Petroleum Geologists. Memoir, 68, 369–380. Saintot, A., Stephenson, R. A., Stovba, S. M., Brunet, M.-F., Yegorova, T. & Starostenko, V. 2006. The evolution of the southern margin of Eastern Europe (Eastern European and Scythian platforms) from the latest Precambrian-Early Palaeozoic to the Early Cretaceous. In: Gee, D. G. & Stephenson, R. A. (eds) European Lithosphere Dynamics. Geological Society of London, Memoir, 32, 481–505. Schellart, W. P. 2004. Kinematics of subduction and subduction-induced flow in the upper mantle. Journal of Geophysical Research, 109, B07401, doi: 10.1029/2004JB002970.

Schellart, W. P. & Lister, G. S. 2005. The role of the East Asian active margin in widespread extensional and strike-slip deformation in East Asia. Journal of the Geological Society, London, 162, 959–972. Schellart, W. P., Jessell, M. W. & Lister, G. S. 2003. Asymmetric deformation in the backarc region of the Kuril arc, northwest Pacific: new insights from analogue modeling. Tectonics, 22, 1047, doi: 10.1029/ 2002TC001473. Schellart, W. P., Lister, G. S. & Jessell, M. W. 2002a. Analogue modelling of asymmetrical backarc extension. Journal of the Virtual Explorer, 7, 25–42. Schellart, W. P., Lister, G. S. & Jessell, M. W. 2002b. Analogue modeling of arc and backarc deformation in the New Hebrides arc and North Fiji Basin. Geology, 30, 311– 314. Schellart, W. P., Freeman, J., Stegman, D. R., Moresi, L. & May, D. 2007. Evolution and diversity of subduction zones controlled by slab width. Nature, 446, 308 –311, doi: 10.1038/nature05615. Shillington, D. J., Scott, C. L., Minshull, T. A., Edwards, R. A., Brown, P. J. & White, N. 2009. Abrupt transition from magma-starved to magma-rich rifting in the eastern Black Sea. Geology, 37, 7 –10. Shillington, D. J., White, N., Minshull, T. A., Edwards, G. R. H., Jones, S. N., Edwards, R. A. & Scott, C. L. 2008. Cenozoic evolution of the eastern Black Sea: a test of depth-dependent stretching models. Earth and Planetary Science Letters, 265, 360–378. Spadini, G., Robinson, A. G. & Cloetingh, S. 1996. Western versus Eastern Black Sea tectonic evolution: pre-rift lithospheric controls on basin formation. Tectonophysics, 266, 139– 154. Spadini, G., Robinson, A. G. & Cloetingh, S. 1997. Thermomechanical modelling of Black Sea basin formation, subsidence and sedimentation. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. American Association of Petroleum Geologists, Memoir, 68, 19–38. Stampfli, G. M. & Kozur, H. W. 2006. Europe from the Variscan to the Alpine cycles. In: Gee, D. G. & Stephenson, R. A. (eds) European Lithosphere Dynamics, Geological Society of London, Memoir, 32, 57–82. Starostenko, V., Buryanov, V. et al. 2004. Topography of the crust-mantle boundary beneath the Black Sea Basin. Tectonophysics, 381, 211–233. Stegman, D. R., Freeman, J., Schellart, W. P., Moresi, L. & May, D. 2006. Influence of trench width on subduction hinge retreat rates in 3-D models of slab rollback. Geochemistry Geophysics Geosystems, 7, Q03012, doi: 10.1029/2005GC001056. Stephenson, R. A., Mart, Y. et al. 2004. TRANSMED Transect VIII: Eastern European Craton to Arabian Craton (Red Star to Red Sea). In: Cavazza, W., Roure, F. M., Spakman, W., Stampfli, G. M. & Ziegler, P. A. (eds) The TRANSMED Atlas – The Mediterranean Region from Crust to Mantle. Springer Verlag, Berlin Heidelberg, CD-ROM.

THE BLACK SEA BACK-ARC BASIN Stephenson, R. A., Yegorova, T. et al. 2006. Late Palaeozoic intra- and pericratonic basins on the East European Craton and its margins. In: Gee, D. G. & Stephenson, R. A. (eds) European Lithosphere Dynamics, Geological Society of London, Memoir, 32, 463–479. Tesauro, M., Kabab, M. & Cloetingh, S. 2009. A new thermal and rheological model of the European lithosphere. Tectonophysics, 476, 478–495. Winchester, J. A., Pharaoh, T. C., Verniers, J., Ioane, D. & Seghedi, A. 2006. Palaeozoic accretion of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories. In: Gee, D. G. & Stephenson, R. A. (eds) European Lithosphere Dynamics. Geological Society of London, Memoir, 32, 323–332.

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Wallace, L. M., McCaffrey, R., Beavan, J. & Ellis, S. 2005. Rapid microplate rotations and backarc rifting at the transition between collision and subduction. Geology, 33, 857–860. Yegorova, T. & Gobarenko, V. 2010. Structure of the Earth’s crust and upper mantle of the West- and East-Black Sea Basins revealed from geophysical data and its tectonic implications. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 23– 42. Zonenshain, L. P. & Le Pichon, X. 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. In: Auboin, J., Le Pichon, X. & Monin, A. S. (eds) Evolution of the Tethys. Tectonophysics, 123, 181–211.

Structure of the Earth’s crust and upper mantle of the West- and East-Black Sea Basins revealed from geophysical data and its tectonic implications TAMARA YEGOROVA* & VALENTINA GOBARENKO Institute of Geophysics, National Academy of Sciences of Ukraine, Pr. Palladina 32, Kiev 03680, Ukraine *Corresponding author (e-mail: [email protected]) Abstract: A back-arc Black Sea Basin consists of two deep sub-basins – the West-Black Sea (WBS) and the East-Black Sea (EBS) – filled with thick sediments (up to 12– 14 km), which are separated by the mid– Black Sea Ridge (MBSR) – a NW trending basement uplift structure. For a better understanding of the lithosphere structure of these two sub-basins, the authors made a comprehensive analysis of the available geological and geophysical data, including carrying out a three-dimensional (3D) gravity back-stripping analysis, a reinterpretation of a number of seismic refraction profiles as well as the re-evaluation of seismological data and local seismic tomography. Inferred differences in the basin architecture and lithosphere structure of the WBS and EBS can be explained by different affinities of the underlying crustal domains and by the peculiarities of their (Cretaceous and younger) rift and post-rift history. Rifting that led to oceanic crust formation in the WBS occurred within the continental crust of the Moesian Platform along Mesozoic sutures with adjoining accreted terranes. The EBS, most probably, formed within the Transcaucasus continental domain due to strike–slip movements along the MBSR. Underthrusting of the EBS oceanic lithosphere beneath the continental domain of the Scythian Platform led to the formation offshore of the Crimean orogen of accretional wedge of Sorokin Trough.

The Black Sea Basin is located between the Ukraine in the north, Russia and Georgia in the NE and east, Turkey on the south and Bulgaria and Romania on the west. Tectonically the Black Sea Basin is situated between the orogenic belts of the Crimea and Great Caucasus on the north and Pontides on the south; on the west the Black Sea cuts the structures of Dobrogea, Moesian Platform and Istrandza (Fig. 1). The Black Sea Basin consists of two deep sub-basins – the West Black Sea Basin (WBS) and the East Black Sea Basin (EBS) filled with Cretaceous and younger sediments as thick as 12– 14 km. These sub-basins are not distinguished in the present-day sea floor bathymetry, which outlines a single flat basin without any topography with the sea floor at a depth of about 2 km (Fig. 1). The WBS and EBS differ in terms of their dimensions and configurations, thicknesses of sediments and general trends of the tectonic units. The rather wide (150 –300 km) WBS runs in a west –east direction in the form of large triangle over a distance of approximately 600 km (Fig. 2). The NW-oriented EBS, parallel with the Caucasus Fold Belt and the Mid–Black Sea Ridge (MBSR), has nearly the same length, but is considerably narrower (up to 100 –150 km). The MBSR, separating the WBS and EBS, is represented in the north by the Andrusov Ridge of up to 200 km length and 20 –80 km

width, and in the south by the Arhangelsky Ridge of 450 km total length and 10–60 km width (Fig. 2). The Black Sea Basin is considered to be a back-arc basin formed in Cretaceous –Palaeogene time due to the subduction of the Neotethys Ocean below the Pontide magmatic arc (Dercourt et al. 1993; Okay et al. 1994; Nikishin et al. 2003). Very thick sediments within the Black Sea Basin cover a heterogeneous basement, which is thought to be represented by a collage of different terranes and microplates accreted during several tectonic events (Pease et al. 2008; Winchester et al. 2006). Much debate has arisen not only with regard to the age of the formation of the Black Sea Basin, but also to the affinity of the crustal domains and lithosphere structure below the Black Sea. Here, a generalization of the geophysical information of the study area is made with the objective of delineating the main lithospherical domains and contributing to an understanding of the tectonic processes and evolution of the whole Black Sea– Caucasus region. Systematic geophysical investigations of the Black Sea region by seismic methods commenced at the end of the 1960s. In the following two decades almost the whole of the Black Sea was covered by differently oriented deep seismic sounding (DSS) profiles (Neprochnov et al. 1970; Malovitsky & Neprochnov 1972; Moskalenko &

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 23– 42. DOI: 10.1144/SP340.3 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Main tectonic units of the Black Sea region shown on the surface topography and bathymetry map. Red lines indicate the DSS profiles shown in Figure 6.

Fig. 2. 3D diagram showing base of Cretaceous successions according to Tugolesov et al. (1985). ANR, Andrusov Ridge; ARR, Arkhangelsky Ridge; ATB, Adjaro-Trialet Belt; BT, Burgas Trough; GT, Gurii Trough; IKT, Indolo-Kuban Trough; KT, Karkinit Trough; KTT, Kerch-Taman Trough; MBSR, Mid– Black Sea Ridge; NT, Niznekamchiisky Trough; SHR, Shatsky Ridge; ST, Sinop Trough; SRT, Sorokin Trough; TT, Tuapse Trough.

STRUCTURE OF THE EARTH’S CRUST

Malovitsky 1974; Belousov & Volvovsky 1989) and deep seismic reflection studies. A set of maps was constructed from these data for the main seismic and stratigraphic horizons within the sedimentary succession of the Black Sea (Tugolesov et al. 1985; Finetti et al. 1988). Thin crust of oceanic/suboceanic type was inferred below the deep-water segments (WBS and EBS) of the Black Sea Basin, with Moho shallows up to a depth of 25 km (Belousov & Volvovsky 1989). Different versions of Moho were shown by Starostenko et al. (2004). The comprehensive results of geophysical investigations of that period have been published in many monographs (Tugolesov et al. 1985; Finetti et al. 1988; Belousov & Volvovsky 1989; Volvovsky & Starostenko 1996). These are complemented by Robinson (1997), which includes a number of geological studies from the viewpoint of oil and gas perspectivity. The Black Sea constitutes only the southeasternmost part of seismic tomography models set up for Europe and therefore has limited useful seismic tomography. In order to fill this gap in studying the lithosphere below the Black Sea we have carried out a local seismic tomography study for the uppermost mantle of the Black Sea. This is complemented by new information on the crustal structure of the Black Sea Basin from the reinterpretation of existed deep seismic sounding (DSS) data using ray-tracing techniques (Yegorova et al. 2010). Interpretation of potential fields (in particular, the gravity field) is widely implemented for the Black Sea area, including three-dimensional (3D) studies incorporating calculations of the gravity signatures of different layers in the crust and upper mantle (Belousov et al. 1988; Yegorova et al. 1996; Starostenko et al. 2004). Here the results of such a gravity back-stripping analysis are integrated with other geophysical data in order to outline better the areal distribution of thin oceanic/sub-oceanic crust beneath the Black Sea. Thus new results, together with the analysis of regional seismicity, available geophysical data from previous studies, as well as geological information, including that obtained during the activity of the Middle East Basin Evolution (MEBE) Programme, allow us to make new insights on the evolution of the Black Sea region.

Main features of the sedimentary cover of the Black Sea Basin The interpretation of huge amounts of deep reflection seismic data (Tugolesov et al. 1985; Finetti et al. 1988) has allowed the compilation of a set of structure maps for the Cenozoic, including

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Anthropogene, Pliocene, Miocene, Maikopian series and Paleocene– Eocene layers (Tugolesov et al. 1985), which in part were reproduced by Starostenko et al. (2004). The lowest layer is represented by the top of the Cretaceous (Fig. 2). A spectacular feature of the eastern part of the Black Sea is the Tuapse Trough, running in a NW direction along the Georgian Black Sea coast and separated from the EBS by the Shatsky Ridge (Figs 2 & 3). The NE-elongated Gurii Trough, running onshore into the Adjaro-Trialet volcanic and flysh foldbelt (ATB in Fig. 2), lies in the south-easternmost termination of the EBS. Two narrow troughs in the WBS (the Niznekamchiisky and Sinop troughs) are characterized by a thickness of Paleocene – Eocene sediments of more than 6 km, greater than anywhere else in the Black Sea; the east –west Karkinit Trough of the northwestern shelf of the WBS is filled with sediments up to 2 km. Paleocene –Eocene sedimentation is represented by rather thick, flat, almost horizontally layered, sedimentary strata, composed by interlayered carbonates and terrigenous rocks (Tugolesov et al. 1985). In the coastal troughs of the Western Caucasus and Pontides they are represented by flysch complexes; volcanic rocks prevail in the Gurii Trough. During Maikopian time (Oligocene–Lower Miocene) a system of narrow troughs (Sorokin and Kerch-Taman Troughs) developed along the Crimea coast. This system can also be traced along the Caucasus coast into the Tuapse Basin. The main subsidence phase and infilling with sediments in these troughs, as well as in the IndoloKuban Trough, occurred in Maikopian time, when up to 5 km of sediments were deposited. The maximum thickness of Maikopian sediments (up to 5 km) is found in the WBS, whereas the EBS is characterized by a smaller thickness (up to 3.5– 4 km) and areal coverage (Fig. 3). The thickness distribution of Miocene sediments indicates a gradual change in the sedimentation environment in the Black Sea Basin: within the WBS their thickness is 1.5–2 times less than in the EBS. The features of the present single Black Sea Basin formed in Anthropogene time are shown in Figure 3 with the accumulation of thick (more than 3 km) deposits within the Danube delta in the west and in the DonKuban delta in the east, over the Kerch-Taman shelf. Reflection seismic data indicates the existence, in the deep part of the basin, of a thick (1.5–4 km) Mesozoic (and possibly older) succession below the Cenozoic strata (Tugolesov et al. 1985; Finetti et al. 1988). The Cretaceous succession is the most investigated among these. Correlation with strata on the flanks of the basin, on the Andrusov and Shatsky Ridges and within the EBS has allowed tentative age estimates to be made. The estimated thickness of Mesozoic sediments in

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Fig. 3. Cross-sections of Cenozoic sedimentary successions constructed according to the data from (Tugolesov et al. 1985).

STRUCTURE OF THE EARTH’S CRUST

the EBS is about 1.7 km (Finetti et al. 1988). The stratigraphy of the lower part of the sedimentary cover is not well constrained and there is some variation in its description and interpretation in the literature (Tugolesov et al. 1985; Finetti et al. 1988; Belousov & Volvovsky 1989; Robinson 1997). Reflection seismic data clearly show two different kinds of deformation processes in the Black Sea Basin. First, a Mesozoic extensional phase that affected the whole Black Sea, is evident in multiple normal faults (in many cases evidently of listric type deepening towards the deep basin), which imply their activity in the Late Cretaceous, frequently continuing into the Paleocene (Finetti et al. 1988). The second tectonic phase, the Cenozoic Alpine compressive phase, evident in the Shatsky Ridge, Dobrogea, Balkanides and Pontides, is related by Finetti et al. (1988) to the first orogenic movements of the Greater Caucasus. Seismic lines crossing the Greater Caucasus show that its compressive deformation continues laterally into the East Black Sea, suggesting the southward propagation of the deformation front (Finetti et al. 1988).

Gravity field of the Black Sea The gravity field of the Black Sea region (Fig. 4a) comprises contributions from various distinct tectonic units in the study region and varies widely in magnitude (from þ140 to 280 mGal). To the north slight negative gravity anomalies (up to 220 mGal) are seen over the Scythian Platform. The regional gravity field of the northeastern frame of the Black Sea is represented by negative Bouguer anomalies (up to 240 mGal) of the Fore-Caucasus region. To the east and SE of the Black Sea strong gravity lows are seen related to the Caucasus and Eastern Pontides orogens. Southwards from the arched gravity low of the Carpathians one can see (Fig. 4a) the west–east oriented strip of slight positive anomalies of Dobrogea, which, in turn, are substituted further south by the negative gravity anomalies of the Moesian Platform. Distinctive gravity feature of the shelf zones linear highs traced along the continental shelf margin (Fig. 4a). The northern one of 20 –80 mGal in amplitude goes from the northwestern shelf through the Crimean orogen up to the Caucasus coast (Fig. 4a). These anomalies have been interpreted as being due to dense mafic bodies intruded into the crust along the margins of the Black Sea (Buryanov & Soloviev 1996; Starostenko et al. 2004). However, keeping in mind that they are located on the margins of the Black Sea and that they, as a rule, are accompanied by a chain of adjacent lows placed towards the basin centre (Fig. 4a), one cannot rule out the possibility that they are

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caused, at least in part, by an edge effect at the basin margin. Against the background of negative Free-Air anomalies within the deep-water part of the Black Sea, two distinct regional highs are situated in the central parts of both the WBS and EBS sub-basins. The former (up to 20 mGal) has an oval shape and runs in a NE direction, whereas the latter (40 mGal) is linear with a NW-orientation (Fig. 4a). These two highs are separated by a wide area characterized by negative anomalies in the central part of the Black Sea. Deep sedimentary troughs along the EBS margins (Tuapse, Sorokin and Indolo-Kuban Troughs) are distinguished by gravity lows with amplitudes ranging from 280 to 260 mGal (Fig. 4a). In addition to two-dimensional (2D) gravity modelling along seismic lines, several attempts have been made to develop a 3D density model for the crust and uppermost mantle of the Black Sea (Belousov et al. 1988; Buryanov et al. 1989; Yegorova et al. 1996; Gordienko 1999; Starostenko et al. 2004). Varying in details, these models have much in common regarding the thickness of sediments in the Black Sea Basin, estimations of the position of the base of the crust (Moho), the presence of high-density thin crust of oceanic/sub-oceanic type, etc. In addition, a mantle gravity component (Buryanov et al. 1987; Belousov et al. 1988) and topography of the Moho (Starostenko et al. 2004) beneath the Black Sea have been estimated. In order to get areal signatures of thin crust of oceanic/suboceanic types we have performed a 3D gravity analysis for described above Cenozoic sedimentary strata (Tugolesov et al. 1985). Residual gravity anomalies, obtained by the removal of the total gravity effect of the seawater layer and four sedimentary layers (Anthropogene, Neogene, Maikopian series, Paleocene– Eocene), are shown in Figure 4b. Two vast high-amplitude (up to 340 mGal) residual anomalies outline the WBS and EBS. They are interpreted by Yegorova et al. (1996) and Starostenko et al. (2004) to have been caused by Moho shallowing up to the depth of 20 –25 km, less than half the average position of the Moho below the adjoining structures. The obtained sediment-free gravity field of the EBS (Fig. 4b) is similar to the Free Air anomaly pattern (Fig. 4a), but larger and more pronounced, spreading out on both sides from the axial anomaly seen in the latter. Thus, it covers the major part of the EBS from the Shatsky Ridge to the Andrusov Ridge; on the north it is cut by the Sorokin Trough low and on the SE it extends towards the Adjaro-Trialet Belt. The WBS residual anomaly seems to be composed by two branches (shown by shading in Fig. 4b) – the northern one of northwestern orientation and the southern

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Fig. 4. Gravity field of the Black Sea and surrounded regions: Free Air anomalies offshore and Bouguer anomalies onshore (a) 4 km data discretization from European database (Wybraniec et al. 1998), and residual (sediment-free) anomalies; (b) dashed line in (a) indicates 200 m isobate of the seawater bathymetry, along which stripe gravity highs are distinguished. For the case (b) calculations were made on 8  8 km grid for the model consisted from five layers: seawater ( r ¼ 1.03  103 kg/m3) and four sedimentary layers in Anthropogene ( r ¼ 2.0  103 kg/m3), Neogene ( r ¼ 2.3  103 kg/m3), Maikopian series ( r ¼ 2.6  103 kg/m3) and Paleocene –Eocene ( r ¼ 2.65  103 kg/m3) sediments. Calculations use anomalous densities derived by the ‘normalizing’ of absolute densities relative to reference density r ¼ 2.85  103 kg/m3. For the abbreviation see explanations to Figure 2; in addition BK, GA, CP indicate the Balkanides, Gubkin anomaly and Carpathian gravity low respectively. Dashed contour in (b) outlines two stripes of sediment-free anomalies within the West Black Sea Basin.

branch of west –east trend. Both branches join together in the central part of the WBS anomaly and follow further eastwards up to the western slope of the Andrusov Ridge as a wide single

anomaly zone. The WBS residual anomaly (Fig. 4b) includes the Gubkin anomaly (GA; Fig. 4), which is the most pronounced gravity high in the northwestern part of the Black Sea.

STRUCTURE OF THE EARTH’S CRUST

The northern margin of the WBS residual anomaly has a steep gradient that can be traced westwards through the Histria Basin on the Romanian shelf to onshore of the Central Dobrogea. The southern branch of the WBS residual anomaly, near its western termination at the Bulgarian shelf, coincides with known onshore structures – the Niznekamchiisky and Burgas Troughs (NT and BT in Fig. 4). To the south the WBS residual anomaly is confined by a zone of high gradients (Fig. 4b).

Magnetic field of the Black Sea The Black Sea Basin is surrounded, offshore and onshore, by tectonic units characterized by different magnetic field patterns (Fig. 5), which are ‘translucent’ through thin sedimentary cover. The magnetic anomaly zones of the highest intensity are seen in the northern parts of both the Odessa Shelf and Azov Sea, whose linear north –south trending anomalies are caused by magnetic sources of the Precambrian basement of the southernmost East European Platform (EEP) that lies to the north of the Black Sea. The southern margin of the EEP is marked by the west –east bending of the southern termination of the Odessa magnetic anomaly (anomaly 1;

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Fig. 5), which is replaced southwards by weak magnetic anomalies of the Dobrogea-Moldova zone (anomaly 2) perhaps related to Carboniferous and Permian extrusive rocks of the Moldova Trough and Palaeozoic intrusions in Northern Dobrogea (Kornev 1982). The Moesian Platform and its offshore continuation are characterized by a smoothed magnetic field, which is related by Kornev (1982) to Riphean methamorphic rock complexes of the basement. The magnetic field of the WBS is dominated by a large positive magnetic anomaly (up to 250 nT) aligned roughly in a west –east direction along the basin axis. Its central, largest amplitude, part coincides with the area of the WBS that lacks a ‘granitic’ layer (Mirlin et al. 1972; Kornev 1982; Shreider et al. 1997). It also corresponds with the central part of the residual gravity anomaly (Fig. 4b). The southern limit of the latter (red dotted line in Fig. 5) coincides with high gradients along the margin of Analtolia (Fig. 5), south of which lays a series of small magnetic highs. The MBSR is distinguished by a negative magnetic field which separates the WBS magnetic anomaly from the magnetic field of the EBS with domination of lineaments of NW orientation.

Fig. 5. Magnetic Field of the Black Sea region (portion of a magnetic field of Europe of 1:5 000 000 scale), isolines in hundred nT. Dotted red lines indicate contours of sediment-free gravity anomalies, solid red line shows the central part of the sediment-free anomaly in the East Black Sea Basin. Main regional anomalies are shown by numbers in circles: (1) Odessa anomaly; (2) Dobrogea-Moldova zone; (3) West Black Sea anomaly; (4) Anatolian anomalous zone; (5) Crimea anomaly zone; (6) Azov Sea (Indolo-Kuban Trough) anomaly; (7) Abkhaz-Kolkhida zone; (8) Adjaro-Trialet anomaly zone; (9) Alushta-Batumi maximum.

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The most spectacular feature of the magnetic field not only of the EBS, but, probably, of the whole Black Sea Basin, is the Alushta-Batumi maximum – a large anomaly (up to 100 km in width) consisting of a series of highs of 250 –400 nT magnitude, which extend in an echelon-like manner for 600 km along the Caucasus coast (anomaly 9; Fig. 5). In its central part the Alushta-Batum high covers the Shatsky Ridge. The coincidence of the southeasternmost part of Alushta-Batumi magnetic high (approximately along its zero isoline) with the northeastern edge of the residual gravity anomaly (Fig. 4b) implies that both types of anomaly (gravity and magnetic) have a common source. We assume that the lineament of tectonic origin may have played a role in the formation of the EBS. Malovitsky et al. (1972) related the Alushta-Batumi magnetic anomaly to a large fault, along which it was supposed that considerable volumes of mafic and ultramafic magma intruded into the crust in a series of parallel bodies. However, the misfit of the EBS magnetic and residual gravity anomalies (see Fig. 5) might be indicative of separate origins from sources at different depths and possibly of different ages. The activation of this fault zone may be related with (Cretaceous –Paleocene) rifting in the EBS. Susequent activity on this fault zone might be related to later (Eocene and younger) orogenic movements in the Great Caucasus. The Crimea anomaly zone (anomaly 5; Fig. 5), seen over the eastern part of the Crimean peninsula, consists of a series of local anomalies related to Middle Jurassic and Early Cretaceous volcanics by Kornev (1982). He interpreted the eastward extension of the Crimea anomaly into the Azov Sea as being caused by highly magnetized Lower Cretaceous volcanic rocks of the Indolo-Kuban Trough. The magnetic anomaly of the Indolo-Kuban Trough (anomaly 6) can be traced southeastwards along the Black Sea eastern coast into the AbkhazKolkhida anomaly zone (anomaly 7; Fig. 5), which can also be related to Middle Jurassic volcanics. The Adjaro-Trialet anomaly can be attributed to a Cenozoic volcanic complex within the southern margin of the Black Sea (Kornev 1982).

Heat flow in the Black Sea The Black Sea Basin, distinguished by very low values of surface heat flow, is surrounded from all sides by tectonic units with high to very high heatflow densities. Here we describe briefly the main features of the pattern of surface heat-flow density in the Black Sea according to the map published by Kutas et al. (1998). From the north and NW the Black Sea Basin is framed by an area characterized

by heat-flow values of 50 –60 mW m22 of Scythian Platform and 60–70 mW m22 of Crimea anomaly. From the NE and east it is surrounded by the region of Forecaucasus and northwestern Caucasus where heat-flow values reach 80 mW m22 (Kutas et al. 1998; Gordienko et al. 2002). The strongest anomaly (up to 100 mW m22) is distinguished offshore of the western extension of the Adjaro-Trialet zone (Kutas et al. 1998). From the south, the Black Sea is confined by the high heat-flow area of Eastern Pontides (80 mW m22, Hurtig et al. 1992) traced northwards offshore of the southern coastline of the EBS by a stripe of 60–70 mW m22 values (Kobzar 1987). In the western part of the Black Sea an area of low heat-flow values (20–35 mW m22) cover the major part of the WBS (Kutas et al. 1998; Verzbitsky et al. 2003; Kutas & Poort 2008). The pattern of the surface heat flow of the EBS looks more heterogeneous and is represented by a number of fragments of heat-flow values ranging from 30 to 57 mW m22 (Kobzar 1987; Kutas et al. 1998; Gordienko et al. 2002; Kutas & Poort 2008). The central part of the Black Sea is distinguished by increased (45– 60 mW m22) heat-flow values above the western slope of the MBSR (Andrusov Ridge). Low heat-flow density in the Black Sea Basin is usually attributed to the very thick Cenozoic sedimentary succession, which can reduce by half the heat flow from the basement. However, both Kutas et al. (1998) and Galushkin et al. (2006) show that the very low surface heat-flow observations of the Black Sea cannot be explained by thick sediments only. After correcting for their effects, the basement heat flow increases to 40 –45 mW m22 and 60 – 70 mW m22 in the deep basins and the central part of the Black Sea respectively (Kobzar 1987; Kutas et al. 1998; Gordienko et al. 2002). The lowest values of basement heat flow (30 mW m22), observed in the WBS, are interpreted to be indicative of the presence of a cold lithospheric domain. Variations of the heat flow in the Black Sea imply distinctions in the temperature regimes of underlying lithospheric domains. Various authors have attempted to determine the age of the Black Sea using the geothermal data. But the significant lateral heterogeneity in observations, the great thicknesses of sediments and the considerable ambiguity of most parameters involved in such modelling, has led to contradictory results. The age of basin formation estimated this way varies from 60 –70 Ma for the whole basin (Verzbitsky et al. 2003) to 130–150 Ma for the EBS (Kutas et al. 1998); 70 –90 Ma age was inferred for the WBS (Kutas et al. 1998). Large variations for lithosphere thickness – from 60 to 65 km for the whole Black Sea (Verzbitsky et al. 2003; Galushkin et al. 2006) to 80–100 km and 120–140 km for the EBS

STRUCTURE OF THE EARTH’S CRUST

and WBS respectively (Kobzar 1987) were also reported.

Structure of the crust below the West- and East-Black Sea Basins according to existing deep seismic sounding (DSS) data The structure of the crystalline crust beneath the Black Sea has been investigated along 21 deep seismic sounding (DSS) profiles, with total length 3800 km, surveys carried out during the 1960s, 1970s and 1980s. The locations of these profiles can be found in Starostenko et al. (2004). The first interpretations of these data revealed three velocity layers in the sedimentary cover of the Black Sea (Neprochnov et al. 1970; Malovitsky & Neprochnov 1972; Moskalenko & Malovitsky 1974). The two deep sub-basins (WBS and the EBS), separated by the MBSR, as described above (Figs 1 & 2), were also revealed by these data. According to the modelled DSS data, the two deep sub-basins are underlain by very a thin high-velocity crystalline crust with the Moho shallowing to a depth of 25 km (Neprochnov et al. 1970; Belousov & Volvovsky 1989). Shallower depths (20 and 22 km) for the Moho beneath the WBS and

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EBS were inferred from the 3D gravity study of Starostenko et al. (2004). Yegorova et al. (2010) reinterpreted some of these DSS data using a ray-tracing technique: two north –south profiles crossing the WBS (Profile 25) and Azov Sea and central part of the Black Sea (Profile 28/29). Their interpretation confirms the presence of two major crustal domains in the Black Sea (oceanic crust below its deep-water part and continental crust beneath the shelf zones) and proposed a crustal velocity structure for each profile (Fig. 6). A key feature of the velocity model of the WBS (seismic line 25, Fig. 6a) is what is interpreted as a high-amplitude normal fault formed at the edge of the continental margin (Scythian Platform), probably a result of rift-related extension in the WBS (Yegorova et al. 2010). The velocity model constructed for the Azov Sea and central part of the Black Sea (Profile 28/29, Fig. 6b) shows, from north to south, thick continental crust of the southern margin of the East European and Scythian platforms below the northern part of Azov Sea, oceanic crust below the EBS of the same type as that below the WBS. The crust of the MSBR represents a bridge of thinned (28 –30 km) continental crust consisting of a series of three blocks, arranged in an en echelon manner, the

Fig. 6. Velocity models along two sections: in the western part of the basin, pr. 25 (a) and along the line crossing the Azov Sea and Central part of the Black Sea, profiles 29 and 28; (b) locations of the profiles are shown in Figure 1. Velocities are in km s21.

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basement of which rises southwards from 10–11 km on the Andrusov Ridge to 6 km on the Arhangelsky Ridge (Fig. 6b).

Seismicity of the Black Sea region The Black Sea is surrounded by young tectonic units displaying active seismicity. Several seismic zones (the Crimea –Caucasus, North Anatolian and Marmora Sea zones) have been defined on the basis of earthquake data collected since 1927 (Chekunov et al. 1996). The strongest earthquakes occur along the southwestern margin of the Black Sea within the North Anatolian zone, where sources with magnitude larger than 7 are common (Chekunov et al. 1996). In order to use the seismicity data for developing a seismic tomography model, 89 earthquakes that occurred within the area limited by 40 –488N latitude and 26–428E longitude between 1970 and 2006, and recorded by seismic stations around the Black Sea (Fig. 7a), have been relocated. The data were taken not only from International Seismological Centre (ISC) Bulletins, but also from bulletins of the regional network of Crimea seismic stations (Pustovitenko 1991, 2005), internet databases (http://www.isc.ac.uk, http://www.sismo. deprem.gov.tr) and other unpublished data from separate local seismic stations. The relocation procedure is described in more detail in Gobarenko & Yegorova (2008). The results are shown in Figure 7b. According to the revised data, the earthquakes recorded within the deep-water part of the WBS have magnitudes M ¼ 3 –4.5 and have focal depths in the lower crust at depths of 15–32 km. The southern peripheral part of the WBS is characterized by moderate seismicity, marked by several shallow earthquakes of low magnitude. Seismicity strongly increases towards the North-Anatolian fault. Near the southern margin of the EEP a number of shallow (0 –15 km) low-magnitude (M , 4) earthquakes were recorded, while at the boundary between the WBS and the Scythian Platform they are deeper in the lower crust– upper mantle. No one earthquake with magnitude M . 3 has been reliably recorded inside the EBS, which is surrounded by zones of high and increased seismicity along the Crimea – Caucasus coast, MBSR and the southern coast. The earthquakes of offshore Crimea occur along a NE-trending Southern Coast fault. Along this fault, detected from DSS studies (Kharechko 1985; Smolyaninova et al. 1996), crustal structure changes dramatically from 25– 30 km thickness below the EBS to 48 –50 km under the Crimean orogen (cf. Starostenko et al.

2004). Here, earthquakes with magnitude M ¼ 4 – 5 and focal depths in the lower crust and upper mantle (H ¼ 20 –46 km) are localized along the northern border of Sorokin Trough, which is perpendicular to the axis of the EBS (Fig. 7b). Focal mechanisms indicate compression directed from the Black Sea towards Crimea (Smolyaninova et al. 1996). In addition, earthquakes related to normal faults and shear zones are also distinguished within this zone. The strongest earthquake (M ¼ 6.8) occurred in 1927 near Yalta on the Crimean coast and was extensional (Pustovitenko et al. 1989). The seismic zones along the Kerch-Taman coast and near the Caucasus coast are characterized by stronger earthquakes (M . 5) with foci in the lower crust and upper mantle (H ¼ 29 –38 km). Focal mechanisms indicate sub-horizontal NE – SW compressional deformations on thrust and strike –slip faults (Pustovitenko 2002). The MBSR is clearly distinguished by a series of NW– SE trending earthquakes with foci either in the upper crust (H , 10 km) or near the crust –mantle boundary (Fig. 7). The southern and southeastern parts of the EBS are distinguished by relatively minor seismicity, with sources in the lower crust and at the crust –mantle boundary.

Seismic tomography study of the Black Sea lithosphere The Black Sea region appears only as a small part of the region covered by published regional seismic tomography models of Europe, the Mediterranean and adjoining parts of Asia (e.g. Spakman et al. 1993; Bijwaard et al. 1998; Marquering & Snieder 1996). The first two of these present P-velocity models from travel–time residuals, while the last shows a 3D S-velocity model obtained by seismic waveform inversion. The distinctions in the P-wave structure of the upper mantle beneath the WBS and EBS are distinguished down to the depth of 150 km (Spakman et al. 1993; Bijwaard et al. 1998; Gobarenko & Nesterov 1996). Figure 8 shows high-velocity anomalies in the WBS, whereas the EBS is covered by an area of negative anomalies. A comparison of the P-wave distribution (Spakman et al. 1993; Bijwaad et al. 1998) with the S-wave pattern (Marquering & Snieder 1996) shows their certain agreement for the WBS and totally different situation in the EBS, where low P-wave velocities contradict to distinct high S-wave velocity anomaly. Available seismic tomography studies show that maximal disagreement between the P-wave (Spakman et al. 1993; Bijwaard et al. 1998) and S-wave (Marquering & Snieder 1996)

STRUCTURE OF THE EARTH’S CRUST

33

Fig. 7. Map of seismicity of the Black Sea region, from (a) ISC data and (b) with corrected parameters (see text), their magnitude and focal depths are shown against the background of the map for top of Cretaceous (Fig. 2). The second data set was used for seismic tomography study (Fig. 9). Black triangles shows seismic stations that recorded the earthquakes. For abbreviations see Figure 2.

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Fig. 8. Black Sea segment of European scale P-wave seismic tomography model at the depth of 95 km (Bijwaard et al. 1998).

models occurs in the depth range of 50 –100 km. This serves the ground for additional study of this interval, undertaken by Gobarenko & Yegorova (2008). The seismic tomography, applied for studying the velocity structure of the Black Sea lithosphere, is based on the Backus –Gilbert approach for 3D cases (Gobarenko et al. 1987). The dataset includes the travel times data of the Pn waves, recorded by the network of seismic stations around the Black Sea (shown in Fig. 7a) from the earthquakes that occurred within the study region. For that purpose only the earthquakes, shown in Fig. 7b, with corrected seismic parameters due to applied relocation procedure were used. In more detail the methodology of applied seismic tomography and technical parameters are presented by Gobarenko & Yegorova (2008). The derived velocity model is shown in Figure 9 by a set of horizontal slices at depths of 40, 45 and 50 km. Lithosphere structure of the WBS at 40– 50 km depth looks rather heterogeneous. Two main domains can be seen here on the upper sections (Fig. 9a, b). First the domain includes an area of high velocities (Vp ¼ 7.9– 8.0 km s21) in western and southwestern parts of the WBS, as well as in the area of Burgas-Niznekamchiisky graben. This domain widens with depth and covers at 50 km

depth the major part of the WBS (Fig. 9c). Second there is a low-velocity domain, covering the northwestern (Odessa) shelf and continental slope at the upper section (Fig. 9a), practically disappears at the depth of 50 km. It includes a linear low-velocity anomaly developing along the northern border of the WBS. Previous study of surface wave azimuth anomalies (Nesterov & Yanovskaya 1988) confirms the existence of low-velocity domain here, which is bent around by surface waves. A number of earthquakes with foci in the crust– mantle transition zone have been recorded in the vicinity of the contact zone of these high- and low-velocity domains of the WBS. Figure 9a shows that at the depth of 40 km the eastern part of the Black Sea Basin is surrounded from the north and east by zones of high-velocity gradient along the Crimea–Caucasus coast, which at greater depths are smoothed out. This highgradient zone confines from the south a low-velocity zone in Figure 9a along the southern Crimea coast (covering the Sorokin Trough and Kerch-Taman Trough), which is explained by a dramatic change in the crustal structure in the transition from the thick continental crust of the Crimean orogen to the thin oceanic/suboceanic crust of the EBS. A spectacular feature of the eastern part of the Black Sea is extended in a NW-direction high-velocity

STRUCTURE OF THE EARTH’S CRUST

Fig. 9. P-wave velocity model constructed by seismic tomography using the data shown in Figure 7b.

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anomaly representing the uppermost mantle below the EBS, which has maximal coverage at 40 km depth (Fig. 9a). A complicated and ambiguous velocity structure was determined below the central part of the Black Sea Basin. A NW–SE low-velocity anomaly with high gradient zone concurring with the Andrusov Ridge is clearly seen at 40 km depth (Fig. 9a). To the west from this low-velocity zone, and in parallel to it, a chain of small high-velocity anomalies have been distinguished (Vp up to 8.0 km s21) traced in a NW direction from the Odessa Shelf to the Central Pontides. This zone of high-velocity anomalies is also distinguished by earthquakes with foci in the uppermost mantle. These features are indicative of the existence of an active tectonic zone separating two lithospheric domains of the Black Sea (the WBS and the EBS). At depths of 45 and 50 km the central part of the basin is featured by a wide (nearly 250 km) lowvelocity zone. The Central Pontides are distinguished by low velocities against the background of high-velocity domains of the Western and Eastern Pontides. Differences in the lithosphere structure below two parts of the Black Sea Basin can be seen in the Figure 10a for dependence of travel–time residuals dtp from the depth of seismic ray dipping, where two sub-basins are featured by different trend of the dtp alteration. These data show that in the WBS the Vp increases with depth, while the EBS, on the contrary, is distinguished by distinct velocity decrease (Fig. 10a). In addition, a combined velocity section for the depth range of 40–200 km has been constructed for both basin parts (Fig. 10b) using the data from previous Black Sea seismic tomography study (Gobarenko & Nesterov 1996).

Discussion and interpretation of the results Recent tectonic setting of the Black Sea is determined by an ensemble of lithospheric plates and microplates of different affinities (fragments of Gondwana and Laurussia), which were amalgamated together by accretion due to closure of the Neotethys Ocean. The WBS constitutes a wide (up to 300 km width and 600 km length) west– east trending sedimentary basin filled with Cretaceous and younger sediments as thick as 14 km. The crust below the WBS belongs to oceanic or suboceanic type with very thick sedimentary strata. Areal coverage of thin crust of oceanic and transition type is contoured by sediment-free gravity anomalies (Fig. 4b), which are caused by dramatic uplift, almost two times, of the Moho below the

Fig. 10. Distribution of time residuals (a) and velocities (b) of P-wave with depth in the western (WBS) and eastern (EBS) parts of the basin.

central part of the Black Sea. Coincidence of west –east oriented WBS magnetic anomaly with central part of the sediment-free gravity anomaly implies that a considerable part of the oceanic crust here has not only high P-wave velocity and density, but also substantial magnetization. As follows from re-evaluation of the DSS data on profile 25 (Yegorova et al. 2010), the crust of the WBS consists mainly of sediments with Vp ¼ 3–4.2 km s21, and thin (about 5 km) highvelocity (Vp ¼ 6.5– 7.2 km s21) crystalline crust underlain by the Moho boundary at the depth of 19 –20 km (Fig. 6a). The velocity model on this profile can be interpreted as showing that rifting occurred in the WBS along the suture between the Moesian and Scythian platforms (Fig. 11). The highamplitude normal fault (half-graben) revealed at the edge of continental margin, belongs to an area of increased seismicity with the earthquake magnitude M ¼ 3 –5 and foci in the lower crust –upper mantle (Fig. 7). This rift-related fault and associated basin structure can be traced westwards through the Histria basin, located on the Romanian shelf of the Black Sea, to the area of NW-oriented

STRUCTURE OF THE EARTH’S CRUST

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Fig. 11. Tectonic scheme of the Black Sea area (after Robinson et al. 1995 with modifications). Red lines indicate seismic profiles of DSS study, shown in Figure 6. Abbreviations COF and PCF shows location of the fault zones confining the Central Dobrogea–Capidava-Ovidiu and Pechenega-Camena fault zones respectively.

Pechenega-Camena and Capidava-Ovidiu Fault zones, which outline the Central Dobrogea and separate the Northern and Southern Dobrogea (Fig. 11). It is generally assumed that in Dobrogea the Pechenega-Camena Fault represents the TornquistTeisseyre Line separating the Moesian Platform from Scythian Platform (Sa˘ndulesku 1978, 1980; Visarion et al. 1990). The Pechenaga-Camena Fault and other NW-trending crustal faults of Dobrogea (Capidava-Ovidiu Fault, Intra-Moesian) played an important role in strike –slip tectonics related to the opening of the western part of the Black Sea (Gradinaru 1984; Sa˘ndulesku 1994; Seghedi & Oaie 1994; Banks & Robinson 1997; Hippolite 2002). Saintot et al. (2006a) consider that the best candidate for the suture zone between the Moesia and the East European platform will be the Capidava-Ovidiu Fault (separating the Moesia plate in two blocks – the Central Dobrogea and the Southern Dobrogea) thus corresponding to the Trans-European Suture Zone. The WBS is thought to form on the crust of the Moesian Platform due to rifting that occurred in Cretaceous (Zonenshain & LePichon 1988; Finetti

et al. 1988; Go¨ru¨r 1988; Okay et al. 1994; Robinson et al. 1996; Banks & Robinson 1997). Rifting may have been localized along the suture zones, formed in the Mesozoic due to accretion/collision of the Moesian Platform with adjoining terranes. The latter include the Scythian Platform on the north and Istanbul and Sakarya zones on the south (Golonka 2004; Saintot et al. 2006a). The lithosphere of the WBS, in comparison with the EBS, is distinguished by higher velocities (Figs 9 & 10; Spakman et al. 1993; Gobarenko & Nesterov 1996; Bijwaard et al. 1998), related probably to higher lithosphere rigidity in the western part of the Black Sea, thus implying the presence of domains of different affinity below the two subbasins. On the other hand, this can also be interpreted in terms of lithosphere thickness, implying thicker lithosphere below the WBS relative to the EBS. This corresponds with the basin modelling of Spadini et al. (1996, 1997), who predict that a cold, thick (200 km) lithosphere was initiated below the West Black Sea, while the Eastern Black Sea is underlain by warm and thinner lithosphere. The existence of two different entities in

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the Black Sea also corresponds with the different types of seismic activity observed in the two parts of the Black Sea. The EBS is characterized by the lack of the earthquakes, whereas in the WBS a number of relatively strong (M  5) earthquakes, with foci reaching the uppermost mantle, have been recorded. A rheologically strong continental lithosphere below the back-arc Black Sea basin is predicted by Stephenson & Schellart (2010) in their geodynamic study of modern analogues. The oceanic domain of the WBS is separated from the EBS by the MBSR – a narrow (80 km width) NW stripe of thin continental crust underlain by the Moho approximately at 29 km depth, which like a bridge joins the continental crust of the Scythian Platform on the north and that of the Pontides on the south (Starostenko et al. 2004; Shillington et al. 2009; Yegorova et al. 2010). The MBSR is represented by an echelon-like system of several (three) uplifted blocks (including the Andrusov and Arkhangelsky Ridges) separated by horst and grabens. Recent tectonic activity within the MBSR is confirmed by earthquakes with M ¼ 4 –5 and foci in the lower crust –upper mantle transition (Fig. 7). The EBS represents narrow NW-trending linear trough underlain by oceanic/sub-oceanic crust, which is squeezed between the continental domains of the MBSR and Shatsky Ridge (Fig. 11). The thickness of sediments in the EBS is 11 –12 km including the Cenozoic successions as thick as 8–10 km (Belousov & Volvovsky 1989; Tugolesov et al. 1985; Yegorova et al. 2010) and several km (up to 3 km) of Mesozoic sediments. Occurrence of Mesozoic sediments in the EBS and surrounding structures is a specific feature of eastern part of the Black Sea (Tugolesov et al. 1985; Finetti et al. 1988). Geological studies and palaeotectonic reconstructions show that the Transcaucasus domain can be a good candidate for the lithospheric block on which the EBS was originated (Adamia et al. 1974; Golonka 2004; Saintot et al. 2006b) during the Latest Cretaceous and the Late Eocene (Finetti et al. 1988; Robinson et al. 1996; Shreider et al. 1997; Golonka 2004), approximately at the same time as the WBS opening. We suggest that the probable mechanism responsible for the opening of the EBS could be strike –slip movements along the MBSR (Yegorova et al. 2010) initiated by moving the Arabian Plate northwards. Genetic linkage of the EBS with the Transcaucasus structures is emphasized by the eastward extension of the EBS into the Adjaro-Trialet Fold Belt (Adamia et al. 1974; Chekunov 1989; Shcherba 1994; Banks et al. 1997). Further moving northwards of Arabia (in Eocene) initiated the first pulses of orogeny in the Caucasus (Ershov et al. 2003) and oblique

rifting in the EBS, which led to ‘basification’ of continental crust of the western part of Trancaucasus domain. Northward spreading of narrow wedge of the EBS oceanic crust terminates abruptly at the edge of the Sorokin Trough developed offshore along the southern Crimea coast (Fig. 11). This trough, together with its northeastern extension into the Kerch-Taman Trough, constitute a system of deep narrow troughs filled with thick Cenozoic sediments (more than 10 km), major part of which belongs to Maikopian strata (Fig. 3). The timing of such events as the formation of the Sorokin and Kerch-Taman Troughs along the Crimea coast and Tuapse Basin along the Caucasus coast in Maikopian time (Oligocene– Lower Miocene), commencement of orogenic pulses in the Caucasus (Eocene) and oblique rifting in the EBS might be indicative of their causal relationship. Therefore a system of linear troughs along the Crimea coast (Sorokin and Kerch-Taman Troughs) might be accretional in origin (Fig. 11) formed by underthrusting of the EBS oceanic plate below the continental Scythian plate. Such an idea was forwarded by Chekunov et al. (1996) in order to explain the origin of the Crimea orogen. This ongoing process is seen in the high seismicity observed offshore along the southeastern Crimea coast where earthquakes with M ¼ 3– 5 and foci in the lower crust –upper mantle have been recorded (Fig. 7). This Crimea zone of high seismicity is traced further southeastwards along the Caucasus coast (Fig. 7). Here, in 1996, near Anapa a strong earthquake with the source in the upper mantle (H ¼ 55 km) and M  6 has been recorded (see also Verzhbitsky et al. 2003). This may point to the Caucasus coastline of distinct NW orientation, which marks the boundary between the Shatsky Ridge offshore and Greater Caucasus onshore, being of tectonic origin similar to that of the Crimea coast. Thus, the zone of high seismicity along the Crimea–Caucasus coast marks an active collision-accretion zone developed by the ensemble of microplates moving northward, including the oceanic EBS domain and continental one of the Shatsky Ridge, below the Scythian Platform and Transcaucasus domain. In contemporary palaeotectonic maps this zone is very often shown as a subduction zone. For example, in the recent publication of Golonka (2004) the Crimea–Caucasus subduction zone, extended southeastwards to the South Caspian, is shown in all Palaeogene and Neogene maps [figs 14– 18 in Golonka (2004)]. Our results definitely indicate that collision – accretion processes along the Crimea–Caucasus coast have been advanced now, at least up to the stage of underthrusting the crust of the East Black

STRUCTURE OF THE EARTH’S CRUST

Sea and Shatsky Ridge below the Scythian Plate and Transcaucasus domain. This was initiated by recent geodynamics of African and Arabian Plates (Reilinger et al. 1997). The velocities of the Arabian Plate are three times more that of the African Plate, reaching 18 and 6 mm/a respectively (LePichon et al. 1995; McClusky et al. 2000). In particular the high geodynamic activity of the Arabian Plate, most likely, was the cause of strengthening the collisional processes in the eastern part of the Black Sea– South Caspian region. As a probable tectonic consequence of this northward movement, in Paleocene the projection of the South-Armenian block wedged into the Transcaucasus basin leading to its splitting into two parts, that initiated the formation of two separate basins – the East Black Sea and South Caspian (Brunet et al. 2003; Golonka 2004; Sebrier pers. comm. 2006). The Transcaucasus sedimentary basin existed in the Jurassic and Cretaceous and extended from Dobrogea through the Black Sea and South-Caspian Depression up to the Kopet-Dag (Khain 1984; Zonenshain & LePichon 1986; Banks et al. 1997; Shreider et al. 1997; Brunet et al. 2003; Golonka 2004).

Summary and conclusions This paper represents a comprehensive analysis of available geological and geophysical information on the Black Sea region accomplished for a better understanding of the basin architecture and lithosphere structure, as well as operated tectonic processes. The analysis also incorporates new data obtained from 3D gravity analysis, local seismic tomography and re-interpretation of existed deep seismic refraction data. The latter are presented in (Yegorova et al. 2010). Practically all of the parameters of the basin architecture and the lithosphere structure, as well as orientation of the main tectonic units and configuration of the Black Sea Basin, show existence of two deep sub-basins in the Black Sea. Local seismic tomography shows that West- and EastBlack Sea Basins have different velocity structure of the lithosphere. These distinctions, traced at least down to the depth of 150 km, could be explained by different thickness of the lithosphere and/or origination of two sub-basins on microplates of different affinities. The high-velocity lithosphere of the WBS could be inferred from its origination on the basement of Moesian Platform due to rifting occurred along the Mesozoic sutures. In this regard the velocity decrease at the depth range of 40 to 45 km relates, most probably, to modification of the matter of the lower crust of Moesian Platform in the course of formation of oceanic crust of the WBS.

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In contrast to the WBS, the EBS is underlain by a lithosphere of lower P-wave velocities (in general on 0.015 km s21) that might be indicative on its origination on different (Trancaucasus) domain and/or on increased temperature regime of the upper mantle. In such a way its development is determined mainly by tectonic evolution of the Caucasus region, which in turn is caused by recent geodynamics of the Arabian Plate. The occurrence of two lithospheric blocks with different rheological properties under the Black Sea corresponds to different character of observed seismicity and different types of unloading the accumulated strength. Accumulated in the EBS, strength is unloaded along the northern seismic zone, distinguished along the Crimea–Caucasus coast, while in the WBS the major seismogenic zone locates along the southern margin within the North-Anatolian seismic zone. A high-seismicity zone along the coast of Southern Crimea and Western Caucasus might be indicative of active underthrusting of the EBS oceanic plate below the Scythian Platform and confirms the accretional origin of Sorokin and Kerch-Taman Troughs. This paper highlights a number of topics, which require further special investigation. Among them the most interesting are supposed to be the study of the cause of extremely low heat flow in the Black Sea, the origin of the giant Alushta-Batumi magnetic anomaly and the reason for strong distinctions between P- and S-wave structures of the Black Sea lithosphere. The scientific investigations leading to this paper were undertaken within the framework of the MEBE programme. We extend our appreciation to the referees, Dr J. Mosar (University of Fribourg, Switzerland) and A. Saintot (Norwegian Geological Survey, Trondheim), whose suggestions were used in improving the manuscript. We thank also R. A. Stephenson for the fruitful discussions and for his help in editing the paper and improving the English.

References Adamia, S., Gamkrelidze, I. P., Zakariadze, G. S. & Lordkipanidze, M. V. 1974. Adjara-Trialeti trough and the problem of the Black Sea origin. Geotectonika, 1, 78–94 (in Russian). Banks, C. J. & Robinson, A. 1997. Mesozoic strike-slip back-arc basins of the western Black Sea region. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 53–62. Banks, C. J., Robinson, A. G. & Williams, M. P. 1997. Structure and regional tectonics of the Achara-Trialet Fold Belt and the adjacent Rioni and Kartli foreland basins, Republic of Georgia. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 331– 345.

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Belousov, V. V. & Volvovsky, B. S. (eds) 1989. Structure and Evolution of the Earth’s Crust and Upper Mantle of the Black Sea. Nauka, Moscow (in Russian). Belousov, V. V., Volvovsky, B. S. et al. 1988. Structure and evolution of the Earth’s crust and upper mantle of the Black Sea. Bolletino di Geofisica Teorica ed Applicata, XXX/117-188, 109–196. Bijwaard, H., Spakman, W. & Engdahl, E. 1998. Closing the gap between regional and global travel time tomography. Journal of Geophysical Research, 103(B12), doi: 10.1029/98JB02467. Brunet, M.-F., Korotaev, M. V., Ershov, A. V. & Nikishin, A. M. 2003. The South-Caspian Basin: a review of its evolution from subsidence modelling. Sedimentary Geology, 156, 119–148. Buryanov, V. B. & Soloviev, V. D. 1996. Study of gravity maxima of the northern and western frame of the Black Sea Depression. In: Volvosky, B. S. & Starostenko, V. I. (eds) Geophysical Parameters of the Lithosphere of the Southern Sector of the Alpine Orogen. Naukova Dumka, Kiev, 111– 119 (in Russian). Buryanov, V. B., Gordienko, V. V. & Rusakov, O. M. 1987. Heterogeneities of the tectonosphere of the Black Sea from geophysical fields study. Geophysical Journal, 3, 34–44 (in Russian). Buryanov, V. B., Gordienko, V. V., Rusakov, O. M., Soloviev, V. D. & Starostenko, V. I. 1989. Gravity field. Geophysical characterization of the Earth’s crust. In: Volvovsky, B. S. & Belousov, V. V. (eds) Structure and Evolution of the Earth’s Crust and Upper Mantle of the Black Sea. Nauka, Moscow, 83– 90 (in Russian). Chekunov, A. V. 1989. Geological problems of the Black Sea basin. In: Belousov, V. V. & Volvovsky, B. S. (eds) Structure and Evolution of the Earth’s Crust and Upper Mantle of the Black Sea. Nauka, Moscow, 110– 120 (in Russian). Chekunov, A. V., Pustovitenko, B. G. & Kulchitsky, V. E. 1996. Depth distribution of foci and energy. Seismicity of the Black Sea region. In: Volvovsky, B. S. & Starostenko, V. I. (eds) Geophysical Parameters of the Lithosphere of the Southern Sector of Alpine Orogen. Naukova Dumka, Kiev, 101–106 (in Russian) Dercourt, J., Gaetani, M. et al. (eds) 1993. Atlas Tethys, Palaeoenviromental Maps. Gauthier-Villars, Paris. Ershov, A. V., Brunet, M.-F., Nikishin, A. M., Bolotov, S. N., Nazarevich, B. P. & Korotaev, M. V. 2003. Northern Caucasus basin: thermal history and synthesis of subsidence models. Sedimentary Geology, 156, 95–118. Finetti, I., Bricchi, G., DEL Ben, A., Pipan, M. & Xuan, Z. 1988. Geophysical study of the Black Sea. Bolletino di Geofisica Teorica ed Applicata, XXX/ 117-118, 197–324. Galushkin, Y. I., Shreider, A. A., Bulychev, A. A. & Shreider, Al. A. 2006. Heat flow and thermal evolution of the lithosphere of the Black Sea basin. Oceanology, 2, 296– 314. Gobarenko, V. S. & Nesterov, A. N. 1996. Volume distribution of velocity parameters of waves from the earthquakes in the lithosphere. Seismicity

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STRUCTURE OF THE EARTH’S CRUST Basin according to hydromagnetic survey. Marine Geology and Geophysics. ONTI VIEMS, Moscow, 3, 12– 21 (in Russian). Marquering, H. & Snieder, R. 1996. Shear-wave velocity structure beneath Europe, the northeastern Atlantic and western Asia from waveform inversions including surface-wave mode coupling. Geophysical Journal International, 127, 283–304. McClusky, S., Balassanian, S. et al. 2000. Global positioning system constraints on plate kinematics and dynamics in the eastern Mediterranean and Caucasus. Journal of Geophysical Research, 105(B3), 5695–5719. Mirlin, E. G., Melikhov, V. R., Mikhailov, O. V. & Neprochnov, Y. P. 1972. On the nature of the magnetic anomalies in the Black Sea Basin. Izvestiya of Academy of Sciences of the USSR, Seriya Geologischeskya, 9, 60–71 (in Russian). Moskalenko, V. N. & Malovitsky, Y. P. 1974. Results of deep seismic sounding on the transmeridional profile along the Azov and Black Seas. Izvestiya of Academy of Sciences of the USSR, Seriya Geologischeskya, 9, 23– 31 (in Russian). Neprochnov, Y. P., Kosminskaya, I. P. & Malovitsky, Ya. P. 1970. Structure of the crust and upper mantle of the Black and Caspian Seas. Tectonophysics, 10, 517–538. Nesterov, A. N. & Yanovskaya, T. B. 1988. Horizontal heterogeneities in the lithosphere of southeastern Europe according to surface wave observations. Izvestiya of Akademy of Sciences of the USSR, Fizika Zemli, 11, 3 –15 (in Russian). Nikishin, A. M., Korotaev, A. M., Ershov, A. V. & Brunet, M.-F. 2003. The Black Sea basin: tectonic history and Neogene-Quaternary rapid subsidence modelling, Sedimentary Geology, 156, 149–168. Okay, A. I., S¸engo¨r, A. M. C. & Go¨ru¨r, N. 1994. Kinematic history of the opening of the Black Sea and its effect on the surrounding regions. Geology, 22, 267–270. Pease, P., Daly, J. S. et al. 2008. Baltica in the Cryogenian, 850–630 Ma. In: Bogdanova, S. V., Li, X. X., Moores, E. M. & Pisarevsky, S. A. (eds) Testing the Rodinia Hypothesis: Records in its Building Blocks, Precambrian Research, 160, 46–65. Pustovitenko, B. G. (ed.) 1991. Seismological Bulletin for the Western ESSN Zone (Crimea-Carpathians), 1970–1990. Naukova Dumka, Kiev (in Russian). Pustovitenko, B. G. 2002. Mechanism of notable earthquakes of Crimea-Black Sea region of last 20 years. In: Pustovitenko, B. G. (ed.) Seismological Bulletin of Ukraine for the 2000 Year. EKOSIGydrofizika, Sevastopol, 59–64 (in Russian). Pustovitenko, B. G. (ed.) 2005. Seismological Bulletin of Ukraine, 1991– 2004. Simferopol (in Russian). Pustovitenko, B. G., Kulchitsky, V. E. & Goryachun, A. V. 1989. The Earthquakes of Crimea-Black Sea Region. Naukova Dumka, Kiev (in Russian). Reilinger, R. E., McClusky, S. L. et al. 1997. Global positioning system measurements of resent-day crustal movements in the Arabia-Africa-Eurasia plate collision zone. Journal of Geophysycal Research, 102, 9983–9999.

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Robinson, A. G. (ed.) 1997. Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68. Robinson, A. G., Rudat, J. H., Banks, C. J. & Wiles, R. L. F. 1996. Petroleum geology of the Black Sea. Marine and Petroleum Geology, 13, 195– 223. Saintot, A., Stephenson, R., Stovba, S., Brunet, M.-F., Yegorova, T. & Starostenko, V. 2006a. The evolution of the southern margin of the Eastern Europe (Eastern European and Scythian platforms) from latest Precambrian-Early Palaeozoic to the Early Cretaceous. In: Gee, D. & Stephenson, R. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 481 –505. Saintot, A., Brunet, M.-F. et al. 2006b. The MesozoicCenozoic evolution of the Greater Caucasus. In: Gee, D. & Stephenson, R. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 277– 289. Sa˘ndulesku, M. 1978. The Moesian Platform and the North Dobrogean orogen. In: Lemoine, M. (ed.) Geological Atlas of Alpine Europe and Adjoining Areas. Elsevier, Amsterdam, 427– 460. Sa˘ndulesku, M. 1980. Analyse ge´otectonique des chaı˘nes alpines situe´es autor de la Mer Noire occidentale. Ann. Inst. Geol. Geofiz., 56, 5 –54. Sa˘ndulesku, M. 1994. Overview on Romanian geology. In: Berza, T. (ed.) Alcapa II Field Guidebook: Geological Evolution of the Alpine-CarpathianPannonian System. Romanian Journal of Tectonics and Regional Geology, 74, 3– 15. Seghedi, A. & Oaie, G. 1994. Tectonic setting of two contrasting types of pre-alpine basement: North versus Central Dobrogea. Romanian Journal of Tectonics and Regional Geology, 75, 56–57. Shcherba, I. G. 1994. Paleogene basin of the Caucasus. Bulletin Moskovskogo Obshchestva Ispytatelei Prirody, otdel Geologicheskiy, 69, 71–80 (in Russian). Shillington, D. J., Scott, C. L., Minshull, T. A., Edwards, R. A., Brown, P. J. & White, N. 2009. Abrupt transition from magma-starved to magma-rich rifting in the eastern Black Sea. Geology, 37, 7 –10. Shreider, A. A., Kazmin, V. G. & Lugin, V. S. 1997. Magnetic anomalies and problem of the age of Black Sea. Geotectonika, 1, 59–70 (in Russian). Smolyaninova, E. I., Mikhailov, V. O. & Lyakhovsky, V. A. 1996. Numerical modelling of regional neotectonic movements in the northern Black Sea. Tectonophysics, 266, 221– 231. Spadini, G., Robinson, A. & Cloetinhg, S. 1996. Western versus Eastern Black Sea tectonic evolution: pre-rift lithospheric controls on basin formation. Tectonophysics, 266, 139– 154. Spadini, G., Robinson, A. G. & Cloetingh, S. A. P. L. 1997. Thermomechanical modelling of Black Sea basin formation, subsidence, and sedimentation. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 19–38. Spakman, W., Van Der Lee, S. & Van Der Hilst, R. 1993. Travel-time tomography of the EuropeanMediterranean mantle down to 1400 km. Physics of the Earth and Planetary Interiors, 79, 3– 74.

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Starostenko, V., Buryanov, V. et al. 2004. Topography of the crust-mantle boundary beneath the Black Sea Basin. Tectonophysics, 381, 211– 233. Stephenson, R. & Schellart, W. P. 2010. The Black Sea back-arc basin: insights on its origin from geodynamic models of modern analogues. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 11–21. Tugolesov, D. A., Gorshkov, A. S. et al. 1985. Tectonics of Mesozoic-Cenozoic Successions of the Black Sea Basin. Nedra, Moscow (in Russian). Verzhbitsky, E., Kuzin, I. & Lobkovsky, L. 2003. Age and thickness of the lithosphere within the Western and Eastern basins of the Black Sea according to geophysical data. Turkish Journal of Earth Sciences, 11, 231–242. Visarion, M., Sa˘ndulesku, M., Ros¸ca, V., Sta˘nica˘, D. & Atanasiu, L. 1990. La Dobrogea dans le cadre l’avant-pays carpathique. Revue Roumaine de Ge´ophysique, 34, 55–65. Volvovsky, B. S. & Starostenko, V. I. (eds) 1996. Geophysical Parameters of the Lithosphere of the Southern Sector of Alpine Orogen. Naukova Dumka, Kiev (in Russian). Winchester, J. A., Pharaoh, T. C., Verniers, J., Ioane, D. & Seghedi, A. 2006. Palaeozoic accretion of

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The crustal structure of the Black Sea from the reinterpretation of deep seismic sounding data acquired in the 1960s TAMARA YEGOROVA*, EKATERINA BARANOVA & VIKTOR OMELCHENKO Institute of Geophysics, National Academy of Sciences of Ukraine, Pr. Palladina 32, Kiev 03680, Ukraine *Corresponding author (e-mail: [email protected]) Abstract: A ray-tracing modelling of seismic refraction data acquired in the 1960s has been undertaken on two north–south lines – Profile 25 in the western part of the Black Sea and Profile 28/29 crossing the Azov Sea and central part of the Black Sea. The velocity model along Profile 25 shows two domains interpreted as thin (5 km) high-velocity (sub-)oceanic crust below the deep-water part of the Western Black Sea (WBS) basin, covered by 12– 13 km of Cretaceous and younger sediments, and a 39 km thick continental domain of the Scythian Platform and southernmost part of the East European Platform. They are separated by a high-amplitude normal fault, interpreted as being related to the opening of the WBS during Late Cretaceous rifting. The velocity model on Profile 28/29 shows what is interpreted as oceanic crust on the northwestern extremity of the Eastern Black Sea Basin (EBS) and thinned continental crust (Moho depths at 29 km) underlying the mid-Black Sea Ridge (MBSR) that separates the EBS and WBS. The basement of the MBSR comprises three units, which in an en echelon-like manner elevate southwards from a depth of 10–11 km beneath the Andrusov Ridge to 6 km on Arkhangelsky Ridge. An inclined seismic boundary at the Moho interface may be related to oblique rifting setting during the initial formation of the EBS.

The Black Sea Basin is an intracontinental sedimentary basin of the Tethyan belt, located between two orogens – the Crimea – Caucasus to the north and the Pontides to the south (Fig. 1a). On its northern margin, the Black Sea is bounded by the southern slope of the Precambrian Eastern European Platform (Ukrainian Shield and Scythian Platform; Fig. 1a), to the west it cuts the Palaeozoic structures of Dobrogea and the Moesian Plate. Much of the present basin floor is a flat abyssal plain lying at a depth of around 2200 m and appears to reflect the presence of a single basin (Fig. 1a). However, deep seismic reflection data (Tugolesov et al. 1985; Finetti et al. 1988) have shown two deep subbasins in the Black Sea – the Western Black Sea Basin (WBS) and the Eastern Black Sea Basin (EBS) filled with Cenozoic sediments as thick as 12– 14 km (Fig. 1b). The two sub-basins are separated by the basement uplift of the mid-Black Sea Ridge (MBSR), which consists of the Andrusov and Arkhangelsky ridges (Fig. 1b). It is generally accepted that the Black Sea Basin formed during the Cretaceous and, possibly, Palaeogene, in a back-arc setting in the hinterland of the Pontide Arc. However, it’s deeper structure and, hence, the processes that actually formed it, remain poorly investigated and there is dearth of modern deep seismic data. In this paper we present new results obtained by re-interpreting deep seismic

sounding (DSS) data (essentially being equivalent to what is now called refraction and wide-angle reflection techniques) collected in the Black Sea during the time of the Soviet Union (1960s– 1980s). Systematic deep seismic investigations (DSS as well as near-vertical reflection common deep point (CDP) studies) were carried out such that the whole of the Black Sea was covered by a set of 21 DSS profiles of different orientation and length (Malovitsky & Neprochnov 1972; Belousov & Volvovsky 1989; Tugolesov et al. 1985). The major orientation of these lines (north– south and east –west) was chosen in order to cross the main tectonic units of the study region. Published interpretations of these data have shown that the deep-water parts of the Black Sea have a thin oceanic crust, lacking a ‘granitic’ layer, with an underlying Moho boundary at depths of 20– 25 km (Neprochnov et al. 1970; Malovitsky & Neprochnov 1972; Moskalenko & Malovitsky 1974; Belousov & Volvovsky 1989). Such velocity models are simple, using the methodologies and interpretation techniques available at that time. Though they constitute the basis for many subsequent geophysical models and tectonic reconstructions of the Black Sea, the actual seismic data have not been reconsidered since. Re-evaluation of these seismic data using raytracing methods may, therefore, be of interest for

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 43– 56. DOI: 10.1144/SP340.4 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Main tectonic elements of the Black Sea Region and location of reinterpreted seismic lines, shown on a (a) topography/bathymetry map and (b) at the base of presumed Cretaceous successions. Red dots in (a) indicate locations of the shot points (red labels) and black numbers refer to distance along the profile in km. The background is digital topographic map from a global topography 2 min database (a) and digitized map of Tugolesov et al. (1985) (b).

the Black Sea region and, here, we present the results of re-interpreting old DSS data from two north–south oriented profiles. One crosses the WBS (Profile 25) and the other (Profile 28/29) combines seismic lines 28 in the Azov Sea and 29 in the central part of the Black Sea. This is the first time that these data have been modelled using computer-based ray-tracing methods, allowing us to take into account simultaneously travel times from all shot points for different types of seismic waves.

Methodology of velocity modelling The reinterpretation of the DSS data was carried out using the travel –time curves (with water depth corrections) constructed on the basis of previously made phase correlations (Malovitsky & Neprochnov 1972; Moskalenko & Malovitsky 1974; Neprochnov et al. 1970). These are shown schematically in Figures 2– 4. These seismic data, acquired more than 50 years ago, have a low signal-to-noise ratio. Accordingly, particular

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Fig. 2. Travel–time curves for the DSS Profile 25 in western part of the Black Sea (East Black Sea Basin) according to (Malovitsky et al. 1969; Neprochnov et al. 1970). For the Profile location see Figure 1. Phases Psd and Pg, shown by thin black lines, correspond to refracted waves from sedimentary layers and basement/upper crystalline crust respectively; Pn (thick black lines) and PmP (thick dotted lines) indicate refracted and reflected waves from the Moho. Numbers indicate apparent velocities in km s21. The travel– time curves used in ray-tracing modelling are shown by black lines, grey lines indicate that not involved in the modelling.

attention has been paid during re-interpretation to careful analysis of the initial seismic data (travel– time curves), the identification of intervals with reliable phase correlation and rejection of unreliable data.

The seismic data were reinterpreted by using a ray-tracing technique (Zelt & Smith 1992) that allows us to take into account travel times from all shotpoints for different types of seismic phases simultaneously. Modern methods of ray-tracing

Fig. 3. Travel– time curves for the DSS line 28 in the Azov Sea. For the location of seismic line see Figure 1. For the identification of seismic phases see Figure 2.

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Fig. 4. Travel–time curves for the DSS line 29 in the central part of the Black Sea in the region of the Mid Black Sea Ridge. For the location of seismic line see Figure 1. For the identification of seismic phases see Figure 2.

modelling allow for converting the kinematic parameters of all wave types into the model structure ˇ erveny´ & Psˇencı´k 1983; Zelt & Smith 1992). (C The velocity modelling is accomplished by iterative calculation of the rays and travel-times for a series of improving velocity models. The procedure is repeated multiple times until a reasonable fit is reached between the observations and computed travel–time curves. The reliability of the final velocity model depends, on one hand, upon the quality of the initial seismic data (accuracy, observation system and wave field characteristics) and, on the other, upon the complexity of the velocity model itself. The accuracy of the constructed velocity models was estimated in the course of the modelling.

Reinterpretation of DSS data in the West Black Sea Basin along Profile 25 Seismic data and wave field description North –south-oriented Profile 25 was acquired in 1966 by a consortium of different industrial and scientific organizations of the former Soviet Union (VNIIMORGEO as the leading research institute) under the leadership of Yu. P. Neprochnov (Malovitsky & Neprochnov 1972). From south to north the profile crosses the WBS (30 –330 km), the Scythian Platform (330 –420 km) and the southernmost part of Eastern European Platform including the Ukrainian Shield (up to 510 km). The thickness of sediments, based on seismic reflection profiling, increases southwards from 3 km on the shelf to 13 km in the deep-water basin (Fig. 1b).

Seismic acquisition was carried out using 20 mobile shotpoints (SP) spaced from 20 to 50 km apart. Shotpoints were ship-based. The observation system is rather irregular (Fig. 2). The southern part of profile (30–240 km) has shotpoint intervals of 40– 50 km with maximum offsets of 80 – 100 km. In the central and northern parts, the SP interval decreases to 15 –30 km and the maximum offsets increase to 160–200 km. Seismic signals were recorded by hydrophones deployed by five ships. Ocean-bottom seismic stations and seismic radio buoys were also utilized. The charges increased in size from 105 kg on the shelf, detonated on the sea bottom, to 135 kg in the deepwater part, where they were detonated at a depth of 90 m. As a result of this DSS study different types of seismic waves, reflected and refracted from boundaries in the crust and uppermost mantle, were recorded. Examples of seismic traces recorded from SP9 and SP8 are shown in Figures 5 and 6. The profile can be divided into two parts: a southern segment within the deep-water basin (30–240 km) and a northern segment, including the continental slope and shelf zone (240–510 km). The wavefield of the deep-water basin was obtained from short travel– time curves from five shotpoints. First arrivals correspond to waves refracted in sediments (Psd) and on the Moho boundary (Pn), where they form two branches with the crossover distance being 35– 40 km, where Psd waves with Vp ¼ 2.0–4.0 km s21 are substituted by upper mantle phases Pn with velocities of about 8.0 km s21 (Fig. 2). From two shotpoints (SP1 and SP3) the Psd phase has been also recorded as a second arrival at 70 –80 km distance in the form of linear branches with Vp ¼ 4.0– 4.5 km s21. This can be

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Fig. 6. Example of seismic recordings from SP 8 of Profile 25 (Malovitsky & Neprochnov 1972).

Fig. 5. Example of seismic recordings from SP 9 of Profile 25. Grey arrows show the first arrivals for Psd, Pg and Pn phases (Malovitsky & Neprochnov 1972).

indicative of the presence of a layer with a lowvelocity gradient. The Pn phase is observed at distances of 60– 110 km from the SP. No reflected waves were recorded in this part of the profile. The lower part of the sedimentary cover and the major part of consolidated crust are not elucidated by first arrivals, forming a so-called ‘hidden layer’. The parameters of this layer were estimated in the model by trying different depths to the top and bottom of the crystalline crust. In particular, the shallowing of the basement leads to the appearance of a Pg phase, which is absent in the recordings, and, vice versa, deepening of the base of the crust generates a Pn phase at greater times than has been recorded. In the northern part of the profile, within the continental slope and shelf zone, the wave field changes substantially. In addition to Psd and Pn phases described above, the first arrivals include the Pg phase, refracted from the basement, seen over the distance 30 –70 km with velocities 5.8– 6.4 km s21 (Fig. 2). The Psd waves are observed as first arrivals with offsets 10 –30 km, while the Pn phase is seen at offsets of 80 to 150 km (Figs 2 & 7). The PmP phase, reflected from the Moho boundary was recorded from several shotpoints as

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Fig. 7. Observed (crossed) and calculated (lines) travel times, and selected model ray paths, for seismic phases observed in the West Black Sea Basin on the seismic line 25, SP 7 (a) and SP 12 (b). For indication of seismic phases see Figure 2. Numbers on the cross-sections indicate modelled velocities in km s21.

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a secondary arrival at distances 70 –130 km (Fig. 7b). The wave field of this part of the profile shows the presence of single intracrustal reflections. Together with the complicated records of the Pg phase, they can be indicative of a complex crustal structure, which cannot be resolved in detail given the poor quality of the seismic data.

Velocity model along Profile 25 The final velocity model along Profile 25 is shown in Figure 8. In its central and southern parts (30 – 320 km) one can see a deep sedimentary basin filled with sediments (Vp ¼ 2.0–4.2 km s21) as thick as 12 –14 km. From the north it is limited by what can be interpreted as a high-amplitude (normal) fault, along which basement deepens abruptly by some 7.5 km (from 3.5 km on the Scythian Platform to 11 km in the WBS). The top of basement is identified as the seismic boundary with Vp ¼ 6.0– 6.2 km s21. The sedimentary basin comprises three layers with velocities varying from 2.0 to 4.2 km s21. Within the southern, deep-water part of the basin (30 –180 km), the basin is made up mainly by the lower layer (Vp ¼ 4.0–4.2 km s21). The two upper layers (Vp ¼ 2.0–3.0 and 3.0–4.0 km s21), with a total thickness up to 9 km, prevail in the northern part of the WBS, forming a sub-basin in the range 200 –320 km. In the northern part of the profile these sediments reach thicknesses of 3.5–4 km and cover the crust of the Scythian Platform and southern slope of the East European Platform. The upper mantle in the model has a velocity of 8.0 km s21, thus defining the Moho. There are dramatic changes in the velocity structure of the overlying crystalline crust. It varies in thickness from

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5 km in the deep-water basin to 39 km beneath the southern slope of the East European Platform (Fig. 8). The Moho is at a depth of 17– 19 km beneath the thin high-velocity (6.5–7.2 km s21) crust of the deep-water basin, which is, accordingly, interpreted as having oceanic affinity. Further north the Moho boundary deepens considerably, in two segments with different slopes. In the interval 160–300 km the Moho gently deepens from 17 to 25 km. Northwards, beneath the continental shelf (300 –360 km), it deepens more steeply down to 39 km, typical for the continental crust of the Scythian Platform and Ukrainian Shield. The geometry of the Moho in this area is controlled by Pn and PmP phases from seven shot points (Fig. 7). The central part of the velocity model (160 –310 km), lying beneath the continental slope corresponds to the transition zone between oceanic and continental crust. The crust in this area has lower velocities (6.2– 6.5 km s21) than the oceanic domain. Where the Moho begins to deepen more steeply beneath the continental shelf (300 km) corresponds to the position of the inferred highamplitude normal fault at the edge of Scythian Platform. In the consolidated crust beneath the Scythian and East European platforms velocities gradually increase from 6.0 km s21 near the surface to 7.2 km s21 at the base of the crust. The accuracy of obtained velocity model, estimated in the course of the modelling, can be characterized by the following parameters. The accuracy of velocity determination in the sedimentary cover on Profile 25 is in the range from +0.2 to +0.3 km s21, while depth to the basement were determined with the accuracy +0.5/1.0 km. In general, the ambiguity of any constructed velocity model increases with depth. In the southern

Fig. 8. Velocity model of the DSS Profile 25 (western part of the Black Sea) crossing the northwestern shelf and west Black Sea Basin. Velocities are in km s21. Thick subhorizontal lines indicate reflectors. Moho interface is shown by dotted line.

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part of profile (30 –160 km) the Moho was modelled using the Pn travel–time curves from several SPs (SP 1– 4). Within the ‘hidden layer’ (mainly in the crystalline crust) the velocities can not be resolved with certainty. Therefore in Figure 8 we show the model with typical oceanic or ‘sub-oceanic’ crustal velocities (Vp ¼ 6.5 –7.2 km s21). The relief of the Moho in the transition from oceanic to continental crust (160 –310 km) has been determined with high reliability since it is controlled by seismic data from many SP: by Pn phase from seven SP and by PmP phase from three SPs (see for example Fig. 7). Thus depths to the Moho were set up with the accuracy 1.5–2 km.

Reinterpretation of DSS data in the Azov Sea and central part of the Black Sea (Profile 28/29) Seismic data and wave field characteristics The northern segment of joint Profile 28/29 (line 28) crosses from north to south of the southeasternmost part of the Eastern European Platform including the southern part of the Azov Massif (Ukrainian Shield), Palaeozoic Scythian Platform and the younger Indolo-Kuban Trough, filled mainly with Cenozoic sediments (see Fig. 1). It lines up to the north with the DOBRE profile (DOBREfraction’99 Working Group 2003) that crosses the Azov Massif itself (Fig. 1a). Seismic lines 28 (246 km length) and 29 (360 km length) were acquired in 1968 by the Interdepartmental Black Sea Expedition (the leading organization was VNIIMORGEO) (Moskalenko & Malovitsky 1974). Trust ‘Dneprgeofizika’ and the Institute of Geophysics of the Academy of Sciences of Ukraine were the Ukrainian partners. Line 28 used five onshore shotpoints, three to the north and two to the south on the easternmost Crimean Peninsula. Seismic signals were recorded by seismic stations on ships and by ocean bottom stations, spaced at 10 –15 km. The observed dataset consists of two groups of travel–time curves separated by the 140 km segment of the Azov Sea (Fig. 3). First arrivals on SP 1 and 2 are phases refracted from within the sedimentary cover (Psd) and have apparent velocities of 2.4– 4.0 km s21. At a distance of 25–35 km these are crossed over by Pg phases with velocities 6.0– 6.4 km s21. Travel– time curves in the north (SP 3–5) are constituted mainly by Pg phases. The Moho refraction Pn phase was not recorded on Profile 28. Reflected phases are visible at offsets of 90– 120 km. These are quite variable and are interpreted to be generated from some common horizon in the lower crust. In general, the accuracy

of seismic data along line 28 is poor because of shortcomings in the observation system, which led to uncertain phase correlations. This was exacerbated by the shallow water conditions of the Azov Sea (maximum depth 14 m) and ‘noise’ inherent to this. The main objective of our reinterpretation here is therefore to derive no more than a generalized model showing the most stable features of the crustal structure. Line 29 crosses the central part of the Black Sea from Feodosiya Bay on the southeastern Crimean Peninsula towards Bafra Bay in Turkey. Much of the present basin floor along the profile is flat abyssal plain lying at a depth of nearly 2000 m (70– 290 km). The northern part of the profile (290 –370 km) crosses continental slope and a narrow shelf zone (Fig. 1a). The shotpoints are spaced very irregularly (Moskalenko & Malovitsky 1974) – from 20 to 90 km in the southern, deepwater, part to 5–7 km within the northern shelf (Fig. 4). The observation system consists of 10 SPs with prevailing counter (coming from the opposite direction) travel –time curves including five SPs with the offsets of 70 –80 km length and five SPs with the 110– 150 km offsets. The present model utilizes the seismic data, represented by travel–time curves, shown in Figure 4; seismic phases are those refracted in sediments (Psd), in the crust (Pg) and in the uppermost mantle, below the Moho (Pn). The Psd phase is seen as first or secondary arrivals along the whole profile. On the deep-water part of the line (70 – 290 km) they form the branch of first arrivals just near the SP up to the offsets of 15–20 km, while within the continental slope and shelf they extend to 30 –40 km. Apparent velocities of Psd range from 2.4 to 4.0–4.5 km s21. From several shotpoints (2–5) the Psd phase is seen as a secondary arrival at offsets of 30– 40 km, with apparent velocities about 2.0, 3.0 and 4.0 km s21 (Figs 4 & 9). Pg waves, penetrating into the basement as deep as 15–18 km, are first arrivals at offsets of 60 – 100 km (Fig. 9). A wide range of apparent velocities (5.6–6.8 km s21) is observed, partly as a result of a low signal-to-noise ratio and basement topography irregularities. This is also related to the change of crustal type from continental (Vp ¼ 6.0– 7.0 km s21) on the southern and central parts of the profile to (sub-)oceanic (Vp ¼ 6.5–7.0 km s21) in the north. The Pn phase, displaying apparent velocities 7.8– 8.2 km s21, is poorly expressed, being recorded only over short (20 –30 km) intervals, with the exception of SP 5 (Fig. 9b).

Velocity model along Profile 28/29 Figure 10 presents our final velocity model along joint Profile 28/29. Its northern segment (line 28)

REINTERPRETATION OF DEEP SEISMIC SOUNDING DATA

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Fig. 9. Observed (crossed) and calculated (lines) travel times, and selected model ray paths, for seismic phases observed in the central part of the Black Sea (Mid–Black Sea Ridge and north-western part of the East Black Sea Basin) on the DSS line 29, SP 2 (a) and SP 5 (b). For the identification of seismic phases see Figure 2. Numbers on the cross-sections indicate modelled velocities in km s21.

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Fig. 10. Velocity model of joint Profile 28/29 (Azov Sea and central part of the Black Sea) crossing onshore the southernmost part of the Ukrainian Shield – Mid– Black Sea Ridge. Velocities are in km s21. Dotted line indicates the Moho position, its probable location under the Sorokin Trough is shown by thick dashed line.

shows crustal structure in the transition from the thick continental crust of the East European Platform and Scythian Platform to Indolo-Kuban Trough, which is the western prolongation of the Greater Caucasus foreland basin (Fore-Caucasus Trough). This is seen in the thickness of the sedimentary cover, which increases from practically zero at the northern end of the line and 2 km beneath the Scythian Platform to about 12 km in the deepest part of the Indolo-Kuban Trough. Sediments in the Indolo-Kuban Trough are mainly of Cenozoic age and have velocities in the range 2.0–4.2 km s21. Crystalline crust in the model displays a velocity increase from 6.0– 6.2 km s21 on the basement surface to 6.7 km s21 in the lower crust. A remarkable feature of the model is a deep reflector seen on all shotpoints. What this reflector represents is unclear, though its coincidence with the 6.7 km s21 velocity isoline suggests that it is the top of a lower crustal layer or possibly the Moho. The southern segment of Profile 28/29 (seismic line 29) images the crustal structure of the central part of the Black Sea from Sorokin Trough to the MBSR. Short offsets and irregularities in Psd phase correlation, given the sparse observations, did not allow definition of lateral velocity variations in the sedimentary cover. The combined thickness of the sedimentary layers increases northwards from 6 km on the Arkhangelsky Ridge of the MBSR to 13 km in the EBS and Sorokin Trough. Most of this, up to 5 km, belongs to the low-velocity (Vp ¼ 2.0– 3.0 km s21) succession. Sediments with higher velocities (Vp ¼ 4.0– 4.5 km s21) fill the lower part of the EBS. The Sorokin Trough has the greatest thickness of these, from a depth of 4 km to about 14 km (Fig. 10). The sediment thickness variations, together with the very irregular morphology of the basement surface, suggest that the basement can be interpreted

as separate blocks of 40 –60 km width with 2 –4 km vertical displacement. The basement horizon is a first-order seismic boundary with a velocity jump of 4.5 –5.0 to 6.0 –6.5 km s21, The top of the basement of the MBSR comprises three structural steps rising from 10 to 11 km depth on the Andrusov Ridge southwards to 6 km on the Arkhangelsky Ridge (Fig. 10). The geometry of the Moho boundary along line 29 is modelled on the basis of a strong Pn phase indicating a velocity jump of 7.0 to 8.0 km s21 and clearly delineates two different crustal domains. The Moho lies at a depth of 19 +1 km beneath the northern domain (250– 360 km), which has a thin (7–10 km) and high-velocity (6.5–7.0 km s21) overlying crustal layer that is interpreted to have oceanic or sub-oceanic affinity. The Moho beneath the southern block lies at depth of 29 +1 km and the crustal layer can be more than 20 km thick and has velocities in the range 6.0–7.0 km s21; it is interpreted to represent thin continental crust underlying the MBSR. These two different crustal domains are separated by an inclined seismic boundary that is interpreted to be the Moho. It deepens southwards below the northern margin of the MBSR by more than 10 km such that it is deeper than the Moho imaged beneath the MSBR itself (Fig. 10). The geometry of this somewhat spectacular feature of the Moho is reliably controlled by refracted Pn phases from several shotpoints (Fig. 9). The accuracy of the obtained velocity model along the southern part of Profile 28/29 (seismic line 29), which was constructed using the refracted waves only, could be characterized by the following. Depths to basement are considered to have an accuracy of +0.3–0.7 km. The structure of the crystalline crust in the model has a variable reliability. Velocities are best constrained in the

REINTERPRETATION OF DEEP SEISMIC SOUNDING DATA

continent to ocean transition zone (180 –250 km). The basement uplift inferred northwards (260– 320 km) is poorly constrained: its gentle southern slope is modelled by observations from one shotpoint only – SP5 (Fig. 9b) and its northern slope has been modelled as a closure of the deep Sorokin Trough. The accuracy in defining the Moho is thought to be in the range +0.5 to +1.0 km. In the northern segment of profile (seismic line 28) crustal velocities were determined with high accuracy (+0.05 km s21).

Discussion Profile 25 crosses the western part of the Black Sea (WBS) in a north–south direction. There is generally good agreement with previously published models (Neprochnov et al. 1970; Malovitsky & Neprochnov 1972; Belousov & Volvovsky 1989), especially regarding thicknesses and velocities in the oceanic segment of the model indicating the ‘lack’ of a ‘granitic layer’ in the central part of the Black Sea. The thickness of crustal layers within the continental domain are revised but, otherwise, the disposition of continental basement and Moho are essentially unchanged. The most important difference between the old and present models, which can be considered as an improvement derived from the better modelling technique being used, lies in the transition zone, in particular the abrupt basin thickening at the margin of the Scythian Platform (310 –330 km) and the reduction of the rate of shallowing of the Moho in the same area. In the new model presented here (Fig. 8), the three main domains (interpreted as continental, transitional and oceanic from north to south) correspond, approximately, to the regional morphology of the sea floor (the shelf, the deep-water basin and the continental slope, respectively). The continental shelf crustal domain has a velocity structure that implies it is of cratonic type, a distinctive feature of which is supposed to be high-velocity lower crust (Pavlenkova 1996). The inferred velocity and thickness of the (deep-water) oceanic crustal domain are typical of young oceanic crust presumably developed during Cretaceous rifting as a result of oceanization/basification of the preexisting continental crust (Yegorova & Gobarenko 2010 and references therein). The abrupt change of basement depth, from 3.5–4.0 km to 11 km at the edge of the Scythian Platform, is interpreted as evidence of a highamplitude normal fault related to WBS rifting. A similar structure has been inferred from reflection seismics in this area (Finetti et al. 1988, fig. 80). The important role of this fault in the formation of the WBS is emphasized by the structure of the

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sedimentary basin of the WBS (Fig. 8, 30 – 320 km), which comprises two sub-basins. To the south, high-velocity sediments (Vp ¼ 4.0–4.2 km s21) prevail, whereas to the north, located mainly within the continental slope, low-velocity sediments (Vp ¼ 2.0–4.0 km s21) occur and are limited to the north (from the Scythian Platform) by the mentioned fault. The internal velocity structure of the sedimentary succession indicates that this normal fault controlled the accumulation of a large thickness of the younger, lighter sediments. The area of this important basin-bounding fault, coinciding with the margin of the northwestern shelf, is characterized by seismicity with earthquake magnitudes around 3–5 and foci located in the lower crust –upper mantle (Yegorova & Gobarenko 2010). Further, the fault zone can be traced westwards through the Histria Basin, located on Romanian shelf, onshore to coincide roughly with major NW-trending faults (Pechenega-Camena and Capidava-Ovidiu faults), that are generally thought to be of crustal-scale and forming part of the ‘Tornquist-Teisseyre Line’ (in fact a zone) in Romania (Sa˘ndulescu 1978, 1980; Visarion et al. 1990). Saintot et al. (2006) considered that the best candidate for the suture between Moesia and the East-European Craton is the Capidava-Ovidiu Fault. All of these inferred major onshore structures played an important role in strike–slip tectonics related to the opening of the western part of the Black Sea (Banks & Robinson 1997; Seghedi & Oaie 1994; Sa˘ndulescu 1994; Hippolite 2002). Thus, the velocity structure of Profile 25 supports the idea that the WBS formed on the crust of Moesian Plate due to rifting occurring along the suture zone (including the Tornquist-Teisseyre Line) between the Moesian and East European crustal domains. The origination of young oceanic crust of the WBS on the thick continental crust is assumed also from the gravity modelling showing the necessity of the uppermost mantle of decreased density below thin high-density oceanic crust (Baranova et al. 2008). Keeping in mind the low-temperature regime of the Black Sea lithosphere (Kobzar 1987; Kutas et al. 1998; Gordienko et al. 2002), this lowdensity sub-crustal layer can hardly be caused by thermal expansion of the upper mantle. More likely, it is caused by modification of the substance of the lower crust –upper mantle interval due to ‘oceanization’ (or ‘basification’) of primary continental crust of Moesian Plate as a result of Upper Cretaceous – Lower Paleocene rifting accompanied by intrusions of basaltic magma into the crust. As a result, depleted low-density sub-crustal layer and high-density basaltic oceanic crust, together with Moho shallowing was originated. The crustal velocity model along Profile 28/29 gives valuable new information in regard to the

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opening of the EBS. From north to south it crosses the southern slope of the East European Craton and Scythian Platform through the Indolo-Kuban Trough and Crimean orogen to the MBSR. It has resolved a rather simple velocity model for the first (platform) segment, which has much in common with the ‘old’ model (Moskalenko & Malovitsky 1974). The latter included two reliable seismic boundaries (the basement surface and Moho boundary) and shows sediment thickness increasing from north to south in the Azov Sea to 15 km, within the Indolo-Kuban Trough. The Moho was interpreted (along the central segment of seismic line 28) to be at a depth of approximately 40 km. In contrast, the present model (Fig. 10) consists of three layers – upper crust (Vp ¼ 6.0– 6.3 km s21), middle crust (6.3–6.5 km s21) and a thick lower crust (6.5–6.7 km s21). It is identical to that of the platform block on the northern part of Profile 25 in the western part of the Black Sea (Fig. 8). There is little difference between the crustal structure of the East European Craton and the Scythian Platform on Profile 28/29, the former is characterized by thinning of both the sedimentary layer and upper crust in comparison with Scythian Platform. New information on the crustal structure of the Azov Sea is imminent from the recently acquired DOBRE-2 seismic profile (Starostenko 2008), which is the southern continuation of the DOBRE line crossing the Donbas Foldbelt (DOBREfraction’99 Working Group 2003; Maystrenko et al. 2003). The crustal structure of the platform segment of Profile 28/29 is also in good agreement with the velocity model of the Azov Massif (Ukrainian Shield) on the DOBRE profile (DOBRE fraction ’99 Working Group 2003), the southern end of which is in close proximity with the northern edge of Profile 28/29 (see Fig. 1a). The transition from the Scythian Platform to the Fore-Caucasus (Indolo-Kuban Trough) is clearly indicated by thickening of sediments. They reach 11 km in thickness below the trough axis under the shallow water of the southernmost part of the Azov Sea. A high amplitude reflector, revealed at the base of the crust and deepening below the basin depocentre, can be interpreted either as the Moho interface or as a boundary within the lower crust. The crustal structure of the Crimean orogen is not illuminated by Profile 28/29 because it lies between the ends of seismic lines 28 and 29 (Fig. 1). However, the crust is likely to be fairly thick in this region, lying as it does between and along strike the thickened crust of the Gornui Crimea to the west and the Greater Caucasus mountains to the east, where Moho depths are thought to reach 50 –55 km (Belousov & Volvosky 1989). The southern segment of Profile 28/29 (seismic line 29) in the central part of the Black Sea

constrains the crustal structure of the poorly investigated MBSR and its transition to the EBS. The new velocity model shows that the basin architecture and the crustal structure of the EBS (Fig. 10) are very similar to that of the WBS, as determined on Profile 25 (Fig. 8). The basement of the MBSR consists of three blocks, arranged in an echelon-like fashion along its strike. The top of basement shallows southwards from 10 –11 km on the Andrusov Ridge to c. 6 km on the Arkhangelsky Ridge. Consolidated crust, consisting of an upper (Vp ¼ 6.0–6.5 km s21) and a lower (6.5– 7.0 km s21) layer, is underlain by Moho at a depth of 29 km. This compares to the old model of Moskalenko & Malovitsky (1974) that simply shows the basement horizon (boundary velocity 5.6– 6.0 km s21) at a depth varying from 4 km on the Andrusov Ridge to 9 km in the Sorokin Trough, underlain by the Moho (constrained only in the central part of the profile as seismic interface with boundary velocity 8.2 km s21) at a constant depth of 20 + 2 km. A peculiar feature of the new model is an oblique seismic boundary at the base of the crust which deepens below the MBSR. We interpret this boundary as the Moho interface at the transition from the thinned continental crust of the MBSR to the oceanic domain of the EBS. Most probably, this seismic boundary, which entrains into the uppermost mantle, has a tectonic origin and, most probably, relates to the opening of the deep EBS in an oblique rifting environment. Available geological and geophysical information suggest that the EBS likely formed within a Transcaucasus terrane (Adamia et al. 1974; Golonka 2004; Saintot et al. 2006). The fragments of this Transcaucasus terrane – the Shatsky Ridge and the MBSR are now intact on either side of the EBS in the eastern and central parts of the Black Sea. As such, the opening of the EBS occurred due to oblique rifting accompanied by strike–slip movements along the MBSR. This formed a narrow NW-trending deep rift basin in the eastern part of the Black Sea, which terminates offshore just at the edge of the very deep and narrow Sorokin Trough. This structure is NE-oriented and is located offshore parallel to the Crimea Mountains (Fig. 1) and, together with its northeastern extension into the Kerch-Taman Trough, can be considered as an accretional trough filled with thick (up to 12 km) Cenozoic sediments. This area is also characterized by active seismicity (Yegorova & Gobarenko 2010).

Conclusion To get new information on the basin architecture and crustal structure of the Black Sea we undertook a reinterpretation of old seismic refraction (DSS)

REINTERPRETATION OF DEEP SEISMIC SOUNDING DATA

data acquired in 1966– 1967 (Neprochnov et al. 1970; Belousov & Volvovsky 1989). This is the first time that these data have been treated using modern computer-based ray-tracing methods. The DSS data were remodelled and reinterpreted on two north–south lines – Profile 25 in the western part of the Black Sea and Profile 28/29 in the Azov Sea and central part of the Black Sea. The crustal structure of the Black Sea as revealed by ray-tracing allows several conclusions that can be summarized as follows. (1)

(2)

(3)

(4)

Deep-water basins in western and eastern parts of the Black Sea are filled with thick (up to 12 –13 km) Cretaceous and younger sediments. These are underlain by a thin (5–7 km) high-velocity crust [Vp ¼ 6.5– 7.0(7.2) km s21] of oceanic or sub-oceanic affinity with the Moho lying at a depth of 19 –20 km. This high-velocity and highdensity oceanic or sub-oceanic crust of the WBS is thought to have been formed by an ‘oceanization’ or ‘basification’ process resulting from the melting of basic and ultrabasic magmas and their intrusion into an extended and thinned lower continental crust during continental rifting. Clear evidence of rifting can be seen in what is interpreted as a high-amplitude normal fault (vertical displacement of basement about 7– 8 km) at the northwestern margin of the shelf zone. The location of this fault just at the edge of the continental domain suggests that it can be considered as a first-order rift controlling structure. It roughly coincides with the suture between the Scythian Platform (or East-European Craton) and Moesian Plate. The crust of Scythian Platform and southern slope of the East European Craton, crossed by the northern segments of Profiles 25 and 28/29, is considered to be indicative of a cratonic type of the crust. It is rather thick (37 –39 km) with P-wave velocity increasing from 6.0 km s21 at the basement surface to 7.2 km s21 at the base of the crust and a relatively thick lower crust with velocities 6.5–7.2 km. No differences were revealed in the crustal structure of the Scythian Platform and the southern part of the East European Craton on either profile, consistent with models in which the former comprises the margin of the latter reworked by Late Proterozoic and younger tectonism (Saintot et al. 2006). The crustal structure and thickness of sediments of the EBS is identical to that of the WBS, within the accuracy of the models. The MBSR comprises a narrow remnant of

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NW-oriented thin continental crust separating the two deep sedimentary basins. It is arranged by a set of blocks, which rises in an echelonlike manner southwards from 6 km on Arkhangelsky Ridge to 11 km on the Andrusov Ridge. An inclined Moho boundary, revealed below the transition zone from thin sub-oceanic domain of the EBS to the thinned continental crust of the MBSR is probably related with strike –slip movements along the latter and might be indicative of the origin (style of opening) of the EBS in an oblique rifting environment within the crust of Transcaucasus terrane. The research described in this paper has been carried out in the framework of the MEBE programme. Our appreciation goes to the referees – Professor M. Grad (University of Warsaw, Poland) and an anonymous one, who made valuable recommendations for improving the previous version of the manuscript. We are thankful also to R. Stephenson (Aberdeen University, UK) for his assistance in editing the style and English of the paper.

References Adamia, S., Gamkrelidze, I. P., Zakariadze, G. S. & Lordkipanidze, M. V. 1974. Adjara-Trialeti trough and the problem of the Black Sea origin. Geotectonika, 1, 78–94 (in Russian). Banks, C. J. & Robinson, A. 1997. Mesozoic strike-slip back-arc basins of the western Black Sea region. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 53–62. Baranova, E. P., Yegorova, T. P. & Omelchenko, V. D. 2008. Reinterpretation of DSS seismic materials and gravity modeling along the profiles 25, 28 and 29 in the Black Sea and the Sea of Azov. Geophysical Journal, 5, 1– 20 (in Russian). Belousov, V. V. & Volvovsky, B. S. (eds) 1989. Structure and Evolution of the Earth Crust and Upper Mantle of the Black Sea. Nauka, Moskow (in Russian). Cˇerveny´, V. & Psˇencı´k, I. 1983. SEIS83-numerical modeling of seismic wave fields in 2-D laterally varying layered structure by the ray method. In: Engdahl, E. R. (ed.) Documentation of Earthquake Algorithm. World Data Center A for Solid Earth Geophysics, Boulder, Report SE, 35, 36– 40. DOBREFRACTION ’99 WORKING GROUP 2003. ‘DOBREfraction’99’, velocity model of the crust and upper mantle beneath the Donbas Foldbelt (east Ukraine). Tectonophysics, 371, 81–110. Finetti, I., Bricchi, G., Pipan, M. & XUAN, Z. 1988. Geophysical study of the Black Sea. Bolletino di Geofisica Teorica ed Applicata, XXX/117-118, 197– 324. Golonka, J. 2004. Plate tectonic evolution of the southern margin of Eurasia in the Mesozoic and Cenozoic. Tectonophysics, 381, 235– 273.

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Gordienko, V. V., Gordienko, I. V., Zavgorodnayay, O. V. & Usenko, O. V. 2002. Heat Flow Field of the Territory of Ukraine. Znanie Ukrainu, Kiev (in Russian). Hippolite, J.-C. 2002. Geodynamics of Dobrogea (Romania): new constraints on the evolution of the Tornquist-Teisseyre Line, the Black Sea and the Carpathians. Tectonophysics, 357, 33–53. Kobzar, V. M. 1987. Heat flow and block structure of the lithosphere of the Black Sea Basin. Geophysical Journal, 4, 89–94 (in Russian). Kutas, R. I., Kobolev, V. P. & Tsvyashchenko, V. A. 1998. Heat flow and geothermal model of the Black Sea depression. Tectonophysics, 291, 91– 100. Malovitsky, Y. P. & Neprochnov, Y. P. (eds) 1972. Structure of the Western Part of the Black Sea Basin. Nauka, Moscow (in Russian). Malovitsky, Y. P., Neprochnov, Yu. P. et al. 1969. The structure of the Earth’s crust in the western part of the Black Sea. Dokllady Akademii Nauk SSSR, 186, 905– 907. Maystrenko, Y., Stovba, S. et al. 2003. Crustal-scale pop-up structure in cratonic lithosphere: DOBRE deep seismic reflection study of the Donbas Foldbelt, Ukraine. Geology, 31, 733–736. Moskalenko, V. N. & Malovitsky, Y. P. 1974. Results of deep seismic sounding along the transmeridional profile across the Azov and Black Seas. Izvestiya of Academy of Sciences of the USSR, Geological Series, 9, 23– 31 (in Russian). Neprochnov, Y. P., Kosminskaya, I. P. & Malovitsky, Y. P. 1970. Structure of the crust and upper mantle of the Black and Caspian Seas. Tectonophysics, 10, 517– 538. Pavlenkova, N. I. 1996. Crust and upper mantle structure in Northern Eurasia from seismic data. Advances in Geophysics, 37, 1 –333. Saintot, A., Stephenson, R., Stovba, S., Brunet, M.-F., Yegorova, T. & Starostenko, V. 2006. The evolution of the southern margin of the Eastern Europe (Eastern European and Scythian platforms) from latest Precambrian –Early Palaeozoic to the Early Cretaceous. In: Gee, D. & Stephenson, R.

(eds) The European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 481– 505. Sa˘ndulescu, M. 1978. The Moesian Platform and the North Dobrogean orogen. In: Lemoine, M. (ed.) Geological Atlas of Alpine Europe and Adjoining Areas. Elsevier, Amsterdam, 427– 460. Sa˘ndulescu, M. 1980. Analyse ge´otectonique des chaı˘nes alpines situe´es autor de la Mer Noire occidentale. Annales de l’Institute de Ge´ologie et Ge´ophysique, 56, 5 –54. Sa˘ndulescu, M. 1994. Overview on Romanian geology. In: Berza, T. (ed.) Alcapa II Field Guidebook: Geological Evolution of the Alpine-Carpathian-Pannonian System. Romanian Journal of Tectonics and Regional Geology, 74, 3– 15. Seghedi, A. & Oaie, G. 1994. Tectonic setting of two contrasting types of pre-alpine basement: North versus Central Dobrogea. Romanian Journal of Tectonics and Regional Geology, 75, 56– 57. Starostenko, V. I. 2008. DOBRE-2: Integrated geophysical studies of the crust and upper mantle on the southern margin of the east European craton (Azov Sea-Crimea-Black Sea area). Abstracts of 33rd International Geological Congress, Oslo (CD-Rom). Tugolesov, D. A., Gorshkov, A. S. et al. 1985. Tectonics of Mesozoic-Cenozoic Successions of the Black Sea Basin. Nedra, Moscow (in Russian). Visarion, M., Sa˘ndulescu, M., Ros¸ca, V., Sta˘nica˘, D. & Atanasiu, L. 1990. La Dobrogea dans le cadre l’ avant-pays carpathique. Revue Roumaine de Ge´ophysique, 34, 55–65. Yegorova, T. & Gobarenko, V. 2010. Structure of the Earth’s crust and upper mantle of West- and EastBlack Sea Basins revealed from geophysical data and its tectonic implications. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 23– 42. Zelt, C. A. & Smith, R. B. 1992. Seismic traveltime inversion for 2-D crustal velocity structure. Geophysical Journal International, 108, 16– 34.

Methane in the northern Black Sea: characterization of its geomorphological and geological environments V. I. STAROSTENKO1, O. M. RUSAKOV1*, E. F. SHNYUKOV2, V. P. KOBOLEV1 & R. I. KUTAS1 1

Institute of Geophysics, National Academy of Sciences of Ukraine, Kiev, 32 Palladin Avenue, 03142 Kiev, Ukraine

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Branch of Marine Geology and Ore Formation, National Academy of Sciences of Ukraine, 55b Gonchara Street, 01054 Kiev, Ukraine *Corresponding author (e-mail: [email protected]) Abstract: Based on hydro-acoustic and geophysical observations, this paper presents an analysis of geomorphological and geological settings of gas methane occurrence on the NW shelf and upper continental slope, in the Sorokin trough and on the Kerch-Taman offshore, in the Black Sea. Gases are associated with seeps, mud volcanoes and gas hydrates. Evidence is given for the thermogenic nature of methane. The gas methane is of mostly abiogenic origin. Small gas releases may be produced by the decomposition of Quaternary organic material near the sea floor through the action of bacteria or biodegradation of redeposited thermogenic hydrocarbons. The origin of carbonate formations is related to degassing sedimentary layers. There is a possible role for deep faults in transporting gas to the sea floor. The gas hydrate stability zone in the Black Sea lies at minimum water depth of 600–650 m with its thickness up to 500 m.

In the northern Black Sea intensive outbursts of methane into the atmosphere were first recorded in the scientific literature as early as 1928, one year after a series of destructive Crimean earthquakes (Popov 1928). Some 60 years later the sources of methane were described when Polikarpov et al. (1989) related acoustic anomalies to the leakage of gas. Recently the Black Sea has become a popular target for interdisciplinary studies on its enormous methane resource. It is estimated that the Earth’s largest anaerobic basin – the Black Sea – contains a total of 96  109 kg of methane (Reeburgh et al. 1991). European River-Ocean Systems project (EROS-2000, EROS-21), CRIMEA project (EC EVK-2-CT-2002-00162) and METROl programme (EC EVK3-CT-2002-00080), the ASSEMBLAGE European Project (EC EVK3-CT-2002-0009), a German research project GHOSTDABS (03G0559A), and an Ukrainian research project ‘Gas volcanism in the sea floor of the Black Sea as prospecting indicator of gas hydrate accumulations and traditional hydrocarbon raw’. These projects have focused on several main topics, namely: (1) mapping venting sites and studying the relationship between spatial distribution of gas releases, sea floor morphology and subsurface features (e.g. Shnyukov & Ziborov 2004; Naudts et al. 2006); (2) examining the geological structure of seepage areas (e.g. Krastel et al. 2003; Lu¨dman et al. 2004; Popescu et al. 2007); and (3) determining the Black Sea’s

contribution of methane to the global ocean and atmosphere (e.g. Schmale et al. 2005; Wallmann et al. 2006). In most publications regarding methane in the Black Sea, its origin is not discussed. Sometimes it is mentioned that it is comes from sedimentary gas reservoirs at emission sites like cold seeps and mud volcanoes where gas is biogenic in origin (e.g. Jørgensen et al. 2001; Durisch-Kaiser et al. 2005). Egorov et al. (2003) presented a comprehensive set of arguments favouring an organic source of methane in the Black Sea. However, these authors dealt only with shallow gas from gravity cores and did not consider higher hydrocarbons from gas seeps that have been described in the literature (e.g. Stadnitskaya & Belen’kaya 2000; Blinova et al. 2003). There have been no attempts to estimate the quantity of organic matter required to generate the evidently enormous accumulation of gas methane in the Black Sea. One issue related to this is the need to assess the potential for abiogenic sources of methane in the Black Sea. A direct relationship between fault distribution with gas release was clearly demonstrated by Kutas et al. (2002, 2004), based on an inventory of the distribution of gas flames in the NW Black Sea. If the faults are assumed to be deep faults cutting the basement, this may imply that original gas methane may be of thermogenic origin. The

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 57– 75. DOI: 10.1144/SP340.5 0305-8719/10/$15.00 # The Geological Society of London 2010.

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aim of this paper is to analyse in a comprehensive fashion the geophysical and isotope data relevant to the distribution of gas seeps, palaeoriver canyons, deep faults, heat flow anomalies and methane composition in the northern Black Sea to shed light on the origin of gas methane.

Gas seepage in the northern Black Sea Numerous gas seeps fields have been found on the shelf of the Black Sea, in the transitional zone between shelf and continental slope, in Sorokin and Kerch-Taman troughs, Shatsky and Arkhangelsky ridges and Giresuan basin (Fig. 1). They are mostly non intensive, their heights above the sea floor reach 100 –130 m but more often they do not exceed 50 –70 m (Fig. 2) At the same time some discrete outbursts are up to 250 –300 m and even 400 –500 m high. The exact number of gas releases is not accurately known because of poor echo sounding coverage. When gas seeps have been studied in detail the results are impressive. For example, in 2003 and 2004 two expeditions (CRIMEA Project) mapped 2778 seeps in an area

of 1540 km2 in the Dnipro palaeo-delta, the NW Black Sea (Naudts et al. 2006). These were all located in the transition zone between the shelf and continental slope at depths from 66 to 825 m. The gas releases are mainly confined to the tops of small uplifts in the sea floor relief (Fig. 3a) though they can also occur on the slopes of uplifts (Fig. 3b) and in areas of low relief lows as well as on the flat sea floor (Fig. 3c). One set of observations in the vicinity of the Danube River canyon, showing that 52 –61% of gas seeps are at depths less than 300 m, about 15% are at depths from 300 to 1000 m and less than 1% are at depths more than 1000 m. Recent hydro-acoustic investigations (Popescu et al. 2007) confirmed large-scale gas seepage within the canyon and submarine alluvial fan of the Danube River revealed earlier (Polikarpov et al. 1989). In some areas (44807, 880 – 30837, 380 ; 44806, 860 – 30838, 980 and others) the gas escapes directly from the neotectonic disruption of the most upper sediments (Shnyukov & Ziborov 2004). Figure 4 illustrates more than 100 gas seeps within the WE belt of 8 –10 miles wide from 358550 to 378E offshore of the Kerch Peninsula

Fig. 1. Location of the mud volcanoes gas seeps and gas hydrates in the Black Sea (C ¸ ifc¸i et al. 2002; Kutas et al. 2002; Vassilev & Dimitrov 2002; Egorov et al. 2003; Shnyukov & Ziborov 2004; Klauke et al 2006; Popescu et al. 2007). Triangles in black, mud volcanoes; circles in red, gas seeps; squares in ice blue, gas hydrate; bold dashed lines in black, shelf edge; bold squared lines, boundaries of tectonic units; filled rectangular in red, Dnipro palaeo-delta area; I, NW Shelf; II, Kerch-Taman trough; III, Sorokin trough; IV, Tuapse trough; V, Shatsky ridge; VI, Andrusov ridge; VII, Arkhangelsky ridge; VIII, Giresuan basin.

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Fig. 4. Gas seeps, local uplifts and submarine canyons offshore of the Kerch Peninsula. Circles in red, gas release; thin lines, isobaths; bold elongated contours with radiating hachures, local uplifts; bold bifurcated lines, submarine canyons; double squared lines, boundary of the continental slope; dashed lines, contours of the burial palaeo-deltaic accumulations; Y-shaped.

Fig. 2. Acoustic images of gas seeps (a) single; (b) groups; (c) massive releases.

Fig. 3. Gas seepage on the tops of small uplifts (a) on the slopes of uplifts; (b) in relief lows; (c) on the flat bottom.

(Shnyukov & Ziborov 2004). The gas emissions on the outer shelf and upper continental slope occur at depths of 115–240 m, with them reaching a height of 60–195 m. The gas releases are mostly located in small submarine-erosional cuttings in a flat sea floor of gently sloping plain. A number of seeps are directly recorded on the canyon slopes, on the tops of submarine mounds, on flanks of small forms of submarine relief (Fig. 4). Generally the belt of gas emissions occurs within the palaeo-delta of the Don River, which forms a large fan in the Black Sea southward from the Kerch Strait. Shimkus et al. (1998) discovered the Quaternary gas-saturated sediments on the Kerch shelf. The gases (mainly methane) are predominantly concentrated in the sandy-siltstone interlayer overlain by clayey pillow in the parallel bedded strata. In sediments, anomalous methane accumulations are distributed in patches. A direct link is not observed between them and small local uplifts (eroded diapirs). On the Kerch-Taman shelf a marine hydro-gas survey recorded the content of hydrocarbon gases in a range from 8  1024 to 340  1024 ml/l in the subsurface water strata (Shimkus et al. 1998). A linkage of large anomalies of hydrocarbon gases to gas seeps of the upper parts of the submarine canyons allows these authors to arrive at the conclusion that gas is of biogenic origin resulting from the anaerobic generation of methane. At the same time they observed the anomalies in local uplifts and fault zones. As on the northwestern shelf, the localization of the gas flames is also controlled by the

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structural-tectonic pattern of the area. The relationship is observed between gas concentration and the uplifts of the Kerch continental slope.

Carbonate formations as derivatives of gas emissions on the northern Black Sea floor In many cases gas emissions are accompanied by genetically related carbonate formations. These authigenic associations are represented by carbonate build-ups, carbonate shelly lithificates in the areas of gas seeps and carbonate crusts on some mud volcanoes. Gevorkyan and his colleagues (1991) were the first to find authigenic carbonates in the areas of gas seeps on the NW Black Sea shelf. Direct observations and sampling from the submersibles ‘Yago’ (Michaelis et al. 2002), ‘Sever-2’ (Shnyukov & Ziborov 2004) and ‘Bentos-300’ (Gevorkyan et al. 1991) discovered the fields of gas chimneys in the NW Black Sea at depths from 230 to 1738 m. The fragments of a large methane chimney were dredged from the Lomonosov submarine massif on the continental slope at the depth of depth 1600 m (Shnyukov & Ziborov 2004). The restoration of the build-up demonstrates that it had a form of a hollow trunk of above 1.5 m high expanded upwards (Fig. 5). The mineralogical composition of the methane chimney shows that its matrix consists of calcite with a small amount of aragonite. Manganese, magnesium, barium and strontium are the impurities in the lattice of the carbonates. The identification of the stable carbon isotopes from different parts of the carbonate build-up and radiocarbon dating were carried out to understand the genesis and to determine the age of the carbonate build-up. The radiocarbon dating shows that the age of the main fragment from the base of the build-up (45 cm long) is of 9050 + 180 years while its upper one is 7540 + 160 years in age (Table 1). It is likely that this fragment has been formed during 1500 years of continuous production of the methane seepage. If we assume a constant rate of the carbonate precipitation one can suggest that 33– 34 years are

Fig. 5. Restoration of a methane chimney. The arrows show the sampling places for isotopic dating.

needed to accumulate 1 cm of build-up. In fact the growth of build-up could be affected by the capacity of the gas releases, possible breaks or variation in the gas intensity, turbidity of water, and finally, by palaeogeographical situation. As only one fragment of build-up was studied one can only assume that build-up on its base is older, but upward it is much younger. Alongside the thoroughly studied bodies of build-ups (derivatives of gas seepage) an important research topic is the carbonate lithificates from the shelf composed of shelly detritus and cemented by carbonate micrite. Lithified carbonate banks within

Table 1. Results of the radiocarbon dating of carbonate formations Sample features Bacterial mats Authiganic carbonates (spherules) Carbonates of the buildup main body, upper part Carbonates of the buildup main body, lower part

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C age, years

240 + 120 5430 + 130 7540 + 160 9050 + 180

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the fields of gas seeps are derivatives of the transformation of hydrocarbon gases upwelling from the deep layers. Carbonate crusts from mud volcanoes represent another type of authigenic carbonates. For example, numerous fragments of flat carbonate crusts were dredged on a depth of 2093 m from the caldera of the mud volcano Dvurechensky in the Sorokin trough (Shnyukov & Ziborov 2004). They are up to 15  20 cm in size and up to 3 cm thick, with one side covered by a mucous bacterial layer up to 1 cm thick. Through small openings, almost round and larger (up to 3 cm) of irregular form were discovered under the bacterial layer. Some openings of the outer side are healed by carbonate knobs producing an irregular surface. The crusts on the sea floor seem to form a continuous cover overlaying the mud sediments. Gas escapes from the underlying gas saturated strata through the openings in the sheet of crusts. Among the other debris there exist tubular structures and their fragments which likely represent the initial stage of growth of tree-like build-ups. The continuous carbonate sheet upon the mud has an age of 5100 + 140 years. The average value of the growth rate of the tubular structure (1 cm per 80 years) differs considerably from that of the above chimney from the rocky base of the Lomonosov submarine massif (1 cm per 34 years). Probably trophic activity of carbonate-creating bacteria is oppressed by turbid water near the sea floor permanently supplying gas through the numerous openings in the sheet of horizontal carbonate crusts. The radiocarbon dating of carbonate formations indicate that they were formed as early as 36 500 years ago (Gulin et al. 2003). Biochemical and chemical methane oxidation explains the genesis of carbonate formations (Egorov et al. 2003). The generation of authiganic carbonates in the zone of the hydrogen sulphide contamination seems to be the result of leakage of carbon dioxide (CO2) from fuel gas as well as still poorly studied processes of microbial oxidation of methane into carbon dioxide.

Geological – geophysical investigations in the areas of gas seeps of the northern Black Sea On the NW Black Sea shelf the geological structure of the gas venting features was studied by seismic-acoustic profiling, which resulted in the detailed signature of sedimentary cover up to 560 ms two-way travel time (TWT) below the sea floor (Korchin et al. 1999; Kutas et al. 2002). The profiles deal with two different geomorphological types of the floor relief where the gas releases

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were registered (for location see Fig. 6a). The first is situated in a shallow area with the flat sea bottom whereas the second is located in the transitional zone between shelf and deep-water basin. Figure 7a demonstrates a typical example of the seismic record along the profile in the area where the depth varies from 20 to 50 m. Three disturbance zones in reflecting horizons are clearly defined on the air-gun seismogram. They are registered in the vicinities of 2, 4.5 and 7 km. The fourth zone is less distinctly seen near 8.5 km. The disturbances are traced down to 300 ms TWT below the sea floor. The structure of the sedimentary nappe in each zone is different. In the first zone some beds are crumpled and slightly uplifted in the centre. The second zone (4.5 km) is characterized by disturbance of sedimentary geometries and fault of its right part. The anticline is observed in the third zone of disturbance (7 km). More detailed characteristic of the uppermost sedimentary cover was obtained with a sparker. Figure 7b presents the fragment of the same profile in the interval of 3.9 –5.1 km. The cone-like body is clearly seen beneath 55 ms TWT. It is bounded by structurally heterogeneous and crumpled deposits that act as a strong reflector. There are relatively homogenous sediments with a strong attenuation of elastic waves in the middle of this area. Such phenomenon is characteristic of mud enriched in gases. Only air gun was used for the second profile (Fig. 8). Here the gas seeps are related to distinctly defined diapir features, which are observed in the bottom relief as alternating cone-shaped knolls and synclinal lows. The height of knolls with symmetrical form is 40–70 m above the sea floor where a water depth varies from 250 to 280 m. Slides, strippings and steps are observed on the slopes of the mounds. The knolls are complicated by vertical disruptions of layer geometries (1–2 km wide) filled by crumpled and destructed mud sediments. The number and intensity of gas seeps in the second site considerably exceed those of the site with smooth relief. They are registered on the flat uplifts and in diapir and synclinal zones. However, independently of the type of relief all of the gas emissions are related to the disturbances of sediments integrity. In both sites the disturbance zones of different scale are vertical channels with minor horizontal fractures. In some cases sites of gas emissions are accompanied by distinct intensive anomalies of heat flow of 20 –200 mW/m2 (Kutas et al. 2004). Figure 9 is an example of such a situation. Maximum local anomaly of heat flow is observed on the rise of a continental slope in the West Crimea fault. A width of intensive anomalies of heat flow rarely exceeds 5–7 km. They are related to heat

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Fig. 6.

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Fig. 7. Examples of the seismic records on the shelf (a) air gun; (b) sparker.

flux of gas and fluid flows along the flank of depression in slopes and in zones of tectonic disruption. Increased values of the heat flow density (55– 65 mW/m2) were registered among the canyons of the Dnieper and Kalanchak palaeorivers (Lu¨dman et al. 2004; Poort et al. 2007) where gas seeps are widespread. Detailed study shows that this gas seep area is characterized by strong small-scale variability of heat flow which is controlled by fluid and gas migration (Kutas & Poort 2008). Here hot gas and fluids are upwelling along present-day active faults delineated by a multi-channel seismic reflection profiling (Lu¨dman et al. 2004). Unfortunately it is not possible to define a real extension depth of the faults due to shallow wave penetration. The above seismic shallow results show clear evidence for a relationship between gas seeps and the tectonic disruptions of the uppermost strata geometries. Really, on the NW shelf the gas releases are situated in areas of deep faults zones. Figure 6a demonstrates deep faults zones identified from interdisciplinary interpretation of geological and geophysical data (Kutas et al.

2004). They can be subdivided into two groups. The first group includes the deep faults, which bound the major tectonic units, the second – deep faults within single features. In Figure 6a the first group includes the fault zone of the southern margin of the East-European Craton (1) and OdessaSinop fault separating the West and East basins (2). The second group is represented by the NorthRomania –Crimea fault (3), Central-Romania – Crimea fault (4) and Western Crimean (5) faults. In the northern Black Sea gas seeps are known to be also related to other major deep faults, such as Peceneaga-Camena (Popescu et al. 2004) and the Kalamitsky Ridge bounding fault zone (Peckmann et al. 2001). Gas seeps are sometimes linked to minor faults, fractures, fissures, and slope slumps, which in turn are derivatives of tectonic activity in the zones of deep faults. This situation is clearly seen in Figure 6b which presents a more detailed map of the Romanian shelf (Popescu et al. 2007). It is worth noting that alignments of seepage distribution in some cases correspond with the directions of minor faults.

Fig. 6. The fault zones and gas seeps in the NW Black Sea. (a) Dashed lines in violet, faults; circles in red, gas seeps; lines with triangles, profiles of seismic-acoustic observations; fine dashed lines, shelf edge; fine dotted lines, avandeltas and canyons. Gas seeps and oil/gas fields in the Danube Canyon area (b) (modified from Popescu et al. 2007). For location see smaller insert in Figure 6. Filled contours in red, oil/gas fields; circles in red, gas seeps; bold lines, active faults; hatched and dotted line, shelf break; PCF, Peceneaga-Camena fault.

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Fig. 8. A fragment of air gun record on the shelf break. Triangles, gas seeps.

Figures 6a and 6b demonstrate that gas seeps are mostly related to palaeorivers, their avandeltas and canyons (Kalanchak, Dnipro, Dnister, Kagulnik, Danube). As deep faults and water flows gave rise to palaeorivers and their canyons (Allison & Palmer 1980), it is clear that tectonic disruptions of sedimentary cover play an important role in

Fig. 9. Heat flow and gas emission in the vicinity of the Western Crimean fault. Parallel lines in red, the Western Crimean fault; lines in red, faults; thin lines in yellow, geological boundaries; circles in red, heat flow values; pillars with upper sharp edges in light blue, gas seeps.

producing vertical pathways for gas migration. A relationship between the gas emissions and zones of deep faults suggests the aboigenic origin of methane. Recent analyses of the chemical and isotope composition of sea floor and deep gases and gas hydrates from the Black Sea provides further evidence that gas is of mostly thermogenic origin. Relative abundance of ethane and higher hydrocarbons are recognized in the sediments from the deep fault in the Sorokin trough (Stadnitskaia & Belen’kaya 2000). Here an anomalously high content of polycyclic aromatic hydrocarbons and the absence of correlation between organic carbon and hydrocarbon compounds is also observed. A contribution of thermogenic gas to methane origin was reported by Peckmann et al. (2001) and Klauke et al. (2006) It has been shown there is a very high concentration of saturated hydrocarbon in mud breccia and absence of any unsaturated homologues in the Dvurechensky volcano (Blinova et al. 2003). There is an absence of any unsaturated homologues and low content of the organic matter (median value is 1.5%) in clays from the fault zone on the continental slope in the NW Black Sea (Kruglyakova et al. 2004). The presence of higher carbons is reported by Mazzini et al. (2004) from a locality on the northern upper slope and from mud the volcaneous in the central Black Sea (Limonov et al. 1997). A composition of gas is the same in the mouth of the prospecting boreholes in the zones of deep faults in the vicinity of gas fields and in the bottom samples from the bays off the Western Crimea (Yurovsky 2001). Thermogenic methane (d13C varying between 236.7‰ and 227.4‰) has been obtained from the Dolgovskoy Mound on the northern Shatsky ridge in the

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western Black Sea (Mazzini et al. 2008). Marine sediments are influenced by localized deep hydrocarbon fluids at the very initial stages of sedimentogenesis and diagenesis that destroy organic material (Dmitrievsky & Valyaev 2002). The western Black Sea regional magnetic anomaly over the OdessaSynop deep fault is caused by authigenic magnetite produced in the16 km sedimentary cover by the reduction of iron hydroxides in the hydrocarbon environment (Kravchenko et al. 2003). The gas and gas condensate fields on the NW shelf (e.g. Golitsynskoe, Yuzhno-Golitsynskoe) are localized in the junctions of the intersection of the OdessaSynop fault with faults of the northeastern strike. There exist gas escapes from the crystalline rocks (plagiogranites, diorites and volcanites) of the Lomonosov submarine massif (Shnyukov & Ziborov 2004). The d13C values from 225 to 227.0‰, and dD values from 2120‰ to 2160‰ of Maikop clays and shales in Black Sea basin (Lukin 2006) are strong evidence for a thermogenic gas source (e.g. Igari 2001; Sassen et al. 2003; Chapman et al. 2004; Milkov et al. 2005; Pohlman et al. 2005). In the mud volcanoes of the Sorokin trough the original source of thermogenic gases is likely to be located below the Maikop shales (Stadnitskaia et al. 2007). As the Maikop series occurs on a 6 km depth in the Sorokin trough (Tugolesov et al. 1985), the source of gas occurs within the sequence where the temperature is about 150 8C which is too high to permit microbial activity (Machel & Foght 2000). However, it does not necessarily imply that the only source of methane seeps in the Black Sea is degassing deep layers. Small amounts of gas in very limited areas may result from the decomposition of shallow organic matter through the action of bacteria (e.g. Egorov et al. 2003; Popescu et al. 2007) or biodegradation of redeposited thermogenic hydrocarbons (Stadnitskaia et al. 2007). Therefore in some localities there exists mixtures of thermogenic and biogenic gases in the northern Black Sea (e.g. Kruglakova et al. 2004; Stadnitskaia et al. 2007). A possible role of deep faults in the formation of gas seeps in the Black Sea is as follows. Gases as a derivative of petroleum are produced in the upper mantle (Mendeleev 1877; Gold 1993). They ascent through faults and accumulate below mud-rich sediments at 10– 20 m beneath the sea floor. The gases escape to seawater when the sedimentary integrity is disrupted due to the gas exceeding the pressure in the capillary system of gas bearing bodies that overburden layers. The crucial condition for the occurrence of gas seeps is pathways for gas migration to shallow depths. Only long-lived active faults in geological horizons can serve as gas conduits. Here tectonic rejuvenation has been

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maintained by the interaction of the African – Arabian and Eurasian plates and the subsidence of the central basin since the Late Cretaceous (e.g. Tugolesov et al. 1985; Okay et al. 1994).

Mud volcanoes of the northern Back Sea Mud volcanoes are important contributors of gas methane in the northern the Black Sea although they are mainly distributed in three separate areas south and east– south of the Crimea and in the area of the Tuapse trough and Shatsky ridge (Fig. 1). About 55 mud volcanoes are known in the northern Black Sea. Based on the study of archival documents of the Hydrographical Survey of the former USSR, we can predict discovery of a minimum of 100 mud volcanoes in future. Mud volcanism is a peculiar manifestation of the diapir tectonics (Shnykov et al. 1971). Gases of deep layers rise upwards under abnormally high pressures of 300–400 105 Pa along disruptions throwing out water, clayey masses and fragments of hard rock. As a result the positive forms of relief are produced on the sea bottom and surrounding land. Submarine mud volcanoes are usually calm. But sometimes catastrophic gas bursts take place. The eruption of the Golubitsky volcano in the Azov Sea in 2000 can serve as an example of a catastrophic explosion in shallow water (Fig. 10). The explosion suddenly happened a half-kilometre along the coast. A small island with salsa in the centre appeared on the sea floor at a depth of 10 m. Over a period of many hours there were huge gas releases, mud breccias, large boulders and rock debris. The mud volcanoes of the south area (Fig. 11) are characterized by a circular or ellipsoidal shape and well developed feeder channels. The height of these mud volcano edifices above their bases is from 20 to 120 m. Geomorphologically, these mud

Fig. 10. Eruption of the Golubitsky Volcano in the Azov Sea.

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Fig. 11. Location of mud volcanoes (triangles in black), gas seeps (circles in red) and gas hydrates (squares in yellow) in the West Black Sea Basin (a) and Sorokin Trough (b) (Meisner et al. 1996; Kruglyakova et al. 2004). MSU, Moscow State University.

uplifts are of a different configuration. The Moscow State University (MSU) mud volcano has complex crater-like form with a distinct flange rim while the Yuzhmorgeologiya volcano is cone-shaped with an almost round base. These mud volcanoes have been formed under stretching conditions of thick sedimentary cover during the long eruption of successive and massive mud flows from the crater along the slopes (Krastel et al. 2005). The Sorokin trough (Fig. 1) is a large structural depression of the present-day deep basin in the Black Sea (Tugolesov et al. 1985). It extends along the southeastern margin of the Crimea Peninsula being 150 km long and 45– 50 km wide. The water

depth is 600–2100 m. The Tetyaev uplift and the Shatsky ridge bound the trough on the south. The Sorokin trough belongs to a system of Oligocene – Miocene troughs of the eastern Black Sea Basin. The most intensive subsidence of the trough occurred during the Neogene –Anthropogene. The north side of the trough has a dip between 30 and 408 and a rupture character; the southern side is gentle The Sorokin trough is characterized by a great thickness of Maikop deposits, up to 5 km thick. The Paleocene –Eocene Sorokin relief depression is supposed to be formed as a trough in the classic sense definitely in the Maikop time. The top of the Maikop sediments is flattened in

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the trough and generally tilts towards the axis of the Black Sea Basin. Brachyanticlinal folds are widespread within the Maikop deposits. Individual groups of folds form several chains of 12 to 40 km long and up to 10 km wide striking along the long axis of the trough from the SW towards the NE. During the Pliocene the Sorokin trough was united with the Tuapse trough within the common Black Sea Basin. Compressive deformation has influenced the growth of diapiric features and facilitated fluid flow to the seafloor and the formation of mud volcanoes (Dvurechensky, Sevastopol, NIOZ, Yalta, Kazakov and Odessa.) above the diapirs (or above their edges) or above the diapir ridges of mainly east –NE strike (Fig. 11b) in the Pliocene and Quaternary deposits. Mud volcanoes are concentrated in three areas in the Sorokin trough. The eastern area is characterized by thick fluid saturated deposits of the Don-Kuban alluvial fan. There exist relatively small cone-like mud volcanoes with two or three individual cones in this area. In the central part large single cone-like mud volcanoes are observed. The Kazakov volcano is the largest among them. The western part includes a number of small cone-like mud volcanoes with destroyed tops. The Dvurechensky mud volcano is characterized by the occurrence of gas releases. During the 57th cruise of the R/V ‘Professor Vodyanitsky’ two large gas fountains were registered within the caldera of this volcano (Shnyukov & Ziborov 2004). Their diameter is 400 m, the height is 850 m above the sea floor where the depth of the sea is about 2000 m (Fig. 12). Their horizontal shape is not round but oval. These gas fountains were not observed here earlier. It is not to be excluded that these fountains are of intermittent origin.

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In the Sorokin trough the high-resolution multichannel seismic observations were conducted to investigate the subsurface structure of the deep seeps at mud volcanoes (Krastel et al. 2003; Wagner-Friedrich et al. 2008). A seismic image of the Dvurechensky volcano is presented in Figure 13. Reflection events were observed within the upper 250 m of the volcano cone. At greater depths they continue in a subsidence structure within the conduit area forming concave-shaped layer geometry. The feeder channel is clearly seen on seismic records. The diameter of the mud volcano on the base is 800–1000 m, and its conduit has the same shape and diameter up to a depth of 600 m below sea floor. The feeder channel penetrates several horizontal sedimentary layers but the stable reflecting boundaries are traced to a depth of 400 m below sea floor. A fault is identified in the centre of the conduit which seems to be a pathway for gas and fluid migration. It is worth noting that there is a great resemblance between the upper part of a conceptual presentation of the mud volcano structure (Shnyukov et al. 1971) and the seismic image of the Dvurechesky mud volcano (Fig. 13). The first three dimensional (3-D) seismic investigations carried out across the Sevastopol mud volcano have made it possible to develop a 3-D morphology of a collapse structured mud volcano and to propose an evolution model (Wagner-Friedrich et al. 2008). The volcano is located above a buried diapiric feature with two ridges. An explosive eruption of overpressure fluids resulted in the collapse depression. Several cones were formed within the depression by subsequent quiet mud extrusions. The mud volcano Sevastopol was supplied with gas and fluids, which migrated along the deep fault of the southern limb of the central diapir. The

Fig. 12. A gas release within the caldera of the Dvurechensky mud volcano.

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Fig. 13. A sub-surface structure of the Dvurechensky volcano (CRIMEA project results 2003).

bright spots were observed either within the anticline features on the diapir tops or along the rise of thinning out beds with diapirs at a relatively constant depth of about 300 ms TWT below the sea floor. They are caused by free gas in the underlying rocks. Gas and fluids migrate predominantly toward the top of diapir where they are accumulated. The bright spots seem to represent the base of the stability zone of gas hydrates that plays the role of a trap.

Gas-hydrate in the northern Black Sea Gas-hydrates are referred to as non-traditional hydrocarbon raw materials. As in the Black Sea their methane reserves are estimated to be 1–5  1012 m3 (Vassilev & Dimitrov 2002), gas production from hydrates may be of great commercial importance in future. For the first time natural gas-hydrates were discovered in a sample from the Danube River fan during the R/V ‘Moscow University’ as early as 1972 (Efremova & Zhizhchenko 1974). Now the area of the Black Sea that is suitable for gas hydrate formation is assessed at 288 100 km2, representing about 68% of the total Black Sea or almost 91% of the deep-water basin (Vassilev & Dimitrov 2002). In the Black Sea gas hydrates have been recovered in specific geological environments on a depth of water that exceeds 700 m namely: in submarine fans, in fractured zones and within mud volcanoes (e.g. Meisner et al. 1996; Bohrmann et al.

2003; Lu¨dmann et al. 2004; Shnyukov & Ziborov 2004; Popescu et al. 2007). Occurrences of gas hydrates in sedimentary cover are mostly inferred from observations of bottom stimulating reflectors (BSRs) which are considered to mark the base of their stability zone. Free gas often occurs beneath the BSR. These situations are exemplified in Figure 14 by a fragment of high-resolution reflection seismic profile from the Danube River fan offshore of Romania (Popescu et al. 2007). BSRs are marked either as a distinct reflection with reversed polarity, or as the top of enhanced reflections. Free gas is indicated by acoustic turbidity and enhanced reflections. BSRs and gas facies are confined inside a channel-levee system, with free gas mainly concentrating at the channel axis. Another area of the existence of BSRs is the vicinity of the Dnipro Canyon within the large deep-sea complex (Lu¨dmann et al. 2004; Zillmer et al. 2005). In both areas BSRs occur in the water depth of 700–1800 m. The localization of BSRs in the fan sediments is a characteristic of the estuaries of the Danube and Dnipro palaeorivers. It is worth mentioning a spatial coincidence of gas seeps and BSRs in the Dnipro-Danube deep-sea fans. Despite sampling gas hydrates from Dvurechesky, Yalta and Odessa mud volcanoes (Bohrmann & Schenk 2002), BSRs were not clearly identified in the Sorokin trough (Krastel et al. 2003). However, pronounced lateral amplitude variations were registered above the diapirs and next to the conduits of mud volcanoes (Fig. 15). They seem to map the base of the gas hydrate stability zone.

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Fig. 14. Bottom-Simulating Reflectors (BSR) offshore of Romania (Popescu et al. 2007).

Gases produce hydrates being enclosed in water molecule cages in specific thermodynamic environments. The physical parameters of a gas hydrate zone depend on gas saturation, gas composition and water chemistry (Trofimuk et al. 1984; Minshull & White 1989; Zatsepina & Buffett 1998). In particular, the presence of heavy hydrocarbons and CO2 increase the temperature of the hydrate

formation while increasing the NaCl concentration decreases the temperature by 1–4 8C (Fig. 16). Such a temperature fluctuation in the sediment changes the thickness of the hydrate zone by 50 –400 m. The phase equilibrium of gas hydrates is very sensitive to the variation of temperature. If the gradient of pressure is 1– 2 MPa/K the temperature

Fig. 15. Amplitude anomalies next to the feeder of a mud volcano in the Sorokin trough (Krastel et al. 2003).

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Fig. 16. Diagram of gas hydrate stability.

variation of 1 8C produces the change in the hydrate zone thickness by 50 –100 m. Even negligible variations of temperature (including variation of temperature with time) should be therefore taken into account in estimating the zone boundaries and its thickness. This variation is caused by a change in the sea floor temperature, conditions of heat exchange in the sub bottom water, sedimentation etc. At the same time the continuous process of sedimentation leads to a decrease in a heat flow in the upper part of the sedimentary strata and its increase in the lower part. When isotherms gradually displace upwards they accelerate the destruction of the gas hydrate layer from the base and build up it in the upper part. Hydrate destruction is accompanied by absorption of heat (400 – 600 KJ/kg in the interval 0–20 8C) and formation of hydrates is accompanied by the emission of heat. During the 61st cruise of the R/V ‘Professor Vodyanitsky’ the temperature of sediments was measured along the whole core with a step of 10 cm directly after uplift of a gravity core. Gas hydrate was clearly registered in some layers by a steep decrease in the temperature of 5–8 8C, with this phenomenon being used as an additional diagnostic indicator of gas hydrate in a core (Fig. 17). Thus, the positive geothermal anomalies should exist above the gas hydrates in a case of their formation, and the negative – above the gas hydrate accumulation, which is subjected to destruction. The state of these opposite processes on the top and the base of gas hydrate layers depends on the rate of sedimentation. A low rate of sedimentation should compensate these processes and the heat released (absorbed) by formation (destruction) of

Fig. 17. A distribution of temperatures in cores of bottom sediments. (a) gas-bearing hydrate deposits (sea depth 2041 m); (b) gas hydrate deposits are absent (sea depth 2098 m).

hydrate does not affect practically on the amount of heat flow. On the contrary, if the rate of sedimentation is high (about 0.1 cm a21) an increase in the temperature retards at the top of the hydrate pool, compensation is disturbed and gas hydrate layer is built up from above. If gas hydrate occupies about half of the pore space, the change of heat flow will be about 5 mWt/m2. In the case of avalanche sedimentary accumulation, which is possible in some alluvial fans of permanent streams or in the zones of intensive development of slides, the rate of growth of the sedimentary layer drastically passes ahead of its temperature change. Therefore, in the newly formed strata or below it favourable environments exist for the increase in gas hydrate thickness or preservation of gas hydrate accumulation. An analysis of the gas hydrates formation (destruction) or their stability requires an account of possible change of thermodynamic conditions with time which can be caused by not only physical-chemical processes in sediments but also a temperature variation of the bottom water layer, sea level change among other factors. If the temperature of water in the Black Sea during the Wu¨rmian glaciation was 4–6 8C lower than now (the present-day value is 9.1 + 0.1 8C) the thickness of the gas hydrate area had to be 100–300 m greater. Since that time the traces of

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glaciation have practically vanished. However, at a depth of 400– 700 m (at the base of the gas hydrate layer) the temperature is still 40–50% lower in comparison with the maximum anomalous effect of cooling. This is 1.5–2 8C if the bottom water temperature decreased to 5 8C and 3 –4 8C if it approached to 0 8C. Figure 18 demonstrates the present-day distribution of temperatures in the sedimentary deposits affected by 5– 9 8C cooling. The rise of temperature has resulted in the decrease in the thickness of the gas hydrate zone by 50– 100 m during the last 7000–8000 years. A huge volume of gas has escaped and this has led to the formation of the sub hydrate deposits. In this area the traps have been formed as elevation of the bases of the gas hydrate layers. Similar elevations are formed even by a negligible rise in the temperature, which can be caused by basement swells, clayey diapirs, faults and other structural elements. The quick formation or destruction of gas hydrates can be a direct reason for originating considerable heat flow anomalies. However, as is seen on the pressure v. temperature stability diagram (Fig. 16), the base and the top of the gas hydrate zone react to the temperature changes in different ways. In this case the base is more sensitive due to vertical (upwards) conductive heat flow of the Earth. Therefore the gas hydrate pool destroying from below is building up in its upper part. This process is accompanied respectively by decreasing and increasing of a heat flow. Based on the stability curves of gas hydrates in the system of methane plus 3.5% solution of NaCl (Fig. 16), an estimatate was obtained for the depth

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and thickness of the hydrate-formation zone in the Western Black Sea sediments (Kutas et al. 1999) The calculations were done along the three profiles across the continental slope of this basin. For evaluating thermal regime in determining the top and base in the zone of possible gas hydrate generation the simplest method is based on the temperature calculation using a determination of heat flow and thermal conductivity. These parameters are employed to define the gradient and the subsequent extrapolation of the temperature in the subsurfice sediments. In this area the observed heat flows are 20–35 mWt/m2, thermal conductivity of the sediments is 0.9–1.1 Wt/ (m K) and geothermal gradients vary in the range of 20 –40 K/km (Kutas & Poort 2008). A decrease in the gradients could be expected with the depth due to the thermal conductivity increase during compaction of the sediments. With the same probability the increase of the gradients can be however expected because of decrease in the thermal conductivity of gas hydrates bearing sediments. The gas hydrates conductivity is known not to be dependent on the temperature and is 0.5 Wt/n . K (Trofimuk et al. 1984; Sloan 1990). Gradients, therefore, could be approximately taken as stable in the upper layer of weakly consolidated sediments of several metres thickness, and could be used for the estimation of a temperature. Comparison of the latter with a phase stability diagram of gas hydrates allow us to determine the hydrate-generated zone. Such methods were used to construct the temperature plots and to eastimate the gas hydrate layer parametres along profiles I and II (Fig. 19).

Fig. 18. Affect of a decrease in the bottom temperature on the geothermal regime of the upper sedimentary strata: 1, stationary temperature; 2, present-day temperature distribution taking into account a decrease in the bottom temperature by DT ¼ 5 8C during the Wu¨rmian glaciation (duration of the cooling period t ¼ 60 000 years); 3, the same by DT ¼ 9 8C; 4 and 5, the temperature distribution at the end of the glaciation respectively by DT ¼ 5 8C and DT ¼ 9 8C; 6 and 7, anomalous temperature respectively by DT ¼ 5 8C and DT ¼ 9 8C.

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Fig. 19. A position of the base of the zone of possible gas hydrate-formation along the profile I (a) and profile II (b). 1 and 2, the values of the geothermal gradient are 20 and 30 8C/km respectively; 3, the geothermal gradient value is 20 8C/km after the end of the glaciation (10 000 years ago).

However, it should be noted that the calculations of the two temperature plots are approximate as they do not take into account the influence of different geological and physico-chemical processes such as heterogeneity of thermal-physical properties of sedimentary deposits, sedimentation process, change of climate, process of formation or dissociation of gas hygrates etc. All of these factors can result in temperature change by several degrees (centigrade scale) in the depth interval of the hydrate-formation that produces the change in the thickness of the hydrate formation layer by 50 – 100 m and more. The temperatures along the seismic profile III were modelled by the finitedifferences method (Fig. 20). The seismic results permitted to adopt the thermal-physical parameters of different horizons in real boundaries. The effects of sedimentation and climate variation were added to the temperatures calculated. Resulting temperature plot was used for determining the geometrical parameters of the hydrate-producing zone. As along this profile the temperatures were calculated more reasonably than along others, the parameters were estimated more accurately for the hydrate-producing zone. The favourable PT conditions occur on a sea depth of 600–650 m to produce methane the hydrates in the Western Black Sea Basin. As sea depth increases the thickness of the hydrate formation zone becomes greater and reaches 350–500 m on the rise slope and in the deep-water. To a great degree the thickness of the

Fig. 20. A distribution of temperatures in the upper sedimentary layer (dash curves) and position of the zone of possible hydrate formation along the profile III (see the profile location in Fig. 4). 1, with regard to the influence of the changes in the bottom temperature, sea level and sedimentation; 2, the same, without regard for the influence of these factors; 3, after the end of the glaciations (10 000 years ago).

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hydrate layer depends on the temperature gradient or heat flow. In the zones of a low heat flow (25–30 mWt/m2) the thickness of the hydrateformation layer is 350–400 m. If a heat flow reaches up to 40 mWt/m2 it decreases to 200– 250 m. A heat flow value of 60 mWt/m2 and more thins it out to dozens of metres. In the zones of anomalous high flows (80– 100 mWt/m2) gas hydrate layer probably is completely absent (Kutas et al. 1999). Thus, in the deep water of the Black Sea the necessary environments exist for accummulation of natural hydrocarbon as solid (gas hydrates) and gas phases, such as rather low temperature of the near-bottom water, appropriate pressure, alternating of porous and clayey deposits, regional abnormal gas saturated water, high rate of sedimantion etc. However, methane of biological origin cannot greatly contribute to the origin of gas hydrate. Hydrocarbon is abundantly formed during diagenesis of the organic material only under a temperature above 80 8C. Such a temperature occurs in the sediments of the Black Sea on depths over 2 km, that is, beyond the zone of hydrate formation. It implies that a volume of gas hydrate seems to be less than it is speculated in theoretically assessing their reserves. This inferences is supported by seismic observations in the northern Black Sea. BSRs are only mapped in the ares of the Dnipro and Danube Canyons (Naudts et al. 2006; Popescu et al. 2007) although subsurface gas hydrates are sampled from many locations.

Conclusion Several thousands of gas seeps are documented in the Black Sea, although their real number is unknown mainly due poor survay coverage. They occur mostly in the shelf –continental slope transitional zone in areas of submarine palaeorivers canyons location of which is controlled by deep regional faults. Gas emissions are often accompanied by carbonate formations such as chimneys, lithificates and crusts, which resulted from degassing deep layers. Mud volcanoes of different morphology and size are directly related to a tectonic activity. Gas seeps and hydrocarbon shows are associated with some mud volcanoes. At a depth of 600 –650 m below the sea floor favourable thermal and pressure conditions exist in a sedimentary layer for originating and stable occurring gas hydrates. Methane is a major compound of gas emissions. The gas is of mostly abiogenic origin although small gas releases may be produced by the decomposition of the near sea surface organic material through action of bacteria or biodegradation of redeposited

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thermogenic hydrocarbons. The deep faults serve as pathways for vertical migration of gas, with a long-time tectonic activity producing such gas conduits. We are very grateful to Prof. J. Cartwright (School of Earth, Ocean and Planetary Sciences, Cardiff University), Dr. R. A. Stephenson (School of Geosciences, King’s College, University of Aberdeen) and anonymous reviewers for critical comments and constructive recommendations to improving the content and clarifying many points of the manuscript.

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METHANE IN THE NORTHERN BLACK SEA northern Cascadia margin as inferred from isotopic (13C/12C and D/H) and molecular composition of hydrate and vent gas. Organic Geochemistry, 36, 703–716. Polikarpov, G. G., Egorov, V. N., Nezhdanov, A. I., Gulin, S. V., Kulev, Yu. D. & Gulin, M. B. 1989. Phenomena of active gas seeps from bottom on the western Black Sea slope. Proceedings of Ukrainian Academy of Sciences, 12, 13–16 (in Russian). Poort, J., Kutas, R. I. et al. 2007. Strong heat flow variability in an active shallow gas environment, Dnepr palaeo-delta, Black Sea. Geo-Marine Letters, 27, 185–195. Popescu, L., Lerocolais, G., Panin, N., Normand, A., Dinu, C. & Le Drezen, E. 2004. The Danube submarine canyon (Black Sea): morphology and sedimentary processes. Marine Geology, 206, 249–265. Popescu, L., Lerocolais, G., Panin, N., De Batist, M. & Gillet, H. 2007. Seismic expression of gas and gas hydrates across the western Black Sea. Geo-Marine Letters, 27, 1– 30. Popov, V. S. 1928. Mud volcanoes. Priroda, 6, 541– 554 (in Russian). Reeburgh, W. S., Ward, B. B., Whalen, S. C., Sandbeck, K. A., Kilpatrick, K. A. & Kerkhof, L. J. 1991. Black Sea methane geochemistry. Deep-Sea Research, 38, Supplement 2: S1189– 1210. Sassen, R., Milkov, A. V. et al. 2003. Gas venting and subsurface charge in the Green Canyon area, Gulf of Mexico continental slope: evidence of a deep bacterial methane source? Organic Geochemistry, 34, 1455–1464. Schmale, O., Greinert, J. & Rehder, G. 2005. Methane emission from high-intensity marine gas seeps in the Black Sea into the atmosphere. Geophysical Research Letters, 32, L07609, doi: 10.1029/2004GL021138. Shimkus, K. M., Moskalenko, V. N., Rayen, V. B. & Pitman, V. K. 1998. Quaternary gas-saturated sediments of the Kerch coastal water part of the Black Sea shelf and theirs relation to littoral facies. Bulletin of Moscow Society of Nature Researchers, Department of Geology, 4, 51– 55 (in Russian). Shnyukov, E. F., Naumenko, P. I. & Lebedev, Y. U. 1971. Mud Volcanism and Ore Formation. Naukova dumka, Kiev (in Russian). Shnyukov, E. F. & Ziborov, A. P. 2004. Mineral Wealth of the Black Sea. National Academy of Sciences of Ukraine, Kiev (in Russian). Sloan, E. D. 1990. Clathrate Hydrates of Natural Gas. Marcel Dekker, New York/Basel. Stadnitskaia, A. N. & Belenkaya, I. Yu. 2000. Composition and origin of hydrocarbon gases and their influence on diagenetic carbonate formation (Sorokin

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Brittle deformation, palaeostress field reconstruction and tectonic evolution of the Eastern Balkanides (Bulgaria) during Mesozoic and Cenozoic times F. BERGERAT1*, D. VANGELOV2 & D. DIMOV2 1

ISTeP, UMR 7193 CNRS-University Pierre & Marie Curie, Case 117, 4 place Jussieu, 75252 Paris cedex 05, France 2

Department of Geology and Geography, Sofia University ‘St. Kliment Ohridsky’, 15 Tzar Osvoboditel, Sofia, Bulgaria *Corresponding author (e-mail: [email protected]) Abstract: Brittle tectonics analysis and stress tensor reconstructions allow us to better depict the Mesozoic and Cenozoic geodynamical evolution of the Eastern Balkanides which is characterized by a series of overimposed basin-systems. The Late Permian–Triassic corresponds to a wide carbonate platform with local embryonic troughs. During the Jurassic–Early Cretaceous period, the area, limited by regional unconformities, was at first dominated by the existence of a deep basin until the latest middle Jurassic, then by its gradual closure during the Late Jurassic– Early Cretaceous. Traces of these tectonic events are found as a result of brittle tectonic data analysis, especially in the Strandja Zone where NE– SW extension and ENE–WSW compression may be related to the Early Alpine phases of rifting and shortening, respectively. The Late Cretaceous– middle Eocene period was governed by the back-arc basin/island-arc system of East Balkan/ Srednogorie zones. The inversion of these extensional zones occurred after the Maastrichtian, then important shearing and thrusting developed during the major shortening in the latest Middle Eocene. These Late Alpine tectonic phases were well characterized by brittle tectonics and the analysis of faulting in terms of stress tensors allows us to identify the main directions of extension of the rifting stage which is north–south to NNE–SSW, and the directions of compression of the Laramian and Illyrian phases, which are NE– SW and north– south, respectively. The relationships between folding and reverse and strike–slip faults are established as well as the occurrence of post-folding normal faults assigned to collapse process following the shortening.

The Bulgarian territory includes two large tectonic units: the Moesian Platform which is the foreland of the orogen and the Alpine fold-and-thrust belt of the Balkanides including the Rhodope Massif (Fig. 1). To the NW (across the territory of Serbia) the Balkanides system joins the Southern Carpathians and to the east it connects to the Pontides and the Caucasus (Bonchev 1930). Since the 1970s, Bulgarian geological literature has been complicated by the mixed use of platetectonics concepts and old geosynclinal nomenclature, resulting in difficulties in interpreting the geology of this region (e.g. Dercourt & Ricou 1987; Bonchev 1988; Ivanov 1988). Whereas new field-based studies have been undertaken in the Rhodopes in order to reinterpret this orogene globally (e.g. Burg et al. 1996; Ricou et al. 1998), the Balkan sensu stricto has not been revisited in terms of global geodynamics. Many of the ideas concerning this orogen, are based on regional studies. Particularly, the Early Alpine structures need to be better understood, correlated and

interpreted. They remain one of the main problems of the Balkanides. Several tectonic phases are classically accepted in the Moesia-Balkan area. Some of them have been described only on the basis of seismic profiles (e.g. Tari et al. 1997a, b; Georgiev et al. 2001), some are demonstrated by clear structural features in the field, others are only suspected on occurrence of unconformities, sequences and facies interpretation and some have been characterized in terms of palaeostress reconstructions in the Bulgarian part of the Moesian platform (Bergerat & Pironkov 1994; Shanov 1990, 2005; Bergerat et al. 1998). The present work concerns principally the eastern part of Balkanides and its aim is twofold: (i) characterizing the tectonic evolution by collecting and revisiting the existing data and by acquisition of new structural observations in the field, and (ii) defining the geometrical and mechanical characters of each successive tectonic event by microstructural measurements in the field, and palaeostress tensor computations.

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 77– 111. DOI: 10.1144/SP340.6 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Schematic structural map of the central part of the Alpine– Himalayan orogenic belt. The dotted rectangle indicates the studied area.

Main features and geodynamical evolution of the Balkanides According to the character, style and date of the orogenic deformations, Ivanov (1988) defined the foreland of the orogen, which is the Moesian Platform, and three tectonic belts within the Balkanides, which are, from north to south: the external, intermediate and internal Balkanides (Fig. 2). North of the Balkanides, the Moesian Platform (Fig. 3) was a wide stable area of shallow marine deposition with significant interruptions of the deposition especially in its eastern part at the end of the Triassic, Lower Cretaceous, Upper Cretaceous, Palaeogene and Miocene. The major unconformities have been related to major compressional events within the Alpine thrust belt (Georgiev and Dabovski 1997). The Early Alpine tectonics is concentrated south of the Moesian Platform, resulting in a gentle folding and subsurface shearing along its southern margin. During the Late Alpine (post-Lutetian) tectonics, part of the external Balkanides was thrusted over the platform southern margin. The external Balkanides, situated south of the Moesian Platform have autochthonous and

para-autochtonous character. They are subdivided into the Kula Zone, West Balkan Zone (both outside of the studied area) and Central Balkan-Forebalkan Zone. The latter consists in several north- and northeast-verging thrusts which took place during the end of Early and Late Alpine phases. The intermediate zone of the Balkanides includes tectonic elements of various ages and different characters. The Early Alpine deformation affected this zone, but the main deformations occurred during the Late Alpine phases: after the Maastrichtian and at the latest middle Eocene (e.g. Bonchev 1910; Kockel 1927; Jaranoff 1960; Bonchev & Karagyuleva 1961; Ivanov 1988). This intermediate zone includes (i) the East Balkan which is thrusted to the north and covers significant parts of the folded system of the Central BalkanForebalkan zone; and (ii) the Srednogorie Zone (Bonchev 1930) which contains fragments of the Early Alpine internal zones (Ivanov 1988), and constitutes a Late Cretaceous island-arc system (e.g. Boccaletti et al. 1974; Hsu et al. 1977) deformed after the uppermost Cretaceous and after the Middle Eocene. Some fragments of that system have been thrusted to the north, over the external zones of the Balkanides.

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Fig. 2. Principal structural zones of the Balkanides in Bulgaria (modified after Ivanov 1988).

The Internal Balkanides occupy the southern areas of the orogen. Single tectonic units of this part of the system were formed in Early Alpine time, during the Late Jurassic and Early Cretaceous. Their Late Alpine tectonic evolution was associated with deformation in the adjacent zones of the Balkanides and is a result of collisional and postcollisional processes during the Late Eocene and Oligocene, even Miocene. The main zones of the internal Balkanides are the Kraishte zone, the Rhodope Massif (both not studied in this work) and the Sakar-Strandja zone (Fig. 2). The Alpine evolution of the Rhodope massif is interpreted as a result of two deformation stages: (i) an intercontinental collision during the latest Early and the beginning of the Late Cretaceous, expressed in regional metamorphism by eclogitic and amphibolitic facies associated with synmetamorphic south-vergent thrusting of large amplitude (Burg et al. 1990); and (ii) an extension and tectonic erosion at the latest Cretaceous and during the Cenozoic, which started with the intrusion of large bodies of syn- to post-metamorphic granitoids, metamorphic core complexes (MCC) exhumation and continued with formation of Late Eocene and Oligocene depressions, filled by turbiditic sediments and products of bi-modal volcanism. The SakarStrandja zone is composed mainly of deformed

Triassic and Jurassic sediments, metamorphosed after the latest Middle Jurassic (Chatalov 1990). This zone includes several thrusts of syn- and post-metamorphic origin. Some depressions filled with thick NeogeneQuaternary terrigeneous sediments are superimposed on Early and Late Alpine tectonic fragments (Ivanov 1988): (i) the so-called Kraishte lineament composed of NNW–SSE and ENE –WSW trending Neogene-Quaternary grabens, located in SW Bulgaria and North Greece; (ii) the so-called Back-Balkan lineament, mainly composed of halfgrabens trending generally east –west, along the southern slopes of the Stara Planina (Srednogorie Zone); and (iii) the Upper Thrace Depression, a wide plain occupying the lowland of Maritsa, Sazliika and Tundja Rivers (Fig. 4). The structural features of the Balkanides cannot be explained by only a simple collision model between Africa and Eurasia during the closure of the Tethyan basin. A relatively deep basin existed during the Early Mesozoic, situated north of the current suture of the Tethys, which can be considered as a western prolongation of the Palaeotethys (Hsu 1977; Sengo¨r et al. 1980, 1984, 1988). There is no existing complete description of the basin evolution yet. Many of the ideas are based on information concerning separated areas. In the

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Fig. 3. Schematic map of the main tectonic zones and units in Eastern Bulgaria (modified after Bokov et al. 1987). 1: Part of the Neogene-Quaternary overimposed basins developed on the top of Balkanides; 2: Post-compressional Priabonian basins developed over the intermediate Balkanides; 3: Moesian Platform (GTS, General Toshevo saddle; BzB, Bezvoditsa block; VB, Vraninski block; SBG, Shabla-Bulgarene graben zone; GCH, Goni-Chiflik horst); 4: Central Balkan-Forebalkan Zone; 5: East Balkan Zone; 6: East Srednogorie unit, 7–9: Central Srednogorie (7: Stara Zagora strip, 8: Shipka-Sliven unit, 9: Tundja-Topolnitsa unit); 10: Sakar-Strandja Zone. CSD, Chudnite Skali Dislocation; NSF, North Strandja Flexure; PL, Ptichevo Line; PSF, Preslav-Skorpilovski Fault; YF, Yantra Fault.

studied area, the existence of an Early Mesozoic deep basin is demonstrated by: (i) a turbiditic and chaotic rock complex (‘East Balkan’ type, Tchoumatchenko et al. 1992) including Triassic and Lower Jurassic large blocks in a mudstone matrix of Lower and Middle Jurassic age (Peybernes et al. 1989; Tchoumatchenko et al. 1992); (ii) the two type of Triassic formations distinguished in the SakarStrandja Zone: shallow and deep marine (Chatalov 1990); the Upper Triassic metamorphosed flyshs of North Strandja are quite similar to those of same age of the East Balkan Zone; (iii) the Palaeozoic and Triassic rocks which constitute the ShipkaSliven unit (Kockel 1927; Ivanov et al. 1974) and can be correlated with a fragment of the SakarStrandja zone. The closure of this basin occurred

during the late Jurassic and the Early Cretaceous. The Balkanides can be considered as a ‘collage orogen’ in the sense of Sengo¨r et al. (1988) with Early Alpine (Cimmerian and Austrian) and Late Alpine (Laramian and Illyrian) tectonic phases. Thus the Alpine evolution of the Balkanides can be summarized as follows. (i) During the Triassic dominated shallow carbonate platform deposition, only in the East Balkan and the allochthonous part of the Strandja unit, Upper Triassic (Norian) calcareous turbiditic sequence was established. (ii) The Jurassic –Lower Cretaceous time is characterized by platform sedimentation on the Moesian platform and shows features of

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Fig. 4. Schematic structural map of the Eastern Balkanides with location of the microtectonic sites. Tectonic zones: CB-F, Central Balkan-Forebalkan; CS, Central Srednogorie; EB-ES, East Balkan-East Srednogorie; MP, Moesian Platform; ND, Neogene depressions; SS, Sakar-Strandja Zone. Main fault zones: BF, Bosna Fault; CSD, Chudnite Skali Dislocation; FM, Maritsa Fault; LF, Lekarnitsa Fault; NSF, North Strandja Flexure (detachment zone?); PL, Ptichevo Line; PSF, Preslav-Skorpilovski Fault; SPT, Stara Planina Thrust; SSBF, Sakar-Strandja boundary fault; SST, Shipka-Sliven Thrust; YF, Yantra Fault; YL, Yambol Line. Important structures and sub-units: BS, Bogdanovo-Slivovo antiform; CS, Central Strandja antiform; ID, Izvor metamorphic dome; SD, Sakar metamorphic dome; SI, Sveti Iliya dome; VS, Volcano-terrigeneous complex of Upper Cretaceous age in Eastern Srednogorie; SS, Shipka-Sliven unit; SZ, Stara Zagora strip. The yellow dots with numbers indicate the brittle tectonic measurement sites. The dotted line near the Black Sea coast indicates the approximate location of the cross-section of Figure 5.

passive continental margin deep-water basin south of it (most likely two basins formed by dextral transtension) until the latest middle Jurassic. The gradual closure of the basin(s) was associated with the enlargement of the Neotethys during the Late Jurassic and Early Cretaceous. The beginning of the Cretaceous is marked by the occurrence of a deep trough (the Nish-Trojan Trough) and laterally, to the south by a narrow margin against the Rhodopes, and to the north by a more stable and wide shallow-water carbonate platform (Moesia) (Minkovska et al. 2002). Fragments of the Central Srednogorie (e.g. Shipka-Sliven unit) and the Sakar-Strandja zone can be considered as elements of the continental margin of the Palaeotethys (thrusting of deep-marine Triassic sequence over the Lower Middle Jurassic rocks in Strandja, ceasing of deposition in the middle Jurassic in the East Balkans, propagating of the basin depocenter northward over the Moesian Platform). (iii) During the Late Cretaceous and the Cenozoic, after the final closure of the basin, the

geodynamical evolution of the Balkanides was controlled by the presence of the northern continental margin of the Neotethys basin and the associated north-dipping subduction zone, giving rise to the development of a back-arc basin (the East Balkan zone) and an island-arc system (the Srednogorie zone), the closure of which occurred during the Maastrichtian – middle Eocene, resulting in shearing and thrusting onto the Moesian platform. The Late Eocene –Oligocene, Miocene and Pliocene –Quaternary basin systems will not be discussed in this paper. The studied area includes the eastern part of Central Balkan-Forebalkan, the Eastern Balkan, Srednogorie and Sakar-Strandja zones (Fig. 4).

Main structural zones and tectonic events Despite the existence of abundant information concerning each of the different structural zones separately, there is no complete coherent description of the tectonics of the Balkan Mountains. During this work, we first collected structural data from the

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literature and then we made new observations in the field, especially near the principal tectonic boundaries. This section describes this (in brief) and discusses the main structures characterizing the major tectonic units and their age, from south to north: the Sakar-Strandja Zone, the Srednogorie Zone, the East Balkan Zone and the Central Balkan-Forebalkan Zone. Their relations with the reconstructed palaeostress fields that prevailed during Mesozoic and Cenozoic times will be discussed in the next sections.

Sakar-Strandja zone This zone classically belongs to the Internal Balkanides and is composed of the Sakar and Strandja units, showing rather different characteristics (Fig. 4). The boundary (Fig. 4) between the Sakar and Strandja units is still questionable, most likely due to normal faulting. The Sakar unit is better investigated and, in general, better explained. It demonstrates Early Alpine evolution with a metamorphic domes exhumation most likely during the Late Jurassic. The erosion and non-deposition on one hand and the overimposed Neogene sediments on the other hand, obscure the Late Alpine evolution. Two metamorphic domes can be recognized (Fig. 4): Izvor and Sakar domes (synmetamorphic and syntectonic granites), with an Upper Palaeozoic –Triassic and Lower-middle (?) Jurassic metamorphosed cover. The metamorphism is high (close to the granitic bodies contacts) to low-grade and lower upward the sequence. The northern boundary of the Sakar dome is a synmetamorphic normal shear zone, while the boundary between the two domes is a synmetamorphic dextral strike–slip zone (Gerdjikov 2005). The Maritsa dextral strike –slip zone (Fig. 4) is accepted to be the southern boundary of the Izvor dome, but its character is still less investigated and less well understood. The Yambol line (discussed below) (Fig. 4) is supposed to constitute the boundary between the Srednogorie and Sakar-Strandja zones. The structure of the Strandja unit is very similar to the Sakar unit, but with different tectonic and metamorphic events. The unit can be subdivided into the Bogdanovo-Slivovo and the Central Strandja antiforms (anticlines), separated by the dextral Bosna strike –slip zone (Fig. 4). Contrary to the Sakar unit the evidence for exhumation of the Strandja ‘domes’ are not so clear and not well investigated yet. The Strandja unit is composed of Lower Palaeozoic, low-grade metamorphites, crosscut by granites and overlain transgressively by Early Triassic to middle Jurassic continental to shallow-marine sequences, known as ‘Balkan’ type, very similar to those in the Moesian platform and Forebalkan.

Over this autochthonous (or para-autochthonous) fragment, were thrusted low-grade metamorphosed, relatively deep-water, Permian(?)–Triassic rocks (so-called ‘Strandja’ type). A Rb–Sr biotite whole rock age dates the regional metamorphism as Late Jurassic (155 Ma) (Okay et al. 2001). This thrusting, which is dated as Late –middle to Late Jurassic is related to Palaeotethyan closure. The metamorphic rocks are uncomformably overlain by shallow marine Cenomanian sansdtones (Okay et al. 2001). The northern ‘domes’ (Bogdanovo-Slivovo and Svety Iliya) (Fig. 4) show some features for secondary Late Cretaceous uplifting and ‘brittle exhumation’. Some of the structural field data indicate reactivating and refolding of the Early Alpine structures during Late Alpine time. On the other hand the offshore seismic data from the Strandja part of the Black Sea may be interpreted as gravity driven thrusts, related to the Western Black sea opening. In summary the Strandja unit is still less understood and interpreted, because of the complicated geology, missing information and, last but not least, the poor exposures. A major unresolved problem remains with regard to the geological relationship between the Strandja Massif and the Rhodope Massif: many similarities but also major differences exist between these two large metamorphic zones (Okay et al. 2001).

Srednogorie zone The main features of the Srednogorie zone are the Late Cretaceous volcanism and synchronous granitoid intrusions, classically interpreted as island-arc/back-arc system, separating the internal from external Balkanides. This zone is characterized by a complicated Early Alpine evolution, including many fragments with specific peculiarities. Based on Re-Os molybdenite ages, Zimmerman et al. (2008) interpreted the width of Late Cretaceous extensional basins and northward younging as the progression of coupled rollbackorogenic collapse. The Late Alpine evolution includes deformation of the ‘island-arc/back-arc system’ at the end of the Late Cretaceous, inversion occurring mainly on previous normal and strike – slip faults. However, the present-day structure of the zone is mainly due to the post-Lutetian compression and the following post-compressional relaxation, dominated by dextral regime. The existing subdivision of the Srednogorie Zone as Western, Central and Eastern units is questionable, mainly because of missing boundaries between the units or depending on criteria used for subdivision (Fig. 3). Generally it is subdivided on the base of different basement characteristics. For example the official boundary between the Central and East Srednogorie units, the Yambol line

BRITTLE TECTONICS OF EASTERN BALKANIDES

(Fig. 4), is only evidenced by geophysics: east of it, the continental crust is rather thinner than to the west. However, on the other hand, it is impossible to separate the depositional pattern in Stara Zagora strip (part of Central Srednogorie) from the one of Emine basin system (Eastern unit and East Balkan). Following the official subdivision, the Eastern unit is represented by a more than 4 km thick, volcano-terrigeneous complex of Late Cretaceous age, including tefro-turbidites and sub-marine volcanites, indicating a relatively shallow palaeoenvironment but fast subsidence during the rifting (Coniacian –Santonian). The Early syn-rift (Cenomanian-Turonian) sequences along the southeastern border of the basin are almost identical to those of northwestern parts of East Balkan and Stara Zagora strip. These sequences overlie transgressively the Early Alpine Strandja unit structure, which is crosscut by numerous Upper Cretaceous volcanites and intrusions. About 5–10 km north of the southern boundary of the Eastern unit (so-called North Strandja flexure) a north-verging fault zone can be traced (Figs 3 & 4), indicating likely reactivation and inversion of the main faults outlining the rift (but back and possibly passive roof thrusting was established in seismic profiles also). The sedimentation in the southern and central parts of the basin ceased during the Maastrichtian and was shifted only on the northern part and on the East Balkan zone, which corresponds to the first contraction events beginning in the Late Campanian to the south, continuing to the middle Eocene. The postcompressional extension resulted in the formation of Late Eocene (Priabonian) basins, including the coal-bearing basin near Burgas. All of these features and some correlations with the surrounding areas, allow this unit (including the Stara Zagora strip) to be interpreted as a dextral pull-apart basin. The Central unit shows a more complicated structure and is itself subdivided, in the studied area, in the Shipka-Sliven, Tundja-Topolnitsa and Stara Zagora units. The Shipka-Sliven unit (Fig. 4) is composed of Palaeozoic basement-metamorphic rocks and intruded granites, covered by Permian volcanites, Triassic dolostones and limestones and Upper Cretaceous sequence. This unit is thrusted over the western part of East Balkan along the Sliven thrust. Its western boundary is not clear but the Yantra sinistral strike– slip zone (Fig. 4), which bounded the East Balkan also, is likely the most logical choice (Vangelov 2006). Its southern boundary is the Stara Planina thrust (Fig. 4). The Stara Zagora strip is characterized by an Upper Cretaceous sequence more or less similar to those of the Shipka-Sliven unit, East Balkan and SE part of the East Srednogorie unit, underlaid by Triassic and Lower Jurassic deposits. The Tundja-Topolnica

83

unit is represented only by Palaeozoic metamorphic rocks and granitoids.

East Balkan zone The East Balkan zone is mainly composed of folded Cenomanian to Eocene turbiditic sequence (Fig. 5), deposited in back-arc setting along the southern margin of the Moesian Platform. The lateral and vertical facies distribution indicates that the basin underwent all stages of basin evolution, from synrift to molassic deposit ramps. Dextral transtension is assumed to be the mechanism of the basin formation. During the Cenomanian to Maastrichtian period the deposition was controlled by extensional tectonics, sea-level fluctuation and thermal subsidence (Late Campanian–Maastrichtian post-rift stage). The Paleocene –middle Eocene sequence is characterized by compressional tectonics and syntectonic sedimentation. The basin closure was likely due to sinistral transpression, involving multiple de´collement levels and lateral ramp development. The transpressional tectonics also could explain the formation and complicated structure of the large anticlines developed in front of the thrust belt by sub-thrust tectonics. The basement of the zone is composed of deformed Triassic and Lower –middle Jurassic sequences, which can represent the Early Alpine fragment (Vangelov 2005). The general trends of the folds in the East Balkan Zone are N908 to N1408. A post-compressional relaxation resulted in half-grabens forming on top of the ramp zones of the major thrusts. The East Balkan Zone is classically subdivided into two units: the Luda Kamchia unit and the Kotel unit (Fig. 5), which comprises the Early Alpine fragment. The Late Alpine evolution of the East Balkan Zone is closely related to the Eastern Srednogorie and Black Sea. Both zones represent the different stages of the evolution of a back-arc basin system and for that reason it is difficult to outline any boundary between them, except in their westernmost part, west of Sliven area, at the junction of three tectonic zones Central BalkanForebalkan, East Balkan and Srednogorie, complicated by the sinistral strike–slip zone between Yantra fault and Ptichevo line (Figs 3 & 4).

Central Balkan – Forebalkan zone Almost the entire zone is essentially constituted by folded Jurassic and Lower Cretaceous formations with dying-out fold intensity toward the Moesian Platform. Some local exposures of Triassic rocks in the cores of anticlines and of Upper Cretaceous – Paleocene in synclines exist. Some east –west trending fold axes are interpreted as Early Alpine, while the NW– SE orientation of the folds and thrusts correspond to the Late Alpine tectonics.

84

F. BERGERAT ET AL.

Fig. 5. Schematic cross-section (without scale) across the coastal part of Balkanides and its foreland. Location on Figure 4.

On the base of the fold style and orientation, the Central Balkan-Forebalkan zone is classically subdivided into two units: the Shipka-Teteven unit in its western part, and the Sevlievo-Preslav unit in the central-eastern parts (Ivanov 1988). In its western part, the Sevlievo-Preslav unit is composed of east– west trending wide anticlines (sometimes changing in direction toward the SW in their western parts) and separated by narrow synclines outlined by the Urgonian formation. A characteristic example of such an anticline is the Elena one, with a core formed by the Tithonian–Berriasian thick turbiditic complex and salt diapirs of Triassic evaporites (proved by well and seismic data), with a north-verging thrust northern boundary (faultpropagation fold). The NE boundary is marked by the Preslav-Skorpilovski fault zone (Figs 3 & 4), narrowing eastward (from 40 km to 2 –3 km); the southern boundary (which represents as well the East Balkan northern boundary) is east– west orientated and marked by the ‘Chudnite Skali’ dislocation (Figs 3 & 4) along which the East Balkan Triassic and Jurassic basement was thrusted over a narrow strip of Upper Cretaceous –Paleocene deposits, representing the foreland of the East Balkan thin-skinned thrust belt. The boundary between central and eastern parts of the SevlievoPreslav unit is the so called Ptichevo line (Figs 3 & 4), a large NE–SW trending sinistral strike–slip zone, crosscutting East Balkan and part of the Srednogorie zones also (see above). The Bjala area (Fig. 5) is supposed to be part of the Preslav unit, however it presents some specific features such as only Upper Cretaceous –Paleocene formations cropping out, as well as facies and style of deformation different from the other parts of the unit. A shallow marine molasse foredeep basin (the so called Lower Kamchia depression) developed, north of Bjala, on the southeastern part of the Moesia since the Late Eocene to Neogene (Figs 3 & 5).

Methods of brittle tectonic analysis Along with the structural observations in the field, we carried out microtectonic measurements in

order to reconstitute palaeostress states relative to the successive tectonic events. The analysis of brittle tectonics in terms of palaeostress reconstruction can be considered now as a routine and efficient tool for investigating the time and space of geodynamical evolution. Therefore the successive palaeostress regimes can be reconstructed quantitatively by the inversion of meso-fault slip data sets collected in the field. Most of the developed inversion methods for the determination of palaeostress tensors are based on stress –shear relationships (e.g. Angelier 1979; Carey 1979; Etchecopar et al. 1981). Because all of these methods have been widely used (e.g. Bergerat 1987) and discussed (e.g. Lisle et al. 1998) in previous papers, they are not further developed herein. In the studied area, 82 sites were investigated: 7 sites in the Moesian Platform; 34 sites in the Central Balkan/Forebalkan Zone; 31 sites in the Srednogorie/East Balkan Zone; and 10 sites in the Sakar/ Strandja Zone. These sites are mainly road cuts, quarries and coastal cliffs of the Black Sea. Microtectonic measurements were carried out in various formations of Triassic (17 sites); Jurassic (7 sites); Lower Cretaceous (17 sites); Upper Cretaceous – Paleocene (40 sites); and Eocene (1 site). Most of the visited sites allowed palaeostress analysis (Tables 1–4). A total of 1940 fault-slip data were measured in order to determine the palaeostress pattern. Because the tectonics in the studied area is highly polyphased, it was necessary to split, for each site, the total population of fault-slip data, into mechanically homogeneous sub-groups as illustrated for example by the sites Bjala Coast (Fig. 6) and Cape Emine (Fig. 7). This separation was made based on mechanical consistencies between brittle structures. Each sub-group of faultslip data is thus consistent with a single stress regime. For each sub-group, the principal axes of the reduced palaeostress tensor have been computed with the INVDIR method (Angelier 1990). Each palaeostress state is characterized by its type (reverse-, normal- or strike –slip type), the attitude of the three stress axes s1, s2 and s3 with s1

Table 1. Palaeostress tensors computed from fault-slip data in Sakar-Strandja Zone Site

Locality

Long. E8

Age1

N2

Type3

s41

s2

s3

F5

a6

Ry 7

T T T T T T T T T J J J T T J J T T T T

8 8 4 14 14 19 17 7 14 12 8 5 19 22 22 24 7 8 7 4

E SS E SS E SS C SS SS E* SS* SS* C E* SS SS E SS SS SS

316 – 71 338 – 2 243 – 49 359 – 21 291 – 64 246 – 32 43 – 5 46 – 73 113 – 37 306 – 50 188 – 6 293 – 7 238 – 5 332 – 74 261 – 12 200 – 41 277 – 66 190 – 12 77 – 37 227 – 0

138– 19 77 – 80 98 – 36 129– 60 121– 26 69 – 58 135– 17 192– 14 270– 51 113– 39 279– 11 180– 72 148– 0 77 – 4 26 – 69 19 – 49 77 – 23 76 – 62 286– 49 317– 44

48 – 1 248 – 10 354 – 18 260 – 21 29 – 4 337 – 1 229 – 72 285 – 9 14 – 11 208 – 7 71 – 77 25 – 16 54 – 85 168 – 16 167 – 17 110 – 1 170 – 7 286 – 24 178 – 14 137 – 46

0.2 0.6 0.8 0.2 0.7 0.6 0.6 0.3 0.1 0.7 0.7 0.2 0.7 0.6 0.7 0.8 0.5 0.6 0.5 0.4

6 18 10 11 12 13 11 8 10 15 11 3 15 14 18 15 8 10 6 7

18 45 28 29 36 29 33 36 45 43 38 27 33 35 43 44 23 33 26 27

170

BOSNA

42.17194

27.43361

171

VELEKA RIVER

42.08667

27.42694

172 –173

STOILOVO

42.01417

27.505

174

BYALA VODA

42.18111

27.44889

175

KONDOLOVO

42.09417

27.68722

176

BLIZNAK

42.15639

27.29972

177

BLUE ROAD

42.17222

27.30833

178

INDJE

42.22111

27.43583

1

Age of the rocks: T, Triassic; J, Jurassic. Number of fault slip data used for calculation. Type of stress state: SS, strike –slip faulting; N, normal faulting; R, reverse faulting. (Asterisk indicates that the calculation has been made after back-tilting.) 4 Maximum, medium and minimum stress axes s1, s2 and s3 (trend and plunge, in degrees). 5 Ratio of stress difference. 6 Average angle a between observed and computed striae (in degrees), acceptable results with a , 308. 7 Ratio upsilon (Ry ): misfit criterion of the INVDIR method (Angelier 1990) in the range 0% (best) – 200% (worst), with acceptable results for RUP , 75%. 8 Chronology relative to the tilted bedding plane: A, ante-tilt; S, syn-tilt; P, post-tilt; Pe, corresponds to possible extrado normal faulting. 2

C8

A A A P A P P

BRITTLE TECTONICS OF EASTERN BALKANIDES

Lat. N8

3

85

86

Table 2. Palaeostress tensors computed from fault-slip data in East-Balkan – Srednogorie Zone Site

Locality

Long. E8

Age1

N2

Type3

s41

s2

s3

F5

a6

Ry 7

T T CS– P CS– P CS CS E E CS CS CS CS– P T T T T CS CS CI P P CI CI P CS CS CS CS CS CI CS CS CS CS CS CS CS CS

13 13 11 12 6 35 13 9 15 11 9 7 4 4 7 11 16 5 12 13 9 5 9 8 7 6 11 11 19 14 14 8 6 9 10 17 4 5

SS SS SS E SS* C E SS SS* E SS SS SS E SS SS SS SS E SS E SS SS SS SS E E SS SS þ C E C SS

163 – 14 230 – 2 31 – 22 3 – 73 213 – 10 189 – 21 260 – 79 220 – 9 15 – 1 105 – 76 78 – 16 213 – 10 337 – 11 159 – 79 20 – 15 283 – 12 72 – 15 108 – 24 293 – 61 64 – 8 66 – 75 286 – 11 185 – 1 323 – 9 209 – 3 48 – 60 24 – 79 345 – 17 27 – 7 12 – 74 18 – 13 43 – 10 140 – 5 295 – 76 162 – 6 217 – 6 134 – 6 208 – 4

62 – 37 107 – 85 176 – 63 252 – 6 319 – 58 87 – 28 33 – 8 348 – 75 107 – 59 329 – 10 325 – 54 75 – 76 197 – 75 302 – 9 142 – 62 82 – 77 302 – 67 299 – 65 94 – 28 321 – 57 315 – 5 29 – 49 282 – 81 64 – 51 300 – 25 273 – 22 283 – 2 165 – 73 296 – 1 103 – 0 108 – 1 267 – 76 232 – 17 113 – 14 276 – 75 344 – 80 244 – 71 299 – 13

271 – 49 321 – 4 295 – 14 161 – 16 117 – 30 310 – 54 125 – 8 128 – 12 284 – 30 237 – 10 178 – 31 304 – 9 69 – 9 33 – 7 283 – 23 192 – 5 167 – 16 200 – 4 188 – 8 159 – 31 224 – 14 187 – 39 95 – 9 225 – 38 112 – 65 175 – 19 193 – 11 75 – 0 197 – 83 193 – 16 203 – 77 135 – 9 35 – 72 203 – 0 70 – 13 126 – 8 42 – 18 103 – 76

0.2 0.1 0.3 0.2 0.6 0.7 0.4 0.0 0.7 0.5 0.8 0.1 0.3 0.6 0.3 0.6 0.4 0.5 0.4 0.0 0.4 0.3 0.6 0.3 0.3 0.8 0.6 0.2 0.4 0.3 0.6 0.8 0.0 0.5 0.5 0.8 0.3 0.4

25 16 14 8 11 6 16 51 14 7 14 12 5 30 20 18 17 6 12 41 14 9 18 15 14 8 14 13 31 17 15 17 28 7 34 32 39 39

52 37 50 25 35 21 32 74 38 25 38 45 36 52 46 33 34 26 20 64 34 29 37 45 51 24 32 48 58 32 32 50 62 19 30 33 34 38

87

SLIVEN-1

42.71667

26.26222

89–90 90 94

SLIVEN-3–4 SLIVEN-4 KOTEL-4

42.74861 42.74861 42.86278

26.43806 26.43806 26.48694

95

AVRAMOV

42.74166

26.63806

96

AJTOS

42.73972

27.24167

97 98

VRESOVO BILKA

42.84611 42.92833

27.18917 27.21667

108

VESELINOVO

42.88083

27.00361

111 111

STARA REKA-2 STARA REKA-2bis

42.85722 42.85722

26.14778 26.14778

112

ELENA-2

42.89056

25.86667

113 114

DRENTA-1 DRENTA-2

42.82222 42.8225

25.79722 25.7725

115

DRENTA-3

42.8225

25.78111

116 117

DRENTA-4 BADEVCI

42.82583 42.88861

25.79528 25.73806

119 160

PCELINOVO CAPE EMINE

42.81528 42.7

25.6525 27.89722

E SS* SS* SS* C

C8

P A P A P P P S P P

P P P P P P

A A A P

F. BERGERAT ET AL.

Lat. N8

TANKOVO

42.70055

27.64417

167

TCHERNOMORETZ

42.44972

27.64111

168

KROUCHEVETZ-1

42.29445

27.495

169

KROUCHEVETZ-2

42.24861

27.47194

179 180

BINKOS CHIVATCHEVO

42.66667 42.68861

26.09639 26.03472

181

BYALA PASS-1

42.72

26.10333

182

BYALA PASS-2

42.69444

26.07972

183

TVARDITZA PASS-1

42.775

25.90806

184

TVARDITZA PASS-2

42.75639

25.89778

185

BORILOVO

42.47583

25.54666

101 þ 161

SOLNIK

42.92083

27.59444

1

CS CS CS CS CS CS CS CS CS T T T T T CS CS T T CS CS T T T CS CS CS CS CS CS

10 7 6 9 5 6 7 28 10 11 31 4 20 25 17 6 19 16 7 20 6 8 6 22 6 8 24 23 28

E E SS C SS* E C SS E E* SS SS* E E E E E SS C E SS E C SS C C C SS E

246 – 83 308 – 80 199 – 5 313 – 7 193 – 21 317 – 55 81 – 27 224 – 3 281 – 88 148 – 76 248 – 1 162 – 6 172 – 60 196 – 78 3 – 75 280 – 79 195 – 79 19 – 21 16 – 3 22 – 78 206 – 9 62 – 75 237 – 11 356 – 12 42 – 0 280 – 21 23 – 7 18 – 0 214 – 83

149 – 3 139 – 10 92 – 72 45 – 10 100 – 7 208 – 13 172 – 2 360 – 86 32 – 1 5 – 11 359 – 86 267 – 67 336 – 29 101 – 1 132 – 10 110 – 11 94 – 2 153 – 61 106 – 6 111 – 5 61 – 79 305 – 7 146 – 4 142 – 76 132 – 22 13-Oct 292 – 4 287 – 78 117 – 1

58 – 6 48 – 2 290 – 17 190 – 78 354 – 67 109 – 32 265 – 63 134 – 3 122 – 2 273 – 8 158 – 3 70 – 23 70 – 7 11 – 12 224 – 12 19 – 2 4 – 11 282 – 20 262 – 83 20 – 11 297 – 6 214 – 13 37 – 78 265 – 8 312 – 68 126 – 67 170 – 82 108 – 12 27 – 7

Age of the rocks: T, Triassic; Cl, Lower Cretaceous; CS, Upper Cretaceous; P, Paleocene; E, Eocene. Number of fault slip data used for calculation. Type of stress state: SS, strike –slip faulting; N, normal faulting; R, reverse faulting. (Asterisk indicates that the calculation has been made after back-tilting.) 4 Maximum, medium and minimum stress axes s1, s2 and s3 (trend and plunge, in degrees). 5 Ratio of stress difference. 6 Average angle a between observed and computed striae (in degrees), acceptable results with a , 308. 7 Ratio upsilon (Ry ): misfit criterion of the INVDIR method (Angelier 1990) in the range 0% (best) – 200% (worst), with acceptable results for RUP , 75%. 8 Chronology relative to the tilted bedding plane: A, ante-tilt; S, syn-tilt; P, post-tilt; Pe, corresponds to possible extrado normal faulting.

0.3 0.4 0.3 0.7 0.5 0.5 0.0 0.4 0.7 0.8 0.4 0.1 0.5 0.5 0.5 1.0 0.8 0.7 0.7 0.5 1.0 0.3 0.6 0.3 0.8 0.6 0.5 0.7 0.4

7 16 20 19 13 14 11 15 15 15 23 15 12 10 11 6 15 12 14 14 17 14 10 15 8 11 24 21 46

18 32 49 44 37 46 35 42 30 36 46 57 26 29 28 20 32 35 25 31 36 33 32 37 29 36 27 36 67

P P P P A

A P A A P A P P P P P P P P P P A

BRITTLE TECTONICS OF EASTERN BALKANIDES

166

2 3

87

88

Table 3. Palaeostress tensors computed from fault-slip data in Western Central Balkan-Forebalkan Zone Site

Locality

Long. E8

Age1

N2

Type3

s41

s2

s3

F5

a6

Ry 7

C8

CI CI CI CI CI CI CI CI CI CI CI CI CI CI CS CS CS CS CS T CS CS CS CS CS CS – P CS – P CS – P CS – P CS CS CI CI

16 4 7 11 11 13 11 10 10 9 16 5 12 18 18 8 12 13 4 10 7 6 7 11 5 35 4 7 6 6 13 6 7

C E SS* C C SS SS þ C C SS E SS SS SS* SS þ C SS SS* E SS C SS SS SS C SS E C E SS SS SS SS SS C

194– 14 229– 83 200– 32 342– 6 15 – 15 5 – 17 188– 14 173– 14 241– 22 9 – 74 18 – 11 285– 16 193– 25 196– 1 348– 22 47 – 12 25 – 74 164– 5 281– 16 286– 5 217– 13 155– 5 300– 1 346– 0 355– 81 209– 0 73 – 70 68 – 36 184– 9 36 – 13 5 – 15 43 – 11 348– 12

102 – 7 101 – 4 74 – 35 250 – 14 106 – 5 199 – 73 287 – 31 263 – 2 7 – 56 267 – 4 247 – 74 131 – 72 16 – 65 101 – 78 163 – 68 299 – 55 181 – 15 19 – 83 189 – 8 146 – 83 73 – 74 250 – 42 210 – 11 77 – 87 247 – 3 119 – 1 324 – 7 288 – 47 286 – 52 253 – 74 142 – 70 252 – 77 257 – 4

346– 74 10– 5 310– 29 94– 75 213– 74 96– 4 77– 55 360– 76 140– 25 176– 16 110– 12 17– 7 283– 1 286– 12 257– 2 145– 32 273– 6 254– 4 72– 72 16– 4 310– 9 59– 47 34– 79 256– 3 157– 9 308– 89 231– 18 174– 21 88– 36 129– 9 271– 13 134– 6 147– 77

0.4 0.4 0.2 0.3 0.4 0.4 0.2 0.4 0.3 0.4 0.5 0.4 0.6 0.3 0.5 0.1 0.3 0.5 0.6 0.5 0.5 0.1 0.6 0.5 0.3 0.7 0.3 0.5 0.4 0.4 0.4 0.1 0.5

11 7 7 16 11 11 13 9 6 10 37 9 16 18 14 4 14 13 11 20 16 13 16 18 24 12 8 10 18 9 5 9 5

26 17 37 42 37 35 39 31 18 29 64 31 42 50 36 30 37 34 36 45 49 38 34 34 50 28 26 33 41 25 10 22 13

S P? A P

1

PRESLAV-1

43.08333

26.78333

80 81

PREOBRAZENSKI SAMOVODENE

43.10695 43.12889

25.60833 25.60722

82 83

DRJANOVO-1 DRJANOVO-2

42.96194 42.94639

25.43861 25.43583

84 85

DRJANOVO-3 ELENA-1

43.03083 42.94667

25.53695 25.90055

86 91

VELIKO TARNOVO KOTEL-1

43.06667 42.91611

25.63806 26.4525

92

KOTEL-2

42.91111

26.4475

93 99

KOTEL-3 G. TRAIJKOV-1

42.9 42.96611

26.45167 27.29528

99

G. TRAIJKOV-2

42.96611

27.29528

102

BJALA

42.87194

27.89417

106

DALGOPOL

43.02861

27.28111

109

TARGOVISTE-1

43.20195

26.52167

P A P P A P P P P P A? P Pe P P

F. BERGERAT ET AL.

Lat. N8

109

TARGOVISTE-2

43.20195

26.52167

VONESTA VODA GABROVO-1

42.86833 42.77306

25.64667 25.41

122 123

GABROVO-2 STOKITE-1

42.78083 42.88972

25.40361 25.07195

124 125 154

STOKITE-2 TROJAN CHIPKA-1

42.88055 42.85194 42.73583

25.07195 24.76889 25.31528

155

CHIPKA-2

42.74778

25.31528

156

CHIPKA-3

42.78

25.31083

157 –158

STRAJA

42.78

25.31056

159

PRESLAV-2

43.13805

26.80806

162

BJALA-1

42.88528

27.90083

4 9 14 11 14 7 8 6 6 8 6 7 9 13 15 5 5 5 7 11 23 7 6 7 16 16 10 34 18 11 7 22 4 8

SS E C SS E C* E SS SS SS C E E SS þ C SS SS* SS E C C* C E SS E C SS* E* E SS C SS E SS C

331– 26 224– 82 19 – 3 192– 21 35 – 73 37 – 7 125– 81 131– 21 303– 18 201– 1 16 – 2 23 – 74 331– 72 9–7 339– 8 171– 11 252– 3 59 – 82 129– 9 328– 4 346– 14 94 – 56 38 – 36 335– 77 352– 19 180– 19 347– 60 141– 26 242– 55 51 – 6 316– 11 99 – 79 33 – 19 199– 1

224 – 30 62 – 8 288 – 10 31 – 68 274 – 9 306 – 5 326 – 9 347 – 65 100 – 70 292 – 44 285 – 11 265 – 8 227 – 4 273 – 41 220 – 74 26 – 77 155 – 65 188 – 5 220 – 7 60 – 24 76 – 1 214 – 18 163 – 39 77 – 3 83 – 2 300 – 55 233 – 14 291 – 4 65 – 35 141 – 1 212 – 49 301 – 10 289 – 34 289 – 12

94– 48 332– 2 127– 79 284– 6 182– 14 179– 81 235– 3 226– 13 211– 7 110– 46 117– 79 173– 14 136– 17 107– 48 71– 14 262– 7 344– 25 278– 6 348– 79 228– 66 172– 76 314– 27 282– 31 168– 12 179– 71 79– 28 136– 27 22– 2 334– 1 245– 84 55– 39 211– 4 146– 49 106– 78

0.4 0.3 0.6 0.5 0.2 0.8 0.3 0.5 0.6 0.4 0.7 0.5 0.3 0.3 0.2 0.4 0.3 0.1 0.5 0.5 0.6 0.4 0 0.6 0.7 0.3 0.8 0.5 0.6 0.7 0.3 0.3 0.4 0.6

6 16 11 14 9 17 16 16 13 16 14 8 8 17 14 18 19 15 8 14 9 15 10 12 15 24 25 19 15 15 46 16 6 6

25 42 27 42 31 45 40 48 38 32 33 21 25 42 35 35 46 40 21 42 15 35 29 29 30 48 53 37 38 42 70 31 32 16

P S Pe P P P P P P P S P S A P P P Pe P A A? P P

BRITTLE TECTONICS OF EASTERN BALKANIDES

118 121

CI CI CI CI J CS CS CS CI J J J CI CI T T T T T T T T T T J J J J – CI J – CI J – CI J – CI CS CS CS

(Continued)

89

90

Table 3. Continued Site

Locality

Long. E8

Age1

N2

Type3

P P P P P CS – P CS – P CS – P CS – P CS – P P P P CS – P

9 12 11 13 5 8 13 6 7 7 21 6 21 8

E SS C SS C SS C SS SS E E SS E E

163

BJALA-2

42.88528

27.90083

164

BJALA-3

42.87806

27.89861

165

BJALA-4

42.88111

27.89833

162 –165

BJALA-1 –4

42.88111

27.89833

1

s41

s2

s3

F5

a6

Ry 7

68 – 75 238– 1 204– 7 2–4 285– 27 58 – 6 17 – 1 171– 7 290– 10 52 – 71 64 – 72 113– 7 64 – 72 6 – 68

285 – 12 332 – 74 114 – 1 258 – 74 34 – 32 301 – 77 286 – 19 65 – 65 185 – 55 278 – 14 279 – 15 346 – 79 279 – 15 210 – 21

194– 9 148– 16 19– 83 93– 16 163– 46 150– 11 110– 71 264– 24 27– 33 185– 13 187– 10 204– 9 187– 10 117– 8

0.4 0.5 0.4 0.3 0.3 0.7 0.3 0.4 0.2 0.2 0.5 0.3 0.5 0.4

10 12 14 15 9 12 12 17 18 11 14 9 14 15

22 35 31 32 45 32 27 35 50 23 32 35 32 46

Age of the rocks: T, Triassic; J, Jurassic; Cl, Lower Cretaceous; CS, Upper Cretaceous; P, Paleocene. Number of fault slip data used for calculation. 3 Type of stress state: SS, strike –slip faulting; N, normal faulting; R, reverse faulting. (Asterisk indicates that the calculation has been made after back-tilting.) 4 Maximum, medium and minimum stress axes s1, s2 and s3 (trend and plunge, in degrees). 5 Ratio of stress difference. 6 Average angle a between observed and computed striae (in degrees), acceptable results with a , 308. 7 Ratio upsilon (Ry ): misfit criterion of the INVDIR method (Angelier 1990) in the range 0% (best) – 200% (worst), with acceptable results for RUP , 75%. 8 Chronology relative to the tilted bedding plane: A, ante-tilt; S, syn-tilt; P, post-tilt; Pe, corresponds to possible extrado normal faulting. 2

C8

F. BERGERAT ET AL.

Lat. N8

Site

Locality

Lat. N8

Long. E8

Age1

N2

Type3

s41

s2

s3

F5

a6

Ry 7

7 4 18 5

E E SS SS Joints Joints SS E C

276– 82 355– 65 359– 3 94– 48

17– 2 192– 24 131– 86 297– 39

0.4 0.3 0.5 0.7

4 20 15 5

8 56 38 7

25– 46 18– 67 67– 7

171– 39 210– 22 158– 11

107 –8 99 –7 269 –3 197 –12 100 175 276 –18 119 –4 305 –77

0.5 0.3 0.6

12 18 8

38 43 23

4 43

SUMEN DEVNJA

43.26667 43.02861

26.9 27.59528

CS CI CI

44 100 103

BELOSLAV PROVADJA GROZDJOVO–1

43.18806 43.15833 43.00139

27.68806 27.43361 27.46667

CS CS CS– P CS– P CS– P

1

9 18 5

Age of the rocks: Cl, Lower Cretaceous; CS, Upper Cretaceous; P, Paleocene. Number of fault slip data used for calculation. Type of stress state: SS, strike –slip faulting; N, normal faulting; R, reverse faulting. (Asterisk indicates that the calculation has been made after back-tilting.) 4 Maximum, medium and minimum stress axes s1, s2 and s3 (trend and plunge, in degrees). 5 Ratio of stress difference. 6 Average angle a between observed and computed striae (in degrees), acceptable results with a , 308. 7 Ratio upsilon (Ry ): misfit criterion of the INVDIR method (Angelier 1990) in the range 0% (best) – 200% (worst), with acceptable results for RUP , 75%. 8 Chronology relative to the tilted bedding plane: A, ante-tilt; S, syn-tilt; P, post-tilt; Pe, corresponds to possible extrado normal faulting. 2 3

C8

A A

BRITTLE TECTONICS OF EASTERN BALKANIDES

Table 4. Palaeostress tensors computed from fault-slip data in Southeastern Moesian Platform Note that only a few new sites have been included herein, detailed study of the Moesian platform can be found in Bergerat & Pironkov (1994) and Bergerat et al. (1998)

91

92

F. BERGERAT ET AL.

Fig. 6. Example of a polyphased site at Bjala, Forebalkan Zone. At least five stress states can be characterized along the cliffs in the marls and micrite limestone Campanian to Paleocene in age [162–165], corresponding to normal (column 1), reverse (column 2), and strike–slip (column 3 –5) regimes. The bedding is slightly tilted (dip 308SW on average, see diagram in the right bottom corner). The stress tensor computations have been carried out separately in the Campanian, Maastrichtian and Paleocene formations in order to verify their presence in all of them and check their homogeneity. Stereoplots are Schmidt projections (lower hemisphere) of the fault populations and computed stress tensors. N is geographic north and M is magnetic north. Fault planes are shown as thin lines; slikenside lineations (striae) are small dots with centrifugal-single, centripetal-single or double thin arrows (mainly normal, reverse or strike–slip, respectively). Three-, four- and five-branch stars are computed axes s3, s2 and s1, respectively. Large solid arrows are directions of horizontal extension and compression. Bedding planes are shown as thin dashed lines. See Table 3 for the detailed stress tensors.

 s2  s3 and the ratio F ¼ (s2 2 s3) / (s1 2 s3) with 0  F 1. These stress states are tentatively attributed to successive tectonic events using three main chronological arguments. (i) Relative chronologies between fault-slip data, such as superimposition of striations on a fault plane indicating fault reactivation with successive slips, or crosscutting relationships between brittle structures, which allow the determination of the succession of stress states. (ii) Collecting faults in strata of various ages which help to give bounds for the age of the calculated stress states, in most cases only indicating the oldest possible age. Finding synsedimentary evidences of faulting or short time span unconformities above faults is ideal, but is rarely the case. (iii) Data collected in tilted sedimentary rocks which permit to date the stress states relatively to the time of tilting.

Stress field analyses carried out in many places during the last 30 years demonstrated that within less than 158 the reconstructed stress axes follow the Andersonian model. That is, one of the principal stress axes is found nearly vertical and the other two axes are nearly horizontal. This effect results from gravity and the Earth’s free surface, which constrains one axis to be vertical, being dependent of rock weight and fluid pressure. Exceptions occur but are rare and correspond to particular situations, such as the vicinity of diapirs or large volcanic centres or in accretionnary prisms where large-scale thrusting occurs (Angelier 1994). For this reason, the backtilt criterion and approach has proved successful and has been used in structural geology for a long time. Reconstructed stress states that occurred after folding in tilted strata obey this criterion. On the contrary, oblique-stress orientations (with two principal stresses in the bedding plane) suggest pre-folding stress states. In this case, the stress state follows Anderson’s model after restoration of the original horizontal

BRITTLE TECTONICS OF EASTERN BALKANIDES

93

Fig. 7. Example of a polyphased site at Cape Emine, East Balkan Zone. Three main stress states can be characterized along the cliffs in the Upper Campanian-Maastrichtian flysch [160], corresponding to four faulting modes: a NE– SW compression marked both by reverse (a) and strike– slip (b) faults, a NNW– SSE compression/SSE– NNW extension illustrated by strike–slip faults (c) and a NE– SW extension (d) marked by normal faults. The bedding is slightly tilted (dip 258NNW on average), the reverse and strike–slip faulting (a–c) are previous to this tilting and herein represented in their back-tilted (initial) positions. The normal faulting (d) is post-tilt. The photograph (e) shows conjugate normal faults (az. 1258, dip 708NE on the right, and az. 1358, dip 758W on the left). These faults belong to the diagram (d). See Figure 6 for the caption of the stereoplots and Table 2 for the detailed stress tensors.

attitude of the bedding. For such pre-folding stress states, the original stress orientations are obtained by computation on the back-tilted faults. Especially significant was the consistency of the principal stress trajectories obtained when backtilted sites (from fold flanks) are mixed with untilted ones (from tabular areas) for a single tectonic event.

Faulting analysis and palaeostress states reconstructions An analysis of brittle deformation and the reconstruction of the successive palaeostress fields has been carried out for the Sakar-Strandja, East Balkan-Srednogorie, and Forebalkan-Central Balkan zones, as well as for a small part of the

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Moesian Platform close to the front of the fold-and-thrust belt. The stress inversion performed on data collected in the 82 studied sites led to the determination of 175 stress states. Among them, 88 are of strike– slip type (919 fault-slip data, 47% of the total faults population), 36 of reverse type (453 fault-slip data, representing 23% of the total faults population) and 51 of normal type (568 faultslip data, representing 30% of the total faults population). Strike –slip and reverse faulting characterizing identical directions of s1 are often associated in a same site and constitute the dominant tectonic regimes (corresponding respectively to 50 and 21% of the computed stress states, respectively) but normal faulting is not negligible (29% of the computed stress states). The results of this stress analysis are presented in Table 1 for the Sakar-Strandja zone, Table 2 for the East Balkan-Srednogorie zone, Table 3 for the Forebalkan-Central Balkan zone, and Table 4 for the small studied part of the Moesian Platform. When certain the ante- or posttilt occurrence of a given stress state is indicated in the tables. Note, however, that (i) in many cases (e.g. sub-horizontal bedding) such relative chronology cannot be established; and (ii) if the age of the main folding events is well known in Bulgaria, it is not always easy to know with certainty to which one a given fold is attributed to. A rose diagram of all of the measured directions of bedding planes is shown in Figure 8a, illustrating the two main phases of folding: the NW– SE trending peak (N 1358) corresponds to the Late Cretaceous compression (i.e. the so called Laramian phase) and the east –west (N80 –958) peak mostly to the compression which occurred at the end of Middle Eocene (i.e. the so called Illyrian phase). Some other small peaks exist, especially one in the direction N1108 for which the age is difficult to establish. It is interesting to note that the NW–SE measured directions are more abundant than the east– west ones in the Srednogorie Zone and that the east –west ones are more abundant than the NW–SE ones in the East Balkan, Central Balkan and Forebalkan (Fig. 8c–e). Note also that the rose diagram of the Sakar-Strandja Zone shows a major peak at N208 which does not exist in the other zones (Fig. 8b); it may correspond to the Early Alpine compression known in this area. Figure 9 shows the computed local trends of the maximal principal stress (s1) for the reverse and strike– slip stress states and of the minimal principal stress (s3) for the normal stress states. The directions of the s1 axes computed from both the strike–slip and reverse faults cover a wide range of directions. The main direction of s1 for reverse faulting is N15 –358and for strike–slip faulting it ranges between N08 and N458, however, some smaller peaks exist (for both strike– slip and

Fig. 8. Rose diagrams of bedding planes in the Eastern Balkanides. Total population of bedding planes measured in the whole studied area (a) 1: directions, 2: dips, and separately for Sakar-Strandja (b), Sredna Gora (c), East Balkan (d), Central-Balkan and Forebalkan (e). The numbers below the diagrams indicate the number of measured bedding planes.

reverse faulting) at N60 –708, N105–1208 and N160–1708. The directions of s3 axes computed from normal faulting are also scattered with two main groups at N170–2358 and N95–1358. Although local stress deviations may occur, the distribution of the azimuths of stress axes rather reflect the occurrence of different tectonic events. Nevertheless as it will be demonstrated below, similar horizontal stress directions do not mean that the stress states have the same age. When available, chronological arguments are then used to determine the age of each stress state. Hereafter we describe the stress states in terms of orientation. For each stress-state orientation, we discuss the possible age of the tectonic event(s) and the relationships between normal-, strike–slipand reverse-stress regimes showing similar

BRITTLE TECTONICS OF EASTERN BALKANIDES

horizontal stress directions, with respect to the possibility of stress permutations between the s1/ s2 or s2/s3 axes (e.g. Angelier & Bergerat 1983). In order to avoid increasing the number of values we group together the less abundant strike –slip and reverse-stress states with identical horizontal direction of s1 (east –west trending s1 and NNW– SSE to NW– SE trending s1), on the contrary for the north–south to NNE –SSW trending s1 and the NE –SW to ENE –WSW trending s1, the strike – slip and reverse-stress states are presented separately for a better clarity of the figures. For the normal-slip stress states, we distinguish four main directions of s3: north –south, NE– SW, east –west and NW– SE. Some sites are presented in detail (diagrams and/or photographs) in order to illustrate each of the main stress states.

Normal faulting and extensional stress states (Fig. 10) Among the 51 extensional stress states calculated, 10 of them (20%) characterize a normal regime with a roughly east –west s3 trend (Fig. 10a). This stress regime is present in the entire studied area except the Sakar-Strandja Zone, affecting Triassic to Late Cretaceous –Paleocene rocks. The majority of the s3 directions range between N708 and N1208 with an average trend of N1008. In two sites, located in Triassic rocks [155, 169], this stress state predates the tilting, meaning that part of these computed stress states may correspond to the initial rifting of the basin during the Jurassic (Fig. 11). The other sites, located in Cretaceous rocks, do not supply any chronological information, except one at Kotel [92] where the normal faults crosscut and displace reverse faulting due to a WNW– ESE compression (Fig. 12). Note that some 15 sites, located in Berriasian to Aptian rocks in the Moesian Platform, also present records of an extensional stress state with a

95

N100–1108 trending s3 (Bergerat et al. 1998). In that case, because of the systematic presence of both strike–slip and normal faulting corresponding to the same direction of s3 in most sites, as well as the systematic absence of relative chronological criteria between these extensional and strike–slip stress states, this extensional stress state has been supposed to be due to a permutation of the stresses s1 and s2. Some of the stress states computed in the chain may likely be related to the same process (see below). A poorly recorded extensional stress trending NW –SE (N125 –1358) is present in six sites (12% of the extensional stresses) (Fig. 10b). This stress state affects rocks from Triassic to Eocene and no relative chronological criteria are available except at Avramov [95], in the Sandstones of middle Eocene age, where it is posterior to the tilting and accompanied by some strike –slip faults characterizing the same direction of s3 (Fig. 13a). It may represent local deviations of the east– west one. Another extensional stress regime is characterized in 14 sites (27% of the extensional stress states) with a NE –SW direction of s3 (N30 –558; Fig. 10c). This stress regime has only been recognized in the Sakar-Strandja Zone where it affects Triassic and Jurassic rocks and in the East BalkanSrednogorie Zone in Late Cretaceous –Paleocene rocks. At Byala Voda [174] in the Strandja Zone, some normal faults related to this stress state and affecting Hettangian-Sinemurian rocks occurred clearly before the probably Early Alpine folding; it is possible to relate them to the Jurassic rifting. An identical state of stress is characterized also by normal faults until the Paleocene that is the case, for example, in the Campanian marls at Bjala [162] (Fig. 14) where the normal faults are slightly tilted with the bedding. These faults most likely correspond to the major extensional tectonics, Upper Cretaceous in age, well known in the Srednogorie Zone. In some other cases the chronological criteria

Fig. 9. Rose diagrams of principal stress axes in the Eastern Balkanides. Trends of s1 associated with reverse stress states (a), trends of s1 associated with strike–slip stress states (b), trends of s3 associated with normal stress states (c).

96 F. BERGERAT ET AL. Fig. 10. Extensional stress states in the Eastern Balkanides. East – west trending extensional stress states (a) NW– SE trending extensional stress states (b) NE– SW trending extensional stress states (c) and north– south trending extensional stress states (d). 1: computed directions of extension (s3) for normal faulting; 2: directions of extension inferred from tectonic joints. See Tables 1– 4 for the detailed stress tensors. B, Burgas; By, Byala; E, Elena; G, Gabrovo; P, Preslav; SZ, Stara Zagora; T, Trojan; V, Varna; VT, Veliko Tarnovo.

BRITTLE TECTONICS OF EASTERN BALKANIDES

97

Fig. 11. Example of polyphased tectonics at Krouchevetz, Srednogorie Zone. Two main tectonic regimes can be characterized along the road, south of Krouchevetz in Upper Triassic turbidites [169]: a strike– slip regime with an ENE–WSW direction of compression (a) and a normal regime with an east–west direction of extension (b). The bedding is tilted to the north (e). The normal faulting is previous to this tilting and herein represented in its back-tilted (initial) position. The strike–slip faults of the diagram (a) belong to both ante- and post-tilt identical stress states and the stress tensor computation has been made on the total strike–slip population. The right-lateral strike–slip faults trend roughly north– south and the pitches of the striation range between 08 and 408; part of the striae is clearly tilted as shown on the photograph (d). Some of the left-lateral strike–slip faults are tilted ones but the most abundant left-lateral motions occurred post-tilting on the tilted bedding planes as shown on the photograph (c). A few right-lateral and left-lateral strike–slip faults are represented in their initial position in (f). The thin black arrows in the photographs indicate the motion of the missing blocks. See Figure 6 for the caption of the stereoplots and Table 1 for the detailed stress tensors.

indicate (i) that the normal faults occurred after folding, especially when the bedding is steeply dipping or even reverse; and (ii) that the directions of the normal faults is the same as those of the fold axes (e.g. site [121], Fig. 13b). This suggests that they can correspond to an extrado extensional process or a post-compressional collapse resulting in half-graben forming at the tops of the ramps. Thus, the NE–SW extensional trend may probably characterize three different tectonic events of Jurassic, Upper Cretaceous and Late Eocene age. The most represented extensional regime is a north–south extension characterized in the whole studied area and corresponding to 41% of the extensional stress states (21 sites). The majority of the s3 directions range between N1708 and N1958 with an average trend of N1828. This stress state affects rocks from Triassic to Paleocene and its main occurrence takes place in the East Balkan and East Srednogorie zones (Fig. 10d). In several places the normal fault systems are obviously tilted (e.g. in Figs 6 & 15b), thus it is likely to link this stress state to the Upper Cretaceous major rifting event,

accompanied by the occurrence of volcanic rocks in Srednogorie zone. Note, however, that, as for the NE –SW extension, some of the north –south extensional stress states are posterior to the main folding (e.g. site [171], Fig. 13c) and may be related to reactivation as normal faults of pre-existing reverse faults, resulting, at a regional scale, in the development of half-grabens on ramp zones. The distribution of the places where the NE–SW and north– south extensional stress states have been recorded deserve consideration (Fig. 10c, d): in the East Balkan Zone it is noticeable that the NE–SW direction is mostly represented in the eastern part of the zone while the north –south one characterize essentially the western part, and moreover among the 35 of sites which have recorded these stress states, only two have recorded both the NE–SW and east –west extensions (sites [111] and [162 –165]), and even in these two cases, the difference in trends correspond to formations of different age in the site (e.g. Campanian v. Maastrichtian and Paleocene at Bjala; Fig. 6). Therefore we can presume that the NE –SW extensional stress state represents

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F. BERGERAT ET AL.

Fig. 12. Example of polyphased tectonics near Kotel. Three faulting regimes are present along a road cut in the sub-horizontal Maastrichtian limestone close to the Kotel Quarry [92] (Central Balkan-Forebalkan Zone): a reverse faulting with a poorly determined WNW–ESE compression (a); a normal faulting showing an east–west extension (b) and a strike–slip regime characterizing a NNW–SSE compression (c). The photograph (d) shows a reverse fault plane (1: az. 308, dip 408W) shifted by a normal fault (2: az. 308, dip 528W). These faults belong to the diagrams (a) and (b), respectively. The thin black arrows in the photograph indicate the motion of the missing blocks. SW of Kotel [94] (East Balkan Zone) the Campanian-Maastrichtian limestone has been strongly folded during the north– south compression. The diagram (f) shows all the measured bedding planes and some fold axes (as white dots). In (e) there is an example of one these small folds. See Figure 6 for the caption of the stereoplots and Table 3 for the detailed stress tensors.

a variation of the north–south one in space (and maybe in time) and that both of them characterize the Upper Cretaceous major extensional event.

Reverse faulting and compressional stress states (Fig. 16) Among the 35 compressional stress states calculated, 8 of them (22%) characterize a compressional

regime with a NE–SW s1 trend (Fig. 16a). The directions of s1 range between N378 and N678 with an average trend of N528. In spite of its slight representation, this stress regime has been recorded in the entire studied area in Triassic to Late Cretaceous –Paleocene rocks. When a chronology relative to folding is available, it indicates that the faulting is generally syn- or post-tilt such as at Kondolovo [175] (Fig. 17), Tvarditza [183 –184]

BRITTLE TECTONICS OF EASTERN BALKANIDES

Fig. 13. Fault populations and stress states of the NW– SE, NE–SW and north– south extensions. The NE– SW trending normal faults are often associated with strike–slip faults, both characterizing a NW –SE direction of extension as shown in the example of Avramov [95] in Middle Eocene sandstones (a1–a2). The examples of Gabrovo [121] in Maastrichtian limestone (b) and Veleka River [171] in Triassic metamorphosed limestone (c); illustrate respectively the NE–SW and north– south extensions, both being characterized by post-tilt normal faults. See Figure 6 for the caption of the stereoplots, and tables 1, 2 and 3 for the detailed stress tensors.

Fig. 14. The NNE– SSW extension in the Campanian marls of the Bjala cliffs, Forebalkan. Numerous east– west to NW– SE trending normal faults characterize this stress state (a). The photograph (b) shows sub-conjugate normal faults with decimetric throws (az. 908, dip 408N on the left, and az. 1108, dip 558S on the right). The bedding planes here trend about N120– 1308 and dip 308SW. See Figure 6 for the caption of the stereoplots and Table 3 for the detailed stress tensors.

99

Fig. 15. Example of polyphased tectonics near the Tvarditza Pass in the Shipka-Sliven unit of the Srednogorie Zone. Two main stress states can be characterized along the road, south of the Tvarditza Pass in the Campanian-Maastrichtian limestone [183] (white diagrams) and in the Middle Triassic dolostone [184] (grey diagrams): a NNE–SSW extension (b, d) and a NNE– SSW to NE-SW compression (c, e). The Cenomanian to Maastrichtian formations are slightly dipping (10– 208NE) and directly overlie the Triassic rocks which are in reverse position and steeply dipping (70– 808S). In the case of the Upper Cretaceous, the normal and reverse faults are obviously neoformed (b –c), contrary to the Triassic where most of the motions, normal and reverse as well, occurred on the bedding planes previously straightened (d –e). Both the Upper Cretaceous and the Triassic have recorded the NNE– SSW trending extension. Because of the small amount of fault slip data, the compression is less determined in trend and it is difficult to know for certain that the two diagrams correspond to the same state of stress. The photograph (a) shows sub-conjugate normal faults (az. 1458, dip 608S on the right and az. 1208, dip 558N on the left) in the Upper Cretaceous limestone. These faults belong to the diagram (b). See Figure 6 for the caption of the stereoplots and Table 2 for the detailed stress tensors.

(Fig. 15) and Drenta [115] (Fig. 18b). Note that this direction of compression NE–SW is also characterized by abundant strike –slip states (see next subsection and Fig. 16b), very often in the same sites (Fig. 7) as for the reverse ones. In several sites, because the reverse faults are generally by far less

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Fig. 16. Compression and strike– slip stress states in the Eastern Balkanides. NE–SW trending compressional stress states (a); NE–SW trending strike–slip stress states (b); east–west to ESE–WNW trending compressional and strike– slip stress states (c); north– south to NNE– SSW trending compressional stress states (d); north–south to

BRITTLE TECTONICS OF EASTERN BALKANIDES

101

Fig. 16. (Continued) NNE– SSW trending strike–slip stress states (e); NW– SE to NNW–SSE trending compressional and strike–slip stress states (f). 1: computed directions of compression (s1) for reverse faulting; 2: computed directions of compression (s1) and extension (s3) for strike– slip faulting. See Tables 1 –4 for the detailed stress tensors. B, Burgas; By, Byala; E, Elena; G, Gabrovo; P, Preslav; SZ, Stara Zagora; T, Trojan; V, Varna; VT, Veliko Tarnovo.

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Fig. 17. The ENE– WSW compression in the Middle-Upper Triassic limestone at Kondolovo, Strandja Zone. The compression is well marked both by folding (a–b) and syn- and post-folding reverse faulting (c). See Figure 6 for the caption of the stereoplots and Table 1 for the detailed stress tensors. Small black dots in (b) represent fold axes.

Fig. 18. Fault populations and stress states of the NE– SW compression. In the entire studied area the NE– SW compression is marked both by strike– slip and reverse faulting indicating the same direction of s1, as shown herein by the examples of Georgi Traijkov Lake site [99] in Maastrichtian limestone, Drenta site [115] in Campanian-Maastrichtian limestone and Cape Emine site [160] in Upper Campanian-Maastrichtian flyshs (a, b and c respectively). The diagrams c1 and c2 represent the present and initial (after back-tilting) attitudes of the faults, respectively. See Figure 6 for the caption of the stereoplots and Tables 2 and 3 for the detailed stress tensors.

abundant than the strike–slip ones, the inversion has been performed on both reverse and strike–slip faults; in that case the resulting tensor is indicated as strike– slip (SS) in the tables and is represented in Figure 16a. Because the folding axes in these sites are in good agreement with those credited to the so called Laramian compression, and because of the close relationship between the reverse faulting and this folding, we propose that this stress state took place at the end of Cretaceous during the first major inversion phase. Another compressional stress trending roughly east –west (N80– 1208) is present in only four sites of Upper Cretaceous age (11% of the extensional stresses) with very few measurements (Fig. 16c). No chronological criteria allows us to propose an age for this fairly recorded stress state, because some 15 sites also show strike–slip faulting with an identical direction of compression, the reverse and strike –slip stress states will be discussed together in the next sub-section. The most represented compressional regime is a north– south to NNE –SSW compression characterized in the northern part of the East Balkan Zone and in the Central Balkan-Forebalkan (Fig. 16d). It corresponds to 40% of the compressional stress states (14 sites). The s1 directions range between N48 and N298 with an average trend of N178. Only in the Preslav site [1] some reverse faults are synfolding (Fig. 19a), but in six other sites they occurred after the tilting (Figs 20c & 15c). In some 30 sites the same direction of compression is indicated by strike –slip faulting (cf. next subsection and Fig. 16e) and in a third of them reverse faulting coexists with strike–slip faulting (Fig. 19). In the latter, in order to check the reality of this association as well as for a better clarity of the diagrams when the number of faults was important, the two fault populations have been processed separately. Only (i) the occurrence of just one of the two regimes (strike–slip or reverse); (ii) the small amount of faults; or (iii) the absence of conjugate faults (e.g. the presence of only left-lateral or rightlateral faults in the strike –slip regime) justified the computation of all of them together (as indicated by SSþC in the tables). Because, in all sites, the small folding axes are in good agreement with the major ones due to the so called Illyrian compression, and because of the close relationship between the reverse faulting and this folding, we propose that this stress regime is related to this major Illyrian inversion phase well marked in front of the thrust belt, at the end of Middle Eocene. Another compressional stress trending roughly NNW– SSE (average trend of s1: N1548) is present in nine sites (26% of the extensional stresses). The same direction of s1 exists for some 20

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Fig. 19. Fault populations and stress states of the north– south to NNE– SSW compression. In the entire studied area the north– south to NNE– SSW compression is marked both by strike–slip and reverse faulting indicating the same direction of s1, as shown herein by the examples of Preslav site [1] in Hauterivian-Valanginian marly limestone, Solnik site [101 þ 161] in Upper Maastrichtian turbidites, Samovodone [81] in Urgonian (Aptian) limestone, and Targoviste site [109-2] in Berriasian-Hauterivian limestone, with separate stress inversions (a1– a2, b1– b2, c1 –c2 and d1– d2, respectively) and of Veliko-Tarnovo site [86] in Urgonian (Aptian) limestone, and Borilovo site [185] in Upper cretaceous limestone, with at once stress inversion (e and f respectively). The stereoplots g1–g2 illustrate the site of Badevci [117] in Maastrichtian limestone, which is the unique site where we can observe both the NNE– SSW and NNW–SSE directions of s1. See Figure 6 for the caption of the stereoplots and Tables 2 and 3 for the detailed stress tensors.

strike– slip stress states (cf. next sub-section and Fig. 16f). Almost all of the sites (for reverse and strike– slip faults) are located close to major thrust: (i) between the Sakar-Strandja zone and the Srednogorie zone; (ii) between East Balkan and Central Balkan-Forebalkan; and (iii) at the boundary with the Moesian Platform. The faults (reverse and strike –slip as well) can be pre-, syn- or post-tilt (Figs 20a, b & 21). Considering the entire population of compressional stresses showing a s1 around north– south (23 sites) only one allows us to distinguish both the directions NNE –SSW and NNW–SSE (Fig. 19g1– g2). In this site the NNW–SSE s1 corresponds to very few fault slip data. Lacking other arguments, the simplest hypothesis is that these stress states should represent local variations of a general north–south strike –slipcompressional regime.

Strike – slip faulting and strike – slip stress states (Fig. 16) Among the 89 strike– slip stress states calculated, 24 of them (27%), present in the entire studied area in Triassic to Late Cretaceous –Paleocene rocks, characterize a strike –slip regime with a NE –SW s1 trend (Fig. 16b). The directions of s1 range between N318 and N698 with an average trend of N448. This stress state has already been discussed above, with the NE –SW reverse regime (cf. previous subsection and Fig. 16a), the only difference with this latter, in addition to its more significant record, is that it can occur before or after folding as well. Figure 18a, c illustrate respectively post- and pre-tilt strike-fault systems characterizing identical directions of compression. Because of their close links with both reverse faulting and

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Fig. 20. The NNW–SSE compression in the Tithonian limestone at Straja, in the Sevlievo-Preslav unit of the Forebalkan. The strike– slip faults are previous to folding, they are presented here in their present-day (a) and initial (b) positions. The reverse faults (c) are post-folding. The photograph shows two of these tilted conjugate strike– slip faults: on the left, a right-lateral one (az. 1408) and on the right, a left-lateral one (az. 108), the thin black arrows indicate the motion of the missing blocks. The bedding plane is trending N858 and dipping 508N (underlined by a white dashed line). See Figure 6 for the caption of the stereoplots and Table 3 for the detailed stress tensors.

Laramian folding, this stress-state may be related to the first major Alpine compressional phase (Latest Cretaceous –Paleocene inversion). Shanov (2005) also recognized this ‘post-Early Cretaceous’ stress field in Aptian limestones of the SE Moesian Platform. An east –west trending strike–slip regime (Fig. 16c) is recorded in 17 sites (19% of the strike– slip stresses) in Triassic to Paleocene rocks; with an average trend of N978 (s1 ranges between N688 and N1238) (Fig. 9). Note that an east –west reverse regime, poorly represented (cf. previous subsection), also shows this direction of compression. As for this latter, the number of faults in each site for the strike –slip regime is generally small and, moreover, no clear chronological criteria exist between the related fault system and the other ones or with the major structures. This is the case, for example, in the Hettangian-Sinemurian rocks at Byala Voda [174] where the east –west strike– slip stress occurred clearly before tilting (Fig. 22a). It is therefore possible that a part of these stress

states (and of the folding as well) should be related to the early stages of Alpine deformation in the Sakar-Strandja area (Jurassic?). The other sites where these stress states are present are located both in the Central-Balkan-Forebalkan and in Eastern Balkan-Srednogorie zones, and, in all cases, very close to the boundary between these two domains, thus it is likely to link them to the normal faulting of the Upper Cretaceous rifting by a permutation of the stresses s1 and s2, the direction of s3 being the same in both cases. Such a permutation has already been suggested in the Moesian Platform (Bergerat & Pironkov 1994). Note, however, that for four of the sites, it postdates different trending folds in rocks from Triassic to Upper cretaceous (Fig. 22b). In conclusion, what the hypotheses concerning this east– west compression may be remains extremely questionable because of the few and inconsistent chronological data. The most represented strike –slip regime shows a north– south to NNE –SSW direction of compression. It is characterized in 29 sites (33% of the entire strike –slip stress states) of the Central BalkanForebalkan and of the East Balkan-Srednogorie zones (Fig. 16e), in Triassic to Late Cretaceous – Paleocene rocks (Fig. 19). The majority of the s1 directions range between N08 and N208 with an average trend of N128. This important (in number of sites as well as in number of faults) strike–slip regime is often associated with the north–south reverse one (cf. previous subsection and Fig. 16d). The main difference with this latter, in addition to its more significant record, is that it can occur before or after folding as well. Some examples of these relations are shown in Figure 19 (see also Fig. 20). Both the strike –slip and reverse faulting may be linked to the Illyrian major inversion phase dated of latest middle Eocene. Another strike–slip stress states is characterized in 19 sites (21% of the strike–slip states) with a NNW– SSE s1 average trend N1598 (Fig. 16f). These stress states have already been discussed with the reverse faulting characterizing the same direction of s1 (cf. previous subsection).

Discussion Early Alpine tectonics The evidence for the Triassic and Jurassic rifting became available at the beginning of the 1990s with the first high-quality seismic reflexion profiles in this area (e.g. Tari et al. 1997a, b; Georgiev et al. 2001) which revealed rift geometry preserved in, and beneath, the thrust sheets. At the southern margin of the Moesian Platform, Georgiev & Botoucharov (2007) also recognized middle Jurassic contourites in cores along syndepositional

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Fig. 21. Fault populations and stress states of the NNW– SSE compression. In the entire studied area the NNW– SSE compression is marked both by strike– slip and reverse faulting indicating the same direction of s1, as shown herein by the example of Drjanovo Monastery site [83] in Aptian limestone, Chipka Pass-2 site [155] in Spathian-Anisian limestone, Chipka-Pass-3 site [156] in Middle-Upper Triassic limestone, and Targoviste site [109-1] in BerriasianHauterivian limestone (a, b, c and d1 –d2, respectively). See Figure 6 for the caption of the stereoplots and Table 3 for the detailed stress tensors.

growth faults. Nevertheless the Triassic to Lower Cretaceous basin systems evolution cannot be explained easily in a whole because surface data display complicated stratigraphy and facies

Fig. 22. Fault populations and stress states of the east– west compression. The two examples illustrate an east– west compression marked by conjugate strike– slip faults occurring ante- (a) or post-tilt (b). In the site of Byala Voda [174] the faults affect Hettangian-Sinemurian sandstones. The diagram (a) shows them in their initial (back tilted) position (reverse bedding: 15–808E). In the site of Kotel [93], in Toarcian sandstones, the faults occurred after the tilting (reverse bedding: 170–608W) (b). Note that in both cases (i) the axis of the folds is trending about north– south; and (ii) the bedding planes are reverse. See Figure 6 for the caption of the stereoplots and Tables 1 and 2 for the detailed stress tensors.

distribution in time and space, and also to missing interaction between them. Because of the important fold-and-thrust deformations which occurred during the Late Alpine tectonics and the middle–Late Jurassic metamorphism in the internal Balkanides, the traces of the Triassic –Jurassic tectonic events have been erased for most of the sites and the microtectonic measurements in the Triassic and Lower Jurassic rocks concern essentially the Late Alpine tectonics (Upper Cretaceous to Eocene in age). Nevertheless we observed some normal faults indicating a NE –SW direction of extension in Triassic and Jurassic rocks of the Strandja unit before a folding attributed to an Early Alpine phase (Fig. 10c). This extension may characterize the so called EoCimmerian phase. In some places, especially in the Strandja unit (Fig. 8b), the fold axes in the Triassic –Jurassic rocks, trend roughly NNE– SSW (which is not a ‘classical’ trend for the Late Alpine features), moreover some traces of an ESE–WNW compression marked by strike –slip as well as reverse faults has been found locally. All of these features should be related to the early stages of Alpine deformation in the Sakar-Strandja area (the so called neoCimmerian or Austrian phases). At the same time fragments of the Early Alpine thrust/fold belt are well recognized in the field

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(e.g. Minkovska et al. 2002) but need to be better characterized, especially for restoration of pre – Late Alpine position and geometry. Note that our propositions concerning both the extension and the compression directions have to be considered cautiously because of the lack of reliable chronological evidence.

Late Alpine tectonics Since the latest Early Cretaceous, the geological history of the Balkanides is more constrained including a period of rift basin evolution, followed by a period of major inversion from the Paleocene and during the Palaeogene (Fig. 23). Syn- and post-rift stages. Based on detailed studies of Tithonian-Valanginian and Hauterivian-Aptian formations in terms of sequence stratigraphy, Nikolov et al. (1998) and Peybernes et al. (1998) characterized the platform-basin-platform pattern of the Central and Eastern Forebalkan during the Early Cretaceous. The extensional regime which prevailed in Srednogorie and East Balkan zones is known to have begun during the Late AlbianCenomanian by a dextral transtension forming segmented systems (an en echelon pull-apart and strike–slip basins) and enlarged during the Turonian and Early Campanian which constitutes the main rifting period. An important volcanic activity occurred in the Srednogorie Zone. In Moesia an asymmetric ‘platform’ basin developed (Fig. 23a, b). Our brittle tectonics analysis not only demonstrates the occurrence of an important extension in the Srednogorie and East Balkan zones (Fig. 10c, d) but also brings arguments for the transtensional process by the presence of a few associated strike– slip faulting indicating identical direction of s3. The Late Campanian-Maastrichtian period is classically considered as a post-rift stage with thermal subsidence in the back-arc basin, a thermal bulge in between the basins (Fig. 23c, d) and the beginning of uplifting in Strandja unit and southern part of Srednogorie Zone. A flooding event covered the region of Forebalkan and Moesia, and then a sea level fall caused the progradation of shallow water sansdstones and carbonates. The sedimentation of Upper-Campanian and Maastrichtian is supposed to be controlled by sea-level fluctuations. Some sites where the sequence is continuous from the Campanian to the Paleocene (e.g. Bjala, Fig. 6) reveal that the same normal faulting (and related palaeostress state) occurred all along this period. Despite the end of the actual rift in the Late Campanian (marked by the end of the volcanic activity), the normal faulting continued to be active up to the Paleocene as demonstrated in many sites (Tables 2 and 3).

This demonstrates that the extensional ambiance persisted at least until the beginning of the Cenozoic. This suggests that the sea level fluctuations maybe not the unique control of the sedimentation but that normal faulting may also have a responsibility in the sedimentary process. Inversion stages. Based on the en echelon distribution of folds and thrusts, and on the undulation of the thrust belt, Doglioni et al. (1996) proposed that the Palaeogene convergence (trending N30 – 458) occurred as a left lateral transpression in the east –west East Balkans, changing to almost pure compression in the NW–SE buried Balkans under the Western Black Sea. We also emphasize the important role of strike–slip faulting all along this period of inversion (Fig. 16), furthermore demonstrating that the stress-field slightly counterclockwise rotated between the Early Paleocene and the latest Lutetian. In the Early Paleocene an important inversion tectonics event occurred producing a separation of the northern part of the Balkanides with a large foredeep basin and the southern part with piggy-back basins onto East Balkan and Srednogorie zones (Fig. 23e). This inversion is well recorded and well characterized by brittle tectonic data (Fig. 16a, b) with both strike –slip and reverse faulting regimes indicating a NE– SW direction of s1 (Fig. 18). In all cases, the faults accompany the folding, the strike–slip faults occurring before, during, and after this event, while the reverse faults occurred during or more generally after the fold process. The available chronological criteria systematically indicate that the faulting related to this compressional event (i) post-dates the normal faults of the Upper Cretaceous rifting and (ii) pre-dates another compressional stresss-state characterized by a north–south direction of s1. It is interesting to note that in some cases, a normal faulting occurred after this compressional episode, especially in the strongly folded areas, corroborating the existence of relaxation processes. At the latest Paleocene and Eocene the foredeep basin expanded quickly over the whole Moesian Platform and coastal part of the Forebalkan. A second inversion phase occurred at the latest Lutetian and most of the important compressional features which are responsible of the present appearance of the Balkanides are related to it. North of the fold-and-thrust belt arose a prograding terrigenious wedge (Fig. 23f). As for the Early Paleocene phase, the strike –slip faulting occurred before, during or after the folding while the reverse faulting is rather syn- or post-folding. The general trend of s1 in the whole area is north–south to NNE–SSW but the direction NNW –SSE is also present

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Fig. 23. Simplified model of evolution for the Central Balkan-Forebalkan zone from Upper Cretaceous to Palaeogene. (a– b) syn-rift stages; (c– d) post-rift stages; (e–f) compressional stages.

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(Fig. 16d –f). The same wide range of trends of s1 has been demonstrated in the Moesian Platform where this phase is the most marked (Bergerat et al. 1998; Bergerat & Pironkov 1994). The chronological criteria do not support, in the Balkanides or in Moesia, the occurrence of two distinct tectonic episodes. Only some local variations in the trend can be considered. A NNE –SSW compression has been characterized in the Pliocene in the northern part of the Moesian Platform (Bergerat et al. 1998); it is impossible to assert that none of the NNE–SSW trend is of the same age. Note also that a strike–slip faulting with a N408 trending s1 exists in the only site located in Eocene rocks ([95], Fig. 13a). It is likely due to a local deviation of the NNE– SSW compressional stress axis. This demonstrates that only a large amount of data in a large number of sites, allow us to justifiably conclude in terms of chronology. In some cases, as for the Laramian phase, relaxation normal faulting occurred after this compressional episode. Post-orogenic stage. After the Illyrian compressional phase a molasse basin developed in Lower Kamchia depression and the coastal part of northeastern Bulgaria during Late Eocene, Oligocene and Miocene, where no microtectonic evidence was found for characterizing the tectonic events. In the East Balkans and Srednogorie Zones a series of post-compressional grabens, infilled by Late Eocene deposits, were formed on top of the main thrusts ramp zones.

Conclusion Despite numerous investigations, and particularly high-quality seismic and well data, the geodynamics of Bulgaria is still not fully understood. Doglioni et al. (1996) already underlined the fact that crosssections are poorly retrodeformable because of transpressional and transtensional components of the Palaeogene and Neogene tectonics, respectively, and the existence of different thickness and facies transitions of the sedimentary cover. Our reconstruction of the different faulting regimes and palaeostress states does not permit directly such retrodeformation, but it allows accurate determination of stress-field evolution in the Balkanides, which cannot be obtained from geometric analyses of seismic profiles only or even from observation of major structural features cropping out. Therefore they provide better understanding of the successive tectonic events, especially in terms of direction of extension and/or compression. Since the Triassic and during all the Mesozoic and Cenozoic times, the tectonic regimes and associated palaeostress states changed many times in the Balkanides, being characterized by a

succession of extensional (basin formation) strike – slip and compressional (inversion) episodes. The stress fields which prevailed during the Early Alpine tectonic period were not easy to characterize, partly because of the metamorphism of middle –Late Jurassic (in the Strandja unit) and mainly because its traces were obliterated by the subsequent Late Alpine tectonics. The sparse stress-state reconstructions indicate the occurrence of a NE –SW direction of extension and an ENE –WSW direction of compression which may correspond, without certainty, to the Eo- and Neo-Cimmerian tectonic phases. In contrast, the succession of stress states related to the Late Alpine tectonic period is better characterized (Fig. 23). It began with the major extension which affected the East Balkan and Srednogorie Zones during the Turonian-Early Campanian. The direction of extension characterizing this rift stage varies from NNW–SSE to NNE–SSW. The rifting classically ends during the Campanian however the brittle tectonic analysis demonstrated that the normal faulting continued until the Paleocene. The first period of inversion followed immediately this extensional stage with a main direction of compression NE–SW. Not only has the direction been characterized by brittle data analysis but also by close association between folding, reverse and strike –slip faulting. The second period of shortening, classically attributed to the latest Lutetian, produced the main tectonic features of the Balkanides as well as the most abundant brittle tectonic data. They characterize a NNW –SSE to NNE – SSW direction of compression, also marked by close relationships between folding and reverse/ strike –slip faulting. Following each of these two compressional phases, normal faulting often occurred, showing that the extension related to collapse did not produce only half-graben at the tops of the ramps, but is a general expression of the post-compressional relaxation. The Pliocene – Quaternary dextral transtension inherited and expanded the graben systems. This study also demonstrates that it is difficult to draw boundaries between the classical existing units of the Central and East Srednogorie Zone and the East Balkan Zone. Multiplying subdivisions is not justified at least since the Upper Cretaceous, all these units belonging to a large island-arc/back-arc basin system with an identical evolution during the Late Alpine history. The Balkanides had a long and complex history, the interpretation of the successive tectonic events is even now debated and their geodynamic significance is still widely discussed. How the evolution of the Balkanides is driven by geodynamic processes at the scale of the whole DinaridicCarpathian-Balkan system (e.g. Kazmin &

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Tikhonova 2006; Schmid et al. 2008; Zimmerman et al. 2008) is out of the scope of this work, but the regional tectonic history that we propose for the Eastern Balkanides is in good agreement with some large-scale geodynamical reconstructions involving the closure of the Vardar Ocean and the formation and closure of marginal sea basins (Karamata 2006), the opening of the Black Sea basins above the northward subduction zone of the Tethyan Ocean (Robinson et al. 1996), the Aegean Cenozoic extensional regime (Burchfiel et al. 2000; Brun & Sokoutis 2007) as well as the MioPliocene subsidence period in the SE Carpathians, driven by slab-pull and intraplate folding due to Quaternary inversion (Matenco et al. 2007). This work was supported by the Middle East Basins Evolution Program (MEBE) and the RILA program. We gladly thank Zˇ. Ivanov for interesting and fruitful discussions on the geology of Bulgaria and V. Vangelova for her help during the field works. We are also grateful to G. Georgiev for exchanges of views on the seismic profiles. The paper has benefited from constructive reviews by Dr. L. Matenco and Dr. S. Vandycke.

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Georgiev, G. & Dabovski, C. 1997. Alpine structure and petroleum geology of Bulgaria. Geology and Mineral Resources, 8– 9, 3– 7. Georgiev, G., Dabovski, C. & Stanisheva-Vassileva, G. 2001. East Srednogorie-Balkan Rift Zone. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & Crasquin-Soleau, S. (eds) Stratigraphy and Evolution of Peri-Tethyan Platforms. Me´moires du Muse´um National Histoire Naturelle, Peri-Tethys Memoir 6, 186, 259–293. Gerdjikov, I. 2005. Alpine metamorphism and granitoid magmatism in the Strandja zone: new data from the Sakar unit, SE Bulgaria. Turkish Journal of Earth Sciences, 14, 167–183. Hsu, K. J., Nachev, I. & Vuchev, V. 1977. The evolution of Bulgaria in light of plate-tectonics. Tectonophysics, 40, 245–256. Ivanov, Zˇ. 1988. Aperc¸u ge´ne´ral sur l’e´volution ge´ologique et structurale du massif des Rhodopes dans le cadre des Balkanides. Bulletin de la Socie´te´ Ge´ologique de France, 4, 227– 240. Ivanov, Zˇ., Moskovski, S. & Koltcheva, K. 1974. Edification ge´ologique d’une partie du noyau de l’anticlinal de Tvarditsa. Annales de l’Universite´ de Sofia, 1, 254– 277. Jaranoff, D. 1960. La tectonique de la Bulgarie. Technica, Sofia (in Bulgarian). Karamata, S. 2006. The geological development of the Balkan Peninsula related to the approach, collision and compression of Gondwanan and Eurasian units. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 155– 178. Kazmin, V. G. & Tikhonova, N. F. 2006. Evolution of Early Mesozoic back-arc basins in the Black SeaCaucasus segment of a Tethyan active margin. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 179–200. Lisle, R. J., Orife, T. O., Arlegui, L., Liasa, C. & Srivastava, D. C. 1998. Favoured states of palaeostress in the Earth’s crust: evidence from fault-slip data. Journal of Structural Geology, 28, 1051–1066. Kockel, C. 1927. Zur Stratigraphie und Tektonik Bulgariens. Geol. Congress, Sofia, 56– 96, Rundschau, 18, 381– 389. Matenco, L., Bertotti, G., Leever, K., Cloething, S., Schmid, S. M., Tara˘poanca˘, M. & Dinu, C. 2007. Large-scale deformation in a locked collisional boundary: interplay between subsidence and uplift, intraplate stress, and inherited lithospheric structure in the late stage of the SE Carpathians evolution. Tectonics, 26, TC4011, doi:10.1029/2006TC001951. Minkovska, V., Peybernes, B. & Nikolov, T. 2002. Paleogeography and geodynamic evolution of the Balkanides and Moesian, microplate (Bulgaria) during the earliest Cretaceous. Cretaceous Research, 23, 37–48. Nikolov, T., Peybernes, B., Ciszak, R. & Ivanov, M. 1998. Enregistrement se´dimentaire de la tectonique extensive et de l’eustatisme dans le Jurassique terminal et le Cre´tace´ basal du Pre´balkan central et oriental (Bulgarie). Comptes-rendus de l‘Acade´mie

des Sciences, Paris, Sciences de la terre et des Plane`tes, 326, 43– 49. Okay, A. I., Satur, M., Tu¨ysu¨z, O., Akyu¨z, S. & Chen, F. 2001. The tectonics of the Strandja Massif: late-Variscan and mid-Mesozoic deformation and metamorphism in the northern Aegean. International Journal of Earth Sciences, 90, 217– 233. Peybernes, B., Ciszak, R., Nikolov, T., Ivanov, M. & Stoykova, K. 1998. Ge´ome´trie des formations et des se´quences de de´poˆt e´ocre´tace´es (intervalle Hauterivien-Aptien) de Bulgarie nord-orientale. Comptes-Rendus de l’Acade´mie des Sciences, Paris, II, 327, 555–562. Peybernes, B., Tchoumachenko, P. et al. 1989. Donne´es nouvelles sur les flyschs jurassiques de la zone de Luda Kamchia (Balkanides orientales, Bulgarie): conse´quences pale´ogeographiques. ComptesRendus de l’Acade´mie des Sciences, Paris, II, 309, 115–124. Ricou, L. E., Burg, J. P., Godfriaux, I. & Ivanov, Zˇ. 1998. Rhodope and Vardar: the metamorphic and the olistostromic paired belts related to the Cretaceous subduction under Europe. Geodinamica Acta, 11, 285–309. Robinson, A. G., Rudat, J. H., Banks, C. J. & Wiles, R. L. F. 1996. Petroleum geology of the Black Sea. Marine and Petroleum Geology, 13, 195– 223. Schmid, S. M., Bernoulli, D. et al. 2008. The Alpine-Carpathian-Dinaridic orogenic system: correlation and evolution of tectonic units. Swiss Journal of Geosciences, 101, 139– 183. Sengo¨r, A., Yilmaz, Y. & Ketin, I. 1980. Remnants of a Pre-Late Jurassic ocean in North Turkey: fragments of Permian-Triassic Paleo-Tethys. Bulletin of the Geological Society of America, 19, 599–699. Sengo¨r, A., Yilmaz, Y. & Sungurlu, O. 1984. Tectonics of the Mediterranean Cimmerides: nature and evolution of the western termination of Paleo-Tethys. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 77–112. Sengo¨r, A., Demir, A., Cin, A., Usyaomer, T. & Hsu, K. 1988. Origin and assembly of the Tethyside orogenic collage at the exprense of Gondwana Land. In: Audley-Charles, M. G. & Hallam, A. (eds) Gondwana and Tethys. Geological Society, Oxford University Press, 119–181. Shanov, S. 1990. Tectonic stress fields in northeast Bulgaria. Geologica Balcanica, 20, 37–47. Shanov, S. 2005. Post-Cretaceous to recent stress fields in the SE Moesian Platform (Bulgaria). Tectonophysics, 410, 217 –233. Tari, G., Dicea, O., Faulkerson, J., Georgiev, G., Popov, S., Stefanescu, M. & Weir, G. 1997a. Cimmerian and Alpine stratigraphy and structural evolution of the Moesian Platform (Romania/ Bulgaria). In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 63–90. Tari, G., Georgiev, G., Hardy, S., Poblet, J. & Stefanescu, M. 1997b. Late Triassic Cimmerian structures beneath the Moesian Platform (Romania/ Bulgaria). The Leading Edge, 16, 1153–1157.

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Dating of the Black Sea Basin: new nannoplankton ages from its inverted margin in the Central Pontides (Turkey) ¨ LLER2, N. KAYMAKCI3 & E. SANGU4 J.-C. HIPPOLYTE1*, C. MU 1

LGCA, University of Savoy, CNRS UMR 5025, 73376 Le Bourget-du-Lac, France Present address: CEREGE, Aix-Marseille III University, CNRS UMR 6635, Europoˆle Me´diterrane´en de L’Arbois, BP 80, 13545 Aix en Provence, France 2

IFP, 6 bis rue Haute 92500 Rueil-Malmaison, France

3

Middle East Technical University, Department of Geological Engineering, 06531-Ankara Turkey; Utrecht University, Forth Hoofddik Paleomag Lab. Budapestlaan 17, 3584 CD Utrecht, the Netherlands 4

Kocaeli University, Department of Geological Engineering, 41100 Kocaeli, Turkey *Corresponding author (e-mail: [email protected]) Abstract: The Eocene uplift and inversion of a part of the Black Sea margin in the Central Pontides, allows us to study the stratigraphic sequence of the Western Black Sea Basin (WBS). The revision of this sequence, with 164 nannoplankton ages, indicates that subsidence and rifting started in the Upper Barremian and accelerated during the Aptian. The rifting of the western Black Sea Basin lasted about 40 Ma (from late Barremian to Coniacian). In the inner, inverted, Black Sea margin, the syn-rift sequence ends up with shallow marine sands. The uppermost Albian to Turonian was a period of erosion or non deposition. This regional mid-Cretaceous stratigraphical gap might result from rift flank uplift, as expected in the case of a thick and cold prerift lithosphere. However, coeval collision of the Kargi Block, along the North Tethyan subduction zone at the southern margin of the Pontides, might also have contributed to this uplift. A rapid thermal post-rift subsidence of the margin occurred during the Coniacian–Santonian. Collision of the Kirs¸ehir continental block commenced in Early Eocene time (zone NP12) giving rise to compressional deformation and sedimentation in piggyback basins in the Central Pontides, whereas the eastern Black Sea was still opening.

It is commonly accepted that the Black Sea Basin opened as a back-arc basin during the Mesozoic, as a consequence of the northward subduction of the Neotethys ocean (Hsu¨ et al. 1977; Letouzey et al. 1977; Zonenshain & Le Pichon 1986). Alternatively, it may have opened under an extensional regime following the Palaeo-Tethyan collision and overthickening of the crust (Yig˘itbas¸ et al. 1999, 2004). However, its precise timing of opening is still under debate (e.g. Nikishin et al. 2003). The eastern Black Sea Basin (EBS) (Fig. 1) is supposed to have rifted in the Upper Paleocene (Robinson et al. 1995; Robinson 1997). This Paleocene (postDanian) age of rifting is supported by the presence of an almost complete Mesozoic to Lower Paleocene series in exploration wells drilled on the Shatsky Ridge (Fig. 1) (e.g. Robinson et al. 1996). The western Black Sea Basin (WBS) (Fig. 1) is generally considered to have rifted during the middle Cretaceous (Late Barremian or Aptian– Albian– Cenomanian; e.g. Finetti et al. 1988; Go¨ru¨r 1988; Manetti et al. 1988; Go¨ru¨r et al. 1993; Robinson

et al. 1996). This age is based on facies and thickness variations in the Cretaceous stratigraphic sequence of the Central Pontides (Go¨ru¨r 1988, 1997; Go¨ru¨r et al. 1993). However, pointing out that arc magmatism started in the Western Pontides only in the Turonian, Tu¨ysu¨z (1999) then Sunal & Tu¨ysu¨z (2002) suggested that the main opening phase had occurred during the Turonian– Maastrichtian. Moreover, based on heat-flow data, Verzhbitsky et al. (2002) obtained a 70 –60 Ma age (Maastrichtian –Danian) for the lithosphere of the western and eastern basins. Surface data concerning the rifting and evolution of the Black Sea can be obtained from the thrust belt of the Pontides, which extends all along its southern margin. The Eocene compression and thrusting have uplifted sediments of the of the Black Sea margin. Therefore, the Cretaceous ‘syn-rift’ sequence can be precisely dated by onshore studies (Fig. 2). We focused our work in the Central Pontides Belt (Fig. 1) where good outcrops of the MesozoicPalaeogene sedimentary sequence are present

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 113–136. DOI: 10.1144/SP340.7 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Location of the arc of Central Pontides between the Western Black Sea Basin and the Kirs¸ehir Block (modified after Robinson et al. 1996; Robinson 1997; Okay & Tu¨ysu¨z 1999).

(Go¨ru¨r 1988; Tu¨ysu¨z 1999), while the Eastern Pontides are mainly constituted by an Upper Cretaceous-Oligocene magmatic arc (e.g. Yılmaz et al. 1997). The Central Pontides Belt results from the inversion of part of the southern margin of the WBS. Thus it may comprise sequences related to the opening of the WBS, and therefore the oldest syn-rift deposits of the Black Sea. In order to constrain the timing of the western Black Sea opening, we have collected 164 samples from the Cretaceous to Palaeogene sedimentary sequences, in 143 localities of the Central Pontides (Table 1). The samples are dated by nannofossils, which provided precise ages for the stratigraphic sequence of the Black Sea margin. The observed main nannofossil assemblages used for each age determination are summarized in Table 2.

Overview of the stratigraphic sequence of the Pontides Owing to facies and thickness variations, the Cretaceous sequences of northern Turkey have been divided into a number of formations with local names that cause a great deal of confusion. Go¨ru¨r

(1997) has proposed a simplified stratigraphic scheme by distinguishing a ‘syn-rift’ sequence of Early Cretaceous age, from a ‘postrift’ sequence (Fig. 3). Lower Cretaceous sediments are generally rare around the Black Sea Basin. They crop out extensively in the Central Pontides (Fig. 2), in particular in the Ulus and the Zonguldak Basins (Fig. 2). It was from stratigraphic studies of these two basins that Go¨ru¨r (1997) proposed that the C ¸ ag˘layan Group (Fig. 3) represents the syn-rift deposits of the western Black Sea. This group is a 200–1300 m thick sequence of grey to black shales, marls and sandstone. Its clastic nature contrasts with the underlying grey to white limestone of the ˙Inaltı Formation (Derman & Sayılı 1995) (Fig. 3). According to Go¨ru¨r et al. (1993) and Go¨ru¨r (1997) these sediments, that are rich in organic matter, are witness for anoxic conditions resulting from restricted water circulations. They proposed that such anoxic conditions resulted from the disintegration of the carbonate platform by normal fault scarps that isolated the western Black Sea rift from the main Tethys Ocean located to the south. The carbonates of the ˙Inaltı Formation are not reliably dated. Locally foraminifers of Late

DATING OF THE BLACK SEA BASIN

Fig. 2. Structural sketch of the Central Pontides arc with location of the studied area (Fig. 4) and the Lower Cretaceous basins. The north Anatolian fault runs through the subduction complexes.

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Table 1. Coordinates and ages of the 164 samples dated with nannoplankton Longitude UTM36

Latitude UTM36

Sample name

458996 456035 435771 373397 392428 412724 411056

4596877 4598317 4562960 4570272 4586435 4581106 4582529

03-1 03-2 03-5 03-7 03-8 03-9 03-10

452433 452433 452433 452433 452433 547642 554238 544191 495198 452595 565862 559198 505499 505499 501158 493185

4621005 4621005 4621005 4621005 4621005 4613153 4617551 4649177 4635760 4620704 4583615 4614685 4633850 4633850 4638416 4634755

03-11 03-12 03-13 03-14 03-15 03-16 03-17 03-21 03-23 03-24 04-2 04-4 04-5 04-6 04-7 04-8

542097 542097 542097 542097 542097 542158 593438 595432

4617526 4617526 4617526 4617526 4617526 4617960 4647043 4646305

04-11 04-12 04-13 04-15 04-16 04-19 04-20 04-21

669099 653700 675286 658851 646635 373427 374757

4630111 4609944 4587100 4594733 4597624 4570281 4569983

04-22 04-23 04-24 04-25 04-26 04-29AB 04-30

378272

4574985

04-31

389917 390201

4584302 4585136

04-32 04-33

443577 452624 452624 453044 443147 435284 445964 450541 460527

4607278 4621116 4621116 4620737 4609861 4563157 4561524 4561488 4555717

04-36 04-41 04-42AB 04-45 04-46 04-47 04-51 04-53 04-55

Formation names Caglayan, Ulus Caglayan, Ulus Akveren–Atbasi Yemislic¸ay, Kale Caglayan, Kilimli Akveren Caglayan, Sapc¸a– Himmetoglu Caglayan, Kilimli Cemaller Kapanbogazi, Basko¨y Yemislic¸ay, Dilence Kapanbogazi, Basko¨y Akveren–Atbasi Caglayan Yemislic¸ay Akveren–Atbasi Kilimli Kusuri Akveren–Atbasi Akveren–Atbasi Akveren–Atbasi Kusuri Kapanbogazi Caglayan Caglayan Kapanbogazi Kapanbogazi Caglayan Akveren–Atbasi Akveren–Atbasi – Kusuri Akveren–Atbasi Caglayan Kusuri Kusuri Kusuri Yemislic¸ay, Kale Yemislisc¸ay, Sarikorkmaz Yemislic¸ay, Red Pellagic L. Inalti Caglayan, Kilimli – Inpiri Kusuri Caglayan Cemaller Caglayan Akveren, Alapli Akveren–Atbasi Caglayan, Ulus Inalti Caglayan, Ulus

Nannoplankton age Early Cretaceous Early Cretaceous Lower Campanian Coniacian– Santonian Early Cretaceous Upper Maastrichtian– Eocene Early Cretaceous Lower Aptian Coniacian Santonian Santonian Coniacian– Santonian Campanian Early Cretaceous Santonian Upper Paleocene Aptian Lower Eocene NP13 Lower Campanian – Maastrichtian Upper Paleocene NP5 Upper Paleocene NP5 Lower Eocene NP13 Upper Valanginian– Lower Barremian Santonian Hauterivian Hauterivian Santonian Santonian Barremian Upper Maastrichtian Lower Eocene NP13 Upper Paleocene NP9 Berriasian –Valanginian Lower Eocene NP13 Middle Eocene NP14 Lower Eocene NP13 Santonian Santonian Late Cretaceous Lower Cretaceous Lower Cretaceous Middle Eocene Lower Aptian Santonian Barremian Upper Campanian Upper Paleocene NP9 Barremian Early Cretaceous Early Cretaceous (Continued)

DATING OF THE BLACK SEA BASIN

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Table 1. Continued Longitude UTM36

Latitude UTM36

Sample name

462152 471258 480964 523991 563169 564605 559672 555747 542097 551523 551523 551523 378854

4556461 4563317 4563642 4564113 4579102 4574754 4609981 4610851 4617526 4611032 4611032 4611032 4575113

04-56 04-58 04-59 04-60 04-61 04-62 04-65 04-66 04-69 04-70 04-70A 04-70C 06-2

391811

4585492

06-4-8

392563 392734 390750 390327 389590 410187 409612 409553 409519 409666 424088 419895 417958 411483 411483

4582373 4582133 4580399 4580791 4581575 4558194 4557625 4557439 4557956 4559005 4571398 4576332 4578064 4582015 4582015

06-11 06-12 06-13 06-14 06-15 06-17 06-19 06-18,20 06-21 06-22 06-23 06-24 06-25 06-27 06-28

411056 411056 410172 392539 392523 393694 394184 392268 391715 452010 452010 452427 452613 452568 453579 453500 454874 454731 446318 446544 443132 430130 419134 404250

4583307 4583307 4584966 4585066 4582627 4581731 4581736 4581087 4581112 4620905 4620905 4620910 4621066 4621172 4620743 4618943 4621982 4621724 4614856 4614643 4609951 4595224 4602847 4594100

06-30 06-31 06-32 06-33 06-34 06-35 06-36 06-38 06-39 06-40 06-41 06-42 06-43,44 06-45 06-49-51 06-52 06-57 06-58 06-59 06-60 06-61 06-62 06-63 06-64

Formation names Kusuri Kusuri Kusuri Kusuri Kusuri Kusuri Kusuri Kusuri Kapanbogazi Kusuri Kusuri Kusuri Yemislic¸ay, Red Pellagic L. Caglayan, Kilimli-Inpiri Caglayan, Tasmaca Caglayan, Tasmaca Caglayan, Cemaller Caglayan, Cemaller Caglayan, Velibey Atbasi Atbasi Atbasi Atbasi Atbasi Atbasi Kusuri Kusuri Caglayan, Tasmaca Yemislic¸ay, Dereko¨y, Cambu Caglayan, Sapc¸a Caglayan, Sapc¸a Caglayan, Sapc¸a Caglayan, Velibey Caglayan, Velibey Caglayan, Tasmaca Caglayan, Tasmaca Caglayan, Cemaller Caglayan, Cemaller Caglayan Cemaller Kapanbogazi Caglayan Caglayan Caglayan Caglayan Caglayan Caglayan Akveren, Alapli Akveren, Alapli Akveren, Alapli Kusuri Yemislic¸ay, Unaz Caglayan, Kilimli – Inpiri

Nannoplankton age Middle Eocene Lower Eocene NP13 Lower Eocene NP13 Lower Eocene NP13 Lower Eocene NP12 Middle Eocene Lower Eocene NP12 Middle Eocene NP14 Santonian Lower Eocene NP12 Lower Eocene NP13 Lower Eocene NP13 Late Cretaceous Barremian Upper Aptian Upper Aptian Upper Albian Early Cretaceous Azoic Uppermost Paleocene NP9 Lower Paleocene NP3 Upper Paleocene NP5 Uppermost Paleocene NP9 Lowermost Eocene NP10 Lower Eocene NP11 Lower Eocene NP12 Lower Eocene NP13 Lower Albian Late Cretaceous Upper Aptian Upper Aptian Lower Albian Azoic Azoic Lower Albian Lower Albian Upper Albian Upper Albian Upper Aptian Coniacian – Santonian Santonian Barremian Barremian Barremian Upper Aptian Upper Aptian Upper Aptian Lower Campanian Lower Campanian Upper Campanian Middle Eocene NP14b Santonian Upper Aptian (Continued)

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Table 1. Continued Longitude UTM36

Latitude UTM36

Sample name

Formation names

404075

4594084

06-66

403348 400957 401409

4594134 4592334 4592715

06-67 06-68 06-72

402512

4593636

06-73

405098 408407 408014 410755 412041 412210 447811 447937 454929

4587045 4586323 4584886 4585914 4586200 4586567 4606314 4605242 4598996

06-74 06-75 06-76 06-77 06-78 06-79-81 06-82 06-83 06-84

Caglayan, Kilimli – Inpiri Caglayan, Kilimli Caglayan, Kilimli Caglayan, Kilimli – Inpiri Caglayan, Kilimli – Inpiri Caglayan, Velibey Caglayan, Sapc¸a Caglayan, Sapc¸a Caglayan, Sapc¸a Caglayan, Sapc¸a Caglayan, Sapc¸a Kusuri Kusuri Akveren–Atbasi

459896 466436 468909 471798 471947 472331 469483 473794 478087 483322 487851 495665 478790 476576 474867 468740 467866 483687 488495 490515 490515 613771 614080 613865 628232 628232

4596892 4598877 4602930 4607246 4608846 4609988 4608745 4604309 4608282 4608514 4611048 4612761 4586589 4584049 4578606 4569277 4564677 4544118 4544328 4544543 4544543 4608680 4608670 4607737 4609604 4609604

06-86 06-87 06-88 06-89 06-90 06-91 06-92-95 06-96-97 06-98 06-99 06-100 06-101-105 06-106 06-107 06-108 06-109 06-114 06-121 06-122 06-124 06-125 06-126 06-127 06-129 06-133 06-134

Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Caglayan, Ulus Kusuri Kusuri Kusuri Kusuri Atbasi Atbasi Paleocene –Eocene Atbasi Atbasi

Oxfordian–Berriasian age were found (Derman & Sayılı 1995). In its stratigraphic log, Go¨ru¨r (1997) considers an Oxfordian –Barremian age for the I˙naltı Formation and an Aptian –Cenomanian age for the uperlying clastic C¸ag˘layan Formation. The C¸ag˘layan Formation is overlain, with a slight angular unconformity, by red to pinkish, thinly bedded pelagic limestones, with volcaniclastic intercalations in its upper part. The basal red pelagic limestone form the Kapanbog˘azı Formation (e.g. Go¨ru¨r et al. 1993) for which an

Nannoplankton age Upper Aptian Aptian Lower Aptian Lower Aptian Lower Aptian Azoic Upper Aptian Upper Aptian Upper Aptian Lower Aptian Lower Aptian Middle Eocene NP15 Upper Eocene NP19-20 Upper Santonian – Lower Campanian Aptian Aptian Aptian Aptian Upper Aptian Upper Aptian Upper Aptian Aptian Barremian Barremian Barremian Barremian Upper Aptian Upper Aptian Upper Aptian Lower Cretaceous Upper Aptian Lower Eocene NP12 Middle Eocene NP16-17 Eocene Middle Eocene NP17 Uppermost Maastrichtian Upper Paleocene NP9 Middle Eocene NP14b Uppermost Maastrichtian Lower Eocene NP13

upper Cenomanian to Campanian age was proposed based on foraminifers (Ketin & Gu¨mus 1963). According to Go¨ru¨r (1997) the drastic change in the style of sedimentation from the dark coloured siliciclastic sediments of the C¸ag˘layan Formations, which accumulated in anoxic conditions, to the overlying red pelagic limestones, resulted from a rapid widening of the rift, end of anoxia, and a regional subsidence. This author interprets the Kapanbog˘azı Formation as a synbreakup succession.

DATING OF THE BLACK SEA BASIN

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Table 2. Nannofossil assemblages recognized for each age determination of Table 1 Stage

Nannoplankton zone

Upper Eocene

NP 19-NP 20 NP17

Middle Eocene

NP 15

NP 14b

NP 13 Lower Eocene

NP 12

Upper Paleocene

NP 10 NP 9 NP 5

Lower Paleocene Upper Maastrichtian

Upper Campanian Lower Campanian Santonian

Coniacian Upper Albian

Lower Albian Upper Aptian

NP 3

Nannofossil assemblages Chiasmolithus oamaruensis, Cycloccolithus formosus, Dictyococcites dictyodus, Discoaster barbadiensis, Ericsqonia subdisticha, Isthmolithus recurvus, Reticulofenestra umbilica Cribrocentrum reticulatum, Cyclococcolithus formosus, Dictyococcites dictyodus, Discoaster barbadiensis, D. saipanensis, D. tani nodifer, Reticulofenestra umbilica, Sphenolithus radians, Zygrhablithus bijugatus Chiasmolithus gigas, C. grandis, C. solithus, Discoaster barbasiesis, Reticulofenestra cf. umbilica (small), Rhabdosphaera gladius, Sphenolithus furcatolithoides, S. pseudoradians, Zygrhablithus bijugatus Chiasmolithus grandis, C. solitus, Cyclococcolithus formosus, Discoaster barbadiensis, D. sublodoensis, Reticulofenestra cf. umbilica (small), Rhabdosphaera inflata, Sphenolithus radians, Zyghrablithus bijugatus Same assemblage as in zone NP 12 but without Mathasterites tribraciatus Campylosphaera dela, Chiasmolithus solitus, Cyclococclithus gammation, C. formosus, Discoaster barbadiensis, D. binodosus, D. lodoensis, Discoasteriodes kuepperi, Marthasterites tribrachiatus, Sphenolithus radians Discoaster binodusus, D. multiradiatus, Marthasterites contortus Coccolithus pelagicus, Discoaster gemmeus, D. multiradiatus, Ellipsolithus macellus, Ericsonia subpertusa, Fasciculithus tympaniformis, Sphenolithus anarophus, Toweius eminens Ellipsolithus macellus, Ericsonia subpertusa, Fasciculithus tympaniformis Chiasmolithus danicus, Coccolithus pelagicus, Cruciplacolithus tenuis, Ericsonia subpertusa, Zygodiscus sigmoides Arkhangelskiella cymbiformis, Ceratolithoides aculeus, Cribrosphaera ehrenbergii, Eiffellithus turriseiffeli, Lithraphidites quadratus, Microrhabdulus decoratus, Micula murus, M. staurophora, Prediscosphaera cretacea, within the latest Maastrichtian occurrence of Micula prinsii Broinsonia parca, Ceratolithoides aculeus, Cribrosphaera ehrenbergii, Eiffellithus eximius, E. turriseiffeli, Lucianorhabdus cayeuxii, Prediscosphaera cretacea, Reinhardtites anthphorus, Quadrum gothicum, Q. trifidum Same assemblage as within the Upper Campanian but without Quadrum gothicum and Q. tifidum Eiffellithus eximius, E. turriseiffeli, Lucianorhabdus cayeuxii, Marthasterites furcatus, Micula staurophora, Prediscosphara cretacea, Reinhardtites anthophorus, within the uppermost part occurrence of Broinsonia parca expansa Same assemblage as in the Santonia but without Reinhardtites anthophorus Eiffellithus turriseiffeli, Eprolithus floralis, Hayesites albiensis, Parhabdolithus angustus, P. embergeri, Prediscosphaera cretacea, Tranolithus orionatus, Zygodiscus diplogrammus, Watznaueria barnesae Ellipsagelosphaera communis, Eprolithus floridanus, Parhabdolithus angustus, P. infinitus, P. embergeri, Prediscosphaera cretacea, Vagalapilla matalosa Chiastozygus litterarius, Coronolithoin achylosus, Ellipsagelosphaera communis, Eprolithus floralis, Nannoconus bucheri, N. circularis, N. elongatus, N. quadriangulus apertus, N. quadriangulus quadriangulus, Parhabdolithus angustus, Rucinolithus irregularis (Continued)

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Table 2. Continued Stage Barremian

Hauterivian

Berriasian – Valanginian

Nannoplankton zone

Nannofossil assemblages Calcicalathina oblongata, Cruciellipsis chiastia, Cyclagelosphaera margerelii, Micrantolithus obtusus, Nannoconus colomii, N. globulus, N. kamptneri, N. steinmannii, N. wassalii, Parhabdolithus asper, Watznaueria barnesae Bipodorhabdus colligatus, Bipodorhabdus colligatus, Calcicalathina oblogata, Cruciellipsis cuvillieri, Cyclagelosphaera margerellii, Ellipsagelosphaera communis, Lithraphidites bollii, Watznaueria barnesae Cyclagelosphaera deflandrei, C. margerelii, Ellipsagelosphaera communis, Nannoconus colomii, Parhbdolithus embergeri, Runcinolithus wisei, Watznaueria barnesae

Tu¨ysu¨z (1999), however, points out that according to a back-arc basin model, the syn-rift formation should include evidence of arc magmatism. He proposes that the older unit showing evidence for arc

magmatism, the Dereko¨y Formation (Fig. 3), is the real syn-rift sequence. This formation, is exposed in the Zonguldak Basin, and consists of thick lavas and carbonates of probable Turonian age. It is

Fig. 3. Comparison of dating of the stratigraphic formations of the Central Pontides (Black Sea coast, Zonguldak Basin and Ulus Basin). CEM, Cemaller Formation; INP, I˙npiri Formation; KAP, Kapanbog˘azı Formation; KIL, Kilimli Formation; SAP, Sapc¸a; TAS, Tasmaca Formation; YEM, Yemis¸lic¸ay Formation; VEL, Velibey Formation.

DATING OF THE BLACK SEA BASIN

noteworthy that, if this interpretation is correct, it would mean that the rifting of the Western Black Sea started in Turonian time and not, as commonly accepted, in Aptian time. Effectively, it is now accepted that no arc magmatism developed in the western Pontides during the Lower Cretaceous (Okay et al. 2006). However, volcanic and volcaniclastic rocks are the main elements of the Upper Cretaceous Black Sea margin sequence. The Kapanbog˘azı Formation conformably passes upwards to the Yemis¸lic¸ay Formation (Go¨ru¨r 1997), which is a thick succession (up to 1500 m) of volcanic rocks and volcaniclastic sediments with intercalations of red pelagic limestones similar to those of the Kapanbog˘azı Formation (Fig. 3). Based on foraminifers, a Turonian to Campanian age was assigned to this formation (Aydın et al. 1986; Tu¨ysu¨z 1999). The Yemis¸lic¸ay Formation is overlain by the Akveren Formation of Maastrichtian age (Ketin & Gu¨mu¨s 1963). This calciturbidite marks the end of magmatic activity in the Maastrichtian (Tu¨ysu¨z 1999; Sunal & Tu¨ysu¨z 2002). It is overlain by the Atbas¸ı Formation of Palaeogene age.

Nannoplankton dating of the Pontide stratigraphic sequence In the following we present the sedimentary units of the three main areas used in previous studies to establish the general stratigraphic sequence of the Central Pontides (Go¨ru¨r et al. 1993; Go¨ru¨r 1997;

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Tu¨ysu¨z 1999): the Black Sea coast, the Zonguldak Basin and the Ulus Basin. Nannoplankton age determinations were made to better constrain the age of these units and their correlations. For reasons of simplicity, we follow the tectono-stratigraphic schema of Go¨ru¨r (1997) that distinguishes the synrift C ¸ ag˘layan Group from the post-rift Upper Cretaceous sequences.

The syn-rift C ¸ ag˘layan Group Black Sea coast Along the Black Sea coast, a 0–200 m thick sequence of dark coloured Cretaceous rocks (sandy or clayey limestones) of the C ¸ ag˘layan Group (Fig. 3), overlays the Upper Jurassic– Lower Cretaceous ˙Inaltı limestones and older rocks. The I˙naltı limestone was interpreted as representing the south facing carbonate platform of the Neotethys Ocean (Koc¸yig˘it & Altıner 2002). The onset of terrigenous sedimentation on the carbonate platform corresponds to a major change that could be related to the opening of the Black Sea rift (Go¨ru¨r 1988). It is therefore crucial to date the oldest deposits of this group. Near Zonguldak, Kilimli and Amasra (Fig. 4) the C¸ag˘layan Group is represented by the Kilimli Formation (Tokay 1952; Go¨ru¨r 1997) (Fig. 3). It is an alternation of limestone, marls and shales that contains ammonites and nannoplanktons indicating an Aptian age (Tokay 1952; Akman 1992). In the

Fig. 4. The Early Cretaceous C ¸ ag˘layan Group and its sites of nannoplankton dating (cf. Table 1).

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Zonguldak area, Tu¨ysu¨z (1999) also distinguished the lower part of the sequence which includes sandstones, sandy limestone and marls of the Late Barremian-Early Aptian, as the ˙Inpiri Formation. However, as the Kilimli and I˙npiri Formations have similar facies and are in the same stratigraphic position, we combine them informally as the Kilimli-I˙npiri Formation (Fig. 5). Nannofossils confirm a Barremian age for the base of the clastic sequence west of Zonguldak (sample 06-4, Fig. 5). Near Amasra, a Barremian age was also determined at the base of the clastic sequence (samples 04-45, 06-43, 44, 45, 06-49, 50, 51, Fig. 6). At Amasra, an Early Aptian age was found in the lower part of the sequence (sample 04-41, Fig. 4) but here most of the C¸ag˘layan sequence have a Late Aptian age (samples 06-40, 06-52, 06-57, 06-58) (Figs 4 & 6). Similarly, at Kilimli, the samples collected along a 3 km long new road cut (06-68, 06-72 and 06-73) (Fig. 5) indicate an Early Aptian age, and the samples collected in the upper part of the sequence East of Kilimli (06-64 and 06-66) indicate a Late Aptian age (Fig. 5). It is concluded that along the Black Sea Coast, the Kilimli-I˙npiri sequence locally started in the Barremian, but most of the sediments were

accumulated during the Late Aptian (Fig. 6). This dating of the first clastic sequence on the platform, together with numerous normal faults observed in the Lower Cretaceous sequence along the Black Sea coast from Zonguldak to Ereg˘li (Fig. 7), suggest that the rifting and breakup of the carbonate platform (Go¨ru¨r 1993) started in the Barremian whereas tectonic activity and subsidence reached its climax during the Aptian.

Zonguldak Basin In the Pontides, the best exposures of the Lower Cretaceous sequence are found in the Zonguldak Basin, immediately SE of the city of Zonguldak (Fig. 5). In this area, the C ¸ ag˘layan Group was previously studied in detail and subdivided into four formations, the Velibey (Fig. 8), Sapc¸a (Fig. 9), Tasmaca and Cemaller Formation (Fig. 10) (Yergo¨k et al. 1989; Go¨ru¨r 1997; Tu¨ysu¨z 1999; Figs 3 & 5). In contrast with the Black Sea coast sections, in most of the northern margin of the Zonguldak Basin, the shelf carbonates of the ˙Inaltı Formation have been eroded before deposition of the Lower Cretaceous detrital sequence. Yellow-orange sands and

Fig. 5. The formations of the Early Cretaceous C¸ag˘layan Group around Zonguldak (Yergo¨k et al. 1989; location on Fig. 2) with sites of nannoplankton dating.

DATING OF THE BLACK SEA BASIN

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Fig. 6. Stratigraphic position of the samples (each dot represents a sample with precise nannoplankton dating) and tentative palaeogeographic interpretation of the facies and age of the C¸ag˘layan Group on a NW–SE section.

well bedded sandstones of the Velibey Formation (Fig. 8) rest directly on the Palaeozoic sequences (Fig. 5). No nannofossils were encountered in the samples collected from the Velibey Formation to constrain its age of deposition. Likewise no palaeontological data have been reported in the previous studies from the sandstones and gravels of this formation (Fig. 8). The sand of the Velibey Formation

consists of 95% quartz. Such an amount of quartz and intense fracturing observed in some outcrops could suggest that some rocks mapped as the Velibey Formation belong to the Pre-Jurassic basement. But its stratigraphic position seems to support an Aptian age. For example, to the NE of Zongulgak, the Velibey formation is underlain by the Kilimli Formation and is overlain by the Sapc¸a

Fig. 7. Stratigraphic contact of the Barremian-Aptian Kilimli-I˙npiri Formation (C ¸ ag˘layan Group) on the I˙naltı Jurassic-Neocomian limestone west of Zonguldak (near site 06-4 of Barremian age, Fig. 5). Along the Black Sea coast, the C¸ag˘layan Formation was cut by numerous normal faults like this one.

Fig. 8. Velibey Formation (C ¸ ag˘layan Group) near Kızılcakilise (Fig. 5). Barren (continental) sandstones and gravels.

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Fig. 9. Sapc¸a Formation (C¸ag˘layan Group) at Sapc¸a (Fig. 5), with an olistolith of Palaeozoic limestone (Zonguldak Formation). Samples 06-78 to 06-81 are Lower Aptian.

Formation. To the west of Zonguldak, it overlies the I˙naltı limestone and is overlain by the Cemaller Formation (Fig. 5). In addition, Tokay (1952) and Derman (1990) reported the presence of rudists, possibly of Late Albian age, in limestones interbedded in its upper part. This allowed Go¨ru¨r (1997) to propose an Early Albian age for these yellow sands. West of Kizilcakilise (site 06-15, Fig. 5), a new road cut allows observation of what is probably a progressive transition between the shelf carbonates (I˙naltı limestone) and the Velibey Formation (Yergo¨k et al. 1989). In this section, white quartz gravels and sandstone layers alternate with dark coal-bearing carbonaceous sandstones. The sample 06-15 collected from this section was barren. However, Late Aptian nannofossils in the

Fig. 10. Cemaller Formation (C¸ag˘layan Group) at site 06-13 between Kizilcakilise and Cemaller (Fig. 5). Marls and sands with thin coal intercalations (in dark). At neighbouring sites 6-38 and 6-39 with olistoliths, nannoplanktons also indicate a Upper Albian (first part) age.

samples collected from marls above it (06-11, 06-12, Fig. 5) suggest a pre– Late Aptian age for this outcrop of the Velibey Formation (Fig. 6). The Velibey Formation is overlain by the marine Sapc¸a Formation (Fig. 9), which is similar in aspect with the Kilimli-I˙npiri Formation, but more sandy (Fig. 7). It is an alternation of dark coloured sandstones with marls or shales rich in organic matter. Its thickness varies from 50 to 450 m and its macrofauna indicates an Albian age (Fig. 3) (Tokay 1952; Aydın et al. 1987; Go¨ru¨r et al. 1993). The Tasmaca Formation, another formation of the C ¸ ag˘layan Group (Fig. 3), is mainly developed near Kızılcakilise (Figs 5 & 11). This formation is a 100–400 m thick succession of poorly bedded organic matter, rich black shales and argillaceous limestones similar to those of the Sapc¸a Formation (Figs 9 & 11). Tokay (1952) proposed that the Tasmaca Formation is of Cenomanian age, based on ammonite fauna. However, nannofossil determinations performed in this study allow precise dating of various levels of the Sapc¸a and Tasmaca Formations. The samples collected from the Sapc¸a formation are dated as Early Aptian (samples 06-78, -81), Late Aptian (samples 06-30, 06-31, 06-75, 06-76, 06-77), and Early Albian (sample 06-32) (Figs 5 & 6). In contrast with the Cenomanian age of Tokay (1952), the nannofossil samples collected from the Tasmaca Formation indicate Late Aptian (samples 06-11, 06-12), and Early Albian ages (sample 06-27, 06-35, 06-36) (Figs 5 & 6). Therefore the Sapc¸a and Tasmaca Formations were contemporaneous during Late Aptian –Early Albian (Fig. 6). We conclude that these formations correspond to local variations in facies and bedding of contemporaneous deposits.

Fig. 11. Tasmaca Formation (C ¸ ag˘layan Group) at Kizilcakilise (Fig. 5). Samples 06-12 and 06-11 (in the village) are Upper Aptian. The top of the formation is Lower Albian (samples 06-35, 36, Fig. 5).

DATING OF THE BLACK SEA BASIN

In the south of the Zonguldak Basin, the Tasmaca Formation is overlain by the Cemaller Formation (Fig. 5). This formation is reported on the MTA 1:100 000 geological map (Yergo¨k et al. 1989) but included in the Senonian units on the MTA 1:500 000 geological map (Aksay et al. 2002), and also considered as part of the Upper Cretaceous series by Tu¨ysu¨z (1999) who describes a ‘shallow marine Cenomanian clastic sequence’. Effectively, the Cemaller Formation does not fit with the deepening character of the basin as indicated by the Sapc¸a and Tasmaca Formations and consists of sands with intercalations of clay and coal (Fig. 10). However, it contains limestone olistoliths similarly to the underlying formations of the C ¸ ag˘layan Group (Figs 5 & 9). Moreover, our three samples (06-13, 06-38, 06-39, Fig. 5) yielded nannoplankton allowing a precise age determination of the first part of the Upper Albian. On one hand, this age is compatible with the Late –Early Albian age of the underlying Tasmaca Formation (Figs 5 & 6). On the other hand, it contrasts with the previously proposed Cenomanian age (Tu¨ysu¨z 1999), and therefore, invalidates the discontinuity in sedimentation between the Cemaller Formation and the underlying formations of the C ¸ ag˘layan Group. The middle Cretaceous unconformity noted by Tu¨ysu¨z (1999) is in fact stratigraphically above the Cemaller Formation. In the Zonguldak Basin, the C¸ag˘layan Group is overlain by the Dereko¨y Formation of probable middle Turonian age (Tu¨ysu¨z 1999). Therefore, the Albian and Cenomanian deposits are missing in this basin (Fig. 6). Note that in the NE of the Zonguldak Basin, the Cemaller Formation was not deposited or was eroded before the middle Turonian. We conclude from these nannoplankton ages that the Velibey, Sapc¸a, Tasmaca and Cemaller Formations of the Zonguldak Basin form a continuous sequence from Late Barremian to the first part of the Late Albian, characterized by non-volcanogenic dark clastic material with limestone olistoliths (Fig. 6). Considering that the C ¸ ag˘layan sequence was interpreted as syn-rift by Go¨ru¨r (1993) our nannoplankton dating would confirm the Aptian– Albian age of rifting (Fig. 3). However, in contrast with Go¨ru¨r’s (1993) rifting model, the syn-rift sequence does not end up with deep deposits, but with shallow marine sands of the Late Albian Cemaller Formation. Furthermore, in the Zonguldak area, the Middle Cretaceous unconformity corresponds to a major gap in sedimentation (Fig. 3).

Ulus Basin The NE– SW trending Ulus Basin is the largest Lower Cretaceous basin of the Pontides (Fig. 2). In contrast to the Zonguldak Basin, the C¸ag˘layan

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Group is described as a single unit: the Ulus Formation (Fig. 3). It starts at the bottom with coarse clastic rocks and grades rapidly into turbiditic sandstones and shales. In the eastern part of the Ulus Basin the flysch deposits are poor in fossils, indicating an Early Cretaceous age (Tu¨ysu¨z 1999). In this study we precisely dated 27 samples from 18 localities in the flysch sequence of the Ulus Basin (Fig. 4). The ages ranged from Hauterivian near Ag˘lıko¨y, (Fig. 4, samples 04-12, 13, Table 1) and Barremian in the centre of the basin (samples 04-51 and 06-98 to 06-105) to Late Aptian (06-90 to 06-95 and 06-106 to 06-108, Fig. 4). These ages are similar to those found along the Black Sea Coast and in the Zonguldak Basin (Fig. 6). Surprisingly the youngest deposits of the Ulus Formation (Late Aptian) were found at the base of the sequence on the northwestern edge of the Ulus Basin (Fig. 4; close to the platform carbonates of the ˙Inaltı Formation). Moreover, samples 06-106 to 06-108 contain reworked species from the Barremian. This reworking and the onlap of the C¸ag˘layan Formation on the surrounding outcrops of the carbonate basement suggest tectonic activity and tilting during sedimentation, since the Barremian. In the Ulus basin, the age of the clastic sequence, is older than along the Black Sea coast (Hauterivian at Ag˘lıko¨y, Fig. 4). However, the geodynamic significance of the age of onset of detritic sedimentation in this basin is not as clear as along the Black Sea coast. Effectively, in this basin, we could not observe large normal faults as along the Black Sea coast (Fig. 7). Moreover, there are conspicuous compressional structures with intensity of deformation increasing toward the south and the east (Ag˘lıko¨y area), that is toward the accreted highpressure –low-temperature complexes of the Early Cretaceous subduction zone (Okay et al. 2006). It is thus possible, that in contrast to the Kilimli-I˙npiri Formation, the Ulus flysch was deposited on the accretionary wedge (Fig. 6). Therefore we will not consider the age of the Ulus Formation as critical for indicating the age of onset of the Black Sea rifting. Near Ag˘lıkoy, in the East of the Ulus Basin (Fig. 4), black shales of Hauterivian age (samples 04-12, 13) are unconformably overlain by the Kapanbog˘azı Formation of Santonian age (samples 04-11, 15, 16, Table 1), and the Barremian, Aptian, Albian, Cenomanian, Turonian and Coniacian are missing (Fig. 12). In the Ulus Basin, the youngest sediments of the Ulus Formation are Late Aptian. Similar to Black Sea coast and the Zonguldak Basin, the Albian, Cenomanian and Turonian deposits are missing in all of the Ulus Basin, which reveals a major gap in sedimentation in the Central Pontides (Fig. 6). This gap indicates erosion or nondeposition in the mid-Cretaceous. In any case this

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Fig. 12. Unconformity of the Kapanbogazy red pelagic limestone (Santonian, samples 04-11, 15, 16) on the C ¸ ag˘layan sandstone (Hauterivian-Barremian, samples 04-12, 13, 19) near Ag˘lıko¨y (Fig. 4).

regional gap was unexpected because according to most of the models (Go¨ru¨r 1988; Okay et al. 1994; Robinson et al. 1996; Banks & Robinson 1997), the WBS was opening at that time (Fig. 3).

Upper Cretaceous – Eocene post-rift sequence The Upper Cretaceous volcanic-sedimentary sequence In contrast to the Lower Cretaceous C ¸ ag˘layan sequence, characterized by rapid facies variations,

a thick sequence of Upper Cretaceous micritic limestone, volcanogenic and volcanic rocks, overlies the Lower Cretaceous and older rocks in most of the Central Pontides (Fig. 2). The limestone layers are mainly present within the lower part of the sequence and are named as the Kapanbog˘azı Formation (Figs 3 & 13) (Go¨ru¨r et al. 1993). They are white to pink (hematite rich) micritic and laminated limestones, in decimetric beds with thin clay intercalations (Fig. 14). They contain foraminifers indicative of a pelagic environment (Go¨ru¨r et al. 1993). Volcaniclastic and volcanic rocks intercalations become dominant upwards and the mainly volcanogenic sequence was called the Yemis¸lic¸ay Formation (Go¨ru¨r 1997) (Figs 3 & 14). The 10– 50 m thick basal ‘red pelagic limestone’ of the Kapanbog˘azı Formation is present over most of the Central Pontides (Fig. 2). It overlies various older rocks including the Lower Cretaceous and Carboniferous. A few kilometres NE of Amasra, a new roadcut allows observation of the unconformity of the Upper Cretaceous rocks with the underlying Lower Cretaceous black shales of the Kilimli-I˙npiri Formation (Yergo¨k et al. 1987) (Fig. 15). It is an angular unconformity of locally up to 508 (Fig. 15). Above the angular unconformity the sequence starts with 5–10 m of yellowish sands (Cemaller Formation, Yergo¨k et al. 1987), with some pebbles at the base locally. It is characterized by abundant burrows, lamellibranches, gastropods, indicating a shallow marine environment, and pieces of coal probably reworked from the Carboniferous basement cropping out nearby (Fig. 15).

Fig. 13. Late Cretaceous Kapanbog˘azı and Yemis¸lic¸ay Formations with their sites of nannoplankton dating (cf. Table 1).

DATING OF THE BLACK SEA BASIN

Fig. 14. Kapanbog˘azı (red pelagic limestone) and Yemis¸lic¸ay (volcaniclastic sediments) Formations at Amasra (Fig. 13). These formations are cut by syndepositional listric normal faults. The layers of the Yemis¸lic¸ay Formation are thicker in the downthrown blocks. A syntectonic wedge of clays of the Yemis¸lic¸ay Formation is dated from the Santonian (03-14). The red limestones are also Santonian (samples 03-13 and 06-42).

Thin sections in the shallow marine sandstones show abundant benthic foraminifers that contrast with the dominantly pelagic foraminifers of the Kapanbog˘azı red limestone immediately above (Fig. 15). The sharp contact between the sandstones and the pelagic limestones implies a sudden deepening of the Black Sea margin (Tu¨ysu¨z 1999).

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Samples collected from the Kilimli-I˙npiri Formation around Amasra contain nannofossils from Barremian to Late Aptian in age (Fig. 6). Above the angular unconformity, the sands are of Coniacian –Santonian age (samples 03-12, 04-42, 06-41). They were mapped as the Cemaller Formation (Yergo¨k et al. 1987), however, our new dating indicates that they are much younger than the Late Albian Cemaller Formation exposed near Zonguldak. We therefore consider that there is no correlation between these sands near Amasra, and the Cemaller Formation exposed near Zonguldak. The Amasra sands belong to the Upper Cretaceous trangressive sequence. In agreement with their Coniacian –Santonian dating, the overlying red pelagic limestone of the Kapanbog˘azı Formation (named Basko¨y Formation on the geological map, Yergo¨k et al. 1987) contains nannofossils of Santonian age (samples 03-13, 15) (Fig. 12). Our nannoplankton ages show that sediments of the Albian, Cenomanian and Turonian are missing in the Amasra stratigraphic sequence (Fig. 6) confirming the Middle Cretaceous gap mentioned above (Figs 6 & 12). Moreover, the observation of an angular unconformity at Amasra demonstrates that the gap in the Cretaceous sequence is at least partly due to erosion (Fig. 15). Note that another angular unconformity, with another dip direction in the Kilimli-I˙npiri Formation, can be observed between this location and the city of Amasra. The

Fig. 15. Angular unconformity of the Late Cretaceous on the Early Cretaceous at Amasra (Fig. 13). (a) The black shales of the Kilimli-I˙npiri Formation (C ¸ ag˘layan Group) are tilted to the West and are overlain by yellowish sands and red limestone of the Kapanbog˘azı Formation. (b) The Kilimli-I˙npiri Formation is dated with nannoplankton from Barremian (06-43, 06-44) to Lower Aptian (under the unconformity, samples 03-11, 04-41). (c, d) The yellow sands are Coniacian –Santonian, and the red limestones are Santonian. Even if the very first layers above the I˙naltı shelf carbonates of the Kilimli-I˙npiri Formation are Barremian, the base of the sequence already contains echinoids, gastropods, belemnites and ammonites (e) of Lower Aptian age. The yellow sands contain clasts of coal from erosion of the nearby Carboniferous sequence (c). Shallow marine environment is indicated by gastropods, burrows and abundant little planktonic foraminifers (f). In contrast, the Kapanbog˘azı limestone is characterized by large planktonic foraminifers (Globotruncanidae) indicative of a much deeper environment (g).

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middle Cretaceous angular unconformity also shows that tectonic deformation occurred before the Coniacian– Santonian transgression. Variations in thickness (hectometres) of the Kilimli-I˙npiri Formation around Amasra, and a local hard ground at the base of this clastic sequence indicate that vertical movements occurred during the deposition of these Barremian–Aptian sediments suggesting extensional block faulting at this time. The Upper Cretaceous section continues with the Kapanbog˘azı red pelagic limestone. Based on planktonic foraminifers, Go¨ru¨r (1997) dated the Kapanbog˘azı Formation as Cenomanian to Campanian (Fig. 3). Such a large time span places this formation as a possible lateral equivalent of the Tasmaca and the Yemis¸lic¸ay Formations. Based on nannofossil study we could precisely date the red pelagic limestone and the Yemis¸lic¸ay Formation at several localities. Around Zonguldak, a red pelagic limestone is named the ˙Ikse Formation of Turonian–Campanian age (Yergo¨k et al. 1987). It corresponds to the formation named Unaz in Tu¨ysu¨z (1999). NE of Zonguldak, near Hisaro¨nu¨ (Fig. 13), nannofossils found in this formation indicate a Santonian age (sample 06-63). Near Ereg˘li (Fig. 13) we also found red pelagic limestones in the ˙Ikse Formation, but we could not find characteristic nannoplankton species for a precise age determination (Upper Cretaceous, sample 06-2, Table 1). However in this area, we found similar pinkish limestones in the Kale Turonian–Campanian Formation (Yergo¨k et al. 1987) lying above the Ikse Formation. The nannofossils also indicate here a Santonian age (samples 03-7, 04-29, Fig. 13). These ages are in agreement with the foraminifer dating of Tu¨ysu¨z (1999) who bracketed the age of the Unaz Formation to the Late Santonian– Campanian (Fig. 3). Considering that our nine ages of the red pelagic limestone in five different localities throughout the Central Pontides (Ereg˘li, Hisaro¨nu¨, Amasra, Doganyurt and Ag˘lıko¨y, Fig. 13) are the same, we conclude that this limestone is Santonian and therefore does not interfinger with the syn-rift Aptian– Albian Tasmaca Formation (Fig. 3, Go¨ru¨r 1997). Furthermore, taking into account the angular unconformity described above, it is clear that a Lower Cretaceous sedimentary cycle has to be distinguished from an Upper Cretaceous cycle and that there is a tectonic event occurring in between. Around Zonguldak, a local formation, comprising clastic, volcanogenic clastic and pyroclastic rocks, is well developed between the Unaz (Kapanbog˘azı) limestone and the formations of the C ¸ ag˘layan Group. Tu¨ysu¨z (1999) named these rocks as the Dereko¨y Formation (Fig. 3). Yergo¨k et al. (1989) had distinguished four formations in this series: the Cemaller sandstone of Albian – Cenomanian age; the Go¨kc¸etepe Formation (lahar

and volcaniclastic sands); the Bas¸ko¨y Formation (volcanogenic clastic and marls); and the Dilence Formation (pyroclastic rocks and tuff ) of Turonian–Campanian age. Tu¨ysu¨z (1999) pointed out that the volcanic rocks of the Dereko¨y Formation represent the onset of arc magmatism in the region, which became more active during the Campanian (Cambu Formation) (Fig. 3). Intercalated in this sequence of pyroclastic and andesiticbasaltic lavas, he found pelagic limestone with foraminifers of middle Turonian age. According to Tu¨ysu¨z (1999) the upper part of the Cemaller Formation, that he considers as Cenomanian, interfingers with the Middle Turonian pyroclastic rocks and lavas. Therefore he proposes that the Dereko¨y Formation is Cenomanian –Turonian in age. However, taking into account our Late Albian nannoplankton dating of the Cemaller Formation near Zonguldak (Fig. 5), and the general gap between the Lower and Upper Cretaceous formations, we cannot follow this interpretation that includes the Cemaller Formation of Zonguldak in the Dereko¨y Formation and we will not retain this global dating of the Dereko¨y Formation. Unfortunately, we do not have any new age determination for the Go¨kc¸etepe, Basko¨y and Dilence Formations around Zongudak because they are mainly volcanogenic rocks, and we could not check the middle Turonian age based on foraminifers for the Dereko¨y Formation (Tu¨ysu¨z 1999). Note, however, that while near Cide, the Cemaller, Basko¨y and Dinlence Formations, included in the Dereko¨y Formation by Tu¨ysu¨z (1999), contain foraminifers suggesting a middle Turonian–Coniacian age for the Dereko¨y Formation (Tu¨ysu¨z 1999), our nannoplankton ages in the same area (Amasra) are more recent, Coniacian –Santonian. Note also that near Amasra, the Basko¨y and Dilence Formations (Yergo¨k et al. 1987), that form the Dereko¨y Formation of Tu¨ysu¨z (1999), correspond on the geological maps to the Kapanbog˘azı and Yemis¸lic¸ay Formations of Go¨ru¨r et al. (1993). The age of the Dereko¨y Formation seems not well constrained. Considering that it corresponds to the lower part of the Yemis¸lic¸ay Formation of Go¨ru¨r (1997) (Fig. 3), it is possible that it is Santonian, like the rest of the volcanic sequence that we dated. Finally we can precisely date the extent of the regional mid Cretaceous stratigraphic gap. Taking into account our lack of ages for the volcanogenic part of the Dereko¨y Formation near Zonguldak, we conclude that in the Western Pontides, deposits from the uppermost Albian to the lower Turonian are missing. Our sections in the Amasra area (samples 03-12, 04-42, 06-41) and in the Ag˘lıko¨y area (samples 04-11, 15, 16) suggest, however, that this stratigraphic gap may extend from the uppermost Albian to the Coniacian.

DATING OF THE BLACK SEA BASIN

The Kapanbog˘azı red pelagic limestone passes upwards into the Yemis¸lic¸ay Formation (Go¨ru¨r et al. 1993) equivalent to the Cambu and Dereko¨y Formations of Tu¨ysu¨z (1999) (Fig. 3). This formation is an up to 1500 m thick widespread series of volcaniclastic sediments and volcanic rocks (andesites and basalts) with local intercalations of red pelagic limestones similar to the Kapanbog˘azı limestones (Ketin & Gu¨mu¨s 1963; Go¨ru¨r 1997) (Fig. 14). It includes the Kazpınar, Liman, Kale and Sarıkorkmaz Formations of Yergo¨k et al. (1989). Based on foraminifers, a Turonian to Campanian age was proposed for this formation (Aydın et al. 1986; Tu¨ysu¨z 1990; Go¨ru¨r et al. 1993) (Fig. 3). However, similar to the Kapanbog˘azı Formation, we have always obtained a Santonian age consistently from several localities in the Central Pontides. For example, at Amasra we dated green and yellow marls intercalated in the lower part of the Yemis¸lic¸ay volcaniclastic rocks (samples 03-14, 06-42) (Figs 13 & 14). Near Ereg˘li, nannofossils found in the calciturbidites of the upper part of the volcanogenic sequence (Kale Formation), confirm this Santonian age (sample 04-30) (Fig. 13). Note that within the volcanogenic sequence, the intercalations of red pelagic limestone also gave a Santonian age at Dog˘anyurt (sample 03-21), at Ereg˘li (samples 03-7, 04-29) and at Hisaro¨nu¨ (sample 06-63, Fig. 13). This age is in agreement with our Santonian dating of the underlying Kapanbog˘azı Formation. Seven kilometres south of Amasra we could date the oldest sediments above the volcanic and volcanogenic sequence of the Yemis¸lic¸ay Formation. These grey

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marls named the Alaplı Formation and equivalent to the Akveren Formation of Go¨ru¨r et al. (1993), contain nannofossils of lower Campanian age (samples 06-59, 60) (Fig. 16). Finally, our nannoplankton dating allows for bracketing the age of the Yemis¸lic¸ay Formation from the Turonian– Coniacian –Santonian– Campanian (e.g. Go¨ru¨r 1997) to the Santonian (Fig. 3). This result is in agreement with the recent age determination by Okay et al. (2006) of a section east of our studied area, near Hano¨nu¨, where the basal and upper part of the Yemis¸lic¸ay Formation contain foraminifers characteristic of the Coniacian –Santonian.

The Upper Cretaceous – Cenozoic sedimentary formations The extensive magmatism ceased after deposition of the Yemis¸lic¸ay Formation (e.g. Tu¨ysu¨z 1999). Whereas Paleocene to Eocene volcanic rocks are well developed in the Eastern Pontides, they are present only locally in the studied area. Sedimentation continues above an unconformity with a 500–3000 m thick mainly turbiditic sequence. While in the southern part of the Pontides, the siliciclastic turbidites of the Gu¨rso¨ku¨ Formation (Ketin & Gu¨mu¨s 1963) are generally interpreted as a Maastrichtian forearc flysch sequences (Go¨ru¨r et al. 1984; Koc¸yig˘it 1991; Okay et al. 2006) in the studied area the Akveren, Atbas¸ı and Kusuri Formations are distinguished in the Maastrichtian to Eocene sequence (Aydın et al. 1986; Go¨ru¨r 1997; Tu¨su¨z 1999, Fig. 3).

Fig. 16. Campanian-Paleocene Akvenren and Atbas¸ı Formations with their sites of nannoplankton dating (cf. Table 1).

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The Akveren Formation consists of carbonates and calciturbidites. In the western Pontides, this formation unconformably overlies the older rocks including the Early Cretaceous Ulus Formation. According to Tu¨ysu¨z (1999) fossils at the base of the Akveren Formation confirm that sedimentation started in a shallow marine environment; then the 10–40 m thick carbonate mudstone of the Atbas¸ı Formation denotes a pelagic environment. The Akveren and Atbas¸ı Formations are followed by the 1000–1500 m thick siliciclactic turbidites of the Kusuri Formation. According to Go¨ru¨r (1997) and Tu¨ysu¨z (1999), the Akveren calciturbidites are Maastrichtian in age and the Atbas¸ı Formation is Paleocene in age based on foraminifers (Fig. 3). Our nannoplankton dating confirms that the calciturbiditic flysch extends into the upper Maastrichtian (samples 03-9 and 04-20) (Fig. 16). But as mentioned above, nannofossils indicate that the oldest sediments resting on the Yemis¸lic¸ay volcanogenic formation, the Alaplı marls, equivalent to the Akveren Formation, are older: Early Campanian (samples 06-59, 60) (Fig. 16). In the upper part of the Akveren-Alaplı Formation we dated thinly bedded limestone of the upper Campanian (samples 04-46, 06-61), thus confirming an age older than estimated before (Go¨ru¨r 1997; Tu¨ysu¨z 1999). We conclude that the Akveren-Alaplı Formation extends from the Lower Campanian to the Upper Maastichtian, which is in agreement with the nannoplankton dating of the underlying Santonian Yemis¸lic¸ay Formation (Fig. 3). Near Ag˘lıko¨y, in the eastern part of the studied area (Fig. 16), we could date as the same age the dark sandstones of the Caylak Formation, (Campanian, sample 03-16 and Early Campanian– Maastichtian, sample 04-4). We infer that this sandstone, that contains echinids and that onlaps the older formations (Akat et al. 1990), represents a shallow marine facies of the Akveren Formation on the southern margin of the Campanian Black Sea. To the east of our mapped area, near Hano¨nu¨ (Fig. 2), a thick sequence of grey marls with 10– 20 m white limestone at its base is transgressional on the Yemis¸lic¸ay-Gu¨rso¨ku Formations. Okay et al. (2006) recently found nannofossils of the Late Paleocene –Lower Eocene. Our samples gave ages ranging from the uppermost Maastrichtian to Middle Eocene (uppermost Maastrichtian for samples 06-126 and 06-133 base of the limestone at the entrance of the village of Sirke and East of Hano¨nu¨; Late Paleocene NP9, Early Eocene NP13 and Middle Eocene NP14b for samples 06-127, 06-134, and 06-129 respectively, taken from the marls above the basal limestone, Table 1). The local presence of uppermost Cretaceous at the base of the transgressional limestone is further

confirmed by the finding of an ammonite near Sirke. Considering the timing of the opening of the Black Sea (Robinson et al. 1995), we propose that this transgression on the accretionary wedge is related to the opening of the EBS. In the Zonguldak-Amasra area, the turbidites reach the upper Eocene in age. The following nannoplankton zones were identified in the Akveren, Atbas¸ı and Kusuri Formations: Paleocene NP3 (sample 06-19) NP5 (samples 06-18, 20), NP9 (samples 04-22, 06-17,21), Earliest Eocene NP10 (samples 06-22) and NP11 (sample 06-23) (Fig. 16), Early Eocene NP13 (sample 06-25), Middle Eocene NP14b (sample 06-62) and NP15 (samples 06-82), Late Eocene NP19-20 (samples 06-83, Fig. 17). In contrast, in the inner part of the Pontide Belt, the Palaeogene sequence fills intramountaineous basins: Karabu¨k Basin, Eflani Basin, Kastamonu Basin, Devrekani Basin (Fig. 17), Boyabat Basin, and Vezirko¨pru Basin (Fig. 2). Intraformational unconformities at the edge of the basins (Fig. 18) show that they are syncompressional piggyback basins formed and filled during the construction of the Pontide Belt, similar to those described in Central Anatolia (Kaymakci 2000). The filling of these intra-mountainous synthrusting basins starts in the lower Eocene (zone NP12, e.g. sample 06-121) and ends in the middle Eocene (zone NP17, e.g. sample 06-125) (Fig. 17). Therefore, even if sedimentation seems continuous in the Zonguldak-Amasra area, on our maps we have distinguished the Paleocene sequence (Akveren and Atbas¸ı Formations, Figs 3 & 16) from the Early-Eocene and Middle-Eocene sediments (C ¸ aycuma and Kusuri Formations, Figs 3 & 17), deposited in a compressional setting. Note that they have a very different geographic distribution, being present in particular inside the Pontide thrust belt (compare Figs 16 & 17).

Geodynamical implications Dating of the stratigraphic sequence of the Black Sea margin in the Central Pontides allows distinguishing two main periods of deposition: Barremian– Albian, and Coniacian– Eocene. It reveals a long mid-Cretaceous period of erosion that contrasts with the classical models of this margin where an Aptian– Albian rifting was immediately followed by rapid Upper Cretaceous thermal subsidence (e.g. Go¨ru¨r et al. 1993).

Barremian to Albian The Barremian –Albian sedimentary cycle starts with shallow marine clastic sediments. The Lower Cretaceous black shales and sandstones were

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Fig. 17. Eocene–Miocene basins with their sites of nannoplankton dating (cf. Table 1). During Eocene piggyback basins are created within the Pontide trust belt. Eocene marine sediments range from NP12 to NP17 in the Eflani, karabu¨k and Kastamonu piggyback basins. Marine sediments up to NP19-20 (Upper Eocene) were found in the Devrek Basin.

interpreted as indicating anoxia during the Black Sea rifting and were related to the opening of the Black Sea as a back-arc basin (Go¨ru¨r et al. 1993). This argument is not conclusive because anoxic events are frequent worldwide during this period. Besides, as noticed by Tu¨ysu¨z (1997) the general absence of subduction-related magmatism during the Early Cretaceous does not support this interpretation. However, there is strong evidence that supports the syn-rift interpretation of the Lower Cretaceous sequence. The arrival of detrital material

Fig. 18. Stratigraphic wedging at the front of a reverse fault along the northern margin of the Kastamonu Eocene basin. Such intra-formational angular unconformities at the border of the Eocene basins demonstrate that they are syncompressional piggyback basins.

on the carbonate platform denotes a major environmental change. At Amasra, Aptian sediments contain abundant clasts of Carboniferous coal attesting for local uplift and erosion during Lower Cretaceous subsidence and sedimentation. We could observe numerous normal faults that control thickness variations in the Lower Cretaceous deposits along the Black Sea coast (Fig. 7). This syntectonic sedimentation is also attested by the presence of olistoliths. Blocks of up to several tens of metres, mostly derived from the Upper Jurassic–Lower Cretaceous ˙Inaltı limestone, have been found in the Kilimli (Go¨ru¨r 1997), Sapc¸a (Derman 1990) and Tasmaca Formations (Siyako et al. 1981). We have also identified a 300 m long olistolith of Palaeozoic limestone within the early Aptian Sapc¸a Formation (Fig. 9). The presence of normal faults, thickness variations, olistoliths and hard grounds in the Lower Cretaceous sequence allows dating the rifting from the Barremian to the Albian. After deposition of hundreds to thousands of metres of sediments, this sedimentary cycle ended up in the upper Albian with sedimentation of shallow marine sands. Nannofossil dating gives evidence for a regional gap ranging from the uppermost Albian to the Turonian/Coniacian. An angular unconformity at Amasra demonstrates that this stratigraphic gap is partly erosional (Figs 12 & 15). Considering that this erosion follows syn-rift sedimentation and subsidence, we propose that it results from a thermally induced uplift of the rift

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shoulders. Such rift flank uplift can be expected during rifting of a thick (cold) lithosphere with high mechanical strength and high depth of necking (level of no vertical motions in the absence of isostatic forces; Fig. 19), which was inferred for the WBS (Robinson et al. 1995; Spadini et al. 1996; Cloetingh et al. 2003). The onset of rifting was characterized by the break of the carbonate platform and the arrival of clastic deposits (Go¨ru¨r 1988). This normal faulting of the carbonate platform is evident all along the

Black Sea coast between Zonguldak and Amasra (Figs 2, 6 & 7). This SW –NE ridge was probably a horst during the Cretaceous. To the SE, Okay et al. (2006) have evidenced SW –NE trending units of Cretaceous high-pressure–low-pressure metamorphic rocks. The Ulus Basin, elongated in the same SW –NE orientation is located between the Zonguldak-Amasra horsts and this Cretaceous accretionary complex. Our dating reveals a diachronous age for the base of this clastic sequence. Clastic sedimentation already existed during the Hauterivian in the eastern areas (near Ag˘lıko¨y, Figs 2 & 12) while carbonate deposition continued in the Zonguldak-Amasra horsts to the West. Likewise, compressional deformation of the Ulus Formation grades rapidly toward the SE and the accretionary complexes (Fig. 2). Taking into account this SW –NE structural trend and this WNW–ESE evolution of the deformation and sedimentation, we propose that the deep depositional environment of the Ulus flysch and its intense deformation are related to its proximity to the Lower Cretaceous accretionary prism (Fig. 6). In this frame, the Hauterivian age of Ag˘lıko¨y might not mean that extension occurred earlier in this area, but that deep marine conditions existed along the active margin (Fig. 6).

Coniacian to Eocene

Fig. 19. The rifting of the western Black Sea Basin in relation with the concept of lithospheric necking. Stratigraphic data show that the rifting started in Late Barremian and was very long at 40 Ma. In the Central Pontides, during the rifting, the change from sedimentation to erosion denotes rift flank uplift starting in Late Albian. These results support the models of rifting of a thick (cold) lithosphere with a large depth of necking (level of no vertical motions in the absence of isostatic forces) [(a) and (b), modified from Spadini et al. 1996]. Note that in the case of the Black sea the uplift of the southern rift shoulders might have been enhanced by the collision of the Kargi continental block. Following the continental breakup the margin subsided and the post-rift deposits onlapped the Central Pontides in Coniacian –Santonian (c).

Following the Turonian erosion, a new sedimentary cycle starts with Coniacian –Santonian shallow marine sands and a thick Santonian volcanic and volcaniclastic sequence with pelagic limestones intercalated. It continues with the deposition of the Akveren-Atbas¸ı flysch sequence which onlaps older rocks (Fig. 3). The distribution of the Senonian deposits along the Black Sea and their north–south variations in facies clearly indicate that they are deposits of the Black Sea margin (Figs 13 & 16). The sharp transition from shallow marine sands to pelagic limestone (Fig. 15) indicates a rapid deepening of the margin that is probably related to a postrift thermally induced subsidence (Go¨ru¨r 1988). The large thickness of the Senonian-Paleocene sequence (up to 3000 m, Fig. 3) and its widespread distribution, support this post-rift interpretation. Consequently the stratigraphic sequence of the Central Pontides allows precise determination of the age of rifting of the WBS: from Late Barremian to Coniacian –Santonian time (Fig. 3). We conclude that the rifting of the western Black Sea was very long: 40 Ma. Compression and uplift occurred from Eocene to present times. In the Pontides, the Eocene sequence was generally deposited within intra-mountainous basins (Fig. 17). Intra-formational angular unconformities (Fig. 18) demonstrate that compression

DATING OF THE BLACK SEA BASIN

was syndepositional. Therefore, the age of the onset of compression could be accurately determined by dating the older syncompressional deposits. In the Central Pontides they are of Early Eocene age (nannoplankton zone NP12). We explain the marine sedimentation in the intra-mountainous piggyback basins by the combined effects of post-rift subsidence of the Black Sea margin and loads of the Pontide thrust piles, in compensating the compressional uplift at its beginning. After the subsidence and filling of the piggyback basins by the end of Middle Eocene (NP 17), compressional deformation continued as shown by the folding of the uppermost Eocene marine sediments. We relate the Eocene onset of compression, to the collision of the Kirs¸ehir continental block, a promontory of the Tauride – Anatolide Block (Fig. 1). This local collision explains the diachronous onset of compression between the Central Pontides (Early Eocene, NP12) and the Greater Caucasus (Late Eocene, e.g. Robinson et al. 1995). The indentation of the Kirs¸ehir Block into the Pontides resulted in the northward convex arc geometry of the Central Pontides (Kaymakci 2000; Kaymakci et al. 2003a, b) and the inversion and uplift of part of the southern Black Sea margin whose sequence is now exposed onshore.

Discussion The Mesozoic –Cenozoic stratigraphy of the Central Pontides shows that the region experienced two main subsidence phases separated by an uplift and erosion during the Cretaceous. The significance of these movements needs to be discussed in the frame of the geodynamic evolution of the Black Sea. It is clear that the Lower Cretaceous represents a period of rifting. However, this rifting was not associated with arc volcanism (e.g. Okay et al. 2006) and according to Tu¨ysu¨z (1999) could predate an Upper Cretaceous rifting and oceanic spreading contemporaneous of arc volcanism. Zonenshain & Le Pichon (1986) proposed that the Black Sea results from back-arc extension during three successive episodes: 1) Early –middle Jurassic (opening of the Great Caucasus Basin); 2) Late Jurassic –beginning of the Cretaceous (opening of the Pre-Black Sea); and 3) end of the Cretaceous –Early Palaeogene. While the second event didn’t lead to complete breakup of the basement, the third episode of extension led to the formation of deep oceanic basins partially closed during the Cenozoic. This model was controversial because during the Neocomian the circum Black Sea region was a shallow shelf (Go¨ru¨r 1988). However, it considers the possibility of pre-Black Sea rifting episodes.

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Accordingly, the Barremian–Albian extensional tectonics (Fig. 7) could be interpreted as a pre-Black Sea rifting that may not have resulted in a complete break-off of the basement. The Ligurian back-arc basin showed such an evolution. Its Provenc¸al margin was cut by NNE –SSW grabens belonging to the Eocene– Oligocene west European intracontinental rift, and then it was broken-off obliquely along the ENE– WSW Late Oligocene Liguro– Provenc¸al back arc rift, that evolved to an oceanic basin (Hippolyte et al. 1993). A supporting evidence for a similar pre-Black Sea rifting unrelated to subduction, could be that along the Black Sea coast subsidence started in Barremian, before the beginning of convergence of Africa with respect to Europe (before 120–83 Ma, Rosenbaum et al. 2002). Such an idea of extensional tectonics unrelated to subduction was already proposed by Yig˘itbas¸ et al. (1999). However, the age of eclogites in the South of the Pontides, shows that, even though there was no arc magmatism in the Pontides during the Late Jurassic –Early Cretaceous (e.g. Okay et al. 2006), subduction and accretion were acting on the northern margin of the Neotethys Ocean during the Albian. Therefore, the Early Cretaceous subsidence and extensional faulting evidenced along the Black Sea coast (Fig. 7) might be related to this subduction. Moreover, in South Dobrogea and in the Moesian platform (Burchfiel 1976; Sandulescu 1978) carbonate deposition was marked by the arrival of abundant terrigenous material during the Aptian –Albian, suggesting that the Barremian– Albian rifting affected the conjugate margins of the western Black Sea. Furthermore, seismic data show that the Karkinit through West of Crimea, opened probably during the mid-Cretaceous and has an Upper Cretaceous –Eocene post-rift sequence (Robinson et al. 1996). Finally, palaeomagnetic analyses of the Kapanbog˘azı Formation in the Central Pontides indicates a palaeolatitude of 21.58N, (Channell et al. 1996) with the implication that the WBS was opened by the Coniacian – Santonian (Okay et al. 2006). In the frame of the Lower Cretaceous rifting, the middle Cretaceous erosion of the Pontides region most probably results from a thermal uplift of the rift shoulders. Seismic data from offshore Romania and Bulgaria show a regional unconformity in agreement with this thermal doming interpretation (Robinson et al. 1996). However, we cannot exclude that a part of the stratigraphic gap identified was related to the evolution of the subduction zone to the south. Effectively, the recent study of Albian eclogites in the accretionary complex south of the studied area (Okay et al. 2006) shows that an up to 11 km thick crustal slice (the Domuzdag complex, Ustao¨mer & Robertson 1997) of the Tethyan oceanic crust was metamorphosed

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at HP-LT at 105 + 5 Ma and exhumated in Turonian–Coniacian times in a fore-arc setting. This exhumation might be the consequence of the collision of the Kargı continental block that occurred just before, along the south facing Tethyan margin of the Pontides (Okay et al. 2006). It is thus possible that this collision participated in uplifting the Central Pontides during the Cenomanian– Turonian. This accretion was followed by the initiation of a new subduction zone to the south (Okay et al. 2006). We infer that this new subduction zone was wider than the Barremian–Albian one that was only related to the opening of the WBS. Effectively this later might have extended all along the Santonian volcanic arc, which is present in all of the Pontides, and was related to the opening of the EBS and the possible reactivation the WBS.

Conclusion Nannofossil investigations provided accurate ages for the sedimentary units of the Central Pontides. That superposed formations dated independently provide compatible ages supports the validity of our age determinations. The rifting of the WBS, that broke up the Upper Jurassic to Lower Cretaceous carbonate platform, started within the Barremian, but the main tectonic activity and subsidence took place during Aptian to Albian times. The syn-rift sequence (C ¸ ag˘layan Formation) is a detritic sequence containing olistoliths. It is characterized by rapid variations in facies and thickness, especially across normal faults. In the inner Black Sea margin, now inverted in the Pontides Belt, sediments of uppermost Albian to Turonian age are missing. This large regional stratigraphic gap, although not clearly identified by means of foraminifers, corresponds to the breakup unconformity of Go¨ru¨r (1997). Although tectonic analysis is necessary to better constrain the origin of the Cretaceous vertical movements, the observation of numerous normal faults in the Early Cretaceous series, and the angular unconformity observed at Amasra, support the idea that erosion occurred during rifting. It is interpreted here as resulting mainly from the thermal uplift of the western Black Sea rift shoulders. Age determinations based on nannofossils show that the post-rift subsidence, which was thought to start in Cenomanian time (Go¨ru¨r 1997), only began in the Coniacian –Santonian. The Cretaceous arc-volcanism that was considered to occur during the Turonian to Campanian, is bracketed to the Santonian in the Central Pontides. Our study confirms that the opening of the Black Sea was diachronous. The rifting of the WBS predates the Paleocene– Eocene rifting of the Eastern

basin. We show that, different to the EBS, the rifting of the WBS was very long (40 Ma) and produced a major uplift of the rift shoulders. These two different characteristics indicate that the WBS, in contrast to the EBS, opened on a thick lithosphere and involved a large depth of necking (about 25 km) as proposed by Spadini et al. (1996) and Cloetingh et al. (2003) based on modelling (Fig. 19). The stratigraphic dating of the Cenozoic sequence also constrains the timing of the Pontides compression. We show that along the southern margin of the Black Sea the orogenic movements are also diachronous. They started in the Central Pontides in the Lower Eocene, with the collision of the Kırs¸ehir block. This work was supported in 2003 and 2004 by the MEBE Programme and in 2006 by TUBITAK (Turkey) (Project No. C¸AYDAG-105Y146) and the University of Savoy (France). We are grateful to A. Arnaud and J. P. Thieuloy for foraminifer and ammonite determinations. Our thanks go also to Professor A. Okay and Professor E. Yig˘itbas¸ for their constructive reviews.

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Cretaceous – Neogene tectonic evolution of the northern margin of the Black Sea from seismic reflection data and tectonic subsidence analysis O. KHRIACHTCHEVSKAIA1*, S. STOVBA1 & R. STEPHENSON2* 1

Naukanaftogaz – Scientific Research Institute of National Joint-Stock Company Naftogaz of Ukraine, Uritskogo Str. 45 Kyiv 03035, Ukraine 2

School of Geosciences, Meston Building, King’s College, University of Aberdeen, Aberdeen AB24 3UE, UK

*Corresponding authors (e-mail: [email protected], [email protected]) Abstract: Three fundamental stages of the Cretaceous–Neogene tectonic evolution of the Odessa Shelf and Azov Sea (northern margins of western and eastern Black Sea basins, respectively) have been documented from the analysis of reinterpreted regional seismic profiling and one-dimensional (1-D) subsidence analysis of 49 wells, for which the stratigraphic interpretation was recently revised. (1) An initial active rifting stage began within the Early Cretaceous (not later than Aptian–Albian times) and continued until the end of the Santonian in the Late Cretaceous (c. 128–83 Ma). A system of half-grabens with mainly south-dipping normal faults developed on the Odessa Shelf at this time. The most profound faulting, accompanied by volcanic activity, occurred in the NE–SW orientated Karkinit-Gubkin rift basin at the boundary between the Eastern European and Scythian platforms. The footwalls of half-grabens were exposed above sea level and subject to erosion at this time. Active extensional processes affected the western part of Azov Sea and, while the onset and cessation of these cannot be tightly constrained, they are compatible with the well constrained results from the Odessa Shelf. (2) The second tectonic stage is one of passive post-rift thermal subsidence that lasted from the Campanian (Late Cretaceous) until the end of the Middle Eocene (83–38.6 Ma). (3) The third stage of basin evolution is one of inversion tectonics in a compressional setting. Discrete inversion events occurred at the end of the Middle Eocene, during the Late Eocene, during the Early Miocene and at Middle Miocene times (c. 38.6 Ma, c. 35.4 Ma, c. 16.3 Ma, c. 10.4 Ma, respectively) and typical inversion structures developed on the Odessa Shelf, some parts of which were uplifted and significantly eroded (down to the Lower Cretaceous succession). The southern part of the Azov Sea, opening into the northernmost eastern Black Sea basin, subsided rapidly during this time; thereafter, until the Quaternary, rapid subsidence was limited to its southeastern part, which was incorporated into the Indolo-Kuban foreland basin of the Greater Caucasus orogen.

The study area includes the northwestern shelf of the Black Sea (Odessa Shelf) and the Azov Sea. The Odessa Shelf is located mainly in the Ukrainian offshore region and is less than 100 m deep over much of its area. The Azov Sea is very shallow (,14 m) and straddles the Ukrainian-Russian border (Fig. 1). Some 15 000 km of regional reflection seismic data have been acquired since 1985 on the Odessa Shelf and within the Azov Sea. About 90 wells, a number of which reach depths .4 km (average depth is 2–3 km, Table 1), have been drilled in the study area for purposes of hydrocarbon exploration during the last four decades. The Mesozoic –Cenozoic stratigraphy, tectonics and evolution of the Black Sea region has been discussed in a number of key papers (Zonenshain & Le Pichon 1986; Finetti et al. 1988; Dercourt et al. 1993; Okay et al. 1994; Robinson et al. 1995;

Jones & Simmons 1997; Banks & Robinson 1997; Spadini et al. 1997; Robinson 1997; Nikishin et al. 1998, 2001, 2003; Morosanu 2002) as well as a number of papers appearing in this volume. However, there is little to no detailed description of the stratigraphy, tectonics and evolution of the Odessa Shelf and Azov Sea published in the international, English-language literature (cf. Stephenson et al. 2004; Stovba & Khriachtchevskaia 2009; Khriashevskaya et al. 2009). A number of one-dimensional (1-D) and twodimensional (2-D) modelling studies of basin subsidence and tectonic evolution have been carried out within the Black Sea basin in the last 20 years (e.g. Zonenshain & Le Pichon 1986; Robinson et al. 1995; Spadini et al. 1996; Cloetingh et al. 2003; Nikishin et al. 2003; Shillington et al. 2008). These were mostly focused on the deep

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 137–157. DOI: 10.1144/SP340.8 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Location of study area and position of (pseudo-)wells and seismic lines used in the study. Bulls-eyes are wells, simple dots are pseudo-wells; tectonic subsidence curves for those in larger font and underlined are shown in Figure 9. Country boundaries are also shown. The well Morskaya-1, mentioned in the text, is on Zmeiniy Island on the westernmost part of the Ukrainian Odessa Shelf.

water of the western and eastern Black Sea subbasins and their southwestern and southeastern margins. This paper presents the first detailed analysis of basin subsidence of the Odessa Shelf and Azov Sea, which allows us to draw conclusions regarding the timing and nature of rifting, post-rift subsidence and the timing and style of compression deformation (basin inversion) that affected these areas during their Cretaceous –Neogene tectonic evolution. It is based on the interpretation of the post-1985 seismic data and on the analysis of tectonic subsidence curves for 35 wells and 14 pseudo-wells and is accompanied by a comprehensive description of the stratigraphy and tectonics of the study area.

Regional tectonic setting of the study area The tectonic units and structure of the northern part of the Black Sea region have been described in a number of key international and local publications, including several in this volume (Tugolesov et al. 1985; Zonenshain & Le Pichon 1986; Go¨ru¨r 1988; Finetti et al. 1988; Beloussov & Volvovsky 1989; Ziegler 1990; Robinson et al. 1996; Banks & Robinson 1997; Robinson 1997; Robinson & Kerusov 1997; Nikishin et al. 1998, 2001, 2003; Stephenson et al. 2004; Stephenson & Schellart 2010; Yegorova & Gobarenko 2010). This section presents a brief review of the main tectonic units of the Odessa Shelf, Azov Sea and the surrounding

onshore regions (Fig. 2), which are mentioned in the paper. The study area occupies the southernmost part of the East European Platform (EEP) and the Scythian Platform (SP) and stretches along the boundary between these two major tectonic units (Fig. 2). The EEP has an Archean–Proterozoic basement. Its southern part is covered in some areas by undeformed Palaeozoic sediments and in others by Jurassic or Cretaceous rocks (Kruglov & Tsypko 1988). The exact locations and nature of the northern and southern boundaries of the SP and its main structural units, the precise age of its basement, and deep structure remain highly debatable issues. The age of the basement is probably Precambrian (e.g. Saintot et al. 2006; Gee & Stephenson 2006) though some authors consider it to be as young as Late Palaeozoic (e.g. Muratov 1968, 1969; Milanovsky 1991; Nikishin et al. 1998, 2001). The Southern Crimea Orogen is located south of, or upon, the basement of the SP and is thought to have been formed by pre-Bajocian, pre-Callovian and intra-Berriasian orogenic events (Nikishin et al. 2001). To the east the study area borders the Greater Caucasus Orogen, which formed as a consequence of shortening and closing of a former Jurassic–Eocene back-arc basin during Late Eocene to recent times (Milanovsky 1991; Ershov et al. 1999, 2003; Nikishin et al. 1998, 2001). The western and eastern Black Sea basins, which were formed as part of a Late Cretaceous –Eocene (?) back-arc oceanic (?) basin system (Nikishin et al.

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Table 1. Wells and pseudo-wells (from depth-converted seismic data) used in the study. An asterisk indicates a well artificially lengthened, with the given depth and age of the deepest strata being from seismic data Well name and abbreviation

Tectonic unit (Fig. 2)

Deepest strata Depth (m)

Age

Azov Sea area Aht-1 Ahtyrskaya-1 An-355 Anastasievsko-Troitskaya-355 Ber-1 Berdyanska-1 Bulg-1 Bulganakska-1 Chk-2 Chkalivska-2 El-1 Elektrorozviduvalna-1 Gen-2 Genicheska-2 Geol-258 Geologicheskaya-258 Griv-15 Grivnenskaya-15 Matr-1 Matroska-1 Mor-2 Morska-2 N.Ker-1 Northern Kerchenska-1 Obit-1 Obitochna-1 Obr-1 Obrucheva-1 Okt-245 Oktyabyrskaya-245 Prib-250 Pribrejnaya-250 S.Siv-1 Southern Sivashska-1 Strel-20 Strelkova-20 Tam-5 Tamanskaya-5 Ust-1 Ust-Salgirska-1 Vor-1 Vorontsovskaya-1

Azov Ridge Indolo-Kuban Depression North Azov Trough Indolo-Kuban Depression North Azov Trough Azov Ridge North Azov Trough Azov Ridge Indolo-Kuban Depression North Azov Trough Azov Ridge Indolo-Kuban Depression North Azov Trough Azov Ridge Azov Ridge Indolo-Kuban Depression Indolo-Kuban Depression Azov Ridge Indolo-Kuban Depression Azov Ridge North Azov Trough

2225 2345 1030 4760 630 668 2755 1325 2135 1908 1053 1255 850 496 1126 2590 666 2000 4300 2815 2180

Cenomanian Middle Miocene Lower Cretaceous Cenomanian Albian Oligocene Lower Cretaceous Middle Miocene Middle Miocene Albian Lower Cretaceous Middle Miocene Lower Cretaceous Oligocene Oligocene Upper Miocene Middle Miocene Albian Cenomanian Lower Cretaceous Lower Cretaceous

Odessa Shelf Arh-1 Bez-2 Des-1 Evp-2 Gamb-2 Gleb-1 Il-2 Kark-1 Odes-1 Odes-2 Olimp-400 Sereb-2 Tend-19 Tuzl-1 P-1 P-2 P-3 P-4 P-5 P-6 P-7 P-8 P-9 P-10 P-11 P-12 P-13 P-14

Karkinit Trough Karkinit Trough Kalamit Ridge Kalamit Ridge Karkinit Trough Karkinit Trough Kalamit Ridge Karkinit Trough East European Platform Karkinit Trough Krayova Step Karkinit Trough East European Platform East European Platform Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Karkinit Trough Gubkin Ridge East European Platform East European Platform East European Platform East European Platform

5217 2185 2515 710 3329 1100 2142 3489 1130 2995 3012 3190 2625 940 3940 4575 4700 5980 5800 3460 4030 3000 1600 1150 2370 3025 2750 1250

Cenomanian* Cenomanian* Cenomanian* Albian Cenomanian Paleocene Cenomanian Cenomanian Cenomanian Cenomanian Aptian Cenomanian Albian Lower Cretaceous Cenomanian Cenomanian Cenomanian Cenomanian Cenomanian Cenomanian Cenomanian Cenomanian Campanian Campanian Campanian Campanian Campanian Lower Paleocene

Arhangelska-1 Bezymyana-2 Desantna-1 Evpatoriyska-2 Gamburceva-2 Glebovska-1 Ilichevska-2 Karkinitska-1 Odesska-1 Odesska-2 Olimpiyska-400 Serebryanska-2 Tendrovska-19 Tuzlovska-1 pseudo-well 1 pseudo-well 2 pseudo-well 3 pseudo-well 4 pseudo-well 5 pseudo-well 6 pseudo-well 7 pseudo-well 8 pseudo-well 9 pseudo-well 10 pseudo-well 11 pseudo-well 12 pseudo-well 13 pseudo-well 14

140 O. KHRIACHTCHEVSKAIA ET AL.

Fig. 2. Major tectonic domains (labelled and coloured) and regional units (labelled and separated by dashed red lines) in the study area and the locations of interpreted seismic reflection profiles shown in Figures 3, 4 and 7. (Pseudo-)wells are labelled as for Figure 1.

CRETACEOUS–NEOGENE TECTONIC EVOLUTION OF THE BLACK SEA

2003; Stephenson & Schellart 2010) are located south of the study area. To the west, the Northern Dobrogea Orogen represents a former PermoTriassic rift basin that is considered to have been deformed by thrusting and folding from the time of the Jurassic/Cretaceous boundary up to the Neocomian (Banks & Robinson 1997; Seghedi 2001; Nikishin et al. 2001; Hippolyte 2002). The orogen is sometimes considered as the westernmost part of a Cimmerian-aged fold belt extending through the Crimea Orogen and onto the Caucasus region (e.g. Sengo¨r 1984; Nikishin et al. 2001). Postorogenic compressional events affected the area in Maastrichtian– Palaeogene times (Hippolyte 2002).

Stratigraphy of the Odessa Shelf and the Azov Sea The stratigraphy of the northern part of the Black Sea region has been described in many Ukrainian and Russian publications and reports (e.g. Astahova et al. 1984; Gojik et al. 2006a, b) and these form the basis of what is summarized here. A growing interest in hydrocarbon exploration in the area stimulated a revision of the existing geological and geophysical data and their interpretations. Recently, the Scientific Research Institute of Oil and Gas Industry – Naukanaftogaz (Kyiv, Ukraine) – together with scientists from National Academy of Science of Ukraine revised the palaeontology of some key wells on the Odessa Shelf and this has allowed a more accurate stratigraphic analysis of the sedimentary sequences penetrated by these wells (Table 2; Khriachtchevskaia et al. 2007, 2009). These new correlations have been used in this paper for calibration of regional seismic reflection sections as well as for construction of tectonic subsidence curves. Unfortunately, the recorrelated stratigraphic data are available only for the Odessa Shelf; the stratigraphy of the sedimentary sucession of the Azov Sea has not been updated since the 1970s. The oldest rocks encountered in several offshore boreholes in the area are Archean –Proterozoic metamorphosed rocks and intrusive complexes of the basement of the East European Platform (EEP) such as exposed further north in the Ukrainian Shield (Kruglov & Tsypko 1988). The EEP basement becomes deeper to the south and is covered unconformably by Cretaceous strata. Upper Proterozoic rocks occur within boreholes on the Simferopol Uplift of Crimea (Fig. 2, Muratov 1969; Kruglov & Tsypko 1988; Nikishin et al. 2001), Central Dobrogea (Chekunov 1994; Bogaec 1976), and Moesian platform (Robinson et al. 1997; Chekunov 1994). Palaeozoic and Triassic –Jurassic rocks occur occasionally within the Ukrainian offshore region.

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Devonian and Silurian sediments were penetrated by the Morskaya-1 borehole on Zmeiniy Island (Figs 1 & 3b) and consist of shaley limestone and marl with beds of limestone and shale (Sulimov 2001). Strata of probably Early Carboniferous age were dragged from the Lomonosov Massif (Fig. 2) on the continental slope to the SW of the Crimea Peninsula (Ivannikov et al. 1999). They are black non-carbonate massive shales. Within the Azov Sea, the Matroska-1 well penetrated strata of probable middle Carboniferous age (Dyakovich et al. 2001). Palaeozoic rocks are known from surrounding onshore regions such as the Predobrogea Depression (Dinu et al. 2002), North-Dobrogea (Catuneanu 1994), Moesian platform (Robinson et al. 1997), northern Crimea (Muratov et al. 1968; Astahova et al. 1984) and the western ForeCaucasus (Letavin 1978). The Upper Triassic–Lower Jurassic complex (Tavrik – or Tauric – series) is encountered in only a few offshore wells and the exact age of this sequence remains controversial. Moreover, none of offshore wells reaches the base of this complex. The Evpatoriyska-2 (Fig. 1) well penetrated 1300 m of Triassic sandstone and siltstone. Triassic –Jurassic siltstone, shale, sandstone and volcanic rocks are also likely present within the Azov Ridge (Figs 2 & 4; Kazmin 1979; Slavin & Khain 1980). Triassic –Jurassic breccia and shales are reported in some Romanian offshore wells (Tambrea et al. 2002). This complex is also widespread on the surrounding onshore areas (Nikishin et al. 2001). The Cretaceous – Cenozoic succession has been drilled by many wells on the Odessa Shelf and within the Azov Sea. It consists of terrigenous and carbonate formations with a thickness from 200 to 3600 m. Volcanic rocks occur within the Lower Cretaceous and lower part of the Upper Cretaceous succession. Some offshore wells have encountered andesites, andesitic tuffs and andesitic porphyrite of Albian age (Nikishin et al. 2001 and references therein). Volcanic rocks are also known to occur in the Cenomanian and Turonian successions of the Kerch peninsula, southeastern Crimea (Nikishin et al. 2003 and references therein). The division of the Cretaceous –Neogene succession into stratigraphic units within the study area is rather variable (Table 2). There is no consistent stratigraphic correlation between the Odessa Shelf and Azov Sea. The main reason is the (variable) sensitivity of the biostratigraphic data, which were used independently in each region. On the Odessa Shelf the Cretaceous –Eocene succession is divided stratigraphically into stages and sometimes even into substages (Table 2). The same sediments within the Azov Sea are only divided into epochs: Lower and Upper Cretaceous, Paleocene, Eocene and Oligocene (Gordievich 1974; Gojik

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Table 2. Cretaceous-Quaternary stratigraphy of the study area (ages from Harland et al. 1990) International stratigraphic scale Epoch

Stage

Pleistocene Piacenzian

Pliocene

Zancleian

Miocene

Messinian U

M

L

Quaternary – Pliocene–Upper Miocene

Serravallian Langhian

Azov Sea

Middle Miocene

Age of base (Ma)

Quaternary

1.64

Pliocene

5.2

Pontian

6.7

Meotian

Tortonian

8.5

Sarmatian

10.4

Konkian

12.0

Karaganian-Chokrakian

16.3

Maykop

35.4

Burdigalian Aquitanian Chattian

Oligocene

Maykop

Rupelian U

Eocene

Odessa Shelf

Holocene

Quaternary

M L

Paleocene

Regional stratigraphy

U L

Cretaceous

U

Priabonian Bartonian Lutetian Ypresian Thanetian Selandian

38.6

Middle Eocene

Eocene

Bahchisarayskiy Upper Paleocene

50.0 56.5

Paleocene

60.5

Danian

Lower Paleocene

65.0

Maastrichtian

Maastrichtian

74.0

Campanian

Campanian

83.0

Santonian

Santonian

Coniacian

Late Cretaceous

86.6

Coniacian-Turonian

90.4

Cenomanian

Cenomanian

97.0

Albian

Albian

112.0

Aptian

Aptian

124.5

Barremian

Upper Barremian

Turonian

L

Alminskiy

Hauterivian Valanginian

Not penetrated

Lower Cretaceous 145.6

Berriasian

et al. 2006a, b). In contrast, the stratigraphy of the Neogene succession of the Azov Sea has been studied in more detail than on the Odessa Shelf. As a result, given stratigraphic units from one region often cannot be fixed within the other.

Lower Cretaceous sediments are found in several offshore wells within the study area and lie with angular unconformity above the preCretaceous succession. According to seismic data they occur irregularly over the region (Fig. 3). The

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Fig. 3. Example interpreted seismic reflection profiles (fragments of regional lines) on the Odessa Shelf, located in Figure 2; fragments (b) and (c) overlap. Well Arh-1 in (a), actually lying some 17 km from the seismic line, illustrates the lengthening of a well used in the subsidence analysis with seismic data. N31-Q, Upper Miocene–Quaternary; N21, middle Miocene; P3-N11, Oligocene– Lower Miocene; P32, Upper Eocene; P22, middle Eocene; P12, Lower Eocene; K2cp-m-P1, Campanian– Maastrichtian –Paleocene; K2cm-st, Cenomanian –Santonian (Upper Cretaceous); K1al, Albian (Lower Cretaceous); J3-K1ap, Upper Jurassic– Albian; S-D, Silurian-Devonian.

(proven) oldest Cretaceous rocks were discovered in the Olimpiyska-400 offshore well (Fig. 1). The well was terminated in Upper Barremian terrigenous strata (siltstone and shale) with an incomplete

thickness of 650 m (Gojik et al. 2006b). Aptian sediments are composed of alternating calcareous shales, siltstones and limestones with a thickness of 402 m in the Olimpiyska-400 well (Gojik et al.

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2006b). A few wells within the Azov Sea penetrated the upper part only of the Aptian succession, comprising sandstones and siltstones up to 100 m in thickness (Gojik et al. 2006a). These sediments are absent on the Azov Ridge (see Fig. 2), as is the whole of the Cretaceous succession, probably as a result of subsequent erosion (Kazmin 1979). An Albian terrigenous and volcanic complex unconformably covers older rocks. Its thickness reaches 220 m in the Olimpiyska-400 well and more than 522 m in the Desantna-1 well (Fig. 1; Gojik et al. 2006b). The average thickness of Albian sediments in the Azov Sea succession is some 300 m (Dyakovich et al. 2001). According to the recent palaeontological data (L. Plotnikova, L. Yakushin 2005, pers. comm.), the Albian stage within the Olimpiyska-400 well is represented by middle and upper substages with the Lower Albian sediments apparently being absent. Upper Cretaceous sediments are widespread but absent in some parts of inverted structures (Azov, Gubkin, Kalamit Ridges; see Figs 2– 4). They unconformably cover the underlying strata and record a change from muddy/sandy Shelf to carbonate sedimentation. The thickness of the Upper Cretaceous carbonate succession within the Azov Sea varies from 200 to 700 m (Gojik et al. 2006a) and is over 2000 m in the axial part of Karkinit Trough (Figs 2 & 3a; Gojik et al. 2006b). The Upper Cretaceous sediments within the Odessa Shelf are subdivided into stages (Table 2). Cenomanian sediments are represented by marls, limestones and sandstones with thicknesses ranging from tens of metres on the northern margin of the Odessa Shelf to 660 m in the axial part of Karkinit Trough, decreasing to 400 m southwards within the Kalamit Ridge (Figs 2 & 3a). Cenomanian sediments are absent in some offshore wells on the Odessa Shelf. Coniacian –Turonian sediments have

a similar carbonate composition and are often considered as a single complex with a thickness up to 700 m (Gojik et al. 2006b). The Santonian unit consists of porcelain-like limestones and has a maximum thickness of 200 m in the axial part of Karkinit Trough. Campanian sediments are composed mostly of clay limestones and marls with beds of clay and sandstone with thicknesses of 400–600 m. In the Olimpiyska wells the thickness of the Campanian is about 100 m (L. Plotnikova & L. Yakushin, pers. comm.). The Maastrichtian unit consists of limestones, rarely of marls and calcareous sandstones and has a thickness of 100–200 m (Gojik et al. 2006b). Paleocene sediments over the area consist of shales, limestones, and marls. They cover the underlying sediments conformably or unconformably (Astahova et al. 1984) and are widespread, except on the Azov Ridge (Figs 2 & 4). In the Azov Sea the Paleocene sediments have thicknesses in the range 100–300 m. On the Odessa Shelf the succession is subdivided into two regional units with total thicknesses up to 350 m (Table 2, Gojik et al. 2006b). Eocene sediments within the study area are represented by shales, siltstones and marls with a thickness of 150–280 m in the Azov Sea and 500–1300 m on the Odessa Shelf. According to seismic and well data the Eocene succession is absent on the Kalamit, Gubkin and Azov Ridges (Figs 2, 3 & 4). Most likely this is because of subsequent erosion. The Eocene succession is divided into three units on the Odessa Shelf (Table 2). The next sediment complex is the widespread Oligocene–Lower Miocene shales, muddy siltstones and sandstones of the so-called Maykopian series. The boundaries of the base and top of the Maykopian are erosive in most of the study area. The thickness of the complex is considerable,

Fig. 4. Seismo-geological profile across the Azov Sea (Gojik et al. 2006a), located in Figure 2. N2-Q, Pliocene-Quaternary; N31m-p, Meotian–Pontian (Upper Miocene); N31s, Sarmatian (Upper Miocene); N21kn, Konkian (Middle Miocene); N21kr-ch, Karaganian –Chokrakian (Middle Miocene); P3-N11, Oligocene–Lower Miocene; P1-2, Paleocene–Eocene; K1-2, Cretaceous; K1, Lower Cretaceous; K2, Upper Cretaceous; T-J-PZ, Triassic–Jurassic– Palaeozoic; AR– PR, Archaean–Proterozoic.

CRETACEOUS–NEOGENE TECTONIC EVOLUTION OF THE BLACK SEA

varying from 400 to 2000 m on the Odessa Shelf, and from 100 m in the north to more than 3000 m in the south of the Azov Sea (Figs 3 & 4). The middle Miocene section is quite thin, not exceeding 300 m in thickness. It consists of sandstones and shales with beds of marls. In some sections of the Azov Sea it is subdivided into the Karaganian-Chokrakian (shales, siltstones, sandstones and limestones) and the Konkian (limestones with beds of shales and siltstones) regional units. The middle Miocene succession has several unconformities (Gojik et al. 2006a, b). The Upper Miocene–Quaternary succession consists of alternating and interleaved shales, siltstones, beds of sandstones and limestones with thicknesses of 300 –700 m. Sarmatian calcareous shales, Meotian limestones and Pontian shales, sandstones, limestones of the Upper Miocene, the Pliocene sands and clays and Quaternary sands and loam can be recognized within the Azov Sea succession (Table 2; Gojik et al. 2006a; Astahova et al. 1984; Gordievich 1974).

Tectonic evolution of the Odessa Shelf and Azov Sea from seismic data The style and age of second order tectonic units recognized within the study area are illustrated by interpreted seismic sections from the Odessa Shelf (Fig. 3) and a cross-section based on seismic data from the Azov Sea (Fig. 4). All of these structures have a roughly east –west trend and, in general, they have been recognized in earlier works, in particular those on the Odessa Shelf (e.g. Tugolesov et al. 1985; Zonenshain & Le Pichon 1986; etc.). However, the new data presented here allow a more detailed assessment of the nature of these structures and the ages at which they formed.

Main tectonic elements of the Odessa Shelf The main tectonic unit of the Odessa Shelf is the Karkinit Trough, which is a part of the same rift basin system as the North-Crimea Trough on the northern Crimea Peninsula (Fig. 2). The Karkinit Trough is a rift basin, presumably formed contemporaneously with the main rift basins of the Black Sea, filled by Upper Jurassic (?)-Aptian (?) pre-rift sediments, Albian–Santonian syn-rift sediments, and Campanian– Pleistocene post-rift sediments with a total thickness up to 8 –9 km (Fig. 3a). The pre-rift sediments lie with angular regional unconformity above older rocks. The age of these underlying rocks is controversial, with ages from Palaeozoic to Triassic –Jurassic having been suggested (Nikishin et al. 2001). It is believed by many authors that pre-rift sediments directly

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overlie the basement of the Scythian Platform (SP). However, no data exist to confirm this. The pre-rift and syn-rift sediments form a system of deep grabens and half-grabens. The Karkinit Trough is separated from the EEP by the Golitsin Fault (Fig. 2). The sediments of the northern flank of the Karkinit Trough onlap the basement of the EEP. To the south, the Karkinit Trough is bordered by the Kalamit Ridge (Fig. 2), which represents the southern margin of the trough as well as perhaps the northern shoulder of the western Black Sea Basin. Relatively thin Mesozoic sediments cover the Upper Triassic –Lower Jurassic succession that crops out widely in the southern Crimea Orogen. The Kalamit Ridge is probably correlative to the Simferopol Uplift on the Crimea Peninsula (Denega et al. 1998). The thinning of the sedimentary succession within the ridge compared to the Karkinit Trough is partly caused by erosion during Cenozoic compressional (basin inversion) events but also reflects non-deposition related to its low rate of post-rift subsidence (Fig. 3a). The Gubkin Ridge is located west the Karkinit Trough and was formed by strong inversion movements along an originally normal syn-rift fault (Fig. 3c). In some places the sedimentary cover was completely eroded and basement crops out beneath the thin Miocene–Pliocene section. Most likely, before inversion, the Gubkin Ridge was the western part of the Karkinit Trough. The Krayova Step, located south of the Gubkin Ridge, has a truncated sedimentary cover consisting of Cretaceous and Palaeogene deposits and a Neogene succession increasing in thickness from 1400 to 2000 m north to south. Before rapid subsidence in the Neogene, the Krayova Step likely represented the western continuation of the Kalamit Ridge and was, as such, part of the southern margin of what can be referred to as a joint Karkinit-Gubkin rift basin. Because of rapid Neogene subsidence, the basement and its Mesozoic cover deepens in a southern direction from a depth of 2500 to 5800 m where the Krayova Step connects with the Histria Depression (Figs 2 & 3c; Denega et al. 1998). The Zmeiniy Uplift (or Kilian-Zmeinian Zone of structural highs) is located north of the Gubkin Ridge and is separated from it by the Sulina Depression (Fig. 3b). The Zmeiniy Uplift most likely developed in a way analogous to the Gubkin Ridge such that before its inversion it also constituted part of the western Karkinit-Gubkin rift basin. Lower Palaeozoic sediments are encountered in the Morskaya-1 well on Zmeiny Island at the top of the uplift (Figs 1 & 3b). The seismic data show that sediments of this age lie beneath a very thin Neogene succession, at least in the vicinity of the island (Fig. 3b). The Sulina Depression

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(Figs 2 & 3c) is also probably a part of the former Karkinit-Gubkin rift basin and, because it is only partially inverted, it still contains a quite thick succession of Mesozoic sediments. The Predobrogea Depression (Fig. 2) is the westernmost segment of the EEP or the SP and is concealed below Cenozoic sediments (e.g. Nikishin et al. 2000; Seghedi 2001); it may be the foreland trough of the Northern Dobrogea zone (e.g. Nikishin et al. 2001).

Main tectonic elements of the Azov Sea The Northern Azov Trough forms the southern part of the EEP in the north and the northern part of the SP with a truncated Cretaceous and Palaeogene sedimentary cover up to 3000 m in thickness (Fig. 4). It is separated from Azov Ridge by the Azov Fault. The Azov Ridge has been generated by the uplift of basement, presumably of Triassic –Jurassic age, and is covered by a thin succession of Oligocene – Pliocene deposits (Fig. 4). A comparison between the structure and formation of the Gubkin Ridge and Azov Ridge suggests that the same compressional events caused formation of the both structures that originally had formed hanging walls of rift (half)grabens. The Indolo-Kuban Depression was formed in the Oligocene –Miocene simultaneously with the uplift of the Greater Caucasus and southern Crimea regions (Fig. 4, Denega et al. 1998). It is a foreland basin (Ershov et al. 1999, 2003).

Age of tectonic events The set of regional seismic reflection profiles from the Odessa Shelf, calibrated with wells (see Fig. 1 for locations), of which those shown in Figures 3 and 4 serve as examples, has allowed three stages of tectonic evolution to be determined (Fig. 5). The data at hand do not allow an independent specification for the Azov Sea; provisionally, it can be assumed that they are, given the proximity of the Azov Sea to the Odessa Shelf and the similarities in their respective tectonic settings, the same. The first tectonic stage is one of active rifting from the Early Cretaceous until the beginning of the Late Cretaceous. Extensional processes were probably active since the Albian, and the seismic data reveal they ceased by the Campanian. This is supported by the presence of volcanic rocks in the Albian and lower part of the Upper Cretaceous successions. Figure 3a demonstrates how pre-rift sediments were disrupted into extensional fault blocks leading to the development of a system of syn-rift grabens and half-grabens, mainly with southdipping normal faults. The so-called Shtormovaya

Fig. 5. Main stages of the tectonic evolution of the Odessa Shelf, determined from the seismic interpretations. N31-N2-Q, Upper Miocene –Pliocene– Quaternary; N21, Middle Miocene; P3-N11, Oligocene– Lower Miocene; P32, Upper Eocene; P21-P1-2 2 , Upper Paleocene– Lower-Middle Eocene; P11, Lower Paleocene; K2cp-m, Campanian– Maastrichtian (Upper Cretaceous); K2cm-t-cn-st, Cenomanian –Turonian– Coniacian– Santonian (Upper Cretaceous); K1, Lower Cretaceous; T-J, Triassic –Jurassic.

‘half-graben’ (Robinson & Kerusov 1997; Dinu et al. 2005) is a good example of this kind of extension structure. During the rifting stage the half-graben developed along the major SulinaTarkhankut Fault, which is fixed in seismic sections along the axis of the Karkinit Trough (Fig. 3a). The most widespread faulting, accompanied by volcanic activity, occurred in the Karkinit-Gubkin rift (Figs 2 & 3a). This structure is separated from the southern margin of the EEP by the regional Golitsin Fault (Fig. 2). The bottom of the pre-rift and syn-rift sediments can be recognized in seismic sections almost everywhere except when they are completely absent (Fig. 3). The uplifted footwalls of some of the half-grabens remained above sea level and were eroded during almost the whole period of active rifting (Fig. 6a). Sediments accumulated mainly in the subsided parts of these halfgrabens (Fig. 3). Sedimentation on the footwalls of most of these structures began only at the end of Albian. Extension occurred in north –south orientation given that the extensional structures are aligned roughly east –west. Rifting and formation of (half)grabens also occurred in the areas of the Gubkin Ridge and Krayova Step (Figs 3b & 3c). This is clearly indicated by the seismic data showing a thickening of syn-rift sediments from the Krayova Step towards the Gubkin Ridge (Fig. 3c). The observed stratigraphic geometries are typical

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Fig. 6. Schematic cross-sections illustrating the evolution of the Karkinit Trough, Odessa Shelf: (a) rift stage; (b) post-rift passive thermal subsidence stage. K1-K2st, Lower Cretaceous–Santonian (Upper Cretaceous); K2cp-m, Campanian– Maastrichtian (Upper Cretaceous); P1-P1-2 2 , Paleocene–Lower-middle Eocene.

for a half-graben. It is evident that Sulina-Tarhankut and Golitsin Faults were normal faults during the rift stage, but were subsequently reactivated as reverse faults, as discussed below. Thus, before basin inversion, the Gubkin graben (Fig. 3c) as well as the areas now occupied by the Zmeiniy Uplift and Sulina Depression (Fig. 3b) represented the western prolongation of what has here been called the Karkinit-Gubkin rift. The second tectonic stage was one of post-rift thermal subsidence from the middle Late Cretaceous (Campanian) until the middle Eocene. The greatest thickness of Upper Cretaceous sediments occurs within the axial part of the Karkinit-Gubkin rift basin and gradually decreases towards its flanks (Figs 3a & 6b). Within the present-day Kalamit Ridge and Krayova Step, the thickness of Cretaceous and Palaeogene sediments is insignificant in comparison with the central part of the Karkinit Trough because of their location on the southern flank of the joint Karkinit-Gubkin post-rift basin. The third tectonic stage was one of basin inversion. Four distinct compressional phases can be identified: at the end of the middle Eocene, at the end of the Late Eocene, at the end of the Early Miocene and at the end of the middle Miocene

(Stovba & Khriachtchevskaia 2009). Compression was, like the earlier extensional stress regime, northdirected, resulting in inversion movements along the pre-existing, major extensional faults and caused the formation of high amplitude positive structures (Fig. 7). Similar folds have also been documented in the adjacent Romanian offshore region (Morosanu 2002). The Shtormovaya ‘halfgraben’ mentioned above was affected at this time and is at present an inversion structure. The timing of formation of the structure shown in Figure 7 is constrained by the observation that the Upper Eocene sediments on its limbs have an onlapping geometry and that Middle–Lower Eocene strata are absent as a result of erosion on the top of the developing anticlines. The strongest compressional events occurred at the end of Late Eocene, at the end of Early and Middle Miocene as can be recognized from the same seismic section. The most profound deformations occurred in the area of the present-day Gubkin Ridge and Zmeiniy Uplift (Fig. 3b, c) where the vertical component of the inversion displacement could be more than 3 km. Most of the Cretaceous –Miocene succession has been eroded in these areas as a result. The geometry of inversion structures as documented on the Odessa Shelf is compatible with the

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Fig. 7. Interpreted seismic profiles showing inversion structures on the Odessa Shelf, located in Figure 2: (a) Gamburceva structure; (b) Zahidnokreydyana structure.

conceptual model of their formation proposed by Cooper et al. (1989).

Tectonic evolution of the Odessa Shelf and Azov Sea from 1-D subsidence analysis The tectonic subsidence of the Odessa Shelf and Azov Sea has been calculated based on standard techniques (e.g. Sclater & Christie 1980). ‘Backstripping’, as such, allows the ‘air-loaded’ position of the basement surface through time to be estimated from the stratigraphy (and formation lithologies) by correcting for the effects of sediment compaction and the local isostatic response to sediment loading. The distribution of wells and pseudo-wells used in the backstripping (tectonic subsidence) analysis is shown in Figure 1. There are 49 in total, including nine wells from the Odessa Shelf, 14 pseudo-wells created from depth-converted seismic data on the Odessa Shelf, 10 wells from the Azov Sea, and 16

onshore near coastal wells. Lithological parameters for the pseudo-wells were obtained from the nearby boreholes. Further, three wells (Arh-1, Bez-2 and Des-1; see Fig. 3a) were artificially deepened through to Late Cretaceous based on correlation with regional seismic lines. Table 1 provides more information about the 49 wells and pseudowells that were used, including the depth and age of the deepest strata penetrated. On the Odessa Shelf, the oldest strata encountered – but never fully penetrated – was Barremian. The deepest well (stratigraphically) used in this study is Olimpiyska-400 (see location in Figs 1 & 2), which ends at 3012 m in Aptian strata. A large number of stratigraphic horizons can be correlated throughout the study area though, as mentioned earlier, the correlation between the Odessa Shelf and Azov Sea is incomplete. Absolute stratigraphic age was derived by applying the timescale of Harland et al. (1990) to the local stratigraphic scheme (Table 2). Regarding palaeobathymetry,

CRETACEOUS–NEOGENE TECTONIC EVOLUTION OF THE BLACK SEA

sedimentary facies and faunal studies suggest that depositional water depth never exceeded 200 m, with much shallower water depth generally prevailing. In the backstripping, an average palaeobathymetry according to the lithological composition of sequences was used. Errors in estimating water depth in the backstripping analysis are likely to be smaller than 50 m.

Odessa Shelf Typical air-loaded tectonic subsidence curves from the Odessa Shelf, resulting from the backstripping calculations, are shown in Figure 8a. One example from each of the tectonic units delineated in Figure 2 is plotted. For any given tectonic unit, all subsidence curves among the total of 28 that were computed, closely conform to one another; there are no notable anomalies. Also shown in Figure 8 are the three tectonic phases identified from the seismic data: syn-rift, post-rift and inversion. In general, there is a close correspondence between the subsidence curves and the tectonic phases identified from the seismic data (Fig. 5). Tectonic subsidence is rapid during the syn-rift phase, slows down during the post-rift phase, and accelerates (or is punctuated) during the inversion phase. A partial exception is Olimpiyska-400, the only well in which Aptian and Albian sediments have been penetrated, which indicates quiet tectonic conditions during the syn-rift phase. It is likely that the Karkinit-Gubkin rift already existed, however, as shown above with the Krayova Step (where the Olimpiyska-400 well is located) as part of its southern boundary. In order to illustrate the regional evolution of the Odessa Shelf, a series of six maps of tectonic subsidence rate was constructed (Fig. 9): one for the Early Cretaceous to early Late Cretaceous rifting phase, one for the thermal subsidence phase beginning in the Late Cretaceous and four for the inversion phase. These correspond to the compressional events identified on the seismic profiles (viz., end of the Middle Eocene, c. 38.6 Ma; Late Eocene, c. 35.4 Ma; Early Miocene, c. 16.3 Ma; and Middle Miocene, c. 10.4 Ma). The syn-rift phase is characterized by rapid tectonic subsidence rates (Figs 8a & 9a), with maxima of 43–52 m/Ma occurring along the SulinaTarhankut Fault. Away from the fault they decrease to 10 m/Ma and less. It is likely that this fault is located near the axis of the Karkinit-Gubkin rift zone whereas the Evpatoriyska-2 and Olimpiyska wells, where subsidence rates are less, probably were closer to the southern flank of the rift. The age of the onset of rifting is not defined by the subsidence data given the lack of pre-Late Cretaceous data. It is likely, however, that what is seen is the

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prolongation of a rifting stage that started in the middle of Early Cretaceous related to the opening of the western Black Sea Basin in the Cenomanian and Coniacian (Nikishin et al. 2003). Tectonic subsidence rates during the Late Cretaceous – Middle Eocene phase are reduced compared to the syn-rift phase, as would be expected with post-rift thermal subsidence (Figs 8a & 9b). The average rate of tectonic subsidence at this time was about 5–10 m/Ma. Maximum subsidence rates are observed along the axis of present-day Karkinit Trough with the fastest observed value being 13 m/Ma at the Arhangelskaya-1 well. The Karkinit Trough and Gubkin Ridge developed jointly during this tectonic phase, which lasted 44 Ma, forming the single tectonic unit referred to earlier as the Karkinit-Gubkin rift. The first part of the inversion phase (Late Eocene) is very well defined on all of the tectonic subsidence curves (Figs 8a & 9c). Most of these indicate a drastic increase in subsidence rate at this time, to an average of about 30 m/Ma. The maximum tectonic subsidence rates (up to 43 m/Ma) are mainly on the most buried part of the present-day Karkinit Trough in close proximity to the Sulina-Tarhankut and Golitsin faults. The acceleration of subsidence within the axis of the Karkinit Trough is related to the onset of compression and the first phase of formation of inverted structures (basin inversion). The tectonic subsidence curves document an uplift within the present-day areas of the Krayova Step, the Kalamit Ridge (Olimp-400, Des-1, Il-2 wells) and the southern part of the onshore North-Crimea Trough (Gleb-1, Sereb-2 wells). It is inferred that compression at this time led to the uplift of the whole of the southern part of Odessa Shelf resulting in the cessation of sedimentation and the onset of erosion within the uplifted areas of the Kalamit Ridge and Krayova Step (Fig. 3) from the end of the Late Eocene through to the beginning of the Oligocene. This erosional event was widespread and is also noted to the SW within the Romanian offshore region (Tambrea et al. 2002). The second part of the inversion phase (Oligocene –Early Miocene) is associated with low tectonic subsidence rates and little differentiation within the study area of the Odessa Shelf (Figs 8a & 9d). It is interpreted to represent, at least in part, the continuation of post-rift thermal subsidence that began in the Late Cretaceous but was interrupted by the previous compressional phase. The average tectonic subsidence rate at that time was 5 –10 m/Ma. It was a time of isolation of the present-day Gubkin Ridge, Sulina Depression and Zmeiniy Uplift. These areas exhibit an insignificant subsidence (2– 3 m/Ma) in contrast to the Karkinit Trough with tectonic subsidence rates of

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Fig. 8. Air-loaded tectonic subsidence curves, typical for regional tectonic units (Fig. 2) of (a) the Odessa Shelf; (b) the Azov Sea, interpreted for syn-rift, post-rift and inversion stages of tectonic evolution (Fig. 5). N2-Q, Pliocene– 2 1 Quaternary; N2-3 1 , Middle– Upper Miocene; N1-Q, Middle Miocene– Quaternary; P3-N1, Oligocene–Lower Miocene;

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Fig. 9. Maps of tectonic subsidence rates, based on backstripping of the wells as shown, for the Odessa Shelf during: (a) Cenomanian– Santonian (syn-rift); (b) Campanian– Middle Eocene (post-rift); (c) Late Eocene (beginning of inversion); (d) Oligocene–Early Miocene (inversion); (e) Middle Miocene (inversion); (f) Late Miocene –Quaternary (post-inversion).

Fig. 8. (Continued) P1-2, Paleocene–Eocene; P32, Upper Eocene; P22, middle Eocene; P12, Lower Eocene; P1, Paleocene; K2, Upper Cretaceous; K2cp-m, Campanian– Maastrichtian (Upper Cretaceous); K2cm-st, Cenomanian –Santonian (Upper Cretaceous); K1, Lower Cretaceous; K1al, Albian (Lower Cretaceous); K1ap, Aptian (Lower Cretaceous). Wells are located in Figures 1 and 2 (underlined) and in the lower panel.

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8–12 m/Ma. The thick Maykopian succession was deposited during this phase within the Karkinit Trough and on the southern slope of the Gubkin Ridge. A considerable part, and in some most uplifted areas even the whole section, of Eocene, Paleocene and Cretaceous were eroded within the Gubkin Ridge, Krayova Step, Sulina Depression and Zmeiniy Uplift. The third part of the inversion phase (middle Miocene) is associated with possible uplift of some areas of the Odessa Shelf (Figs 8a & 9e), including the Tarhankut Peninsula, Kalamit Ridge, Odessa High and the northern part of the Odessa Shelf. Only the SW part of the Odessa Shelf displays weak tectonic subsidence (up to 12 m/Ma), where intensive sedimentation took place since the beginning of middle Miocene (Fig. 3c). Finally, tectonic subsidence rates decline during the Late Miocene –Quaternary time interval (Figs 8a & 9f), with an average of 2–5 m/Ma, increasing southward up to 10 m/Ma within Krayova Step. At the same time some uplifted areas are observed: the Crimea Peninsula and Odessa Shelf area.

Azov Sea Typical air-loaded tectonic subsidence curves from the Azov Sea, resulting from the backstripping calculations, are shown in Figure 8b. One example from each of the tectonic units delineated in Figure 2 is plotted. For any given tectonic unit, as for the Odessa Shelf, all subsidence curves among the total of 21 that were computed, closely conform to one another; there are no notable anomalies. And, similarly, there is generally a close correspondence between the subsidence curves and the tectonic phases identified from the seismic data although the reduced stratigraphic resolution in some wells makes this less obvious than for the Odessa Shelf. The regional evolution of the Azov Sea basin is summarized in a series of five maps of tectonic subsidence rate (Fig. 10): one for the rifting phase, one straddling the break between rifting and post-rift thermal subsidence phases and three for the thermal subsidence and inversion phases, one less than for the Odessa Shelf because of the more limited stratigraphic resolution. In contrast to the Odessa Shelf, the character of syn-rift subsidence during the Albian is available for the Azov Sea because from ten wells located in its northwestern and northern segments (Figs 8b & 10a). The highest tectonic subsidence rates at this time (.20 m/Ma) are observed in the western part of the Azov Sea. It is possible that this area represents the eastern continuation of the Karkinit-Gubkin rift as defined on the Odessa

Shelf. The lowest subsidence rates are observed in the northern and northeastern parts of the Azov Sea, indicating that this area was probably marginal to the Early Cretaceous rift basin. There is a general lack of detailed chronostratigraphic constraints for the Azov Sea during its post-rift thermal subsidence phase, with the whole of the Late Cretaceous (which probably includes some of the syn-rift period as well) as a single time slice in most of the wells (Figs 8b & 10b). The one exception is the Matroska-1 well, which demonstrates an accelerated subsidence during the time interval 97– 83 Ma. This suggests that the evolution of the Azov Sea may correspond to that of the Odessa Shelf, with the active syn-rift stage of evolution terminating at this time. All of the other wells suggest an average Late Cretaceous tectonic subsidence rate that is quite low, increasing moderately to the SE in the Azov Sea. The highest values (12 m/Ma) are observed in the Tamanskaya-5 well (Fig. 10b). Subsidence rates during the Paleocene – Late Eocene part of the thermal subsidence phase are mapped separately (Fig. 10c) because of the increased number of datapoints available at this time. Subsidence rates during this time remain slow, with an average rate of tectonic subsidence of 3–5 m/Ma (Fig. 10c). However, there are two local areas with relatively higher rates (7–8 m/Ma), seen in the southwestern and eastern parts of the Azov Sea. There is a sharp increase in tectonic subsidence rate at the beginning of the Oligocene, through the Early Miocene (Figs 8b & 10d) with values up to 37 m/Ma, towards the southern part of the Azov Sea. Rates in the northern part of the Azov Sea are much less, ,5 m/Ma. The accelerated tectonic subsidence of the southern part of the Azov Sea at this time is interpreted as being related to the onset of formation of the Indolo-Kuban foreland basin of the Greater Caucasus. The thick Maykopian succession (up to 3000 m) was deposited during this time. In the subsequent time slice (Middle Miocene –Quaternary), rapid subsidence rates were maintained (Figs 8b & 10e), with rates .20 m/Ma in the southeastern part of the Azov Sea area. Elsewhere, they were relatively low (5–10 m/Ma).

Discussion Timing of rifting in the Black Sea The seismic interpretations and tectonic subsidence curves presented in this paper provide robust constraints on the timing of rifting and other tectonic events on the northern margin of the Black Sea. Based on a history of studies on the margins of

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Fig. 10. Maps of tectonic subsidence rates, based on backstripping of the wells as shown, for the Azov Sea during: (a) Albian (pre- to syn-rift?) (b) Late Cretaceous (syn- to post-rift); (c) Paleocene– Eocene (post-rift); (d) Oligocene– Early Miocene (inversion); (e) Middle Miocene– Quaternary (inversion to greater Caucasus foreland).

other basins, this can also be taken as direct evidence about the timing of rifting in the main rift structures of the deeper basins of the Black Sea. Since there are few to no reliable ages associated with the deeply buried stratigraphy of the deeper basins of the Black Sea, this means that the present results form an important basis for defining the timing of rifting in the Black Sea, which has been a matter of some debate. Observations from the Odessa Shelf, where stratigraphy has been newly recorrelated and where seismic interpretations are well constrained, show

quite unequivocally that active rifting, accompanied by volcanism, was underway in the Early Cretaceous (Aptian–Albian) and continued until the end of the Santonian in the Late Cretaceous (c. 128– 83 Ma; Fig. 8). Tectonic subsidence rates in the NE–SW oriented Karkinit-Gubkin rift basin were as high as 30– 50 m/Ma during this time (Fig. 9). This structure formed in the vicinity of the boundary between the East European and Scythian platforms and it involved the present-day Gubkin Ridge to the west and continued to the east across the northern Crimea Peninsula into the

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Azov Sea (Fig. 2). Horsts, grabens and half-grabens were generated during rifting and the seismic reflection data show that syn-rift erosion occurred on some uplifted footwall blocks (Fig. 3). Constraints on the timing of rifting in the Azov Sea are much more ambiguous than for the Odessa Shelf but the available evidence (the seismo-geological profile, Fig. 4, and the tectonic subsidence curves, Fig. 8) are consistent with synchronous syn-rift activity in the Early to Late Cretaceous. Contiguous areas reportedly show the same pattern. The southwestern Crimea Peninsula is characterized by rapid Early Cretaceous subsidence (Nikishin et al. 1998, 2001) as is the Indolo-Kuban Depression in the southern Azov Sea and Greater Caucasus foreland (Fig. 2), particularly during Aptian –Albian times when it was accompanied by basaltic and andesitic volcanism (Bolotov 2006). The time of cessation of active rifting and onset of post-rift thermal subsidence is unambiguous in the Odessa Shelf seismic and subsidence data. It occurred at the beginning of the Campanian (Late Cretaceous) and passive post-rift subsidence continued uninterrupted until the end of the middle Eocene (83–38.6 Ma). The Karkinit-Gubkin trough clearly developed as a single post-rift sag basin during this time (Fig. 3). Tectonic subsidence rates (5–10 m/Ma, Fig. 9) are consistent with a passive thermal subsidence mechanism and accommodated the deposition of a thick sequence (up to 2000 m) of limestones and marls in the KarkinitGubkin post-rift basin during this phase of basin development. A similar regime has been described for the Romanian shelf (Morosanu 2002). The backstripping data for the Azov Sea (Figs 8 & 10) show that the sedimentary basin in this area, on the northern margin of the eastern Black Sea, was likely passively subsiding during at least part of the Late Cretaceous but most certainly during much of the Palaeogene. Although differences in stratigraphic correlation and resolution between the Odessa Shelf and Azov Sea make a precise comparison difficult, the results suggest that both areas were subject to the same tectonic regimes of (1) active (syn)rifting from sometime in the Early Cretaceous (probably Aptian –Albian) until the end of the Santonian; and (2) passive post-rift thermal subsidence from the Campanian throughout much of the Palaeogene. There is no evidence to suggest, in subsidence patterns or rates, that any part of the Azov Sea area was subject to active rifting during the Paleocene. Given the location of the Odessa Shelf on the margin of the western Black Sea basin and the Azov Sea on the subsequently deformed (with the development of the Greater Caucasus-Crimea mountain belt) margin of the eastern Black Sea basin, it seems reasonable to suggest that the timing of rifting in the main Black Sea basins also

occurred, more or less synchronously, from the Early Cretaceous until the end of the Santonian, followed by passive thermal subsidence until basin inversion processes intervened.

Timing of inversion in the Black Sea Inversion of the Karkinit-Gubkin rift basin on the Odessa Shelf and its western prolongation or associated basin in the Azov Sea occurred abruptly at the end of the Middle Eocene. Vertical movements associated with inversion would have been superimposed upon the effects of ongoing passive thermal subsidence derived from the earlier Cretaceous rifting phase. The seismic reflection data from the Odessa Shelf show that there were several discrete phases of compression: the initial one (at the end of the Middle Eocene; c. 38.6 Ma) followed by later phases in the Late Eocene, Early Miocene and middle Miocene (c. 35.4 Ma, c. 16.3 Ma, c. 10.4 Ma respectively). These inversion events caused uplift of parts of the former basin. The Late Eocene event produced uplift of the southern part of the Odessa Shelf (Kalamit Ridge and Krayova Step) and the Azov Sea (Azov Ridge), resulting in widespread erosion, down to the upper part of Upper Cretaceous succession, in these areas. The Early Miocene event led to uplift and widespread erosion in the western part of Odessa Shelf (Fig. 3). Initially extensional faults such as the Golitsin and SulinaTarhankut faults were inverted during these events, so that they now display reverse kinematic relationships (Fig. 3) and control inversion structures as seen on the Odessa Shelf (Fig. 7). Similar inversion structures occurred within the Romanian offshore region (Morosanu 2002). Rapid subsidence of the southern part of the Azov Sea occurred during Oligocene– Early Miocene times and, thereafter, even more rapid subsidence of its southeasternmost segment is observed (Fig. 10). This subsidence is linked with the development of the Indolo-Kuban Basin, in the foreland of the western Greater Caucasus. The timing of compressional events constrained in this study, from reflection seismic data and subsidence analysis on the Odessa Shelf and in the Azov Sea, is compatible with those determined regionally, which are generally ascribed to Eurasian – Arabian plate convergence and collision (Dercourt et al. 1993; Robinson et al. 1995; Nikishin et al. 2003). The resolution afforded by the seismic data on the Odessa Shelf, however, lead to a much more detailed definition of the timing and magnitude of compressional (basin inversion) events there than available regionally and suggest that inversion ended, or at least was suspended, by the Late Miocene. In contrast, both Nikishin et al. (2003) and Robinson et al. (1995) inferred, from

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(more poorly constrained) stratigraphic data, a period of rapid tectonic subsidence starting in the Pliocene or Late Miocene (respectively) and continuing until the present. No such event is documented in the present results, from the northern margin of the Black Sea, where only insignificant tectonic subsidence is observed (excluding the IndoloKuban Basin. Figs 9 & 10). This implies that Late Miocene and younger rapid sedimentation in the deeper Black Sea is not strongly controlled by regional tectonic processes but could be related to other processes such as sediment supply and transport mechanisms.

Summary and conclusions Industry seismic reflection and well data from the Odessa Shelf and the Azov Sea, on the northern margin of the Black Sea, have been interpreted and analysed to elucidate the timing and style of tectonic events shaping the architecture of the Black Sea. Age interpretations have been facilitated by revised stratigraphic correlations newly available for the sedimentary succession of the Odessa Shelf and structural interpretations by a large set of recently acquired seismic reflection data. The results show clearly that extension (rifting) on the Odessa Shelf was active in the Early Cretaceous (by Aptian –Albian times) and continued until the end of the Santonian in the Late Cretaceous (c. 128 –83 Ma) when it stopped. The NE –SW orientated Karkinit-Gubkin rift basin, comprising the present-day Karkinit Trough as well as other structures that were tectonically inverted since Cretaceous rifting, formed at this time near or on the boundary between the Eastern European and Scythian platforms. The uplifted footwalls of associated half-grabens were exposed above sea level and subject to erosion during syn-rift tectonics. The onset of rifting is poorly to not resolved by the Azov Sea data, which are, however, wholly consistent with the Aptian –Albian onset observed on the Odessa Shelf. The cessation of active rifting in the Azov Sea is also not as clearly displayed as on the Odessa Shelf but the available observations suggest a similar, end of Santonian, date. There is no evidence in the documented observations to suggest any active tectonic process, in particular active extension (rifting) on the Odessa Shelf or in the Azov Sea from the end of the Santonian (beginning of Campanian) until the Middle Eocene. Rather, basin evolution during this time is consistent with passive thermal subsidence, occurring in a post-rift setting subsequent to the earlier, documented Cretaceous rifting stage. Given the proximity of the study areas to the western and eastern Black Sea basins, it is concluded that the latter are also unlikely to have been affected by active processes during this time.

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The present results lend support to a model of Black Sea formation in which the opening of its western and eastern segments was essentially synchronous, initiated and terminated in the Cretaceous (Aptian –Albian up to and including the Santonian). Typical basin inversion processes began to affect the initially extensional structures on the Odessa Shelf in the late Middle Eocene. These occurred several discrete phases: the initial one (at the end of the Middle Eocene; c. 38.6 Ma) followed by later phases in the Late Eocene, Early Miocene and Middle Miocene (c. 35.4 Ma, c. 16.3 Ma and c. 10.4 Ma respectively). The Late Eocene event was most profound and led to widespread erosion on the Odessa Shelf. No active inversion on the Odessa Shelf is observed since the Middle Miocene event. Rapid post-Miocene subsidence in the southeastern Azov Sea is related to the formation of the Indolo-Kuban Basin, in the foreland of the western Greater Caucasus. This research was funded through the Ukrainian national scientific project by NAK ‘Naftogaz of Ukraine’ and the MEBE project. Thorough reviews of an earlier version of the manuscript by M.-F. Brunet (Paris) and S. Egan (Keele) are much appreciated. The School of Geosciences of the University of Aberdeen is thanked for facilitating a visit by O. Khriachtchevskaia and S. Stovba to prepare the (near) final version of the manuscript.

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CRETACEOUS–NEOGENE TECTONIC EVOLUTION OF THE BLACK SEA Nikishin, A. M., Seghedi, A., Bolotov, S. & Stephenson, R. A. 2000. Crimea and Dobrogea: a comparison of their Mesozoic geological histories. Geophysical Journal, 4, 114– 116 (in Russian). Nikishin, A., Ziegler, P. A. et al. 2001. Mesozoic and Cenozoic evolution of the Scythian Platform – Black Sea – Caucasus domain. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & CrasquinSoleau, S. (eds) Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basins and Passive Margins. Me´moires du Muse´um National d’Histoire Naturelle, Paris, 186, 295–346. Okay, A. I., Sengo¨r, A. M. C. & Go¨ru¨r, N. 1994. Kinematic history of the opening of the Black Sea and its effect on the surrounding regions. Geology, 22, 267–270. Robinson, A. G. 1997. Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68. Robinson, A. G. & Kerusov, E. 1997. Stratigraphic and structural development of the Gulf of Odessa, Ukrainian Black Sea: implications for petroleum explorations. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 369– 380. Robinson, A. G., Rudat, J. H., Banks, C. J. & Wiles, R. L. F. 1996. Petroleum geology of the Black Sea. Marine and Petroleum Geology, 13, 195– 223. Robinson, A. G., Spadini, G. & Cloetingh, S. 1995. Stratigraphic evolution of the Black Sea: inferences from basin modelling. Marine and Petroleum Geology, 12, 821– 836. Saintot, A., Stephenson, R. A., Stovba, S. M., Brunet, M.-F., Yegorova, T. & Starostenko, V. 2006. The evolution of the southern margin of Eastern Europe (Eastern European and Scythian platforms) from the latest Precambrian-Early Paleozoic to the Early Cretaceous. In: Gee, D. G. & Stephenson, R. A. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 481–505. Sclater, J. G. & Christie, P. A. E. 1980. Continental stretching: an explanation of the post-mid-Cretaceous subsidence of the central North Sea Basin. Journal of Geophysical Research, 85, 3711–3739. Seghedi, A. 2001. The North Dobrogea orogenic belt (Romania): a review. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & Crasquin-Soleau, S. (eds) Peri-Tethys Memoir 6: Peri-Tethyan Rift/ Wrench Basins and Passive Margins. Me´moires du Muse´um National d’Histoire Naturelle, Paris, 186, 237–257. Sengo¨r, A. M. C. 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia. Geological Society of America, Special Paper, 195, 81. Shillington, D. J., White, N., Minshull, T. A., Edwards, G. R. H., Jones, S. M., Edwards, R. A. & Scott, C. L. 2008. Cenozoic evolution of the eastern Black Sea: a test of depth-dependent stretching models. Earth and Planetary Science Letters, 265, 360–378. Slavin, V. I. & Khain, V. E. 1980. Early Cimmerian geosynclinal trough of north of central part Mediterranean

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Seismic sequence stratigraphy, structure and subsidence history of the Romanian Black Sea shelf CATALINA KONERDING1,2, CORNELIU DINU3* & HOW KIN WONG1 1

Institute of Biogeochemistry and Marine Chemistry, Hamburg University, Germany 2

Present address: G&W Systems, Corp., Houston, USA

3

Faculty of Geology and Geophysics, Bucharest University, Romania

*Corresponding author (e-mail: [email protected], [email protected]) Abstract: The Black Sea is an extensional back-arc basin developed along the northern active margin of the Tethys Ocean which was subducted northward from the Triassic to Miocene times. The Romanian Black Sea shelf is dominated by mid-Cretaceous extensional structures and their sedimentary cover, subsequently affected by Cenozoic compression. Here we analyse the post-Oligocene structural and sedimentological evolution of the shelf, based on Romanian oil industry data consisting of (1) 5300 line-km reflection seismic profiles situated on the shelf and continental slope; and (2) depth and lithostratigraphic information from 60 boreholes on the shelf. Our study provides evidence for a changing evolution of the shelf during a relatively short period of time directly related to the pre-Miocene period and the evolution of the Romanian onshore. Mio-Pleistocene subsidence of the Romanian Black Sea shelf is highly variable and is directly dependent on sediment input, tectonic activity as well as water-level fluctuations. Subsidence during the Badenian– Sarmatian and the Dacian– Quaternary was limited. In contrast, during the Pontian, shelf subsidence was progressively faster in the basinward direction. Subsidence on the outer shelf was much more significant than elsewhere on the shelf. Tectonically, the most active period during the Mio-Pleistocene was the Pontian. The Badenian–Sarmatian was largely quiescent and the Dacian–Quaternary saw a decrease in the Pontian tectonic activity, coming possibly even to a halt. From a sequence-stratigraphical point of view, eight systems tracts were identified for the Mio-Pleistocene sedimentary section. The Badenian-Sarmatian unit was attributed to a HST (highstand systems tract). The Pontian unit was subdivided into P1, P2, P3 and P4. Subunit P1, which was laid down on the slope at the time of deposition, is progradational and attributable to the lowstand wedge of a LST (lowstand systems tract). Subunit P2 is likewise also attributed to a LST, having the continental slope and the deep basin as palaeo-depositional environments. The reflection terminations and the wedge-shape of P3 suggest that it was deposited in the deep basin during a sea level lowstand. The next transgressive systems tract (TST) and HST developed during the deposition of P4. The boundary between P4 and the Dacian is represented by an erosional hiatus, which comprises the LST that follows the formation of the sequence boundary at the end of the Pontian. During the Dacian–Quaternary, the subsequent TST and HST were deposited on the inner and middle shelves. Sedimentation on the Romanian shelf during the Mio-Pleistocene period was thus strongly influenced by sediment input and subsidence, while sea level fluctuations played a lesser role. As sediment input is related to the evolution of the adjacent land and subsidence is dependent on sediment supply, tectonic activity and sea level fluctuations, these two factors are not totally independent.

The Black Sea is located between Romania, the Ukraine, Russia, Georgia, Turkey and Bulgaria. It is one of the largest enclosed marine seas with an area of 423 000 km2, a volume of 534 000 km3 and a maximum water depth of 2206 m (Ross et al. 1978). The Black Sea basin came into existence about 200 million years ago and over 15 km of sediment have accumulated in some parts (Neprochnov & Ross 1978). It is an extensional back-arc basin developed along the northern active margin of the Tethys Ocean which was subducted northward from the Triassic to the Miocene. It consists of two parts: the western Black Sea which is underlain by oceanic to suboceanic crust, and the

eastern Black Sea underlain by continental crust. The two basins are separated by a strike –slip fault system along the Mid–Black Sea/Andrusov Ridge, which comprises continental crust. The Romanian sector of the Black Sea lies to the east and SE of the Danube Delta and has an area of over 35 000 km2. The Romanian shelf with an area of 20 000 km2 is up to 140 km wide and usually has a water depth of less than 120– 140 m (Popescu et al. 2004). This area is dominated by mid-Cretaceous extensional structures and their sedimentary cover, and was subsequently affected by Cenozoic compression (Robinson et al. 1996).

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 159–180. DOI: 10.1144/SP340.9 0305-8719/10/$15.00 # The Geological Society of London 2010.

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The present paper is based on reflection seismic profiles and borehole data made available by the Romanian oil industry (Fig. 1). The seismic profiles were recorded and processed by the Romanian company Prospect¸iuni S.A. between 1980 and 1994. About 70 profiles situated on the Romanian shelf and on the continental slope with a total length of 5300 line-km were used for interpretation of the Mio-Pleistocene sequences. Velocity information from processing was used for depth conversion. In addition to seismic data, depth and lithostratigraphic information from 60 boreholes drilled on the Romanian shelf were available. Most of these boreholes are located on the inner shelf, particularly on the offshore prolongation of the North Dobrogea Orogen. The purpose of this paper is to map the hitherto unknown distribution and structure of the MioPleistocene succession on the Romanian Black Sea shelf in time slices, as well as to reconstruct (1) the structural evolution of the shelf; and (2) the sequence stratigraphic framework for the Miocene and the post-Miocene in relation to the evolution of the transition zone from the Carpathians to the Black Sea basin.

Structural characteristics of the Romanian shelf Pre-Miocene structures Previous seismic studies show a continuation of the main structural features and major faults in the Dobrogea region (onshore Romania) onto the Romanian Black Sea shelf (Dinu et al. 2002, 2005). This includes the Sf. Gheorghe Fault, which separates the Pre-Dobrogea Depression from the North Dobrogea Orogen, the Peceneaga-Camena Fault, which separates the North Dobrogea Orogen from the Moesian Platform (identified on the shelf for over 60 km) and the Capidava-Ovidiu Fault separating central and south Dobrogea from the Moesian Platform (Fig. 2). The seaward extension of these land structures dips to the east and is covered by Eocene, Oligocene and Neogene deposits. Beginning with the Oligocene, the geological evolution of the Romanian shelf was decoupled from that of the land (Dinu et al. 2002, 2005). According to Moros¸anu (2002), three main tectonic stages can be distinguished on the Romanian shelf.

Fig. 1. Location of the seismic lines and boreholes used in the present paper. Lines 1 and 2 show the idealized sections along which the subsidence history was reconstructed.

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Fig. 2. The major Romanian onshore tectonic structures and their offshore prolongations (from Dinu et al. 2002).

(1)

(2)

(3)

During the first stage, which began in the Jurassic and continued until the Albian, extensional tectonics took place with the opening of a NW– SE oriented rift. This rift developed in direct connection with the opening of the western Black Sea Basin. In the Albian, the Histria Depression and its landward continuation, the Babadag Syncline, were formed (Dinu et al. 2002, 2005; Fig. 2). Extension was very active in the Albian and continued until the Upper Cretaceous with reduced activity. During the second stage that took place from the Upper Cretaceous to the Eocene, rifting ceased and the depositional rate increased. The depression was filled with Upper Cretaceous deposits, Eocene carbonates covering only the highest areas. The third stage developed from the Late Eocene to the Lower Oligocene. It is characterized by inversion tectonics, going from an extensional to a compressional regime.

The Oligocene succession was deposited in a deep, subsiding basin and Mio-Pliocene sediments filled the entire depositional area. Subsidence decreased

in the Miocene and the Pliocene and very thick terrigenous deposits accumulated on the eastern part of the Romanian shelf. Pre-Miocene faulting on the shelf led to: (i) pre-Albian structures, which represent the offshore continuation of the tectonic units that existed in the Dobrogea region before the opening of the Western Black Sea Basin during the Lower Cretaceous; and (ii) extensional structures, directly associated with the opening of the western Black Sea Basin (Dinu et al. 2005). Eocene and Oligocene sedimentation in the Histria Depression was controlled by the Laramian and Pyrenean orogenic phases when inverted structures came into being. The reverse faults on the northern flank of the Histria Depression were presumably also formed during this period (Dinu et al. 2002). An important tectonic element on the Romanian Black Sea shelf is the Histria Depression, the offshore prolongation of the Babadag Syncline, which represents the southern part of the North Dobrogea Orogen (Fig. 2). This depression is bounded by a major structural feature, the ‘Euxinian Threshold’, a zone in which Palaeogene deposits subsided strongly (Fig. 2; Ionescu 2000; Pa˘trut¸

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1975). The Euxinian Threshold was a continental palaeo-slope developed during the Upper Eocene, but may also mark the limit of the shelf deposits during the Upper Cretaceous (Ionescu 2000). The Histria Depression is the subsided block that developed between the central and the northern segments of the Euxinian Threshold. Mio-Pleistocene sedimentation and tectonic processes in the study area are closely related to the development of this depression.

Mio-Pleistocene structures Our study complements previous knowledge on the pre-Miocene structures with new information on the post-Miocene evolution of the Romanian shelf. The Mio-Pleistocene units have been identified on seismic data and chronostratigraphically calibrated using borehole data. From the beginning of the Miocene to the present, these units are: Badenian–Sarmatian, Pontian (subdivided into P1, P2, P3 and P4), Dacian and Romanian–Quaternary (Fig. 3). The Badenian and Sarmatian are undifferentiated on the Romanian shelf because of their limited thickness. Our seismic data suggest that they are present only in two distinct areas: in the southern part of the middle shelf in the area surrounding the Delfin wells, and on the northwestern inner and middle shelf as well as on the palaeo-slope around the Histria Depression (Fig. 4). Borehole data suggest their presence also at the 1 Ovidiu and 12 Midia wells, but the reduced vertical resolution makes their identification on adjacent seismic lines difficult. The thickness distribution of these deposits varies from 0 m at the periphery

to 245 m in the internal part of the northwestern area. In the southern area, the distribution pattern is similar but thickness values of only about 100 m are reached. Pontian deposits have been identified in all wells drilled on the Romanian shelf (Fig. 5). Their thickness varies from 0 m on the inner shelf to 4000 m on the outer shelf in the eastern Coba˘lcescu area. This variation is due primarily to an increase in subsidence of the base of the Pontian to the east, non-uniform sedimentation rates on the shelf, and complex tectonics in the Coba˘lcescu area. The Dacian unit occurs over the entire study area and increases in thickness from the inner to the outer shelf, from values of 0 m in the vicinity of the coast to 1150 m in the east and southeast (Fig. 6). The Romanian–Quaternary section occurs over the entire study area as well. It shows only small variations in thickness from 0 m in the coastal area to 575 m on the outer shelf and the continental slope (Fig. 7). The boundary between the Dacian and the Romanian–Quaternary deposits is conformable. Therefore, their separation was based solely on borehole information. These Mio-Pleistocene units are separated by the following erosional unconformities (Fig. 3): † the BBU, which represents the unconformity at the base of the Badenian –Sarmatian deposits; † the BPU, the unconformity at the base of the Pontian; † IPU 1, the unconformity that separates the P1 and P2 subunits, occurs only locally on the inner shelf; † IPU 2, separating the P2 and P3 subunits; † IPU 3, separating the P3 and P4 subunits; and

Fig. 3. Interpreted seismic line showing the distribution of the Mio-Pleistocene seismo-stratigraphic units: Badenian– Sarmatian, Pontian (1, 2, 3 and 4), Dacian and Romanian–Quaternary, and the main erosional unconformities: the BBU (the unconformity at the base of the Badenian), the BPU (the unconformity at the base of the Pontian), the IPU 1, IPU 2, IPU 3 (intra-Pontian unconformities) and the PDU (Pontian– Dacian unconformity) on the Romanian shelf. TWT, two-way travel.

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Fig. 4. Isopach map of the Badenian-Sarmatian section showing only local faulting during this period. The location of the Histria Depression as well as the positions of the major faults (Peceneaga-Camena and Capidava-Ovidiu) are also given.

† BDU, the unconformity near the top of the Pontian as well as near the base of the Dacian.

Faulting activity and its classification Faulting on the Romanian shelf evolved significantly during the Mio-Pleistocene. It was most active in the Pontian. In contrast, the Badenian –Sarmatian and the Dacian –Quaternary are considerably more quiescent. The Euxinian Threshold separates two distinct zones on the shelf: the northwestern zone which is practically unaffected by tectonic processes, and the southeastern zone which is tectonically very active in connection with the development of the Histria Depression. The Badenian–Sarmatian was a period of tectonic quiescence; only local extensional structures identifiable on our seismic data developed (Fig. 4). The vertical offset of such structures is typically about 350 m. The Pontian was a time of intense tectonic activity on the Romanian shelf. These activities did not affect the entire shelf, but were concentrated in the area basinward of the palaeo-slope (the Histria Depression) as well as in a small depression around

the Delfin wells (Fig. 5). The structures that developed are mostly NE–SW trending grabens, horsts and flower structures (Figs 8 & 9). Faulting that affected the Pontian section can be classified in two ways: temporal and spatial. From the temporal point of view, we distinguish between faults that originated in the pre-Oligocene and continued their activities into the Pontian, faults that developed during the Oligocene–Pontian, and faults that were active only during the Pontian. Spatially, faulting can be related to a NE–SW depression (which we call the ‘Pontian Depression’) with its depocentre linking the 1 Ovidiu and 75 Cobaˇlcescu wells (Fig. 5), and gravitational faulting at the shelf-break. Temporal classification of the Pontian faults. The pre-Oligocene faults are related to the formation of the Histria Depression (Figs 8 & 9). Presumably these faults were reactivated during the Lower Pontian (T¸ambrea et al. 2000) in an extensional regime characterized by NE–SW trending normal faults and grabens because of large sediment loading and rapid subsidence (Fig. 8). Vertical offsets are in the order of a few metres up to about 800 m. They are mostly present in the central and

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Fig. 5. Isopach map of the Pontian deposits showing the general NE– SW pattern of faulting during this period of time. The wells used for calibration are also shown.

northern Histria Depression (around and NE of the well 75 Cobaˇlcescu in Fig. 5), but some of them developed also on the outer shelf (Fig. 9). The second category of faults that started their activity in the Oligocene and continued into the Pontian are gravitational faults that developed in similar sedimentary facies existing both during Oligocene and Pontian times (Figs 8 & 9). The central and northern Histria Depression is an extensional zone characterized by NE–SW trending normal faults and grabens that extend northeastward. The vertical offsets vary from a few metres to almost 1000 m. To the south, extensional faulting with offsets of less than 100 m occurs. Because of the lack of seismic data, however, it is not possible to follow these faults to the NE. Transfer faulting between these faults and the northern Histria Depression has not been observed. The northeastern part of the outermost shelf is characterized by NE–SW oriented negative flower structures, a horst and normal faults over a distance of about 20 km (Fig. 9). The horst is developed along the shelf-break and has a lateral extent of 13 km in the northeast, increasing to 20 km in the SW. Vertical offsets are in the order of a few metres up to about

700 m. A NW– SE oriented transfer fault separates this complex from the southern fault system. On the southeastern and southern outer shelves, NE– SW striking normal faults, grabens and negative flower structures developed locally. They are bounded by transfer normal and thrust faults with a NW –SE trend. The Delfin area is less affected by faulting (Fig. 5); normal faults with vertical offsets of only a few metres have been mapped. The last category consists of gravitational faults that developed only during the Pontian. Some of them were active only in the Lower Pontian in relation to the extensional regime (Fig. 8), whereas others continued their activity in the Upper Pontian and in the Dacian –Quaternary (Fig. 9). In the Lower Pontian, roll-over structures characteristic of an extensional regime formed (Figs 8 & 9). Spatial classification of the Pontian faults. The first category consists of faults related to the development of the Pontian Depression. They strike both NW– SE (parallel to the Histria Depression), and NE–SW (crossing the wells 1 Ovidiu and 75 Cobaˇlcescu; Fig. 5). The NE– SW striking faults were developed during the Pontian extension.

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Fig. 6. Isopach map of the Dacian section showing the faults that were active during this period.

The second class of faults formed along the shelf-break under a gravitational regime. Some of them originated in the Oligocene while others developed during the Pontian. These faults were active till the Quaternary. Their vertical offsets are in the order of a few metres. Basinward of the shelfbreak, the Pontian sediments were deposited in a tectonically quiescent regime. The Dacian –Quaternary section is less faulted than the Pontian sequences. Most of the faults observed here are Oligocene in age or older, and were reactivated during the Pontian (Fig. 9). Vertical offsets are in the order of a few metres to tens of metres. The fault systems trend NE –SW and are separated by NW–SE oriented transfer faults. Only in the west does the fault trend change to east –west.

Mio-Pleistocene subsidence history of the Romanian shelf The subsidence of a passive continental margin is controlled by a number of factors, of which sediment loading and thermal cooling are the most important. Our subsidence analysis is based on

lithostratigraphic information from 11 wells located on the Romanian shelf (Figs 10 & 11). The wells were projected along two lines across the shelf, one striking WNW –ESE and the other NE–SW (Fig. 1). The present-day stratigraphic thickness values were corrected for compaction, palaeobathymetry and absolute water-level fluctuations. The palaeo-water depth was estimated from palaeontological information (T¸ambrea et al. 2000). The occurrence of a relatively rich association of ostracods indicates a brackish-to-fresh water environment with a palaeo-water depth estimated at about 20 m. Information on water-level fluctuations was taken from a water-level curve estimated for the Romanian shelf (Konerding 2006). The unconformities at the base of the Badenian (c. 16–15.8 Ma) and at the base of the Pontian (c. 9–7.3 Ma) were taken into account (Fig. 9). We assume that during the 16– 15.8 Ma and 9 –7.3 Ma time periods only erosion took place on the Romanian Black Sea shelf. The subsidence was computed over the time interval 16 Ma (beginning of the Miocene) to the present. The subsidence history of lines 1 and 2 is shown in Figures 12 and 13 respectively. The history of tectonic subsidence and the history of total subsidence

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Fig. 7. Isopach map of the Romanian–Quaternary section. Note that faulting is predominantly NE– SW.

are similar along lines 1 and 2. From the beginning of the Miocene to the end of the Sarmatian, subsidence was slow. It increased abruptly during the Pontian, slowed down during the Dacian, and

almost came to a halt in the Romanian–Quaternary. Along line 1, the rate of total subsidence during the Badenian– Sarmatian was small (Fig. 12). Subsidence increased from 27 m in the well 814 Leba˘da

Fig. 8. Seismic line crossing the Romanian shelf in a WNW– ESE direction. Note the intense Pontian faulting activity in the Histria Depression. TWT, two way travel time.

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Fig. 9. Seismic line crossing the Romanian outer shelf from NW to SE. TWT, two way travel time.

(on the inner shelf) to 125 m in the well 1 Ovidiu (on the outer shelf). By the end of the Pontian, the total subsidence amounted to 584 m on the inner shelf and 2803 m on the outer shelf, implying an increase of 557 m and 2678 m in these two areas respectively. At the end of the Dacian, the total subsidence reached 733 m on the inner shelf and 3428 m on the outer shelf. This suggests a much smaller increase in subsidence during the Dacian, namely 149 m on the inner shelf and 625 m on the outer shelf. The corresponding values for the Romanian –Quaternary are 175 m (inner shelf) and 242 m (outer shelf) respectively. Thus, the Romanian shelf was largely stable during the Badenian–Sarmatian and the Dacian – Quaternary, but was strongly subsiding during the Pontian. This large increase in subsidence, both in the vertical (for each borehole) and lateral (basinward) directions was a result of intense tectonic activities and high sediment supply during this time period. Subsidence along line 2 (NE–SW; Fig. 13) is significantly lower than along line 1 (WNW–ESE; Fig. 12). Line 2 was stable during the Badenian– Sarmatian, with a total subsidence of 9 m except for the well 18 Lotus, where 11 m of uplift took place. During the Pontian, the total subsidence increased considerably in the southern area of wells 10 Tomis and 6 Delfin. In the Leba˘da and Lotus boreholes, the total subsidence was on average about 400 m higher than during the Badenian–Sarmatian, while for the Tomis and Delfin boreholes, this value is over 1100 m. This

difference is attributed to the different positions of the boreholes on the shelf. Leba˘da and 18 Lotus wells are located on the inner shelf, whereas 10 Tomis and 6 Delfin lie farther basinward where the base of the Pontian began to subside strongly. The Dacian and the Romanian–Quaternary deposits were less affected by subsidence. During the Dacian, the total subsidence decreased from 208 m (at the end of the Pontian) in the well 25 Leba˘da to 150 m in the well 18 Lotus and increased again to 223 m in the well 6 Delfin. The Romanian– Quaternary is marked by very small changes in subsidence, from 148 m in the north to 142 m in the south. We assume that the amount of sediment delivered to the Black Sea during Mio-Pleistocene time was related to uplift of the Carpathians as well as to a possible Messinian drop of the Black Sea level and rapid erosion of the exposed areas (correlated with the intra-Pontian IPU2; Gillet et al. 2007). The distribution pattern of Mio-Pleistocene sediments in the Romanian foreland basin (the Dacic Basin; Jipa 1997; Gillet et al. 2007) correlates negatively with that on the Black Sea shelf. During the Badenian–Sarmatian when uplift of the Carpathians started, a large amount of sediment was delivered to and deposited in the foreland (.5 km; Ta˘ra˘poanca˘ 2004). In contrast, on the Black Sea shelf, Badenian–Sarmatian sediments are present only locally and are very thin (,250 m). Upper Miocene (Meotian) deposits reach a thickness of up to 1.6 km in the foreland basin, while on the

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Fig. 10. Lithological description of the wells projected on line 1, the WNW–ESE profile.

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Fig. 11. Lithological description of the wells projected onto line 2, the NE– SW profile.

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Fig. 12. Subsidence history along line 1 from the Miocene to the present. Zero denotes present-day water level of the Black Sea.

Black Sea shelf they are absent. In the uppermost Miocene (Pontian), a sediment layer of 1.6 km maximum thickness was laid down in the Dacic Basin. A major part of the sediment supplied from the Carpathians, together with erosional products from the subaerial parts of the shelf and slope exposed by a possible Messinian draw-down of the sea level, was transported into the Black Sea, where the Pontian reached a thickness of 4 km in the central Histria Depression. During the Dacian – Quaternary, up to 4.5 km of sediment were deposited in the Dacic Basin, while on the Black Sea shelf, thicknesses of only about 1.5 km were reached. The tectonic subsidence curves, obtained by removing the effect of sediment loading from the total subsidence, show a trend similar to that of the total subsidence curves and suggest that the total subsidence is controlled more by sediment loading than by vertical tectonics. Tectonic subsidence of the Badenian–Sarmatian deposits is observed only in the well 13 Heraclea, where it reached 19 m at the beginning of the Badenian and

28 m at the end of the Sarmatian. In all of the other wells, uplift of a few metres took place. During the Pontian, tectonic subsidence along line 1 (WNW –ESE) increased basinward from an average of 190 m on the inner shelf to 325 m at the well 13 Heraclea and 751 m on the outer shelf. Thus, tectonics contributed significantly to total subsidence on the outer shelf, where intense tectonic activity took place during the Pontian. In the northeastern segment of line 2 (NE– SW), the average tectonic subsidence was 130 m except for the well 18 Lotus, where a value of only 67 m was deduced. In the SW, it increased to 300 m. During the Dacian, the tectonic subsidence increased about 90 m along line 1 as compared to the end of the Pontian, except for the well 1 Ovidiu, where the increase was 205 m. Along line 2 the corresponding increase was 90 –100 m. For the Romanian– Quaternary, the tectonic subsidence amounted to 80 –90 m. To summarize, vertical tectonics had less influence on the Mio-Pleistocene subsidence than

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Fig. 13. Subsidence history along line 2 from the Miocene to the present. Zero denotes present-day water level of the Black Sea.

sediment loading. Only on the outer shelf where intense tectonic activities occurred during the Pontian and in part during the Dacian did vertical tectonics play an important role.

Mio-Pleistocene seismic sequence stratigraphy of the northwestern Black Sea shelf A sea level curve for the Black Sea shelf during the Mio-Pleistocene time span was estimated (Fig. 14). This was done using the global sea level curve (Haq et al. 1987), an unpublished sea-level curve based on seismic data from the northwestern Black Sea shelf (T¸ambrea et al. 2000; Konerding 2006; T¸ambrea, pers. comm.), and salinity information (Jones & Simmons 1997). Because the unpublished sea-level curve for the northwestern Black Sea shelf gives only relative sea levels, it was correlated with the sea-level curve of Haq et al. (1987), to convert it into a curve with absolute values for the

present study. This correlation was carried out by choosing four points on the curve of Haq et al. (1987), during a period when the Black Sea was connected to the global oceans (Fig. 14). To identify these points, information on palaeo-salinity was used, assuming that salinities of 20 –30‰ indicate marine conditions. At the four points chosen, the absolute sea levels on the two curves were identical. This provides the necessary calibration of the absolute sea-level scale. In order to estimate the depth of the Badenian– Sarmatian sequence at the time of deposition, the influence of subsidence from the Badenian to the present was taken into account. In the northwestern area, this total subsidence is 38 m, suggesting that the Badenian–Sarmatian section was located 38 m higher than the present-day seafloor at the time of deposition. Borehole information points to a palaeowater depth of 50 m in this area (Konerding 2006; T¸ambrea, pers. comm.), implying that the section was located 12 m below the present-day sea level on the inner shelf at the time of deposition. In the

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Fig. 14. Chronostratigraphic calibration of the Mio-Pleistocene seismic units from the Romanian Black Sea shelf. PDU, Pontian-Dacian unconformity; IPU 1, IPU 2, IPU 3, intra-Pontian unconformities; BPU, erosional unconformity at the base of the Pontian; BBU, erosional unconformity at the base of the Badenian; IPU 1 occurs only locally in a limited area; MFS, maximum flooding surface; TS, transgressive surface; SB, sequence boundary. Sea-level curve is relative to the present-day sea level (0 m). The interval in black on the water-level curve was estimated for the Romanian shelf, while the grey portion of the curve was taken from Haq et al. (1987). The geologic periods shown in grey illustrate the Mio-Pleistocene erosional unconformities.

south, the total subsidence is 35 m and the palaeowater depth was about 65 m, yielding a location 30 m below the present-day sea level on the inner shelf at the time of deposition. The internal configuration of the unit (Fig. 15) and its inner shelf position suggest that it was deposited during a late highstand when the sea level began to fall slowly after reaching a maximum. At that time, the rate of deposition was higher than that of sea level fall; the parasequences prograded basinward and downlapped onto the lower sequence boundary. Where Pontian 1 (P1) is developed, the total subsidence is 89 m from the beginning of the Pontian to the present. This suggests a position of P1 89 m below the present-day seafloor, which was 50 m below the present-day sea level. This in turn implies a location on the slope at the time of deposition, since the Black Sea level at P1 time was about 80 m higher than that of today

(Fig. 14). The onlap terminations against the lower sequence boundary coupled with the progradational nature of the subunit (Fig. 16) suggests that it was deposited during a late lowstand, when sea level began to rise very slowly. The high sediment supply and slow sea level rise led to a progradational depositional pattern, corresponding to the lowstand wedge of a lowstand systems tract (LST). Subunit Pontian 2 (P2) exhibits two different seismic facies patterns. On the western inner shelf, it is thin and is characterized by divergent, discontinuous, low-amplitude reflectors that onlap the lower boundary, while the upper boundary is represented by an erosional truncation (Fig. 17). On the middle and outer shelves, its thickness increases rapidly eastwards and it is characterized either by medium amplitude reflectors parallel to the lower boundary, or occasionally by chaotic reflections (Fig. 17). Both the lower and the upper boundaries

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Fig. 15. Detailed view of the Badenian– Sarmatian section from the northwestern middle shelf; grey arrows mark downlap and toplap terminations respectively.

are erosional. The internal configuration and external geometry of subunit P2 suggests that at the time of deposition, it straddled the continental slope and the deep basin. Since basinal deposition is possible only during a lowstand when the shelf is exposed and the sediment bypasses the shelf to be laid down on the slope and in the basin, this unit is attributed to a LST. For the subunit Pontian 3 (P3), only an approximate subsidence correction is possible because of insufficient borehole information. On the outer shelf, a difference of .200 m between the present-day

depth of P3 and the total subsidence was estimated. At the time of deposition, P3 was over 300 m below the present-day sea level. The reflection terminations and the wedge-shape (Fig. 18) suggest that this subunit was deposited during a sea level lowstand. The sea level fall that produced the lower subunit boundary was beyond the shelf-break; the shelf became subaerial and rivers incised into the exposed shelf. The rate of sea level fall exceeded the rate of subsidence, and a large sediment supply produced basinward progradation of sediments deposited on the slope.

Fig. 16. Detailed view of the section P1; grey arrows indicate onlap terminations against the lower subunit boundary (also a sequence boundary).

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Fig. 17. Details of subunit P2 showing its characteristics on the inner shelf (above) and on the middle and outer shelves (below). In the upper figure grey arrows indicate onlap terminations against the lower subunit boundary. Note the erosional truncations at both the upper and lower boundaries (below).

For the subunit Pontian 4 (P4), a subsidence correction is not possible because of the lack of borehole information for this interval. Based on reflection terminations and the seismic facies characteristics (Fig. 19), we speculate that P4 was located on the shelf at the time of deposition. The lower part of the subunit is attributed to a transgressive systems tract (TST), which formed as sea level began to rise and reached the shelf-edge. Thereafter, accommodation space was created at a rate faster than it could be filled with sediment and a retrogradational pattern marked by onlap terminations developed. At the end of the TST, the sea level rose at a rate faster than the rate of sediment supply to the basin and a maximum flooding

surface formed. Thereafter, sea-level rise slowed down, while the rate of sediment supply increased; an aggradational pattern typical for the HST developed. The subsequent sea-level fall led to the formation of a sequence boundary. The external geometry of the Dacian unit (Fig. 20) suggests that it was located on the shelf at the time of deposition. After the formation of the previous sequence boundary, sea level fell below the shelf-break. The continental shelf was exposed and retrogressive subaerial erosion of the shelf-break took place, leading to c. 300 m deep, canyon-like incisions (Gillet et al. 2007). Subsequent to the lowstand, the incised canyons were filled transgressively during Dacian time.

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Fig. 18. Details of the subunit P3. Grey arrows indicate downlap terminations against the lower subunit boundary.

We presume that the Romanian–Quaternary section (Fig. 20) was deposited after the transgressive phase of the Dacian during a sea level highstand. As the rate of sea-level rise slowed down after reaching a maximum, sediment supply increased and a maximum flooding surface formed. Sediment continued to be delivered at a faster rate than that of sea-level rise and typical aggradational parasequences developed.

Stratigraphic modelling The results obtained here were used to create a sedimentation model that reproduces the findings described above. The model was generated in

order to determine how different parameters controlled the depositional processes. The position of the modelled profile was constructed on the basis of six wells distributed from the inner to the outer shelf, following the direction of line 1 from the subsidence analysis (Fig. 1). The modelled profile has a length of 71 km, and the simulation started at the base of the Badenian (16 Ma BP). Lithological information from boreholes suggests that only a marginal amount of carbonates are present within the basin; almost all sediments are clastic. The direction of the simulated deposition is from NW to SE, namely from the left. Modelling was carried out taking into account the effect of isostasy during deposition. Hereby the

Fig. 19. Details of subunit P4. Grey arrows indicate onlap terminations against the lower subunit boundary.

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Fig. 20. Details of the Dacian, Romanian and Quaternary section. Incised valley of about 300 m depth occurs at the base of the Dacian incised into the Pontian.

lithosphere was assumed to behave as an elastic beam with a flexural rigidity of 1023. Information on the water-level curve and subsidence history is based on this study. The compaction parameters for this area are from Ionescu (2000). The best-fit model of the simulation is shown in Figure 21 (upper panel), while the corresponding Wheeler diagram is shown in the lower panel. For comparison, a seismic profile that crosses the shelf in a NW–SE direction close to the modelled line is shown in Figure 22. Simulation began with the erosional unconformity that marks the base of the Badenian–Sarmatian unit (BBU). The final model shows that the Badenian–Sarmatian formation extends to the SE on the outer shelf, while this extension is not obvious in the seismic data because of their low vertical resolution. The second unconformity at the base of Pontian (BPU) could also be simulated over the entire profile. However, only two of the intra-Pontian unconformities, namely IPU 2 and IPU 3, could be simulated. IPU 1 occurs only locally in the seismic data and did not appear in the simulation. The Pontian deposits could be simulated satisfactorily, but this is not the case for the position of IPU 2 and IPU 3 because of insufficient data. While the shapes of these unconformities are well modelled on the inner shelf (left part of Fig. 21), they are less precise basinward, where the dip of IPU 2 is too large and the position of IPU 3 is too high. We attribute this to a lack of lithological information between the wells 13 Heracleea and 40 Albatros as well as between 40 Albatros and 1 Ovidiu. A second problem concerns the uncertainties in the age of the formations bounding IPU 2 and IPU 3 since these intra-Pontian unconformities are not seen in the borehole data. In addition, the strong influence of tectonics on Pontian

sedimentation in this area could not be taken into account during the simulation. The internal seismic characteristics of the intra-Pontian units P2 and P3 could be well modelled. The Lower Pontian comprises two different seismic facies: thin divergent reflectors on the inner shelf and parallel reflecting packages that thicken basinward on the middle and outer shelves. P3 (the intra-Pontian unit between IPU 2 and IPU 3) has a modeled thickness larger than that deduced seismically but has the required internal pattern, namely wedge-shaped reflectors that downlap at the lower boundary and are erosionally truncated at the top. They dip steeply in the west but are almost horizontal in the east. Above IPU 3, parallel reflectors characteristic for the Upper Pontian –Quaternary are modelled. The Wheeler diagram in the lower panel of Figure 21 shows the corresponding chronostratigraphic development. During the first 7 Ma of simulation, the Badenian–Sarmatian unit, which has a non-depositional hiatus in the middle part of the profile, was deposited. Two periods of nondeposition (3– 3.4 Ma and 4–5 Ma simulation time, ST) occurred in the proximal section. The Badenian– Sarmatian deposits are thin, reaching only a maximum of 250 m on the inner shelf. This low sedimentation rate was controlled largely by low sediment input (most of the sediments eroded from the Carpathians were deposited in the Dacic Basin) and slow subsidence, with sea level fluctuations playing a minor role. At the end of the Badenian– Sarmatian (c. 8.5 Ma ST), the Wheeler diagram shows another hiatus on the inner shelf (left side of the diagram) which corresponds to the sequence boundary (BPU) that formed in response to sea level fall. Basinward the hiatus pinches out and the water depth increases abruptly from about

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Fig. 21. Modelled profile for the simulation with isostatic compensation; the flexural rigidity was chosen to be 1023. Upper panel: the simulated model with colour-coded water depths at the time of deposition of the different units. Lower panel: the corresponding Wheeler diagram. For this panel, the vertical scale is simulation time in years, 0 marks the start of simulation. The Black Sea water-level curve is plotted in light blue on the right, the vertical violet line represents present-day sea level. The light blue line to the left of the violet line gives the water-level scale. The water depth at the time of deposition is colour-coded. Light and dark grey fields mark the areas of zero accumulation: the former corresponds to locations where sediment was deposited but was subsequently eroded, and the latter locations where deposition never occurred (hiatuses).

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Fig. 22. Seismic profile crossing the shelf in a NW– SE direction, chosen for comparison (see Fig. 1 for profile location).

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200 to 800 m. Above the BPU, the palaeo-water depths varies from c. 50 m on the present-day inner shelf to .1000 m on the present-day outer shelf. The unconformity IPU 2 (c. 10 –10.5 Ma ST) could be simulated only in three areas, elsewhere there is an abrupt increase in water depth. IPU 2 corresponds to a transgression and the subsequent highstand, as indicated by the sea level curve. Above IPU 2, subunit P3 developed at water depths of c. 50 m on the present-day inner shelf to 1700 m on the present-day outer shelf. That is, the depositional environment was a shelf proximally and a basin distally. IPU 3 (c. 11.5– 11.75 Ma ST) is represented on the Wheeler diagram only in areas where borehole data are available. The thick, wedge-shaped Pontian section deposited only in about 2 Ma marks a time when the Dacic Basin was filled and sediments were laid down on the continental shelf and beyond. The youngest (Romanian, c. 1.3 Ma ST) unconformity visible in the Wheeler diagram could not be identified in the seismic data. Our simulation shows that Romanian– Quaternary sedimentation after the formation of this unconformity was quiet.

Conclusion Structural and sedimentological analyses of the Mio-Pleistocene deposits on the Romanian Black Sea shelf using seismic and borehole data show a changing evolution during a relatively short period of time. This evolution is directly related to the pre-Miocene evolution of the Romanian shelf. The available seismic data, which is relatively uniformly distributed over the study area, yielded important information on the tectonic evolution of these deposits. Mio-Pleistocene subsidence of the Romanian Black Sea shelf is directly dependent on sediment input, tectonic activity as well as water-level fluctuations and is highly variable. Subsidence during the Badenian–Sarmatian and the Dacian –Quaternary was limited. The amount of subsidence calculated along the WNW–ESE profile increases slowly from the inner to the outer shelf. In the NW– SE direction, subsidence was slow and constant. In contrast, during the Pontian, shelf subsidence was progressively faster in the basinward direction. Subsidence on the outer shelf was much more significant than elsewhere on the shelf. This is the result of a large sediment input to the basin and of significant tectonic activity in the Histria Depression. The most tectonically active period during the Mio-Pleistocene was the Pontian. The Badenian– Sarmatian was largely quiescent and the Dacianto-Quaternary saw a decrease in the Pontian tectonic activity, coming possibly even to a halt. We

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classified the faulting that affected the Pontian sedimentary section both temporally and spatially. From the temporal point of view, we distinguish between (1) faults that originated in the Pre-Oligocene and remained active up to the Pontian; (2) faults that developed during the Oligocene– Pontian; and (3) faults that were active only during the Pontian. The first are supposedly reactivated faults that developed during the formation of the Histria Depression, the second are gravitational faults that developed in sediments of a similar facies, and the last are gravitational faults related to the extensional phase in the Pontian. Spatially, we distinguish between (1) extensional faults that accompanied the development of the Pontian Depression; and (2) gravitational faults at the shelf-break. Based on the available seismic data, the reconstructed subsidence history and the water level curve, the Mio-Pleistocene sedimentary section was divided into eight systems tracts. The Badenian– Sarmatian unit was attributed to a HST (highstand systems tract), being deposited in a late highstand when the sea level began to fall slowly after reaching a maximum. The next cycle cannot be observed in the seismic data, because it corresponds to sedimentation on the Romanian shelf during Meotian time, and was subsequently completely eroded. The Pontian subunit P1, which was laid down on the slope at the time of deposition, is progradational and attributable to the lowstand wedge of a LST. Subunit P2 is likewise also attributed to a LST, having the continental slope and the deep basin as palaeo-depositional environments. Because the boundary between units P1 and P2 occurs only locally and could not be correlated with our estimated sea-level curve, it was presumed to have formed during a higher order sea-level cycle. Subunit P2 is followed by a non-depositional hiatus that separates P2 from P3. On the sea level curve, it corresponds to the HST and TST that followed the LST of P2 and ends with a sequence boundary. The reflection terminations and the wedge-shape of P3 suggest that it was deposited in the deep basin during a sea level lowstand. The succeeding TS (transgressive surface) corresponds on the seismic data to the erosional unconformity between the subunits P3 and P4. This TS is followed by the next TST and HST that developed during the deposition of P4, but a MFS (maximum flooding surface) is missing in the seismic data because of insufficient vertical resolution. Subunit P4 was deposited on the inner and middle shelves, as deduced from its reflection terminations and seismic facies characteristics. The boundary between P4 and the Dacian is represented by another erosional hiatus, which comprises the LST that follows the formation of the sequence boundary at the end of the Pontian. During the Dacian –Quaternary, the subsequent TST

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and HST were deposited on the inner and middle shelves. The MFS that separates the TST and the HST could not be observed in the available data. The stratigraphic model obtained by simulation of the seismic section, in which the subsidence history and the Black Sea water-level curve were taken into account, shows a strong influence of sediment input and subsidence on sedimentation, while sea level fluctuations played a lesser role. The sediment input is related to the evolution of the adjacent land, while subsidence is dependent on sediment supply, tectonic activity and sea level fluctuations. Thus, these two factors are not totally independent. Gillet et al. (2007) and Hsu¨ & Giovanoli (1979) argued convincingly for a temporary Messinian evaporative draw-down of the Black Sea from 6.25 to 5.2 Ma contemporary with the Messinian salinity crisis of the Mediterranean. Our proposed sequence stratigraphic framework is consistent with their scenario. The draw-down event is marked by our IPU 2, the equivalent of their IPU. However, our proposed subsidence history as well as our reconstructed palaeo-water-level curve does not have a resolution high enough to verify their hypothesis. The authors would like to thank the National Agency for Mineral Resources of Romania and PETROM Oil Company which gave us access to seismic and well data. Thanks are also due to Dr I. Popescu for her constructive review of our manuscript. For one of the authors (CD) this work was supported by CNCSIS project number IDEI 960/2007.

References Dinu, C., Wong, H. K. & T¸ambrea, D. 2002. Stratigraphic and tectonic syntheses of the Romanian Black Sea shelf and correlation with major land structures. Bucharest Geoscience Forum, Special Volume, 2, 101– 117. Dinu, C., Wong, H. K., T¸ambrea, D. & Matenco, L. 2005. Stratigraphic and structural characteristics of the Romanian Black Sea shelf. Tectonophysics, 410, 417– 435. Gillet, H., Lericolais, G. & Rehault, J. P. 2007. Messinian event in the Black Sea: evidence of a Messinian erosional surface. Marine Geology, 244, 142– 165. Haq, B. U., Hardenbol, J. & Vail, P. R. 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235, 1156– 1167.

Hsu¨, K. J. & Giovanoli, F. 1979. Messinian event in the Black Sea. Palaeogeography, Palaeoclimatology, Palaeoecology, 29, 75– 93. Ionescu, G. 2000. Facies models of the Paleogene formations on the northwestern Black Sea shelf. PhD thesis, Bucharest University, Faculty of Geology and Geophysics. Jipa, D. 1997. Late Neogene – Quaternary evolution of Dacic Basin (Romania). An analysis of sediment thickness pattern. Geoecomarina, 2, 127–134. Jones, R. W. & Simmons, M. D. 1997. A review of the stratigraphy of Eastern Paratethys (OligoceneHolocene), with particular emphasis on the Black Sea. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoir, 68, 39– 52. Konerding, C. 2006. Mio-Pleistocene sedimentation and structure of the Romanian shelf, northwestern Black Sea. PhD thesis, University of Hamburg, Institute of Biogeochemistry and Marine Chemistry. Moros¸anu, I. 2002. Inverted structural types of the Romanian offshore Black Sea. In: Dinu, C. & Mocanu, V. (eds) Geology and Tectonics of the Romanian Black Sea Shelf and its Hydrocarbon Potential. BGF Special Volume, 2, 21–28. Neprochnov, Y. P. & Ross, D. A. 1978. Black Sea Geophysical Framework: Initial Reports of the Deep Sea Drilling Project, 42(2), US Govt. Printing Office, Washington, DC, 1043–1055. Pa˘trut¸, I. 1975. Investigations of the geological structure of the Black Sea continental platform and the position of the first drilling operations. Internal report, Ministry of Petroleum, Bucharest. Popescu, I., Lericolais, G., Panin, N., Normand, A., Dinu, C. & Le Dresen, E. 2004. The Danube submarine canyon (Black Sea): morphology and sedimentary processes. Marine Geology, 206, 249–265. Robinson, A. G., Rudat, J. H., Banks, C. J. & Wiles, R. L. F. 1996. Petroleum geology of the Black Sea. Marine and Petroleum Geology, 2, 195–223. Ross, D. A., Stoffers, P. & Trimonis, E. S. 1978. Black Sea Sedimentary Framework: Initial Reports of the Deep Sea Drilling Project, 42(2), US Govt. Printing Office, Washington, DC, 359–372. Ta˘ra˘poanca˘, M. 2004. Architecture, 3D geometry and tectonic evolution of the Carpathians foreland basin. PhD thesis, Netherlands Research School of Sedimentary Geology (NSG), Publication No. 2004.01.03. T¸ambrea, D., S¸indilar, V. & Olaru, R. 2000. The Pontian from the Romanian continental shelf of the Black Sea. Romanian Oil Journal, 7, 9 –20.

The Early Mesozoic evolution of the Western Greater Caucasus (Russia): Triassic– Jurassic sedimentary and magmatic history T. MCCANN1*, F. CHALOT-PRAT2 & A. SAINTOT3 1

Geologisches Institut, Bonn University, Nußallee 8, 53115 Bonn, Germany

2

CRPG-CNRS/Nancy Universite´, BP 20, 54501 Vandoeuvre-Les-Nancy Cedex, France 3

Geological Survey of Norway, NO-7491 Trondheim, Norway *Corresponding author (e-mail: [email protected])

Abstract: The Greater Caucasus (GC) forms a high Alpine fold-and-thrust belt on the southern margin of the East European Platform (EEP). The Triassic, and particularly, the Jurassic history of the Western Greater Caucasus region is important for our understanding of the palaeogeographic and tectonic evolution of the western Tethys area. In order to better constrain the nature and relevance of these events in the evolution of the region, which are classically described as the Late Triassic to Late Jurassic Cimmerian events, a field campaign in the Western Greater Caucasus was undertaken. Analysis of structural, sedimentological and petrological data from 41 sites in the ForeCaucasus (Malaya Laba, Mount Tkhach-Belaya River), the Central Greater Caucasus (Georgievskoye, Otdaleni) and Southern Slope (Krasnaya Poliana) areas of the Western Greater Caucasus revealed that a broad asymmetric basin, with associated emergent volcanic islands, formed in the area in Jurassic times. Incipient back-arc rifting in Pliensbachian times was coeval with similar rifting episodes in the Pontides and South Caspian Sea areas. The synchroneity of these events may have been related to the renewal of the Tethys subduction to the south of the Eo-Cimmerian accretionary belt. Rift reactivation, with significant thinning of the continental lithosphere, occurred in the Aalenian. Despite the strong Alpine tectonic overprinting, some structural data confirms that the extension trend was east– west (almost parallel to the active margin) resulting in the formation of a series of pull-apart basins in the GC and the South Caspian region behind the Eastern Pontides–Lesser Caucasus subduction-related volcanic belt. In Bajocian times, subduction-related volcanic activity largely expanded from the Eastern Pontides– Lesser Caucasus to encompass the Transcaucasus, the southern part of GC and the Crimea region. Such widening of the volcanic arc was probably due to a shallowing of the northward subducting slab. In the back-arc GC region, this signalled the onset of the post-rift stage. The return of the slab to normal steepness resulted in subsidence in the back-arc region and in the GC with extensive accommodation space creation. This was subsequently filled by the Late Jurassic, Cretaceous and Cenozoic sedimentary successions.

The Greater Caucasus (GC) belt, situated on the southern margin of the East European Platform (EEP) extends from the northern margin of the eastern Black Sea Basin to the western margin of the South Caspian Basin and forms a morphological barrier along the southern margin of the Scythian Platform (Figs 1 & 2). The GC has been studied for over 150 years and there is a significant volume of Soviet, Russian and Georgian literature dealing with the geological evolution of the GC belt, much of which is not easily accessible (see Saintot et al. 2006a for details). The formation of the GC fold-and-thrust belt resulted from several phases of deformation during Phanerozoic times (Milanovsky & Khain 1963; Razvetaev 1977, 1989; Adamia et al. 1977, 1981; Milanovsky et al. 1984; Muratov et al. 1984; Khain 1984; Gamkrelidze 1986; Dotduyev 1989; Philip et al. 1989; Zonenshain et al. 1990; Milanovsky 1991; Nikishin et al. 1998a, b, 2001) related to accretion (Gondwanan, Tethyan terranes) along the southernmost edge of the East European

Platform (EEP) (see discussions by Sengo¨r 1984; Zonenshain et al. 1990; Dercourt et al. 1993, 2000). A broad overview of the Jurassic succession of the Caucasus region has been provided by Rostovtsev et al. (1992; stratigraphy subsequently updated by Ruban 2003). Two main sedimentary successions have been recognized, an older (Sinemurian –Bathonian) complex predominantly clastic, and which is up to 10 000 m thick and overlain by a carbonate-dominated complex (Callovian– Tithonian; up to 3000 m; Akmedov et al. 2003; Ruban 2006); evaporites are present in the uppermost part (Tsejsler 1977; Rostovtsev et al. 1992). In some areas (e.g. the Lesser Caucasus), volcanics and volcaniclastics are abundant. In Jurassic times, the Caucasus region was located on the northern active margin of the NeoTethys Ocean (Gamkrelidze 1986; Dercourt et al. 2000; Stampfli & Borel 2002; Golonka 2004; Barrier & Vrielynck 2008). The area formed part of a widespread system of marine basins, and

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 181–238. DOI: 10.1144/SP340.10 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Simplified geological map of the Greater Caucasus and Lesser Caucasus areas (after Milanovsky & Khain 1963; Belusov et al. 1988). MCT, Main Caucasian Thrust; PTF, Pshekish-Tyrnauz Fault; RLFZ, Racha-Lechkhumy Fault Zone. Greater Caucasus: 1, Main Range; 2, Peredovoy Range; 3, Betcha Anticline; 4, Svanetia Anticline; 5, Laba-Malka Monocline; 6, Dagestan Folded Zone; 7, SE Greater Caucasus Flysch Zone; 8, NW Greater Caucasus Flysch Zone; 9, Kuban syn-orogenic basin; 10, Terek-Caspian syn-orogenic basin; 11, Kusar-Divichi syn-orogenic basin. Lesser Caucasus: 12, Somketo-Karabakh Zone; 13, Artvin-Bolnisi Zone; 14, Adzharo-Trialet; 15, Talesh; 16, Sevan–Akera; 17, Kafan; 18, Vedin; 19, Zangezur; 20, Mishkhan-Zangezur Massif; 21, Ararat-Djulfa Massif; 22, Araks Basin. Trans Caucasus & Black Sea intramontane zones: 23, Rioni Basin; 24, Kura Basin; 25, Dzirula Massif; 26, Tuapse Basin; 27, Shatsky Ridge.

Fig. 2. Generalized cross-section across the central part of the Greater Caucasus showing the southward vergent thrusting of the belt over the Transcaucasus (after Dotduyev 1987, 1989). Location indicated on Figure 1.

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associated emergent volcanic islands, that developed along the southern margin of the Scythian Platform, behind the main Tethys subduction zone (e.g. Ershov et al. 2003; Efendiyeva & Ruban 2005). However, the precise geometries of these basins changed during the Jurassic (Ruban 2006), and these changes have not yet been evaluated. Furthermore, the tectonic history of the Caucasus region needs to be integrated within a global palaeotectonic framework (Stampfli & Borel 2002; Stampfli et al. 2002; von Raumer et al. 2003) in order to understand the evolution of the adjacent domains of Europe and the Middle East. The main aim of the present contribution is to examine the Jurassic history of the region of the Western Greater Caucasus – a critical time in its evolution since the Cimmerian orogenic phases (in terms of their extent and age) are not well understood. Within the framework of the MEBE (Middle East Basin Evolution) programme, fieldwork was carried out in the western GC and 41 sites were visited (Fig. 3 and Table 1). Fieldwork was concentrated in four areas (Georgievskoye – sites 1 to 11, 18, Krasnaya Poliana – sites 12 to 17, Otdaleni/Pchekha Pshekha River – sites 19 to 28 and Kamennomostsky/Belaya River – sites 29 to 41). In the Belaya River area, the Triassic

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succession was also examined (cf. Gaetani et al. 2005). The Western Greater Caucasus region is large, complex, poorly exposed, relatively inaccessible and overprinted by penetrative Alpine folding and thrusting. All of these factors contribute to our poor understanding of the dynamics of the region. Published stratigraphic sections are generally composite sections obtained through the interpretation and correlation of separate outcrops. This present study, while examining sedimentary and volcanic successions from a series of Jurassic-age outcrops across the Western Greater Caucasus, is neither a revision nor an upgrading of the present stratigraphy (such a study would necessitate extensive fieldwork). Rather, our aim was to characterize the pattern of both sedimentation and volcanic activity, in terms of areal and stratigraphic distribution, for the Western Greater Caucasus region in Jurassic times. In addition, geochemical (major & trace elements) and isotopic (Nd–Sm, Rb–Sr) studies of the volcanic associations were carried out in order to determine the locations of the magma sources within the lithosphere and asthenosphere. This multi-disciplinary approach thus provides us with a more precise evaluation of the plate tectonic evolution and geodynamic history of the Greater Caucasus region.

Fig. 3. Locations of sampling sites from the southwestern part of the Greater Caucasus region. (Location indicated on Fig. 1, after the Geological Map of the Caucasus 1964.)

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Table 1. List of locations, Western Greater Caucasus Site number

Latitude

Longitude

1 2 3 4 5 6 7 8 9 10 11 12

44880 59.4100 N 44860 46.9000 N 44860 51.5900 N 44890 53.2300 N 448110 18.1500 N 448150 10.1600 N 448190 20.2500 N 448190 14.5200 N 448190 27.6200 N 448190 41.1900 N 448140 12.5800 N 438400 53.6800 N

398160 42.0200 E 398180 23.1200 E 398180 14.0000 E 398160 22.1600 E 398140 31.9100 E 398150 1.4400 E 398170 20.2000 E 398200 48.4500 E 398210 36.8200 E 398210 2.0300 E 398170 44.1700 E 408160 25.7200 E

13

438410 20.8300 N

408140 54.9600 E

14

438410 13.0400 N

408140 5.2000 E

15

438410 4.1800 N

408140 4.9000 E

16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34

438400 4.0300 N 438380 0.1700 N 448210 50.1000 N 44800 0.7800 N 44800 31.4200 N 44820 16.4300 N 44840 44.6200 N 44840 34.2300 N 44830 40.9000 N 44840 24.6100 N 44890 37.5200 N 448100 12.7700 N 448130 3.7900 N 448160 38.3800 N 448150 16.0900 N 44800 7.9100 N 44800 30.6300 N 44870 36.0400 N 44860 19.9900 N

408110 7.1300 E 408170 14.6800 E 398180 33.9000 E 398400 43.1300 E 398410 17.6800 E 398420 14.4000 E 398410 8.0800 E 398390 55.2700 E 398360 31.4200 E 398380 46.6200 E 398390 10.5200 E 398390 7.4800 E 398420 47.1300 E 408110 3.2200 E 408120 6.5800 E 40880 3.4100 E 40880 28.5600 E 40870 28.3700 E 40800 48.0400 E

35 36 37 38 39 40

44840 56.8200 N 44860 42.4800 N 448120 38.9900 N 448100 14.7200 N 44890 32.1600 N 44870 46.8400 N

40800 21.2400 E 40810 11.2100 E 408150 57.4400 E 408170 51.9800 E 408180 0.1700 E 408190 0.3500 E

41

44870 57.6000 N

408180 54.8200 E

Formation Makarov Tchataltapa Makarov Makarov Mount Indyuk Mount Indyuk/Pchich Mount Indyuk Pchich Pchich Tchataltapa/Pchich Mount Indyuk/ Makarov Turovoy/Estasadok/ Tchvejips/Illaryonov/ Anchkhoi/Atchichkh Estasadok Estasadok/Illaryonov/ Anchkhoi Estasadok Porphyritovaya Series Porphyritovaya Series Pikhtar Tchataltapa/Makarov Bzerpia Tchataltapa Tchataltapa Bzerpia Tchataltapa Turovoy Zeitun Yablon Pikhtar Kamennomostsky Tkhach Group Psebai Psebai Djangur Psebai/ Kamennomostsky/ Gerpegem Gerpegem/Mezmai Kamennomostsky Psebai Bugunja Yatyrgvarta (Tkhach Group) Maliy Tkhach (Tkhach Group) and Khodz Group Maliy Tkhach (Tkhach Group) and Khodz Group

Structure of the Greater Caucasus region and Early Mesozoic Greater Caucasus basin development The present-day Greater Caucasus south-vergent fold-and-thrust belt comprises a basement core

Age Aalenian Aalenian Aalenian Aalenian Aalenian Aalenian – Bajocian Aalenian Aalenian – Bajocian Aalenian – Bajocian Aalenian Aalenian Pliensbachian – Bajocian Sinemurian – middle Pliensbachian Sinemurian – Aalenian Sinemurian – middle Pliensbachian Bajocian Bajocian Oxfordian– Kimmeridgian Aalenian Toarcian Aalenian Aalenian Toarcian Aalenian Pliensbachian – Toarcian Late Aalenian – Bajocian post-Callovian Oxfordian– Kimmeridgian Callovian– Oxfordian Early and Middle Triassic Toarcian– Aalenian Toarcian– Aalenian Bajocian– Oxfordian Toarcian– Callovian Oxfordian– Thitonian Callovian Toarcian– Aalenian Sinemurian – Pliensbachian Early Triassic Middle and Late Triassic Middle and Late Triassic

containing strata as old as Proterozoic (the Main Range and the Peredovoy Range) with Jurassic to Eocene-age units located along its flanks (Fig. 1). The deep crustal structure of the GC and Transcaucasus are not well known (cf. Saintot et al. 2006a, b). Within the Greater Caucasus three distinct tectonic areas can be noted – the southernmost

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Southern Slope, the Central Greater Caucasus (¼ the Main Range plus the Peredovoy or Fore Range), and the Fore-Caucasus to the north (Zonenshain et al. 1990) (Fig. 1). The Central Greater Caucasus comprises the Proterozoic- and Palaeozoic-age, crystalline and metamorphic core of the GC (Main Range and Peredovoy Range), and the Late Palaeozoic to Middle Jurassic mainly sedimentary succession. Along the Main Caucasian Thrust sensu stricto (MCT) the Main Range thrusted southward over the Lower and Middle Jurassic units. The Southern Slope is separated from the Central Greater Caucasus by a southern branch of the Main Caucasian Thrust (these two branches form the MCT sensu lato). A variety of Palaeozoic sedimentary complexes are exposed in the central part of the Greater Caucasus, where the Greater Caucasus Terrane was identified by Gamkrelidze (1997), Tawadros et al. (2006) and Ruban et al. (2007). The Southern Slope comprises an intensely deformed south-vergent fold-and-thrust belt which thrusts over the Transcaucasus along the RachaLechkhumy Fault Zone (Fig. 1). The Southern Slope is separated from the Main Range of the GC by the MCT. The Pshekish-Tyrnauz Fault separates the Main Range and the Peredovoy Range of the Greater Caucasus (Fig. 1). A large structure also separates the Peredovoy Range from the ForeCaucasus extending far to the north (Fig. 1). As noted above, the structural setting of the GC is complex, and not fully understood. The most appropriate tectonic framework was proposed by Khain (1975), Gamkrelidze (1986), Gamkrelidze & Giorgobiani (1990) and Ershov et al. (2003) who interpreted this domain as an Alpine WNW– ESE-oriented fold-and-thrust belt disrupted by parallel trans-regional shear zones with strike –slip displacement along the central part of the Greater Caucasus. The present structure of the Greater Caucasus was created in the Cenozoic (Ershov et al. 2003). However, compressional and extensional phases in both Palaeozoic and Mesozoic times have been recognized. For our purpose, in Early Mesozoic times, three main tectonic phases can be recognized. In Late Triassic times (1) a compressive event (Eo-Cimmerian tectonic phase) resulted in the inversion of the Permo(?)-Triassic continental rift basins (Nikishin et al. 1998a, b, 2001; Gaetani et al. 2005). Renewed rifting in the Early Jurassic (2) accompanied by both marine and continental sedimentation and volcanic eruptions (Sinemurian– Pliensbachian according to Lordkipanidze et al. 1989) affected the whole region behind the Tethys subduction trench. North –south-striking Aalenian-age synsedimentary normal faulting suggests that there was a major phase of renewed rifting (3) under a dextral transtensional regime along the basin margin (Se´brier et al.

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1997; Saintot et al. 2004). This rifting episode was very significant, as evidenced by the eruption of MOR-basalts, suggesting that the asthenosphere was close to the surface as a result of the marked lithospheric thinning. As part of our fieldwork, detailed structural analysis was carried out in order to better constrain the tectonic regime during the Jurassic rifting episode(s), that is, to identify the rare structures related to the various rifting episodes (and not destroyed by subsequent Alpine compressional overprinting events). The stress inversion confirms the highly oblique trend of extension (relative to the strike of the belt) that affected the Lower to Middle Jurassic rocks (Fig. 4). Therefore, the structural setting of the sedimentation and volcanism (as described in this paper) was most probably within a system of pull-apart basins formed along large WNW –ESE-trending dextral faults and approximately north –south-trending normal faults. (Note that these large WNW–ESE dextral faults provide the main structural grain of the Alpine GC fold-and-thrust belt since they were probably reactivated during later tectonic inversion to form the main Alpine thrusts.) An angular unconformity present at the base of the Callovian marks a major event in the evolution of the GC. However, its significance and extent remains unclear. The unconformity may be related to local orogenic events (Ershov et al. 2003) or it may have resulted from a fall in sea level (Ruban 2007). Whatever the nature of the pre-Callovian event, the Late Jurassic to Late Eocene depositional history of the GC was mostly dominated by the infill of a thick succession (6 –8 km) of deep-marine sediments (cf. Saintot et al. 2006a). Ongoing Alpine compression resulted in the formation of the southvergent GC fold-and-thrust belt.

Stratigraphic succession within the Greater Caucasus area The area of the Greater Caucasus can be subdivided into three main regions, namely, the Fore-Caucasus, the Central Greater Caucasus (sensu stricto) and the Southern Slope areas. Within these areas, the stratigraphic successions of the Western Greater Caucasus are distributed in a number of what are termed (in the local literature) ‘zones’ (see below, Geological Map of the Russian Federation 2001, 2002; Figs 5–8). These zones are defined in terms of their tectonic history or the facies present (see fig. 3 in Ruban 2006 for map of zones). As such they are not terranes in the classical sense (i.e. faultbounded entities of regional extent, characterized by a geological history which differs from that of the adjacent terranes and where the boundaries

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Fig. 4. Trends of extension observed in Lower and Middle Jurassic rocks of the western Greater Caucasus that confirm the development of pull-apart basins along WNW– ESE strike– slip and roughly north– south normal faults. Synsedimentary normal faults constrain an Aalenian age of the event (Se´brier et al. 1997; Saintot et al. 2006a). Site numbers as on Figure 3 and Table 1, this study (except site 143 from Saintot & Angelier 2002). Method for stress tensor reconstruction is described in Angelier 1990, 1994). Key for stereoplots: Schmidt’s projection, lower hemisphere; bedding planes as broken lines, tension gashes as dotted lines, fault planes as thin lines, striae as small arrows (outward-directed, normal, couple of thin arrows, strike–slip). BR as back-rotated stress state (it means that the structures developed prior to any tilting/folding of the beds, that is, in agreement with a Middle Jurassic age of faulting prior to any compressional tectonic phase, cf. discussion in Saintot et al. 2006a). RUP, criterion of quality ranging from 0% (maximum calculated shear parallel to actual striae with the same sense) to 200% (maximum calculated shear parallel to striae but opposite in sense), acceptable results with RUP ,75%. a, average angle between observed striae and computed shear, in degrees (acceptable with a , 22.58).

may be transcurrent faults, thrust faults or a me´lange, being cryptic or polygenetic; Oliver et al. 1993). In the Western Greater Caucasus, the boundaries between the individual zones do not

always crop out, and it is difficult to understand the precise relationship between one zone and those adjacent to it. In some cases, the differences appear to be facies related (e.g. proximal-distal changes)

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to consider the individual zones as separate, but related, entities. The precise interpretation of these zones would require a great deal of basic mapping, which was beyond the scope of our study. However, our fieldwork included the areas of Malaya Laba and Mount Tkhach-Belaya River (Fore-Caucasus), Georgievskoye and Otdaleni (Central Greater Caucasus) and Krasnaya Poliana (Southern Slope). The chosen areas thus provided a broad regional base to trace the evolution of the Jurassic succession (Fig. 3). Additionally, in the northern part of the region, a selection of Triassic outcrops was also examined and sampled. The various areas will be discussed in terms of their lithologies, facies, chronology, volcanic eruption dynamics, sedimentary and magmatic petrology and environmental interpretations. One further point, each zone comprises a stratigraphic succession of units that are described by the Russian term ‘svita’. Although the concept is slightly different, we have chosen to transcribe ‘svita’ by the term ‘formation’ (fm).

Fore-Caucasus area: Mount Tkhach-Belaya River/Malaya Laba (Triassic, Fig. 5; Sites 29, 30, 38, 39, 40, 41; Fig. 3, Table 1)

Fig. 5. Stratigraphy of the Triassic of the Mount Tkhach/Belaya River area (Fore-Caucasus) (after Geological Map of the Russian Federation 2002 and Gaetani et al. 2005).

whereas in other situations, the differences are more fundamental (e.g. differences in age). In terms of our interpretations, therefore, we have chosen to group the various zones into the Fore-Caucasus, Central Greater Caucasus, and Southern Slope areas (because each of these areas represents a well defined tectonic entity clearly separated by major thrusts, and thus provide a firm basis for our study; see above), and within each of these areas,

The Triassic succession is mainly confined to the northern part of the Western Greater Caucasus region (i.e. the Fore-Caucasus area). Here it crops out in two locations, the Mount Tkhach-Belaya River, and the Malaya Laba areas which were examined in order to provide a basis for our more detailed analysis of the Jurassic succession. The Triassic succession of the Mount TkhachBelaya River area, which is the most complete in the region, does not crop out continuously, and published sections are based on composites of the three main groups, from base to top, the Tkhach, Sakhrai and Khodz groups (Fig. 5; the Tkhach Group being divided from bottom to top into three formations namely the Yatyrgvarta, Maliy Tkhach and Acheshbok formations; cf. Gaetani et al. 2005). Dating of the various groups is problematic and thus a precise correlation of the various formations is not always possible. Depositional environments in the Tkhach Group (sites 30, 39, 40) are mainly low energy, possibly in a distal shelf setting. The sediments of the overlying Sakhrai Group (site 29 – dated as Ladinian in Gaetani et al. 2005; and sites 38, 41) were deposited in a continental environment, which included both alluvial and deltaic settings. However, in the Sakhrai River area there is evidence of turbiditic sediments, which would be indicative of deeper marine conditions; this would suggest that there are significant problems with regard to the definition of these various units. The uppermost Khodz Group (sites 40, 41) comprises mainly

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Fig. 6. Stratigraphy of the Jurassic of the Laba-Malka Zone (Fore-Caucasus) (after Geological Map of the Russian Federation 2001).

Fig. 7. Stratigraphy of the Jurassic of the Central Greater Caucasus: Psekhako-Berezov Zone, Goitkh Zone (Key: 1, Novorossiisk-Lazarev Subzone; 2, Grachev Subzone; 3, Bezeps-Pchich Subzone; 4, Nevieb Subzone), Abino-Gunai Zone. (After Geological Map of the Russian Federation 2001, 2002.)

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Fig. 8. Stratigraphy of the Jurassic of the Krasnaya Poliana area (Krasnopoliana, Abkhazo-Racha and Tchvejips zones of the Southern Slope and Psekhako-Berezov Zone of the Central GC). (After Geological Map of the Russian Federation 2001, 2002.)

carbonates, which were deposited in a shelf environment. At site 41 there is an unconformity between intensely-folded Middle Triassic rocks (Sakhrai Group) and the Upper Triassic-age muddy conglomerates (debris flows; Khodz Group). The Triassic succession of the Malaya Laba area comprises mainly Lower Triassic and Anisian rocks (cf. Gaetani et al. 2005). The succession consists mainly of fine-grained, thinly-bedded mudstonewackestone units, indicative of a low-energy depositional setting deposited below fair-weather wave base. Repeated episodes of terrigeneous flooding also occurred. Locally, alluvial facies prograde onto shore facies. Of interest, is the presence of an Olenekian conglomeratic unit, which includes serpentine schist pebbles (Gaetani et al. 2005). Such sediments are only found in modern sands supplied by Alpine-type subducted ophiolites with a strong metamorphic overprint (e.g. Garzanti et al. 1998).

Fore-Caucasus area: Mount Tkhach-Belaya River (Jurassic, Fig. 6, Sites 31, 32, 33; Fig. 3, Table 1) The Jurassic succession of the northern Caucasus area crops out in two zones, the Labina-Malka Zone which contains a succession ranging in age from mid-Sinemurian to Bajocian and the overlying

upper Callovian– Tithonian succession in the Labina Zone. The lowermost succession (i.e. the Laba-Malka Zone), which has a total thickness of 2800 m, commences with the Bugunja Formation which is overlain by the Psebai and Djangur formations (Fig. 6). The adjacent overlying Labina Zone also contains three formations, although the contacts between the various formations are not exposed (Fig. 6).

Laba-Malka Zone: Bugunja Formation (Fig. 6 – Site 38; Fig. 3, Table 1) Description. This formation is of Sinemurian– Pliensbachian age and comprises mainly sandstones with rarer conglomerates. The sandstones are mainly quartz rich, and crop out as both thin (,1.0 m) and thick (up to 4.0 m) beds which are laterally continuous over a range of metres. Internally, the beds are parallel laminated to massive, although the thinner beds also show evidence of trough and parallel cross lamination. Some of the thicker beds appear to have a channel-type form, cutting down into the underlying finer-grained sediments. Additionally, some of the thicker sandstones show evidence of loading at the base. Associated conglomerates are mainly clast supported, with a range of clasts including sedimentary (sandstones, siltstones, mudstones),

190

T. MCCANN ET AL.

metamorphic (mica schists) and magmatic (granites, volcanics) lithic fragments, in a patchy carbonate cement. Cements in the sandstones are also calcareous (including dolomite). Rare fossils (e.g. bryozoa) are found in the sandstones. Plant fragments are common throughout the siliciclastic sediments, being concentrated both at the bed bases and also as discrete layers within the beds, providing evidence of bed amalgamation and/or variations in flow energy within individual flows. Interpretation. The relatively coarse-grained nature of the sediments would suggest that they were generally deposited in a high energy setting, probably in a coastal environment. The thick channel sandstones are interpreted as part of a (possibly multistorey) channel complex deposited in a deltaic setting. The thinner cross-bedded sandstones may indicate periods of wave flooding of the delta surface. Such an interpretation would correlate with the lateral outcrops of the Bugunja Fm which, to the east of the present area, are predominantly continental. The succession here would thus represent fluvial or fluvial-wave interaction deltaic setting.

Labina-Malka Zone: Psebai Formation (Fig. 6 – Sites 31, 32, 34, 37; Fig. 3, Table 1) Description. This Pliensbachian –Aalenian-age formation comprises a mixed siliciclastic-calcareous succession of mainly mudstones, siltstones, sandstones and limestones. The mudstones and siltstones mainly occur as thin and very thin (sensu Stow 2006), laterally continuous beds. Mud pellets and possible shelly fragments (now diagenetically altered) are present within the beds. Thin sandstones range in thickness up to 5.0 cm and are internally trough and parallel laminated, with possible sequence development (analogous to the Bouma sequence). In one location, where little mudstone is present, the sandstones are quartz rich, and micaceous. Some loading was observed at the base of one parallel-laminated bed. The sandstones contain a range of lithic fragments, including volcanics (e.g. rhyolites, basalts), sediments and rare, disaggregated fossils (algae, shelly material). The markedly subordinate calcareous material is mainly micritic limestone (biomicrite to sparse biomicrite) and dolomite. Crinoidal limestones/micrites were also noted. Interpretation. The sediments from this formation are predominantly fine-grained mudstones, occurring in laterally continuous beds (over a scale of metres), suggesting that deposition occurred in a low-energy setting, possibly distal shelf. The rare thin sandstones with internal structures analogous

to Bouma sequences may have been deposited as a result of storm surges on the shelf, resulting in the transport of coarser-grained material down slope. Similar sands have been interpreted as deposits of major storm surges (e.g. Nelson 1982). Such an interpretation would agree with the predominantly clastic nature of the formation. The rare carbonates (mainly micritic) may be related to periods of little or no clastic input resulting in the occasional superposition of a carbonate facies mosaic on the region. The presence of crinoidal limestone may suggest a depositional setting located between shallow-marine deposits and pelagic carbonates (e.g. Jach 2005).

Labina-Malka Zone: Djangur Formation (Fig. 6 – Site 33; Fig. 3, Table 1) Description. This Aalenian-age formation comprises a mixed clastic/calcareous succession where carbonates are initially predominant, and the succession becomes increasingly siliciclastic towards the top. The basal carbonates contain a range of fossil material (e.g. shells, echinoderms, corals, bryozoa, algae, forams) as well as ooids and intraclasts. Lithic fragments are also present and include sediments (siltstone, mudstone), volcanics (rhyolite?, tuff ?) and individual crystals (including embayed monocrystalline quartz). The upper part of the succession comprises very thick (up to 2.7 m), graded sandstone beds which are parallel laminated and contain weakly-developed dish structures. The bed bases (i.e. lowermost 10 – 15 cm) are sometimes grooved and may be microconglomeratic. The sandstones contain mainly magmatic lithic fragments (granitoids, volcanics) as well as rare fossils (e.g. echinoderms). Interpretation. Deposition of these sediments presumably took place in a shelf setting analogous to that of the underlying Psebai Fm, but in a more proximal setting, as suggested by the coarse-grained nature of the rocks. Individual beds contain a mixture of clastic and carbonate components; the former are angular in form, while the latter are disaggregated. This material was transported as mass flows from shallower parts of the shelf into deeper areas, possibly as result of storm activity (although the role of tectonics and/or volcanic activity cannot be ruled out).

Labina Zone: Kamennomostsky Formation (Fig. 6 – Sites 34, 36; Fig. 3, Table 1) Description. The succession in the Labina Zone differs markedly from that of the subadjacent LabaMalka Zone in that it is predominantly calcareous.

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

The unit commences with the 25 m thick Callovianage Kamennomostsky Fm which is transgressive and unconformable on the underlying lithologies (i.e. Psebai Fm). The contact with the Psebai Fm at site 34 is marked by the presence of a soil horizon with evidence of root structures. The Kamennomostsky Fm is calcareous with a variety of limestones present, including, intrabiomicrites, biomicrites, micrites and dolomites. The material ranges from fossil rich (e.g. packed biomicrites, with echinoderms, shells, corals, algae, bryozoa) with intraclasts to fossil poor (e.g. intra-biopelmicrites, where peloids are more common). Ooids are also present. Rare thin (centimetres), siliciclastic interbeds are present, while some limestones contain much siliciclastic material including quartz (10–15%), rare plagioclase or zircon grains. The individual limestones are parallel bedded, while internally they are graded or, more rarely, laminated. Interpretation. The predominance of limestones in this formation suggests that deposition occurred in a shallow-marine ramp or platform setting. A more precise definition (i.e. ramp or platform) is difficult to determine given the lack of outcrop. The nature of the sediments suggests a relatively lowenergy setting, probably a carbonate-dominated shelf. The predominance of micrite in some parts of the section suggests that energy levels varied within the environment.

Labina Zone: Gerpegem and Mezmai formations (Fig. 6 – Sites 34, 35; Fig. 3, Table 1) Description. Two Late Jurassic-age formations terminate the succession of the Labina Zone. The 147 m thick Gerpegem Fm (Sites 34, 35, Fig. 6) comprises Oxfordian-Kimmeridgian limestones, which may be dolomitized. The carbonates include both crystalline (grainstone) forms as well as micritic sediments with patchy spar cements. The micritic parts of the formation include a range of fossils (forams, bivalves, echinoderms?) as well as micritic intraclasts and peloids(?). The overlying Mezmai Fm (Site 35, Fig. 6) comprises a 200– 400 m thick succession of Tithonian-age crystalline limestones (grainstones, and recrystallized limestones). It is notable that both formations lack any terrigeneous input. Interpretation. Deposition of these two formations occurred in a low-energy carbonate rich setting. The lack of any sedimentary structures, together with the predominance of micrites and rare fossils suggests that the sedimentary environment was most probably a shallow-marine shelf setting. The

191

absence of terrigeneous input in both formations (unlike in the underlying formations) reflects the lack of source areas (low relief?, carbonate coast?, tectonic quiescence?).

Fore-Caucasus: summary The succession in the Fore-Caucasus region can be subdivided into two periods – Sinemurian – Aalenian and Callovian– Tithonian. The first of these periods is characterized by clastic shelf/ coastal sedimentation, while the second is marked by the predominance of carbonate deposition in a shelf setting. Volcanic activity was completely absent from the Fore-Caucasus area during both Jurassic periods. There is no evidence of volcanic activity in the Fore-Caucasus at this time.

Central Greater Caucasus: Psekhako-Berezov Zone (Figs 7 & 8) The rocks of this zone are Jurassic in age, and comprise three formations, namely Turovoy, Bzerpia and Tchataltapa. This succession crops out to the south of the Main Range and north of the Southern Slope (the south-vergent branches of the MCT sensu lato. being the limits between these three tectonic units; Fig. 1). It extends from the northwestern part to the southeastern parts (from sites 2 & 10 to site 12) of our investigated area. The broad extent of this zone provided us with an opportunity to examine lateral trends within the various formations. Lateral facies changes were observed, for example, in the Turovoy Fm which crops out in two widely spaced locations (i.e. Krasnaya Poliana and Otdaleni/Pshekha River areas; sites 12 & 25, respectively). (It should be noted that the stratigraphy within this zone is particularly problematic. In the study area, the formations present are as described below. However, elsewhere in this zone the formations are not so clearly defined, and additionally, locally, restricted formations are also present.)

Psekhako-Berezov Zone: Turovoy Formation (Figs 7 & 8; Sites 12, 25; Fig. 3, Table 1) Description. The lowermost unit, the Pliensbachian– Toarcian-age Turovoy Fm (c. 400 m at site 25 but thickening to 1000 m towards the SE, near site 12) is predominantly mudstone (c. 90%) with rare siltstones, fine-grained sandstones and limestones. The coarser siliciclastic beds are up to 20 cm thick, but are generally much thinner (centimetres), and tend to be lenticular and discontinuous. Internally, they are graded (sometimes with mudstone clasts), with some internal parallel- or crosslamination. The beds have sharp bases and rippled

192

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tops and are generally lenticular and discontinuous. The rare limestones are finely laminated with rare intraclasts. There are clear lateral variations within this formation. The sediments deposited at site 12 (i.e. close to the Southern Slope region) are mudstone dominated with rare silts. However, at site 25 (i.e. close to the presumed basin depocentre), the sediments contain more silts and sandstones with evidence of Bouma sequence development (Tcde, Tce). Palaeocurrents would suggest derivation from the NE. The thickness increase in the formation towards the SE (Krasnaya Poliana Region; compare Figs 7 & 8) is mainly due to the interstratified volcanics. Pliensbachian-age interstratified volcanics have been only observed at site 12 (close to the Southern Slope region; Fig. 3; Table 1) and comprise wellstratified and fine-grained pumice-rich ash-fall tuffs, mostly rhyolitic but sometimes trachy-andesitic (Table 2). The rhyolitic phenoclasts comprise quartz, biotite and alkali feldspar while the trachyandesitic phenocrysts are plagioclase, alkali feldspar and augite.

Transport was to the north–NNE. The individual beds, which are internally structureless, are both scoured and show evidence of channelling and amalgamation. Thinner sandstones and siltstones are lenticular (both continuous and discontinuous) with parallel- and cross-lamination, convolute lamination, and loading at the base. Lithic fragments present within the sandstones include volcanics (basalt, dolerite), metamorphics (mica schist) and granitoids as well as abundant siliciclastic clasts. Rare fossils (algae) are also present. The calcareous sandstones are rarer and mainly comprise micrite and dolomite with c. 50 –60% detrital siliciclastic fragments. The carbonates are probably cements. Rare dolomite/sparry calcite beds with reduced clastic contents (c. 30%) are also present. The uppermost part of the formation is muddy and shows evidence of slumping within the thin, fine-grained sandstones with a variety of internal structures (parallel and cross-lamination) and basal grooves. The ichnofossils Planolites, Helminthopsis and Cosmorhaphe have been noted. Volcanics are absent.

Interpretation. The predominantly fine-grained nature of the deposits from the Turovoy Fm, together with the laminated/massive nature of the mudstones suggests that deposition took place in a low-energy setting. The rare lenticular sandstones, which are rippled and internally parallel- or cross-laminated, provide evidence of occasional current activity. Deposition probably occurred in a shelf setting below storm wave base, with possible evidence of deeper-basin conditions in the main part of the depocentre. The rare limestones may be evidence of climatic variations, or simply related to changes in clastic input. It is also possible that limestone formation was confined to basement highs (cf. Devonian of northern Germany, McCann 1999). The volcanic events were explosive. The abundance of pumice shards in the matrix of the volcanic fall tuffs suggests subaerial eruptions or very shallow water deposition without any significant reworking of the fall deposits. The presence of ash fall only suggests that the subaerial eruptive centres were distal (hundreds of kms?) to the depositional sites. A possible source was the Lesser Caucasus Arc to the south.

Interpretation. The change between the mudstonedominated Turovoy Fm and the overlying clasticdominated Bzerpia Fm is marked. Again, deposition probably occurred in a shelf setting, although the presence of the ichnofossil Cosmorhaphe suggests a deeper-water environment which would support the idea of deepening as observed between the two Turovoy Fm sites. However, Hansen & MacEachern (2005) have reported the occurrence of Cosmorhaphe from deltaic settings (Cruziana ichnofacies), which would support the idea of a shelfal depositional environment, although, as noted above, it is generally considered to be indicative of the Nereites ichnofacies (e.g. Ksia˛z˙kiewicz 1977; McCann 1990; Fu¨rsich et al. 2007). Thus, the precise depositional setting is unclear, extending from distal shelf through to a deep-marine setting. The Toarcian was a period of volcano-tectonic quiescence in the Psekhako-Berezov Zone.

Psekhako-Berezov Zone: Bzerpia Formation (Fig. 7; Sites 20, 23; Fig. 3, Table 1) Description. The overlying Bzerpia Fm (760 m) is Toarcian in age and predominantly clastic, with rarer calcareous sandstones. The siliciclastics are mainly sandstones, siltstones and mudstones. Individual sandstone beds are very thick (up to 2.5 m), with bases which exhibit groove and flute casts.

Psekhako-Berezov Zone: Tchataltapa Formation (Fig. 7; Sites 2, 10, 19, 21, 22, 24; Fig. 3, Table 1) Description. The uppermost formation present in this zone is the Aalenian-age Tchataltapa Fm (510 m). It mainly consists of a thick (510 m) volcanic succession, the sampling sites of which are rather numerous and across the study area (Fig. 3). The volcanics (Table 2) are mostly lava flows with their related flow-breccias, or may form pillow lava edifices (site 10 only). Occasionally, extrusion breccias are present (site 24). These are more or less

Table 2. Petrography of the Jurassic-age volcanic rocks from the Western Greater Caucasus. Rock names are based on petrographic and major element data (Table 3). The geochemical signatures of the basic to igneous rocks are based on the trace element data (Table 3) Sample number

Formation

Age

03RU01

Tchataltapa

Aalenian

basanitic lava – E-MORB

03RU02A

Makarov

Early Aalenian

rhyolitic or trachytic ignimbrite

03RU02B

Makarov

Early Aalenian

03RU02C

Makarov

Early Aalenian

rhyolitic pumice ash fall tuf with accretionary lapillis rhyolitic pumice cinerite

03RU02D

Makarov

Early Aalenian

rhyolitic pumice cinerite

03RU03

Makarov

Early Aalenian

rhyolitic ignimbrite – very alterated

03RU04

Makarov

Early Aalenian

rhyolitic ignimbrite

03RU05A

Makarov

Early Aalenian

rhyolitic lapilli tuf

03RU05B 03RU06

Makarov Tchataltapa

Early Aalenian Aalenian

rhyolitic lapilli and ash tuf basaltic flow breccia – alkaline

03RU07

Tchataltapa

Aalenian

pillow margin breccia of andesitic basalt

03RU08

Mount Indyuk

Aalenian

rhyolitic lapilli tuf

Petrographic descriptions

Sampling site number (as on Fig. 3)

Porphyric, microlitic with chloritized matrix, 10% phenocrysts: (calcitized) olivine þ augite; 60% microcrysts: plagio (chloritized) Porphyric, finely vitroclastic with pumiceous frgts, 10 – 15% phenocrysts: Qz* þ AFd† þ Plag‡ Finely vitroclastic matrix with lapillis of aphyric lava and accretionnary ash lapillis (4mm dia) Finely vitroclastic matrix with lapillis of aphyric lava and accretionnary ash lapillis (4mm dia) Finely vitroclastic matrix with lapillis of aphyric lava and accretionary ash lapillis (4mm dia) ‘Matrix supported’ vitroclastic texture with numerous pumice clasts and accretionnary ash lapillis; phenoclasts: Qz þ Afd þ Biotite Vitroclastic texture with numerous pumice clasts and accretionnary ash lapillis: ‘matrix supported’/phenoclasts: Qz-rich þ well preserved ‘AFd þ Plag þ Biotite’ Pyroclastic texture with crystals (Qz þ Afd þ Plag þ Biotite) and pumice fragments in an ashy matrix Contact between ash and lapilli tufs Clastic texture within an horizon above a lava flow; pyroclasts not sorted (up to 2 cm), with shreded contours scattered in a fine matrix; clasts with a trachytic porphyritic texture; plag as pheno- and micro-crysts Brecciated pillow margin; texture and mineralogy are similar to this of sample 03RU06 Very fine and well sorted vitroclastic matrix with pumice fragments and Qz þ Afd þ Plag þ Biotite phenoclasts

Site 2

Site 3 Site 3 Site 3 Site 3 Site 3 Site 3

Site 4 Site 4 Site 10

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

Rock name (composition, eruptive dynamic) – trace element geochemical signature

Site 10 Site 11 193 (Continued)

194

Table 2. Continued Sample number

Formation

Age

Rock name (composition, eruptive dynamic) – trace element geochemical signature

Tchataltapa

Aalenian

basaltic lava (alterated), alkaline

03RU10

Makarov

Early Aalenian

rhyolitic ash fall tuf

03RU11

Atchichkh

Bajocian

trachy-basaltic lava, alkaline

03RU12A 03RU12B

Turovoy Turovoy

Pliensbachian Pliensbachian

rhyolitic cinerite rhyolitic fine and coarse ash tuf

03RU12C

Turovoy

Pliensbachian

trachy-andesitic pumice cinerite – calc-alkaline

03RU12D

Turovoy

Pliensbachian

rhyolitic fine and coarse ash tuf

03RU12E

Turovoy

Pliensbachian

rhyolitic fine and coarse ash tuf

03RU13

Porphyritovaya Series

Bajocian

basaltic lapilli tuffite

03RU14A

Porphyritovaya Series Porphyritovaya Series Porphyritovaya Series

Bajocian

basaltic ash tuf

Bajocian

trachy-andesitic ash tuf – calc-alkaline plagio & augite trachy-basalt – E-MORB

03RU14B 03RU15

Bajocian

Sampling site number (as on Fig. 3)

Porphyric microlitic with a chloritized/ calcitized matrix; phenocrysts (1 –5 mm) of sometimes fresh olivine þ numerous more or less preserved plag þ rare but preserved augite; found as a pebble on the site 11; analog to the Tchetaltapa Fm 03RU01 or 03RU06 Very fine and well sorted vitroclastic matrix with pumice fragments and Qz þ preserved ‘Afd þ Plag þ Biotite’ phenoclasts Porphyric microlitic with a chloritized/ calcitized matrix; phenocrysts (1 –5 mm) of numerous more or less preserved plag þ often preserved augite þ altered olivine; microcrysts of augite Fine to very fine ash matrix Fine to very fine ash matrix with numerous pumice clasts and phenoclasts of Qz þ Fds§ þ Biotite Fine to very fine ash matrix with numerous pumice clasts and phenoclasts of Plag þ Afd þ augite; important schistosity Fine to very fine ash matrix with numerous pumice clasts and phenoclasts of Qz þ Fds þ Biotite; important schistosity Fine to very fine ash matrix with numerous pumice clasts and phenoclasts of Qz þ Fds þ Biotite; important schistosity Pyroclasts of lava and Plag crystals within a sparitic carbonated matrix ¼ ash and lapilli fall within a sediment Pyroclasts of Plag crystals within a ashy matrix

Site 11

Ultra fine and laminated ashy matrix

Site 16

Porphyric poecillitic to microlitic texture; phenocrysts of plag and magnetite; poecilitic augite; chloritized matrix by place

Site 16

Site 11 Site 12

Site 12 Site 12 Site 12 Site 12 Site 12 Site 16 Site 16

T. MCCANN ET AL.

03RU09

Petrographic descriptions

Porphyritovaya Series

Bajocian

plagio & augite basaltic trachy-andesite – E-MORB

03RU16B

Porphyritovaya Series Porphyritovaya Series Porphyritovaya Series Porphyritovaya Series Porphyritovaya Series

Bajocian

16 & 15 contact

Bajocian Bajocian

olivine, augite, plagio basaltic trachy-andesite – calc-alkaline rhyolitic pumice ash fall

Bajocian

basaltic pumice and crystal ash fall

Bajocian

basaltic pumice and crystal lapilli fall basaltic trachy-andesitic pumice and crystal coarse ash fall – calc-alkaline trachy-andesitic ash to lapilli tuf – calc-alkaline

03RU17 03RU18A 03RU18B 03RU18C 03RU19

Porphyritovaya Series

Bajocian

03RU20

Porphyritovaya Series

Bajocian

03RU21

Porphyritovaya Series Porphyritovaya Series

Bajocian Bajocian

trachytic pumice ash fall – calc-alkaline augite basalt

Porphyritovaya Series Porphyritovaya Series Porphyritovaya Series

Bajocian

basaltic ash fall tuf

Bajocian

basaltic ash fall tuf

Bajocian

olivine & augite pillow basalt

03RU22

03RU23A 03RU23B 03RU24A

Porphyritic texture in a chloritized matrix; phenocrysts of plag (4 – 5 mm) with always double twins þ rare augite and skeletal magnetite Contact between poecilitic lava. and porphyric lava Porphyritic texture; order of crystallization: olivine, then augite, then plag Only very fine vitroclasts of rhyolitic pumice Very fine vitroclasts of basaltic pumice and phenoclasts of plag, olivine and augite Lapilli size vitroclasts of basaltic pumice and phenoclasts of plag, olivine and augite, without any matrix; subaerial deposits Very fine vitroclasts of basaltic pumice and phenoclasts of plag, olivine and augite Lapilli size vitroclasts þ accretionary lapilli of basaltic pumice and phenoclasts of plag, olivine and augite Only very fine vitroclasts of basaltic pumice, well stratified Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of numerous often preserved augite þ rare altered olivine; microcrysts of augite; vesicules filled with calcite Only very fine vitroclasts of basaltic pumice, well stratified Only very fine vitroclasts of basaltic pumice, well stratified Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours

Site 16

Site 16 Site 16 Site 17 Site 17 Site 17 Site 17 Site 17 Site 17 Site 17

Site 17 Site 17 Site 17

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

03RU16A

(Continued) 195

196

Table 2. Continued Sample number

Formation

Age

Rock name (composition, eruptive dynamic) – trace element geochemical signature

Porphyritovaya Series

Bajocian

olivine & augite pillow basanite – alterated – calc-alkaline

03RU24C

Porphyritovaya Series

Bajocian

olivine & augite pillow basanite – alterated – calc-alkaline

03RU25

Porphyritovaya Series

Bajocian

olivine & augite pillow basalt

03RU26

Porphyritovaya Series

Bajocian

olivine & augite pillow basalt

03RU27

Porphyritovaya Series

Bajocian

olivine & augite pillow basanite – alterated – calc-alkaline

03RU28

Porphyritovaya Series Tchataltapa

Bajocian

basaltic hyaloclastites

Aalenian

plagio, olivine & augite basalt – N-MORB

03RU29

Sampling site number (as on Fig. 3)

Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Porphyric microlitic with a more or less vesiculated chloritized matrix; phenocrysts (1 –5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Hyaloclastitic texture; numerous ‘vitreous’ flams between crystals Porphyritic intersertal to poecilitic texture; phenocrysts of plag, rare olivine, xenomorph preserved augite; intertitial matrix from place to place

Site 17

Site 17

Site 17

Site 17

Site 17

Site 17 Site 19

T. MCCANN ET AL.

03RU24B

Petrographic descriptions

Tchataltapa

Aalenian

plagio, olivine & augite basalt – N-MORB

03RU31

Tchataltapa

Aalenian

03RU32

Makarov

Early Aalenian

plagio basaltic trachy-andesitic explosive breccia – N-MORB þ Ta anomaly rhyolite lava

03RU33

Makarov

Early Aalenian

rhyolitic flow breccia

03RU34

Makarov

Early Aalenian

dacitic lava

03RU35

Tchataltapa

Aalenian

trachy-dacitic lava flow breccia

03RU36A

Tchataltapa

Aalenian

andesitic lava

03RU36B

Tchataltapa

Aalenian

dacitic lava, calc-alkaline

03RU37

Tchataltapa

Aalenian

basaltic trachy-andesitic lava, N-MORB

03RU39

Tchataltapa

Aalenian

andesitic extrusion breccia, N-MORB þ Ta anomaly

03RU40

Tchataltapa

Aalenian

andesitic extrusion breccia

Site 19

Site 19 Site 19 Site 19 Site 19 Site 21 Site 21

Site 21

Site 22 Site 24

Site 24

197

Porphyritic intersertal to poecilitic texture; phenocrysts of plag, rare olivine, xenomorph preserved augite; intertitial matrix from place to place Fragment of basalt with plagio in a porphyritic and microlitic and more or less vesiculated texture Porphyritic and vesiculated texture; matrix previously vitreous; phenocrysts of Qz, altered Fds Pyroclastic texture; not sorted; flow breccia of the lava flow 32 Porphyritic and vesiculated texture; matrix previously vitreous; phenocrysts of Qz, altered Fds; no garnet as seen at sample scale Pyroclastic texture; not sorted; flow breccia of the lava flow 37 Porphyric (30%) microlitic (rare) with a more or less vesiculated chloritized matrix; phenocrysts (1 – 5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Porphyric (30%) microlitic (rare) with a more or less vesiculated chloritized matrix; phenocrysts (1 – 5 mm) of chloritized olivine (5%), preserved augite (15%)þ; microcrysts of augite; vesicules filled with calcite and hematite on their contours Intersertal texture; phenocrysts of plag þ magnetite (60%); interstitial chloritized matrix (40%) Pyroclastic texture; pyroclasts not sorted (1 –8 mm) with shreded contours scattered in a finegrained microcristalline siliceous matrix; clasts with a microlitic and vesiculated texture; plag as micro-crysts Pyroclastic texture; pyroclasts not sorted (1 –8 mm) with shreded contours scattered in a finegrained microcristalline calcitic and siliceous matrix (hydrothermalism); clasts with a microlitic and vesiculated texture; plag as micro-crysts

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03RU30

(Continued)

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Table 2. Continued Sample number

Formation

Age

Rock name (composition, eruptive dynamic) – trace element geochemical signature

Tchataltapa

Aalenian

andesitic extrusion breccia

03RU42

Tchataltapa

Aalenian

clast from a plagio & olivine trachy-andesitic lava – E-MORB

03RU43

Tchataltapa

Aalenian

andesitic extrusion breccia – E-MORB

*Qz, Quartz. † Afd, Alkali feldspar. ‡ Plag, Plagioclase. § Fds, Alkali feldspar & Plagioclase.

Sampling site number (as on Fig. 3)

Pyroclastic texture; pyroclasts not sorted (1 –8 mm) with shreded contours scattered in a finegrained microcristalline hematitized matrix (hydrothermalism); clasts with a microlitic and vesiculated texture; plag as pheno- and micro-crysts Fragment of basaltic lava flow within the extrusion breccia; porphyric and vesiculated (calcite inside) texture; phenocryst of plag and microcrysts of olivine; crystals surrounded with a hematite crust Pyroclastic texture; pyroclasts not sorted (1 –8 mm) with shreded contours scattered in a finegrained microcristalline calcitic and siliceous matrix (hydrothermalism); clasts with a microlitic and vesiculated texture; plag as micro-crysts

Site 24

Site 24

Site 24

T. MCCANN ET AL.

03RU41

Petrographic descriptions

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

vesiculated, mostly porphyritic (plagioclase, olivine, augite) and microlitic. Their interstitial matrix is sometimes chloritized or/and calcitized (sites 2, 10, 24). They are compositionally variable (Table 2) depending on the sampling site, and may be basaltic to andesitic basaltic (sites 2, 10, 11, 19), mainly andesitic to trachyandesitic (sites 22, 24), or andesitic to trachydacitic and dacitic (site 21). Associated with the volcanic rocks are great thicknesses of dark grey to black mudstones (up to tens of metres thick, e.g. site 24). These mudstones are massive to laminar. Contacts with the volcanics may be baked (as seen at sites 2, 21). Interpretation. The Aalenian was a period of marked eruptive activity in the Psekhako-Berezov Zone. The volcanism was effusive, mainly subaerial and sometimes subaqueous. In the latter, alteration of the interstitial glass was most likely due to interaction with sea water. The presence of extrusive breccias provides evidence of the explosive character of the andesitic magmas. The absence of any significant sedimentary units within the volcanic succession may suggest that volcano-tectonic activity was dominant (although it may have been relatively short in duration). While eruptions of the basic magmas predominated, those of acid magmas were also important. The mudstones were deposited in a low-energy setting, possibly deeper marine (e.g. outer shelf ). Such an interpretation would agree with the idea of marine extrusion of the lavas (i.e. presence of pillow basalts). However, the idea that some of the lavas were subaerially extruded would suggest that at least part of the mudstone-rich formations were deposited close to, or indeed, occasionally above, sea level, possibly in a lagoonal setting.

Central Greater Caucasus: Goitkh Zone (Fig. 7) The Goitkh Zone, which is one of the most complex zones in the area, crops out in the area located to the north of Tuapse. The succession is a relatively restricted one, comprising only the Aalenian and Bajocian. The uppermost part of the Goitkh Zone is separated from the lower part by an unconformity which is either tectonic or stratigraphic in nature. Six formations are present in this area (Makarov, Naujni, Mount Indyuk, Pchich, Sosnov, Kochkanov), four of which were examined.

Goitkh Zone: Makarov Formation (Fig. 7; Sites 1, 3, 4, 11, 19; Fig. 3, Table 1) Description. This formation (1930 m), which is Aalenian in age, comprises a succession of

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interbedded sediments and volcanic rocks. The former are predominantly mudstones (sometimes silty), with rare siltstones, and siderite nodules. Rare limestone lenses, up to 60 cm thick, were noted which may contain a fauna of ammonites, pelycypods and bivalves, as well as micritic intraclasts. The lack of detailed sedimentary structures is largely a function of the poor preservation of the section. Thick volcanic piles form spectacular cliffs (several decimetres high) across the landscape. The volcanics are rhyolitic to dacitic, either aphyric or porphyritic with phenoclasts of quartz, alkali feldspar, plagioclase, and biotite (Table 2). Both pyroclastic deposits (sites 3, 4 and 11) and lava flows with their related flow breccias (site 19) have been recognized. The pyroclastic rocks include numerous pumiceous fragments and shards, the internal structure of which is typical of subaerial pyroclastic flows or related ash-fall deposits (Table 2). Numerous accretionary ash lapillis (up to 4 mm) are included in the fall deposits. Interpretation. The lower Aalenian-age rhyolites of the Makarov Formation were emplaced as subaerial ‘nue´es ardentes’ (sites 3, 4, 11) or as subaerial lava flows (site 19). Evidence for subaerial deposition of the pyroclastics is provided by the observed structures within the pyroclastic flows as well as the presence of accretionary ash lapillis in the fall deposits and the abundance of pumice shards present in their matrix. However, it is also possible that at least part of the pumice flow and fall deposits were deposited in very shallow waters where there was no significant reworking of the deposits. Evidence for the subaerial origin of the lava flows is provided by their association with flow breccias which formed a cooling surrounding cortex (5– 20 cm in thickness) at the base and top of the flows. The fine-grained nature of the sediments together with the relative absence of structures suggestive of wave or current activity would suggest deposition in a low-energy setting, possibly below storm-wave base. However, as previously noted a subaerial (or possibly shallow water) environment is suggested by the volcanics. The sediments were deposited in a low-energy, mudstone-dominated setting – possibly coastal flats or a mud-rich fluvial floodplain. In such a setting the rare carbonates may represent marine incursions.

Goitkh Zone: Naujni Formation (Fig. 7) Description. The overlying Aalenian-age Naujni Fm (430 m) did not crop out in the study area. However, the succession is predominantly siliciclastic, comprising sandstones and mudstones, with rare interbedded rhyodacites (Geological Map of the Russian Federation 2001).

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Interpretation. The Naujni Fm was possibly deposited in a marine setting.

Goitkh Zone: Mount Indyuk Formation (Fig. 7; Sites 5, 6, 7, 11; Fig. 3, Table 1)

related to the increased volcanic activity in the area. The ash fall rhyolitic tuffs interstratified with the sediments could be subaerial in origin or they were emplaced in very shallow water; in both cases deposition occurred distal to the eruptive centres.

Description. This volcano-sedimentary formation (1600 m) is late Aalenian in age. The sediments present include mudstones, siltstones, sandstones and rare conglomerates. The mudstones, which dominate the lower part of the formation, are mainly structureless, although microscopically they sometimes appear pelleted, while rare siderite nodules are also noted. Interbedded silt- and fine-grained sandstones occur within the mudstone-dominated part of the Mount Indyuk Fm. The individual sandstone beds are lenticular and c. 15 –20 cm thick. The sandstones and conglomerates contain a variety of lithic fragments including sediments, volcanics (basalt, dolerite, rhyolite, tuff ), metamorphics (mica schist, slate) and rare granites, as well as fossils (bryozoa, algae, coral(?), crinoids). Sandstones are concentrated in the lower upper part of the formation, where they are up to c. 80 cm thick. Sandstone beds, which are mainly sharp based, have a variety of bottom structures, including loads, grooves, flute marks (palaeocurrent directions to the east) while fossils (e.g. bivalves, cephalopods, plants) and ichnofossils (e.g. Palaeophycus, Thalassinoides) are also present. Internally, the sandstones are parallel- to cross-laminated (sometimes in couplets with siderite) with some possible hummocky crossstratification. The upper parts of some beds may be convolute laminated. Rhyolitic very fine grained ash fall tuffs (Table 2) are interstratified within the sedimentary sequence of the Mount Indyuk Fm.

Goitkh Zone: Pchich Formation (Fig. 7; Sites 6, 8, 9; Fig. 3, Table 1)

Interpretation. The sedimentary succession of the Mount Indyuk Fm is dominated by fine-grained clastics which would suggest that the depositional setting was a low-energy one. The presence of possible hummocky cross stratification within the sediments of the Mount Indyuk Fm would suggest that deposition occurred between the storm and the fairweather wave bases in a shelf setting. Such a setting would be supported by the presence of siderite. Similar occurrences of siderite have been reported from offshore shelf sediments deposited under aerobic conditions (cf. McKay et al. 2006). The geometry of the sandstone lenses (i.e. the discontinuous nature of the individual beds) as well as their internal structure suggests that they may have been deposited as sand lobes on top, or adjacent to, the dominant shelf mudstones. Deposition was from traction currents and/or mass flows that may have been generated as a result of storm events on the shelf. It is also possible that some of these mass flows may have had a tectonic component

Goitkh Zone: Sosnov & Kochkanov formations (Fig. 7)

Description. The overlying Pchich Fm (1280 m) straddles the Aalenian-Bajocian boundary and comprises mainly siliciclastics (mudstone, sandstone, conglomerate) as well as limestones. The sandstones, which are present within the very thick mudstone units (up to 8.0 m) tend to be thin (,10 cm) and are parallel to cross-laminated. Individual sandstones beds may show evidence of amalgamation (e.g. concentrations of mudstone rip-up clasts, or mudstone partings). Lithic fragments within the coarser sandstones include volcanics (basalt, dolerite), sediments, metamorphics (slate) and granites. Fossils, including algae, forams, and bivalves, are also present. The mudstones show some evidence of parallel lamination. Rare limestones are mainly micritic and intraclasts may be present (intramicrite), although one sample includes volcanic lithic fragments (basalt). Volcanics are absent. Interpretation. The fine-grained nature of the Pchich Fm suggest that deposition occurred in a low-energy environment. The depositional setting was probably similar to that of the underlying Mount Indyuk Fm. Of interest, however, is the presence of rare limestones, which would indicate possible changes in the marine chemistry or temperature, or a reduction in clastic supply to the region.

Description. The uppermost formations of this zone in the Middle Jurassic succession (the Goitkh Zone extends up into the Cretaceous) comprise the Upper Bajocian-age Sosnov & Kochkanov formations (660 m), which were not examined as part of this study. However, they comprise a mixed carbonatesiliciclastic succession including mudstones with siderite, limestones, sandstones and conglomerates (up to 30 m thick) and it is probable that the depositional setting was similar to that of the underlying formations (Fig. 7).

Central Greater Caucasus: Abino-Gunai Zone (Fig. 7) This zone comprises four formations (Zeitun, Yablon, Pshekha, and Pikhtar) which range in age

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

from Bajocian to Tithonian, with a gap in the stratigraphic record coincident with the upper Bathonian (Fig. 7). Volcanics have not been recorded from this zone.

Abino-Gunai Zone: Zeitun Formation (Fig. 7; Site 26; Fig. 3, Table 1) Description. The succession commences with the Zeitun Fm (40–400 m) which is of Bajocian age and comprises sandstones and mudstones. The sandstones range in thickness from 3 to 40 cm and bed bases may be loaded. Internally, the beds may be internally parallel- or cross-laminated. Rare plant fragments are noted, as are microfossils (?forams) and a range of lithic fragments (granites, ?volcanics, sediments). Layers, which include blocks of up to c. 0.5 m diameter (although they can be up to 1.5 m) of siderite and mudstone, are found close to the base of the formation. Four of these layers have been noted, each of which is separated by a mudstone interval. Internally, the blocks are parallel laminated. Interpretation. The fine-grained nature of the Zeitun Fm suggests that deposition took place in a lowenergy setting, possible a distal shelf location. Such a setting would conform with the rare sandstones which are present. The formation containing the siderite nodules probably formed as a result of early cementation processes (although the possibility of additional, localized, transport cannot be ruled out) and the extensive siderite cementation may have been a consequence of the low net accumulation rate of sediment in a possibly transgressive sequence.

Abino-Gunai Zone: Yablon Formation (Fig. 7; Site 27; Fig. 3, Table 1) Description. The post-Callovian unconformity part of the Abino-Gunai Zone comprises two sub-zones (Fig. 9). In the Abino-Gunai Zone, the Callovianage Yablon Fm (440 m) supposedly rests unconformably on the Zeitun Fm. However, structural measurements along the section from the Zeitun Fm through to the Yablon Fm have revealed the absence of an unconformity between the formations. Instead, they both show a common (and constant) strike (N100) with a dip of 358 to the north. The Yablon Fm sediments are limestone-dominated and include peloids, intraclasts and fossils (coral, echinoderms, algae, forams, bivalves, crinoids, ?ooids). The limestones may be classified as wackestones and pel-intrasparite/intra-pelsparite or intrasparite. Interpretation. The marked difference between the Yablon Fm and the underlying Zeitun Fm is the predominance of limestones in the former, which

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resulted due to a major change in the depositional conditions. The superposition of a carbonate mosaic on the clastic shelf setting records the marked reduction in clastic supply to the region as well as possible temperature and/or depth changes. Sedimentation was in a low- to medium-energy setting.

Abino-Gunai Zone: Pshekha and Pikhtar formations (Fig. 7; Sites 18, 28; Fig. 3, Table 1) Description. The uppermost formations in this zone are the Pshekha Fm (230 m) which is Oxfordian – Kimmeridgian in age, and the Pikhtar Fm (800 m) which is, Tithonian– Early Berriasian in age. The formations are both limestone dominated, with mainly intraclasts (including micrite, biomicrite, silt/mud), fossils (algae, bivalves, forams, echinoderms), peloids and ?ooids present. The rocks tend to be predominantly micritic and represent both packed and sparse micrites. Internally, parallel lamination may be observed. The limestones can be classified as mudstones or wackestones (biosparite, intrasparite, pelsparite, intra-biosparite). Rare coarser-grained beds are poorly graded and show some evidence of parallel lamination. Interpretation. The micritic nature of the sediments in this formation would suggest that they were deposited in a shallow-marine, low-energy shelf setting. From exposures it is clear that the finergrained limestones were the background sedimentation (forming beds up to 1.0 m thick), while the coarser material, which is lenticular in form, was probably deposited from mass flows (?storm deposits). One particular bed, close to the top of the exposure, is clearly a debris flow and is overlain by sandy limestones which pass upwards into the sandier intervals which form part of the overlying Berriasian succession.

Central Greater Caucasus: summary In total, 13 formations in three different ‘zones’ extending from the Pliensbachian to the Bajocian were examined across the Central Greater Caucasus. In the Psekhako-Berezov Zone there is evidence of a shallowing-upward succession from Pliensbachian through to Aalenian times. In the adjacent Goitkh Zone the early Aalenian is also coastal/ nearshore, but over time the succession deepens to a shelfal one (Aalenian through to Bajocian). In the Abino-Gunai Zone, the succession begins in Bajocian times in a clastic shelf setting. The overlying Oxfordian to Berriasian-age formations are predominantly carbonates, deposited in a shelf environment. Volcanic activity occurred in the PsekhakoBerezov and Goitkh zones, and was absent in the

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Fig. 9. Sandstone modal data for the Triassic succession, Fore-Caucasus. QFL, quartz-feldspar-lithic fragments; QmFLt, monocrystalline quartz-feldspar-total lithic fragments; SrfCrfFo, siliciclastic lithic fragments-carbonate lithic fragments-fossils. Tectonic discrimination fields defined by Ingersoll & Suczek (1979) and Dickinson et al. (1983).

Abino-Gunai Zone. Magmatic activity commenced in the Pliensbachian with rhyolitic and explosive eruptions in the Psekhako-Berezov Zone. Following a period of quiescence in the Toarcian, activity recommenced in the Aalenian with rhyolitic and explosive eruptions in both the Psekhako-Berezov and Goitkh zones; a large volume of basaltic (sensu lato) lava-flows erupted in the PsekhakoBerezov Zone (Tchataltapa Fm). The Makarov, Mount Indyuk and Tchataltapa formations are coeval and are probably lateral equivalents. Evidence would suggest that the sediments of the Mount Indyuk Fm were derived not only from exhumed basement highs but also from the adjacent basic volcanoes of the Tchataltapa Fm. This might explain the large amount of volcanic pebbles from the Tchataltapa Fm that is observed within the sedimentary succession of the Mount Indyuk Fm (03RU09; Table 2). Volcanic activity ceased at the beginning of the Bajocian across the Central Greater Caucasus.

Southern slope: Krasnaya Poliana area The southernmost examined succession in the Western Greater Caucasus region is located close to the border with Georgia. It was not possible to work in the border region due to the political situation and thus the fieldwork area was restricted. The stratigraphy of the area extends from the Sinemurian through to the Tithonian (Fig. 8). The examined Lower and Middle Jurassic succession in the Krasnaya Poliana area does not crop out continuously, instead being subdivided into three main zones: the Krasnopoliana and AbkhazoRacha zones of the Southern Slope, and the southeastern extension of the previously described Psekhako-Berezov Zone of the Central Greater Caucasus (containing the Turovoy Formation at site 12, see above). Our work was focused on the formations within the zone of the Southern Slope, that is, in the Krasnopoliana (Sites 12, 13, 14, 15; Fig. 3) and Abkhazo-Racha (Sites 16, 17; Fig. 3)

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

zones (Fig. 8). As previously noted the PsekhakoBerezov Zone belongs to the Central Greater Caucasus and was thrust over the Krasnopoliana Zone of the Southern Slope along the southern branch of the MCT. The Bathonian-age Betag Fm (500 m) of the Krasnopoliana Zone, which overlies the Atchichkh Fm, was not examined (Fig. 8).

Krasnopoliana Zone: Estasadok Formation (Fig. 8; Sites 13, 14, 15; Fig. 3, Table 1) Description. The Estasadok Fm (600 m), which is Sinemurian– Middle Pliensbachian in age, comprises siliciclastic sediments, which are the oldest recorded in the Krasnaya Poliana area. These are mainly structureless mudstones which are extensively weathered, and contain thin to very thin, lenticular and laterally discontinuous sandstone beds. Interpretation. During Sinemurian to Middle Pliensbachian times, the predominance of mudstones in the depositional environment, together with the rare occurrences of thin, laterally discontinuous sandstones, suggest that sedimentation occurred in a low energy setting. This probably took place in a sheltered, nearshore, or floodplain setting. The stability of the Krasnopoliana Zone is evidenced by the lack of volcanic activity.

Krasnopoliana Zone: Tchvejips & Illaryonov formations (Fig. 8; Sites 12, 14; Fig. 3, Table 1) Description. The overlying Tchvejips (700 m; Upper Pliensbachian) and Illaryonov (1000 m; Toarcian) formations both closely resemble the underlying Estasadok Fm, in that they are mudstone dominated with rare intercalated thin to very thin, laterally discontinuous sandstone beds. Volcanics are absent. Interpretation. During Upper Pliensbachian to Toarcian times, the depositional setting of both of these formations was probably similar to that of the underlying Estasadok Fm given the similarity of the facies present (i.e. deposited in a low-energy setting), and a nearshore/floodplain setting is probable. The ongoing stability of the area is evidenced by the lack of volcanic activity.

Krasnopoliana Zone: Anchkhoi Formation (Fig. 8; Sites 12, 14) Description. This Aalenian-age formation (700 m) is similar to the other formations within this zone in that mudstones (generally shaly and internally structureless) are predominant and the thin to very thin sandstones are rare and laterally discontinuous.

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The sandstones contain some sedimentary lithic fragments. Volcanics are absent. Interpretation. During the Aalenian times, deposition took place in a low-energy environment, probably nearshore or in a continental floodplain.

Krasnopoliana Zone: Atchichkh Formation (Fig. 8; Site 12; Fig. 3, Table 1) Description. The Bajocian-age Atchichkh Fm (1400 m) consists mainly of siliciclastic rocks, with rare limestones and episodic interstratified lava flows. The mudstones, which locally predominate, are structureless to finely laminated. Interbedded sandstones and siltstones are graded, of variable thickness (up to 1.0 m) and laterally continuous or channellized. Ichnofossils (Planolites, Thalassinoides) are present. Associated limestones are mainly micritic and laminated and include rare fossils (?forams, ?algae, calcispheres). This formation is coeval with the Porphyritovaya Series (2000 m, see below). Interstratified lavas flows are trachy-basaltic (Table 2) and porphyric microlitic. Phenocrysts (1–5 mm) comprise numerous well preserved plagioclases, as well as preserved augites and altered olivines. The microcrysts are augite only. The chloritized/calcitized matrix could be due to sea water alteration of the glassy matrix. This unit is coeval with the Porphyritovaya Series (Abkhazo-Racha Zone; see below). Interpretation. Deposition of the Atchichkh Fm probably occurred in a shallow-marine setting close to shore, but below fairweather wave base. For the first time in the Jurassic there is evidence of volcanism in the Krasnopoliana Zone. Eruptions were effusive but rather limited in terms of volume. Based on the observed alteration of their glassy matrix, the lava flows may have been emplaced in a shallow marine setting suggesting also that the eruptions themselves were also possibly submarine. Compositionally, the volcanics were basaltic (sensu lato) only.

Abkhazo-Racha Zone: Porphyritovaya Series (Fig. 8; Sites 16, 17; Fig. 3, Table 1) Description. The Bajocian-age Porphyritovaya Series (2000 m) is mainly composed of volcanics, in marked contrast to the coeval Atchichkh Fm. It commences with lava flows followed by pyroclastic fall deposits which are interbedded with rare finegrained siliciclastics. Flows (some hundreds of metres in total) were mainly emplaced as pillows, and these pass upwards (and laterally) into lenticular deposits of hyaloclastites. The accumulation of these latter units around the pillow edifices (some

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metres thick) suggests a low-energy subaqueous setting. Lavas are mainly basaltic to andesitic basaltic (Table 2). They are mostly vesiculated and coarsely porphyritic (augite, olivine, plagioclase; Table 2). The overlying fall deposits are broadly basaltic to andesitic and sometimes trachytic or rhyolitic. They comprise successions, hundreds of metres thick, of well stratified and often laminated fine and coarse pumiceous ash and lapilli tuffs including numerous accretionary lapillis. Tuffites are rare. The associated interstratified siliciclastic sediments are finely laminated and relatively thin (millimetres thick), although individual beds may be laterally continuous. The sediments have a marked red colour, suggesting that they may be continental in origin. Interpretation. The Bajocian in the Abkhazo-Racha Zone was a period of very intense volcanic activity, predominantly effusive rather than explosive. The pillow edifices were probably built up in a shallowwater setting. The presence of numerous accretionary lapilli in the overlying fall deposits suggests subaerial explosion, while the abundant pumice shards within an ash matrix demonstrates both the pyroclastic origin of the explosions and the subaerial deposition. Thus, the volcanic evidence would suggest that there was uplift in the area, with the stratigraphically lower pillows being emplaced in a shallow water setting. The original interpretation of the sediments of the Porphyritovaya Series suggested that they were nearshore, but the lack of structures indicative of wave reworking, together with the low-energy level of the environment and the subaerial deposition of great volumes of fall tuffs, suggests that deposition occurred either in a nearshore marine setting, possibly lagoonal, or in a lacustrine or floodplain setting.

Tchvejips Zone: Aibga & Agepst formations (Fig. 8; Site 17; Fig. 3, Table 1) Description. These two formations crop out in the Tchvejips Zone, but the outcrops are extremely poor. The lowermost of the two formations (Aibga Fm) is Oxfordian-Kimmeridgian in age and comprises mainly micritic limestones. This is overlain by the Tithonian-age Agepst Fm which is calcareous in nature. The limestones are micritic and often laminated. Secondary dolomites also occur. Volcanics are absent. Interpretation. The calcareous nature of these two formations, together with the fine-grained nature of the sediments suggests that deposition occurred in a low-energy setting, probably a shelf from Oxfordian to Tithonian times. The often finely

laminated nature of the micrites would support such a setting. Of interest, are the lack of bioturbation and the rare presence of fossils, which may be evidence of increased salinities.

Southern slope: summary A total of eight different formations from three different ‘zones’ were examined in the Southern Slope region, the majority of which were in the Krasnopoliana Zone. Here the succession, extending from Sinemurian through to Bajocian, is mainly a coastal one, where the sediments were deposited in a nearshore/coastal plain setting. The adjacent Abkhazo-Racha Zone (Bajocian) was deposited in a similar setting, while the sediments from the Tchvejips Zone (Oxfordian–Tithonian) were deposited in a carbonate shelf or nearshore/brackish environment. Volcanic activity was confined to the Bajocian in both the Porphyritovaya Series (Abkhazo-Racha Zone) and the Atchichkh Fm (Krasnopoliana Zone). Both of these units differ in terms of the relative amounts of volcanics and sediments present. In the Porphyritovaya Series volcanic activity was marked, with both effusive and explosive eruptions, while in the Atchichkh Fm, the activity was less intense, with only effusive eruptions occurring. In contrast, the sediments of the former were predominantly continental (where present), whereas in the latter they are generally shallow marine (and much more common). Compositionally, the volcanics were mostly basic, and rarely acidic.

Petrography and modal composition of the siliciclastic formations The petrography of 227 sandstones and carbonates from the three main areas of the Western Greater Caucasus was analysed: samples were collected from 41 Triassic- to Jurassic-age outcrops throughout the belt (Fig. 3, Table 1) The sample selection criterion was to obtain representative samples from each stratigraphic formation, with the aim of sampling the coarsest grained sandstones available. Care was taken to avoid collecting weathered samples. After preliminary observations, 70 of the sandstones were selected for detailed modal analysis following the procedures of Valloni (1985). Samples with a medium grain size of ,125 mm were discarded [counting 300–500 points per thin section]. Medium- to coarse-grained sandstones were analysed. Grain parameters are those of Ingersoll & Suczek (1979), albeit modified. For sandstone classification purposes, the schemes of Folk et al. (1970) and Zuffa (1980) were used.

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

The majority of the samples were medium to well sorted, with subangular to subrounded individual grains. The matrix is generally abundant in the sandstones (range: 0.5–55%). It occurs as crushed lithic grains, small quartz grains, and phyllosilicates (primary matrix, pseudomatrix, cf. Cox & Lowe 1996). Poikilotopic, pore-filling and patchy carbonate cements (sparry calcite, micrite, dolomite) are abundant in some samples (range: 1– 55%). Cementation by quartz was less common (range: 0–14%).

Fore-Caucasus area: Triassic and Jurassic The Triassic-age samples from the Mount TkhachBelaya River area were restricted to the sandstones and conglomerates of the Maliy Tkhach Fm (Tkhach Group) (n ¼ 4) and the Sakhrai Group (n ¼ 3). The presence of quartz (Q), feldspars (F), fine-grained lithics (L) and phyllosilicates characterize these sandstones; grains are predominantly subrounded to subangular in shape. Quartz occurs as monocrystalline (Qm) and fine- to coarse-grained polycrystalline grains. Polycrystalline quartz (Qp) is subordinate (Qm . Qp). Both undulose (.58) and non-undulose varieties of Qm are present, with nonundulatory monocrystalline quartz being the most common quartz type (typically .73% of the total quartz grains). The individual grains do not, however, show a common orientation, suggesting that strain occurred in the source area. Qm grains may also be polycomposite, while some rare Qm grains show evidence of a volcanic origin (embayments). Most of the polycrystalline (Qp) grains present consist of more than three crystals. The contacts between the individual sub-grains are straight to sutured, the latter occurs more commonly. The sub-grain size is variable, even within a single composite grain of Qp. Chert (microcrystalline Qp) is commonly recognized. Inclusions are present within Qm grains and commonly include zircon and tourmaline. Plagioclase and K-feldspar are both subordinate to the quartz grains, and generally do not exceed more than 18% of the total detrital components. The plagioclase to K-feldspar (P/K) ratio is 0.55. Plagioclase grains range from large, euhedral, compositionally zoned crystals to subangular grains. Twinning is common. Feldspars may be fresh and unaltered, but more commonly, they are replaced by carbonate or altered to sericite and clay minerals. Muscovites and biotites represent the majority of the phyllosilicates, whereas chlorite is rarely present. Lithic fragments present include plutonics (granitoids), volcanics (?basalt, rhyolite, dolerite and tuffs), low-grade metamorphic clasts (e.g. quartzmica phyllites) as well as higher-grade clasts (e.g. mica schist), and sediments (both clastic and carbonate, including shales, siltstones and cherts).

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The sandstones have a carbonate cement (sometimes dolomitic). Accessory minerals include epidote, while in some samples heavy minerals (e.g. zircon) are observed. The Triassic-age sandstones of the Fore-Caucasus area can be classified as sub-arkose arenites with a mean of Q72.75%F16.09% L10.53%. The Jurassic-age samples from this region are confined to the Bugunja (n ¼ 4), Psebai (n ¼ 5) and Djangur (n ¼ 5) formations. As with the underlying Triassic-age samples, these sandstones are characterized by the presence of quartz (Qm . Qp), feldspars, and fine-grained lithic fragments and phyllosilicates. Individual grains are predominantly subrounded to subangular. Long grain contacts are common. Both undulose and non-undulose varieties of Qm are present, although the latter predominate, and no common orientation was observed. Embayed Qm has been recorded from the Bugunja Fm. Qm grains may also contain inclusions, including microcrystals of tourmaline and zircon (Bugunja Fm). Individual Qm grains may also exhibit quartz overgrowths (Psebai Fm). Qp is present in all of the samples, although Qp , Qm. Most of the Qp present consists of more than three subcrystals, where the contacts between the individual sub-grains are straight to sutured. Chert is also recorded. The Jurassic-age sandstones of the Fore-Caucasus area can be classified as subarkose to arkose arenites with rare feldspathic litharenites (Bugunja Fm). The Psebai Fm samples are richer in quartz (i.e. quartz arenites, subarkose and sublitharenites) while the Djangur Fm samples range from subarkose to lithic subarkose, as well as litharenites. The mean QFL values for the samples are Q65.42%F23.37%L11.2% (Bugunja Fm), Q89.95%F4.75%L5.3% (Psebai Fm) and Q53.93%F7.87%L38.2% (Djangur Fm). Feldspars are subordinate to the quartz grains (4.7% of detrital fraction). K-feldspar is generally the most abundant feldspar present, although this varies between the formations (e.g. Bugunja Fm, P/K ¼ 0.8; Psebai Fm, P/K ¼ 0.3; Djangur Fm, P/K ¼ 1.1). Plagioclase grains are generally euhedral in form. Microcline is also recorded (Bugunja Fm), as is spindle texture (Psebai Fm). Alteration and replacement textures are common, particularly in the K-feldspars. Carbonate cements (sometimes dolomitic) are most common, and are usually patchily distributed. Lithic fragments are predominantly sedimentary (sandstones, siltstones, mudstones), magmatics (?granitoids), volcanics (?rhyolite, tuff, ?basalt, dolerite; fragments with devitrified textures), and metamorphics (mica schist, phyllite). Fossil fragments are also present in all three formations, with the percentage of organisms increasing steadily up profile. The lowermost Bugunja Fm contains just

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bryozoa. Interestingly, these are the only fossil fragments present, suggesting that the sands were derived from an area rich in such material. In the overlying Psebai Fm the range of fossil material increases and includes algae and shelly material. The uppermost formation shows a marked increase in the percentage of fossil material present (up to 50% of the clasts present are carbonates). Fossils present include echinoderms, ?algae, corals, bryozoa, forams and shelly fragments. Ooids and intraclasts are also present. Accessory minerals in the Bugunja Fm include muscovite, zircon, epidote, and tourmaline as well as granophyre and ?myrmekite. The overlying Psebai Fm contains biotite, muscovite, chert, and chalcedony. The Djangur Fm contains muscovite, biotite, chlorite and zircon.

Central Greater Caucasus A total of 31 sandstone samples of Jurassic age were analysed from the Central Greater Caucasus area. These samples were derived from five different formations distributed across three different zones, namely, the Psekhako-Berezov (Turovoy, n ¼ 2 and Bzerpia, n ¼ 8 formations), Goitkh (Mount Indyuk, n ¼ 8 and Pchich, n ¼ 7 formations), and Abino-Gunai (Zeitun Fm, n ¼ 6) Zones. The Jurassic-age sandstones of the Central Greater Caucasus region can be classified as subarkose to lithic subarkose arenites (Psekhako-Berezov Zone), subarkose to sublitharenite with rare lithic subarkose arenites (Goitkh Zone), and subarkose to quartz arenites (Abino-Gunai Zone). The mean QFL values for the samples are Q77.38%F14.88%L7.74% (Psekhako-Berezov Zone), Q74.65%F8.75%L16.6% (Goitkh Zone) and Q90.95%F6.79%L2.2% (AbinoGunai Zone). The various sandstones are characterized by the presence of quartz (Qm . Qp); indeed, quartz is the predominant detrital clast present in the samples. Individual grains are mainly subrounded to subangular. Both undulose and non-undulose varieties of Qm are present, although the latter predominate. As with other samples from the region, no common orientation of undulosity was observed. Qm grains may contain microliths of tourmaline and rutile (Psekhako-Berezov Zone). Qp is present in all of the samples, although Qp , Qm. Most of the Qp present consists of more than three subcrystals, where the contacts between the individual sub-grains are straight to sutured. Feldspars (K . P), fine-grained lithic fragments and phyllosilicates are also present. The feldspars are subordinate to the quartz grains (7– 15% of detrital fraction). K-feldspar is generally the most abundant feldspar present, although this varies between the formations (e.g. Psekhako-Berezov

Zone, P/K ¼ 0.83; Goitkh Zone, P/K ¼ 1.1; Abino-Gunai Zone, P/K ¼ 0.7). Plagioclase grains are generally euhedral in form. Cements are generally patchy carbonates (Psekhako-Berezov Zone), although they may be more extensive (Abino-Gunai Zone), while chert cements are also observed (Goitkh Zone). The carbonates are also sometimes found as replacements (e.g. K-feldspar, Abino-Gunai Zone). Matrix is generally abundant in the sandstones. It occurs as crushed lithic grains, small quartz grains, and phyllosilicates (particularly sericite, pseudomatrix), and as epimatrix and orthomatrix. Poikilotopic, pore-filling and patchy carbonate (sparry calcite, micrite) is abundant in some samples. Lithic fragments present in the PsekhakoBerezov Zone include sediments, magmatics (granites), volcanics (basalt, dolerite, ?diorite, glassy fragments showing evidence of devitrification) and metamorphics (mica schist, phyllite, gneiss). Accessory minerals include biotite, muscovite, zircon, chlorite, pyroxene (titanaugite), chert, chalcedony, augite and rare biogenic material (algae). In the Goitkh Zone the fragments include metamorphics (mica schist, low-grade metamorphics, e.g. slate), volcanics (basalt, dolerite, trachyte), sediments and rare granites. Accessory minerals include biotite, muscovite, tourmaline, ?olivine, hornblende, granophyre, chlorite, zircon, chert, chalcedony, glauconite and rare fossil material (bryozoa, algae, coral, shelly fragments). Finally, in the Abino-Gunai Zone sediments are the most common lithic fragments followed by rare granites, ?volcanic lithics and metamorphics. Accessory minerals include biotite, muscovite, chlorite, chalcedony, chert and zircon. Rare biogenic fragments (forams) are also present. The fact that biotite is such a common accessory mineral implies a certain proximity to the source area and/or rapid deposition.

Southern slope: Krasnaya Poliana area (3 samples) The clastic samples from the Krasnaya Poliana area were taken from the lower part of the succession (in contrast, the upper part of the succession, i.e. the Tchvejips Zone, is carbonate dominated). The sampled successions include the Illaryonov (n ¼ 1), Anchkhoi (n ¼ 1), and Atchichkh (n ¼ 1) formations. The sandstones present are quartz dominated (Qm . Qp), where the individual grains are mainly subrounded to subangular. Both undulose and non-undulose varieties of Qm are found, although the latter is predominant. As in the other areas, no common orientation of undulosity was observed, confirming that all observed strain (i.e. as recorded by the undulose extinction) occurred in the source regions rather than in the depositional basins.

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

Qp is present in all of the samples, although Qp , Qm. The sandstones are subarkosic arenites. The mean QFL values for the Krasnopoliana Zone are Q90.74%F7.16%L2.1%. Feldspars and lithic fragments are also present. The feldspars are subordinate to the quartz grains (7.16% of detrital fraction); P/K ¼ 0.74. Plagioclase grains are generally euhedral. Both quartz and carbonate cements are present. Lithic grains include sediments, volcanics (basalt-both vesicular and amygdaloidal, dolerite, glassy fragments, tuff, and pumice) and igneous fragments. Accessory minerals in the Illaryonov Fm include biotite, muscovite, zircon, and pyroxene (titanaugite). The overlying formations contain less accessory minerals, mainly muscovite, biotite and zircon.

Petrofacies A series of ternary plots (cf. Dickinson et al. 1983) have been used to compare the sandstones of the Triassic and Jurassic succession in the Western Greater Caucasus. On a standard QFL petrography diagram no clear compositional trend could be observed, nor was it possible to clearly distinguish the various sand-rich depositional formations. An initial QFL of the Triassic indicates that the data plots either in the Recycled Orogenic or the Transitional Continental fields (i.e. quartz rich) (Fig. 9). This pattern is confirmed by the QmFLt plot. The predominance of quartz in the samples may be indicative of severe chemical weathering (and resultant quartz enrichment), or possibly a long transport distance. Much of the additional scatter in the data results from differences in the populations of quartz and feldspar detrital grains. Despite the degree of dispersion, an uplifted continental fragment can be suggested as the likely source area (cf. Cingolani et al. 2003). The analytical data from the Jurassic sandstones was subdivided into three time slices (Sinemurian – Toarcian, Aalenian, Bajocian –Callovian) in order to ascertain the degree of stratigraphic differentiation across the region, and whether any particular patterns could be observed. On the first series of plots (Sinemurian–Toarcian), the data plot mainly in the Recycled Orogenic field and this is confirmed by the QmFLt plot (Fig. 10). These plots also suggest that there is slight trend to an increasingly quartz-rich signature from north–south. From the QpLvmLsm plot there is a suggestion of a source in a collision orogen from the northernmost succession. Lithic fragments are predominantly sedimentary, both siliciclastic and carbonate (see below; Fig. 10). For the Aalenian, the data plots mainly in the Craton Interior field, with some indication of a more lithic-rich source in the formations located to

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the south (Fig. 11). A similar pattern (i.e. quartzose with lithics) is seen on the QmFLt plot, which also suggests a degree of arc signature for the Tchataltapa Fm. However, a degree of caution is advised here; Butler & Woronow (1986) suggested that the subdivision of the magmatic-arc province may be an artefact of the constant-sum of compositional data. On the QpLvmLsm plot the data plots mainly in the Collision Orogen source field, with little indication of change over time (although the northernmost formations show a defined pattern (Fig. 11). Lithic fragments are predominantly volcanic and sedimentary. However, in the Mount Indyuk Fm they are mainly carbonates with fossils, indicating a source rich in carbonates. The final set of plots (Bajocian –Callovian) again show a Recycled Orogen to Craton Interior trend, with the northernmost formations being more lithic rich (Fig. 12). This trend is confirmed by the QmFLt plot, which indicates that the northernmost formations are derived from Lithic Recycled sources, becoming increasingly quartz-rich towards the south. The samples were predominantly derived from a Collision Orogen source. Lithic fragments are predominantly sedimentary, with a predominance of carbonates and fossils in the northernmost formation (Djangur Fm) (Fig. 12). Overall, the siliciclastic fraction alone is insufficient to clearly distinguish between the different depositional systems and recognize any compositional trends. All of the depositional systems in the region generally have a plagioclase-to-K-feldspar ratio of ,1. A low P/K ratio has long been assumed to indicate a polycyclic origin for the clastic sediments (Blatt 1967), because of the lower chemical stability of calcium-bearing plagioclase with respect to potassium feldspar (Nesbitt et al. 1997). The relatively high proportion of monocrystalline quartz grains may be attributed to the igneous origin of the quartz grains (cf. Dal Gupta & Pickering 2008). However, there are numerous, potentially interlinked factors which may contribute to the formation of quartz-rich sandstones, including sourcearea characteristics, chemical weathering, climate, topography and orogenesis, multicycling, sediment transport and storage pathways, and diagenesis and/or leaching (see Smyth et al. 2008 for references), and so care must be taken in assigning a particular source region. The relative proportions of the detrital components indicate that the sandstones of the Western Greater Caucasus region were predominantly derived from quartz-rich sources, with subordinate magmatic and sedimentary input. Lithic grains comprise, in decreasing abundance: † Fore-Caucasus – siliciclastics (sandstones, siltstones, mudstones, carbonates), magmatics

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Fig. 10. Sandstone modal data for the Sinemurian-Toarcian succession, Western Greater Caucasus. QFL, quartzfeldspar-lithic fragments; QmFLt, monocrystalline quartz-feldspar-total lithic fragments; VrfIrfSrf, volcanic rock fragment-igneous rock fragment-sedimentary rock fragment; SrfCrfFo, siliciclastic lithic fragments-carbonate lithic fragments-fossils. Tectonic discrimination fields defined by Ingersoll & Suczek (1979) and Dickinson et al. (1983).

(granitoids, rhyolite, ?basalt, dolerite, tuff ), metamorphics (mica schist, phyllite). † Central Greater Caucasus – siliciclastics (sandstones, siltstones), magmatics (basalt, dolerite, trachyte, granite), metamorphics (mica schists, phyllite). † Southern Slope – siliciclastics (sandstones, siltstone, shale), magmatics (basalt, dolerite, glass, tuff, pumice). Thus the content of lithic fragments in the sandstones of the region is predominantly siliciclastic, with volcanics, metamorphics and plutonic rocks being secondary. The detrital components of the sandstones from the Western Greater Caucasus region indicate that they were mainly derived from a quartz-rich source with subordinate sedimentary and volcanic areas. The samples can thus be

termed sedimentoclastic sandstones (i.e. higher in sedimentary lithics) and can be classified as predominantly quartzose (mainly Qm with minor Qp and F). Quartzose sandstones are predominantly derived from deeply weathered cratonic landmasses or recycled sediments while quartzolithic can be derived from uplifted strata of fold-and-thrust belts (Dickinson 1988). For QFL and QmFLt plots to succeed in establishing the provenance and tectonic setting, they require a single homogenous source terrane (Cawood 1991). Furthermore, these plots do not categorize all types of tectonic regimes, in particular, those in strike –slip regimes where the resultant sandstones often show complex compositions due to the juxtaposition of different source area terranes (Miall 1990; Asiedu et al. 2000). As noted above, the structural setting of the Western Greater

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

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Fig. 11. Sandstone modal data for the Aalenian succession, Western Greater Caucasus. QFL, quartz-feldsparlithic fragments; QmFLt, monocrystalline quartz-feldspar-total lithic fragments; QpLvmLsm, polycrystalline quartz-volcanic & metavolcanic lithic fragments-sedimentary & metasedimentary lithic fragments; VrfIrfSrf, volcanic rock fragment-igneous rock fragment-sedimentary rock fragment; SrfCrfFo, siliciclastic lithic fragments-carbonate lithic fragments-fossils. Tectonic discrimination fields defined by Ingersoll & Suczek (1979) and Dickinson et al. (1983).

Caucasus region is characterized by strike –slip activity. This, together with the relative lack of continuity within, and between, the various formations results in a problematic basis for any detailed petrographic work. However, the results obtained here, together with the information derived from the detailed analysis of the facies and depositional settings, when integrated with the structural and other data, provides an overview of the petrographic evolution of the region. Overall, the Triassic and Jurassic sandstones of the Western Greater Caucasus are generally quartz rich and derived largely from cratonic or recycled sources. The predominant petrofacies are quartzose (dominantly Qm) derived from deeply weathered cratonic landmasses, or recycled sediments, and to a lesser extent, volcaniclastic (dominantly Lv), derived from volcanic fields of active magmatic activity. However, there are some minor petrographic differences between the Triassic and Jurassic successions. In general, the sandstones were derived from sedimentary successions (quartz recycling) supplemented by mainly granitic or gneissic exposures. Volcanic rocks, as source rocks, were generally subordinate. Additional petrographic discrimination, based on lithic fragment populations, indicates that most of

the samples plot in the Quartzose Recycled/ Craton Interior fields on standard ternary tectonic discrimination diagrams (see above). The dominance of volcanic and sedimentary rock fragments would suggest that sources were either rift or arc related. As such they could have been associated with initial rift events in the region, or with arc volcanism related to the final stages prior to collision.

Chemical and isotopic composition of the volcanics Both major and trace elements were analysed using ICP-AES and ICP-MS, respectively (SARM; CRPG-CNRS Nancy). Sr and Nd isotope data were obtained following the use of standard chemical separation techniques on a Finnigan MAT 262 (CNRS-CRPG; Nancy).

Major element compositions Major element analysis of 32 volcanic rocks (Table 3) provides a basis to classify them and to characterize both their common features as well as their differences. Such an approach allowed the

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Fig. 12. Sandstone modal data for the Bajocian-Callovian succession, Western Greater Caucasus. QFL, quartz-feldspar-lithic fragments; QmFLt, monocrystalline quartz-feldspar-total lithic fragments; QpLvmLsm, polycrystalline quartz-volcanic & metavolcanic lithic fragments-sedimentary & metasedimentary lithic fragments; VrfIrfSrf, volcanic rock fragment-igneous rock fragment-sedimentary rock fragment; SrfCrfFo, siliciclastic lithic fragments-carbonate lithic fragments-fossils. Tectonic discrimination fields defined by Ingersoll & Suczek (1979) and Dickinson et al. (1983).

possible genetic relationships in space and time to be analysed in detail. Observations. Mafic to intermediate volcanics were emplaced in Aalenian times in the Central Greater Caucasus (Tchataltapa Fm – Psekhako-Berezov Zone), and Bajocian times in the Southern Slope (Porphyritovaya Series – Abkhazo-Racha Zone). Eruptions of acid volcanics occurred sporadically from Pliensbachian through to Bajocian times (with the exception of the Toarcian). Prior to any detailed analysis of the compositions, it should be noted that the loss on ignition (LOI; Table 3) of the Jurassic rocks varies from 0.8 to 15.8%; the highest values measured were those of rocks with the least silica. There are significant losses in SiO2, Na2O and K2O, and gains in CaO for a LOI  8% (7 mafic volcanics) (Fig. 13). Only those data which were recalculated on an anhydrous basis were plotted in the binary diagrams (Fig. 14). According to the Le Maitre classification scheme (1989), the observed compositions range from basanites to rhyolites (Fig. 14). The datapoints plotted mostly within the silica-saturated ‘central’ field, less often in the silica-oversaturated field, and rarely in the silica-undersaturated field

(i.e. the most basic samples). The total of alkali is on average higher, and thus the silica-saturated feature is more pronounced in the Porphyritovaya Series than in the Tchataltapa Fm. Additionally, the volcanic succession is bimodal since there is a compositional gap in SiO2 values (between 65 and 70%) between the basic to intermediate volcanics and the acid ones. In more detail, the SiO2 content is observed to correlate negatively with the MgO, CaO and Fe2O3 contents. The volcanics are generally more sodic (up to 7%) than potassic (up to 4.2%), except for a highly potassic trachyte (Porphyritovaya Series) and a trachyandesite (Turovoy Fm). The percentage of Na2O broadly increases with SiO2 from basic to intermediate (1– 7%) rocks, whereas it remains lower (2–5%) in the acid rocks. K2O values are 1% in the Tchataltapa Fm, but somewhat higher (0.5–4.2%) in the samples from both the Porphyritovaya Series and the rhyolitic group. This somewhat heterogeneous rhyolitic group of samples is always TiO2-poor and more sodic than potassic. The basic sample group is rather heterogeneous. With a Mg# 67 (up to 72), basanites and trachybasalts of the Porphyritovaya Series can be

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

distinguished from the picrobasalt, basalts and trachybasalts of the Tchataltapa Fm, the Mg# of which only ranges up to 56. It should be noted that the basanites of the Porphyritovaya Series are very rich in olivine and pyroxene (Table 2), which could be a cumulative feature and would have increased the Mg# of the whole rocks. TiO2 contents can be used to distinguish 2 groups of basalts (sensu lato), a Ti-poor one (0.5 to 1%; basanites/ basaltic trachyandesites/ trachyandesites of the Porphyritovaya Series; picrobasalt/andesites of the Tchataltapa Fm) and a Ti-rich one (1– 2.5%; trachybasalts of the Porphyritovaya Series; basalts/trachybasalts of the Tchataltapa Fm). The Ti-rich basalts (sensu lato) are more sodic (Na2O% 3) than the Ti-poor ones. Interpretation. The bimodal aspect of the volcanics from the Western Greater Caucasus suggests that the rocks of the rhyolitic group are not genetically linked with the basic to intermediate volcanics. This is true for all of the analysed samples, whether of Pliensbachian, Aalenian or Bajocian age. The rhyolites were probably derived from partial melting of a continental crustal source. This is supported by the fact that the rhyolitic eruptions are either predominant (Pliensbachian times) or that they predated the basaltic to intermediate events (Aalenian times) without evidence of any transitional magmatism. In contrast, there is evidence of a regular and continuous trend between the basic to intermediate rocks in the Tchataltapa Fm and the Porphyritovaya Series, and this is probably due to fractional crystallization. With a Mg# content of 67, the basanites and trachybasalts of the Porphyritovaya Series maintained primitive magma features. In turn, the picrobasalts, basalts and trachybasalts of the Tchataltapa Fm are representative of more differentiated magmas as suggested by their lower Mg# (up to 56). Nevertheless the existence of two trends (TiO2 v. SiO2) excludes any genetic relationship between both of the basic sets independent of whether they are Aalenian or Bajocian in age. Indeed, the observed differences both in TiO2 and in Na2O and K2O contents between the two basic sets within the Tchataltapa Fm and the Porphyritovaya Series may reflect compositional differences between the mantle sources involved. The basalts and trachybasalts of the Tchataltapa Fm were derived from K2O-poor, Na2O-rich and TiO2-poor or TiO2-rich mantle sources. The mantle sources of the trachybasalts of the Porphyritovaya Series were richer in K2O than the TiO2-rich ones of the Tchataltapa Fm. The basanites of the Porphyritovaya Series had mantle sources rich in K2O but poor in TiO2 and Na2O.

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Trace element compositions The aim of the trace element study was to characterize the magmatic signatures of the mafic and acid eruptive magmas in order to determine the compositions of their mantle and crustal sources. Observations. As noted above, the magmatic rocks of the region are bimodal. Chondrite-normalized Rare Earth element (REEN) and Primitive Mantlenormalized Trace Element (TEN) spidergrams (Fig. 15) of the Western Greater Caucasus volcanics can be used to separate the basic (%SiO2  55) to intermediate (%SiO2  60) rocks from the trachy-dacitic (60  %SiO2  65) and rhyolitic (%SiO2  70) rocks. Of interest is the great diversity of the spidergram patterns, and by extension the magmatic signatures, among the basic to intermediate volcanics and to a lesser extent among the acid ones (Figs 15 & 16a, b), not only stratigraphically (in Jurassic times) but also across the region (Fig. 16c). Among the basic to intermediate rocks, four patterns representative of different magmatic signatures can be identified: (1) (2) (3) (4)

Normal-Mid Oceanic Ridge Basalts or N-MORB, in the Aalenian (sites 19, 22); Enriched-Mid Oceanic Ridge Basalts or E-MORB, in the Aalenian (sites 2, 24) and Bajocian (site 16); Alkaline, in the Aalenian (sites 10, 11) and Bajocian (site 12A); and Calc-Alkaline, in the Pliensbachian (site 12) and Bajocian (sites 16, 17). N-MORB type REEN patterns are rather flat and are always (to a greater/lesser degree) depleted in Light REEN (LaN to EuN) (cf. recent N-MORB patterns from various oceans; Fig. 15).

The REEN contents (down to 4 and up to 30 times chondritic values) are also comparable to those of recent oceans (ref. http://earthref.org/GERM). Similarly, the TEN patterns are rather flat and slightly depleted in the most incompatible elements (Th/TaN and La/YbN  1; Fig. 16a, b). Other features such as the high contents of Rb and Ba and the slightly negative anomalies in Ta and Nb are also observed on N-MORB patterns from recent oceans as well as from late Jurassic MORB (ChalotPrat 2005). In comparison, the REEN patterns of E-MORB type are slightly enriched in Light REEN and are comparable to E-MORB from recent oceans (Fig. 15). The REEN contents (down to 7 and up to 105) are similar or slightly below those of the E-MORB patterns of recent oceans. The TEN patterns are rather flat but always slightly enriched in the most incompatible elements with slightly

Table 3. Major and trace element data of selected volcanic from the Western Greater Caucasus Stratigraphic age

Aalenian

Early Aalenian

Aalenian

Middle Aalenian

Aalenian

Early Aalenian

Formation

Tch

Mak

Tch

Ind

Tch

Mak

Rock name

basanitic lava

rhyolitic ignimbrite

basaltic flow breccia

rhyolitic lapilli tuf

basaltic lava

rhyolitic ash fall tuf

03RU01

03RU04

03RU06

03RU08

03RU09

03RU10

71.46 0.28 11.26 1.85 0.16 0.71 3.88 2.33 1.57 0.13 6.92 100.5

46.20 1.47 17.39 7.51 0.10 4.75 7.94 3.14 0.46 0.29 10.51 99.8

69.79 0.39 12.29 3.43 0.14 0.82 4.80 2.07 2.10 0.09 4.41 100.3

43.73 2.34 16.23 11.38 0.55 4.38 7.67 4.49 0.53 0.69 7.64 99.6

73.54 0.47 13.08 2.25 0.04 0.70 2.08 2.80 3.17 0.10 2.13 100.4

44.64 0.88 17.87 11.89 0.41 6.72 14.60 2.57 0.30 0.12 53

76.32 0.30 12.02 1.98 0.17 0.76 4.14 2.49 1.68 0.14

51.77 1.64 19.49 8.42 0.11 5.32 8.89 3.52 0.51 0.33 56

72.75 0.41 12.81 3.58 0.15 0.86 5.00 2.16 2.19 0.09

47.54 2.54 17.64 12.37 0.60 4.76 8.34 4.89 0.58 0.75 43

74.86 0.48 13.32 2.29 0.04 0.71 2.12 2.85 3.23 0.10

,L.D. 282 0.573 ,L.D. ,L.D. 9.41 43.04 368 0.788 134 1.99 1.14 0.686 15.8 1.82 1.64 1.30 0.404 ,L.D. 3.88 0.180 ,L.D. 1.440 6.05 286 2.39 1.31 4.62 ,L.D. 1.65 0.529 307 0.134 0.308 0.960 0.175 0.203 182 ,L.D. 11.16 1.174 71.1 44.2

7.56 303 1.892 0.485 ,L.D. 52.2 3.58 20.3 18.5 10.6 6.51 3.81 0.447 13.2 5.53 1.59 4.86 1.310 0.076 22.91 0.560 ,L.D. 7.12 23.26 9.10 25.11 6.12 51.49 0.870 5.62 5.79 130 0.966 1.015 12.92 0.582 5.310 19.3 3.19 41.09 3.79 37.1 173

,L.D. 223 1.038 ,L.D. ,L.D. 23.1 23.84 133.9 2.63 41 3.43 1.89 1.353 17.4 3.54 1.70 2.79 0.671 ,L.D. 10.00 0.285 ,L.D. 10.35 13.60 26.06 1.48 3.11 14.57 ,L.D. 3.43 1.175 261 0.815 0.561 0.988 0.279 0.375 168 ,L.D. 18.86 1.861 92.2 133

6.87 458 1.162 0.217 ,L.D. 52.8 5.71 32.2 3.21 18.4 3.88 2.02 1.146 15.4 4.20 1.75 3.99 0.733 ,L.D. 24.28 0.306 0.675 7.47 22.95 15.63 22.44 6.01 68.40 0.454 4.82 3.05 118 0.701 0.653 9.66 0.310 2.107 39.7 1.487 21.66 2.03 48.3 152

,L.D. 322 1.382 ,L.D. ,L.D. 42.8 21.70 27.0 0.72 50 5.38 2.79 2.478 20.2 5.75 2.02 3.21 1.025 0.070 19.35 0.377 1.607 18.03 23.43 15.56 1.91 5.46 7.32 0.107 5.55 2.77 354 1.417 0.895 1.791 0.397 0.560 211 0.150 29.46 2.49 92.4 155

7.07 438 2.076 0.276 ,L.D. 63.2 5.08 29.9 12.3 14.6 5.12 2.79 0.873 15.7 5.47 1.51 4.68 0.991 ,L.D. 29.19 0.386 0.511 9.39 27.90 12.91 23.31 7.37 128 0.677 6.09 4.20 104 0.881 0.860 13.89 0.401 3.26 32.2 1.675 29.37 2.57 49.2 177

SiO2 37.68 TiO2 0.74 Al2O3 15.08 Fe2O3 10.03 MnO 0.35 MgO 5.67 CaO 12.32 Na2O 2.17 K2O 0.25 P2O5 0.10 LOI 15.28 Total 99.7 Recalculated on an anhydrous basis SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 Mg# (SiO2  55%) As Ba Be Bi Cd Ce Co Cr Cs Cu Dy Er Eu Ga Gd Ge Hf Ho In La Lu Mo Nb Nd Ni Pb Pr Rb Sb Sm Sn Sr Ta Tb Th Tm U V W Y Yb Zn Zr

Bajocian

Pliens-bachian

Pliens-bachian

Bajocian

Bajocian

Bajocian

Bajocian

Por

Tur

Tur

Por

Por

Por

Por

trachy-basaltic lava

rhyolitic cinerite

trachy-andesitic pumice cinerite

trachy-andesitic ash tuf

trachybasaltic lava

basaltic trachyandesitic lava

basaltic trachyandesitic lava

03RU11

03RU12A

03RU12C

03RU14B

03RU15

03RU16A

03RU17

47.20 2.31 18.15 8.55 0.15 5.15 8.37 3.37 2.62 0.65 3.23 99.7

82.88 0.16 9.19 0.86 0.00 0.00 1.60 3.67 0.67 0.06 0.78 99.9

57.61 0.90 16.72 8.93 0.40 2.81 1.69 1.85 5.86 0.19 2.96 99.9

58.33 0.82 17.73 8.03 0.26 2.77 1.21 4.91 2.61 0.16 3.58 100.4

46.18 1.79 16.96 11.14 0.17 5.41 9.95 3.41 1.04 0.28 3.39 99.7

53.03 2.08 15.66 11.05 0.23 2.07 5.60 5.99 1.02 0.75 2.15 99.6

50.25 0.60 14.39 7.67 0.17 7.91 9.67 3.08 2.69 0.24 3.14 99.8

48.90 2.39 18.80 8.86 0.16 5.33 8.67 3.49 2.71 0.67 54

83.64 0.16 9.27 0.86 0.00 0.00 1.61 3.71 0.68 0.06

59.41 0.93 17.25 9.21 0.41 2.90 1.75 1.91 6.04 0.19

60.23 0.84 18.31 8.29 0.27 2.86 1.25 5.07 2.70 0.17

47.94 1.86 17.61 11.57 0.18 5.62 10.33 3.54 1.08 0.29 49

54.41 2.13 16.07 11.34 0.24 2.12 5.74 6.14 1.05 0.77 27

51.98 0.62 14.89 7.94 0.17 8.18 10.00 3.19 2.78 0.25 67

1.17 593 1.734 ,L.D. ,L.D. 73.5 27.49 94.0 1.48 84 4.95 2.59 1.992 20.4 5.54 1.61 4.50 0.940 0.074 39.52 0.375 ,L.D. 59.97 30.55 43.70 5.42 8.03 52.78 1.060 6.09 5.64 432 4.47 0.844 6.63 0.371 1.580 248 0.498 26.79 2.44 82.8 214

5.84 108 1.991 0.411 ,L.D. 23.0 1.380 13.3 1.31 11.5 3.34 1.98 0.194 8.4 2.67 1.49 2.50 0.681 ,L.D. 10.00 0.293 ,L.D. 5.23 10.58 5.50 20.60 2.75 30.81 0.455 2.67 5.44 65 0.709 0.507 7.45 0.302 4.90 10.3 2.04 20.17 1.991 20.5 83.4

13.41 653 1.203 ,L.D. ,L.D. 32.3 20.33 57.7 6.63 94 5.14 3.14 1.203 18.9 4.77 3.01 3.18 1.091 0.074 15.29 0.521 ,L.D. 4.04 17.46 28.00 8.72 4.05 165 0.591 4.44 2.33 150 0.360 0.811 3.84 0.484 1.132 199 3.44 31.07 3.30 94 117

11.40 199 1.860 0.488 ,L.D. 49.7 21.76 71.8 4.24 116 4.70 2.78 1.304 21.0 4.56 1.71 3.35 0.968 0.097 26.69 0.471 ,L.D. 7.25 21.68 47.17 37.66 5.55 109 0.599 5.02 2.98 136 0.637 0.767 8.83 0.440 1.642 219 0.658 26.56 3.00 114 122

,L.D. 273 0.714 ,L.D. ,L.D. 23.7 37.23 132 1.50 108 4.16 2.30 1.513 19.7 4.18 1.40 2.52 0.836 0.062 9.44 0.345 ,L.D. 4.11 15.43 44.89 3.29 3.33 16.15 0.114 3.96 2.23 832 0.328 0.684 0.831 0.344 0.227 277 ,L.D. 23.58 2.24 93.0 103

,L.D. 289 1.898 ,L.D. ,L.D. 72.8 15.19 ,L.D. 1.17 8.5 10.18 5.79 3.532 25.3 10.22 1.85 6.85 2.061 0.107 30.07 0.903 1.031 13.11 43.07 ,L.D. 6.44 9.81 9.57 ,L.D. 10.28 2.98 382 0.999 1.675 2.616 0.859 0.763 85.7 0.227 57.87 5.75 138 295

,L.D. 821 1.315 ,L.D. ,L.D. 25.2 33.80 312 1.36 100 2.55 1.42 0.916 10.7 2.93 1.12 1.14 0.501 0.076 10.48 0.223 0.531 1.92 13.91 89 1.96 3.20 54.9 0.168 3.28 0.807 388 0.135 0.442 3.18 0.214 0.930 205 0.298 14.10 1.400 64.1 39.7 (Continued)

Table 3. Continued Stratigraphic age

Bajocian

Bajocian

Bajocian

Bajocian

Bajocian

Formation

Por

Por

Por

Por

Por

Rock name

rhyolitic pumice cinerite

basaltic trachyandesitic pumice cinerite

trachy-andesitic lapilli tuf

trachytic pumice cinerite

basanitic lava

03RU18A

03RU19

03RU20

03RU21

03RU24B

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total

72.47 0.36 11.94 4.39 0.17 1.22 1.35 5.36 1.27 0.08 1.21 99.8

51.11 1.10 16.62 10.71 0.26 3.31 3.66 5.74 2.25 0.39 4.86 100.0

55.85 0.72 18.04 7.50 0.19 2.32 4.25 4.82 4.21 0.27 2.43 100.6

59.31 0.64 17.55 5.47 0.16 1.55 1.29 3.68 8.21 0.27 1.96 100.1

35.52 0.52 9.59 8.42 0.11 9.15 18.02 1.34 2.68 0.21 14.01 99.6

Recalculated on an anhydrous basis SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5

73.50 0.36 12.11 4.45 0.17 1.24 1.37 5.44 1.28 0.08

53.72 1.15 17.47 11.26 0.27 3.48 3.84 6.03 2.37 0.41

56.90 0.73 18.38 7.64 0.20 2.36 4.33 4.91 4.29 0.27

60.44 0.65 17.88 5.57 0.17 1.58 1.31 3.75 8.37 0.28

41.52 0.60 11.21 9.84 0.13 10.69 21.06 1.57 3.13 0.25

3.13 88 1.315 ,L.D. ,L.D. 26.9 7.47 10.700 ,L.D. 41.8 6.09 3.90 0.876 15.2 4.80 0.90 5.29 1.299 0.076 10.26 0.653 ,L.D. 5.46 15.61 7.79 11.41 3.63 23.44 0.511 4.53 1.652 162 0.477 0.915 5.88 0.610 1.746 44.8 0.298 36.68 4.15 59.9 195

38 2.02 189 1.935 ,L.D. ,L.D. 48.3 26.61 24.7 7.88 104 5.81 3.38 1.835 18.5 5.96 1.46 3.42 1.164 0.081 23.29 0.537 ,L.D. 6.17 25.46 14.77 11.20 6.00 72.5 0.138 6.07 1.586 454 0.498 0.966 5.85 0.509 1.987 228 0.854 32.88 3.47 120 123

3.39 524 1.959 ,L.D. ,L.D. 42.4 17.40 15.7 0.142 91 4.54 2.69 1.403 18.7 4.71 1.43 3.05 0.924 0.081 21.04 0.453 ,L.D. 6.07 21.16 11.90 10.70 5.24 98 ,L.D. 4.81 1.365 911 0.506 0.744 5.90 0.411 1.889 133 0.851 27.62 2.82 91.5 119

3.58 2174 0.843 ,L.D. ,L.D. 39.2 9.74 4.76 0.234 28.4 3.97 2.49 1.254 10.7 3.82 0.748 3.61 0.827 ,L.D. 21.8 0.461 0.720 7.94 18.7 5.96 14.0 4.62 262 0.248 4.05 1.85 411 0.646 0.633 5.53 0.401 2.11 69 1.52 22.8 2.93 68.7 133

68 ,L.D. 336 0.645 ,L.D. ,L.D. 24.7 46.4 901 0.766 36.9 2.72 1.52 0.902 9.62 3.32 1.097 1.07 0.542 ,L.D. 14.4 0.223 ,L.D. 1.62 16.3 244 5.11 3.81 46 0.125 3.71 1.78 250 0.132 0.475 3.42 0.215 0.710 200 0.285 16.8 1.47 67.2 35.5

Mg# (SiO2  55%) As Ba Be Bi Cd Ce Co Cr Cs Cu Dy Er Eu Ga Gd Ge Hf Ho In La Lu Mo Nb Nd Ni Pb Pr Rb Sb Sm Sn Sr Ta Tb Th Tm U V W Y Yb Zn Zr

Bajocian

Bajocian

Aalenian

Aalenian

Aalenian

Early Aalenian

Early Aalenian

Early Aalenian

Por

Por

Tch

Tch

Tch

Mak

Mak

Mak

basanitic lava

basanitic lava

basaltic lava

basaltic lava

basaltic trachyandesitic explosive breccia

rhyolitic lava

rhyolitic flow breccia

dacitic lava

03RU24C

03RU27

03RU29

03RU30

03RU31

03RU32

03RU33

03RU34

40.69 0.62 10.97 10.64 0.14 13.80 10.56 1.13 2.31 0.23 8.48 99.6

37.42 0.49 8.98 8.46 0.13 11.05 18.63 1.80 0.90 0.18 12.19 100.2

47.75 1.83 14.28 12.26 0.22 6.95 9.88 2.94 0.28 0.18 2.72 99.3

47.84 1.89 13.91 12.91 0.23 6.55 8.99 3.84 0.33 0.21 2.67 99.3

52.98 0.48 19.82 6.27 0.14 3.06 4.88 4.92 0.80 0.06 5.80 99.2

68.26 0.60 14.58 4.79 0.06 1.34 0.31 3.41 3.54 0.12 2.81 99.8

70.50 0.53 14.24 4.11 0.05 1.42 0.25 4.09 2.03 0.11 2.48 99.8

67.29 0.55 14.39 4.47 0.09 1.55 1.16 4.49 3.24 0.11 2.18 99.5

44.67 0.69 12.04 11.68 0.15 15.16 11.59 1.24 2.54 0.25

42.51 0.56 10.20 9.61 0.15 12.55 21.16 2.04 1.02 0.20

49.44 1.90 14.79 12.69 0.23 7.19 10.23 3.05 0.29 0.19

49.48 1.95 14.39 13.35 0.23 6.77 9.29 3.97 0.34 0.21

56.72 0.51 21.22 6.71 0.15 3.28 5.23 5.26 0.86 0.06

70.36 0.62 15.03 4.94 0.06 1.38 0.32 3.51 3.65 0.12

72.43 0.54 14.64 4.22 0.05 1.46 0.26 4.20 2.09 0.11

69.13 0.56 14.79 4.59 0.09 1.59 1.19 4.61 3.32 0.12

72 1.49 318 0.588 ,L.D. ,L.D. 30.6 68.5 1257 0.981 51.7 2.91 1.62 1.234 12.7 3.70 1.241 1.29 0.584 ,L.D. 15.0 0.235 ,L.D. 1.92 18.3 371 4.81 4.19 45 0.109 4.18 1.22 184 0.151 0.533 4.12 0.232 0.725 238 0.326 17.0 1.54 96.5 42.3

72 1.12 165 0.628 ,L.D. ,L.D. 21.2 47.4 929 0.383 30.9 2.16 1.20 0.764 8.99 2.57 1.305 1.00 0.424 ,L.D. 8.97 0.177 ,L.D. 1.45 11.8 396 3.45 2.70 16.12 0.121 2.86 0.718 242 0.112 0.380 3.31 0.174 0.698 190 0.252 11.6 1.16 58.1 32.9

53 ,L.D. 141 0.619 ,L.D. ,L.D. 15.3 41.1 152 0.380 95.5 6.64 4.04 1.676 19.2 5.60 1.730 3.29 1.401 0.089 5.14 0.627 ,L.D. 3.02 13.2 56 3.43 2.49 4.99 0.239 4.38 4.14 186 0.283 1.018 0.313 0.601 0.114 322 ,L.D. 38.5 4.09 102.0 117

50 ,L.D. 174 0.681 ,L.D. ,L.D. 17.1 41.5 107 0.481 70.9 7.27 4.39 1.757 19.4 6.11 1.660 3.73 1.526 0.094 5.82 0.672 ,L.D. 3.39 14.6 48 2.13 2.80 6.63 0.138 4.83 2.41 279 0.325 1.093 0.344 0.656 0.130 336 ,L.D. 41.8 4.42 99.4 133

,L.D. 206 ,L.D. ,L.D. ,L.D. 2.7 13.8 13.71 0.789 35.5 2.33 1.43 0.528 14.6 1.91 0.938 1.023 0.498 ,L.D. 1.28 0.221 ,L.D. 0.299 3.28 15.6 2.04 0.57 14.46 ,L.D. 1.33 0.314 361 0.028 0.347 0.146 0.212 0.027 183 ,L.D. 14.3 1.43 71.2 27.9

8.43 445 1.194 ,L.D. ,L.D. 53.7 8.60 44.97 4.89 17.3 5.15 2.96 0.893 16.9 5.09 1.599 5.35 1.036 0.062 25.9 0.477 ,L.D. 8.23 25.0 16.0 25.03 6.47 128 0.454 5.53 2.50 38 0.824 0.844 11.45 0.452 2.90 70 1.402 29.9 3.07 77.1 188

7.66 262 1.379 ,L.D. ,L.D. 56.3 7.31 40.01 3.11 7.20 5.42 3.09 0.913 17.0 5.58 1.883 5.37 1.087 ,L.D. 27.7 0.484 ,L.D. 7.82 27.2 14.6 9.80 6.99 82 0.299 5.94 2.80 40 0.794 0.897 12.32 0.464 3.05 62 1.375 32.0 3.10 37.7 189

2.95 516 1.385 ,L.D. ,L.D. 52.3 8.86 46.01 2.55 8.07 5.22 3.04 0.844 16.6 5.05 0.938 5.17 1.046 ,L.D. 26.7 0.478 ,L.D. 7.98 25.1 17.0 21.9 6.52 122 0.134 5.48 1.79 143 0.809 0.836 11.66 0.467 2.99 72 1.157 31.3 3.12 77.0 183 (Continued )

Table 3. Continued Stratigraphic age

Aalenian

Aalenian

Aalenian

Aalenian

Aalenian

Aalenian

Formation

Tch

Tch

Tch

Tch

Tch

Tch

Rock name

trachy-dacitic lava flow breccia

dacitic lava

basaltic trachy-andesitic lava

andesitic extrusion breccia

fragment of trachy-andesitic lava

andesitic extrusion breccia

03RU35

03RU36B

03RU37

03RU39

03RU42

03RU43

60.05 0.51 16.84 4.65 0.12 4.06 1.81 6.72 0.43 0.14 4.03 99.4

60.23 0.47 15.23 5.25 0.25 4.03 2.96 5.28 0.45 0.29 5.24 99.7

52.13 1.34 15.39 9.81 0.26 4.27 3.93 6.30 0.00 0.16 5.72 99.3

54.19 0.80 17.57 9.37 0.14 3.88 1.84 2.54 0.16 0.06 8.18 98.7

57.85 0.62 15.92 7.00 0.09 3.85 2.41 6.64 0.14 0.18 3.48 98.2

51.91 0.74 16.77 7.32 0.07 2.95 6.89 2.75 0.18 0.05 8.66 98.3

63.78 0.50 16.13 5.56 0.26 4.26 3.13 5.59 0.48 0.31

55.70 1.43 16.44 10.48 0.28 4.56 4.20 6.73 0.00 0.17

59.85 0.88 19.40 10.35 0.15 4.29 2.03 2.80 0.17 0.07

61.08 0.65 16.81 7.39 0.10 4.07 2.55 7.01 0.15 0.19

57.92 0.82 18.71 8.16 0.08 3.30 7.69 3.06 0.20 0.06

2.48 291 0.621 ,L.D. ,L.D. 21.8 12.6 36.80 0.378 22.6 2.14 1.39 0.578 12.0 2.22 1.125 2.46 0.461 ,L.D. 13.3 0.256 ,L.D. 3.46 9.89 61 5.48 2.55 7.28 0.161 2.13 1.37 268 0.296 0.338 3.19 0.220 0.822 51 0.773 15.8 1.56 54.2 97

46 ,L.D. 49 ,L.D. ,L.D. ,L.D. 13.4 23.8 4.40 ,L.D. 15.35 4.79 2.88 1.261 17.2 4.00 1.374 2.57 1.001 0.068 4.59 0.446 ,L.D. 1.28 10.9 4.19 1.65 2.16 ,L.D. ,L.D. 3.38 1.28 105 0.127 0.716 0.451 0.434 0.147 240 0.238 28.1 2.89 78.3 96

45 ,L.D. 48 ,L.D. ,L.D. ,L.D. 5.51 20.9 10.80 0.613 137.7 3.28 2.07 0.738 19.3 2.68 2.527 1.76 0.706 0.066 2.80 0.347 ,L.D. 0.727 5.69 19.8 7.02 1.09 2.48 0.427 2.02 5.30 36 0.070 0.489 0.357 0.327 0.112 233 0.207 19.9 2.23 93.7 56.8

1.83 37 ,L.D. ,L.D. ,L.D. 8.24 11.4 7.04 ,L.D. 23.9 3.67 2.37 0.658 12.5 2.87 0.880 1.58 0.792 ,L.D. 3.95 0.367 ,L.D. 0.807 6.11 8.80 2.23 1.20 1.27 ,L.D. 2.11 0.704 139 0.072 0.533 0.411 0.362 0.181 108 0.258 23.9 2.36 78.3 42.4

,L.D. 66 ,L.D. ,L.D. ,L.D. 6.87 12.8 10.72 0.624 29.9 2.49 1.69 0.582 19.0 1.90 1.291 1.95 0.543 ,L.D. 3.01 0.342 ,L.D. 1.54 4.55 11.4 2.74 0.97 3.18 0.132 1.56 0.797 187 0.131 0.363 1.12 0.287 0.274 188 0.183 14.5 2.03 65.1 56.4

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total

Recalculated on an anhydrous basis SiO2 62.99 TiO2 0.54 Al2O3 17.67 Fe2O3 4.88 MnO 0.13 MgO 4.26 CaO 1.90 Na2O 7.05 K2O 0.45 P2O5 0.14 Mg# (SiO2  55%) As Ba Be Bi Cd Ce Co Cr Cs Cu Dy Er Eu Ga Gd Ge Hf Ho In La Lu Mo Nb Nd Ni Pb Pr Rb Sb Sm Sn Sr Ta Tb Th Tm U V W Y Yb Zn Zr

5.83 368 0.720 ,L.D. ,L.D. 20.8 12.2 43.44 0.223 11.31 2.09 1.23 0.643 15.2 1.97 0.893 2.66 0.425 ,L.D. 11.6 0.213 ,L.D. 3.74 8.74 63 2.45 2.31 6.89 ,L.D. 1.90 0.923 115 0.317 0.329 3.55 0.194 0.822 73 0.794 12.5 1.32 47.9 106

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

217

Fig. 13. Plot of the variations in SiO2, CaO, Na2O, K2O and total alkali with Loss on Ignition (LOI) on whole-rock samples (raw data) of volcanics from the Western Greater Caucasus region. Pliensbachian (Turovoy Fm): blue diamond; early Aalenian (Makarov Fm): yellow circle; middle to late Aalenian (Mount Indyuk Unit): green triangle; Aalenian (Tchataltapa Fm): black square; Bajocian (Porphyritovaya Series): red star.

negative anomalies of Ta and Nb (1 , Th/TaN  5 and 1 , La/YbN  4; Fig. 16). Of note is the fact that the total TE content is higher in Bajocian-age E-MOR basalts and andesites than in Aalenianage ones. Alkaline type REEN patterns are strongly enriched in Light REEN (to a greater degree in Bajocian rather than Aalenian basalts). The TEN patterns are regular and strongly enriched in the most incompatible elements with slight positive anomalies in Ta and Nb (Th/TaN  1 and La/YbN  1), all of which are characteristic trends of alkaline mafic rocks (Wilson 1989; Best & Christiansen 2001). Calc-alkaline type REEN patterns are also strongly enriched in Light REEN, and cannot be distinguished from alkaline ones. The TEN patterns are irregular and enriched in the most incompatible elements, with significant negative anomalies in Ta and Nb (Th/TaN  5 and La/YbN  4; Fig. 16),

typical of calc-alkaline mafic rocks (Wilson 1989; Best & Christiansen 2001). Spidergrams of the basalts show strong negative anomalies in Zr, Hf and Ti none of which (with the exception of Ti) are present in the andesites. Both negative and positive anomalies in Sr are present in both rock types. Of note is the rather low value of La/YbN (3.3) in the Pliensbachian-age samples which differs from the pattern observed in the Bajocian ones. Despite the various magmatic signatures, there are no systematic differences between the basalts and the andesites. The REEN and TEN patterns are either parallel or superposed, while the total TE contents of the andesites (sensu lato) may be below or above those of the basalts (sensu lato). The TiO2-poor and TiO2-rich basic (%SiO2  55) samples can be grouped into four and two sets, respectively, each of which has different magmatic signatures (Fig. 17). The TiO2-poor (0.5–1%)

218

T. MCCANN ET AL.

Fig. 14. Classification of volcanics rocks after Le Maitre (1989) and variations of major element oxide percentages v. silica (data recalculated on an anhydrous basis). For key to symbols, see Figure 13. 1, picrobasalt; 2, tephrite-basanite; 3, basalt; 4, trachybasalt; 5, basaltic andesite; 6, basaltic trachyandesite; 7, andesite; 8, trachyandesite; 9, dacite; 10, trachyte; 11, rhyolite.

group includes the Aalenian E-MORB set and the Bajocian calc-alkaline set. The TiO2-rich (1.5– 2.5%) group contains together both the Aalenian and Bajocian alkaline sets, the Aalenian N-MORB set and the Bajocian E-MORB set. There is a

relationship between the high or low TiO2% contents and the absence or presence of Ti negative anomalies in the alkaline and calc-alkaline trace element spidergrams, respectively. However, this is not verified for the TiO2-rich N-MORB, the

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

Fig. 15. Primitive mantle-normalized trace element, chondrite-normalized Rare Earth Element spidergrams and magmatic signatures of the western Greater Caucasus volcanics in Pliensbachian, Aalenian and Bajocian times: (a) basic (SiO2%  55; sample number in bold) to intermediate (SiO2%  60);

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Fig. 15. (Continued ) (b) acid (rhyolite SiO2%  70; sample number in bold; trachyte-dacite 60  SiO2%  65). Formation name and site number (located on Fig. 3) in green bold; magmatic affinity in red block letters. Reference trace element data of N-MORB and E-MORB from Atlantic (Atl), Pacific (Pac) and Indian (Ind) oceans from http://earthref. org/GERM.

spidergrams of which show slightly negative Ti anomalies as is usual for N-MORB (Chalot-Prat 2005 and references therein) and the relatively Ti-poor and Ti-rich E-MORB sets, the spidergrams of which do not show any Ti anomaly. With regard to the acid rocks, all of the REEN and TEN patterns (with one exception: rhyolite 03RU18A), show the characteristic features of basic calc-alkaline rocks, with strong enrichments in the most incompatible elements and with marked Ta, Nb and Ti anomalies (Th/TaN  4 and La/YbN  3; Fig. 16). The similarity of the rhyolite (Mount Indyuk Fm) and the dacite/trachydacite (Tchataltapa Fm) patterns is noticeable, although these dacite/trachydacite patterns differ from those of the basic to intermediate rocks of the Tchataltapa Fm. These patterns resulted in the inclusion of the dacite/trachydacite rocks within the acid group. In contrast, the spidergram of the Bajocian rhyolite (03RU18A) has a magmatic

signature close to E-MORB-type ones for basic volcanics with a rather flat REEN pattern (La/ YbN , 2) and a negative Ta anomaly (Th/ TaN  6). Thus, the acid volcanics have similar Th/TaN ratios (4 to 7), but different La/YbN (2 to 9) and therefore, show a horizontal trend (Fig. 16a, b), while the basic calc-alkaline rocks show a vertical trend. This trend separation clearly demonstrates the bimodal nature of the volcanic rocks in the area. Interpretation. The bimodal aspect of the volcanic rocks in the region is confirmed by the trace element analysis. Since basic volcanic eruptions were absent, or not significant, during periods (Pliensbachian; Early to Middle Aalenian) of major rhyolitic eruptions, a crustal origin for the acid group is probable. The involved crust was likely a heterogeneous basic lower crust, showing a calc-alkaline signature. The E-MORB signature

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Fig. 16. Trace element ratios (N: normalized to the Primitive Mantle) and related magmatic signatures of the Western Greater Caucasus (a) basalts and andesites; and (b) rhyolites/dacite/trachyte/trachydacite at Pliensbachian (Turovoy Fm as diamond, blue: andesite/red: rhyolite), Early Aalenian (Makarov Fm: rhyolite as yellow circle), Middle Aalenian (Mount Indyuk Fm: rhyolite as green triangle), Middle to Late Aalenian (Tchataltapa Fm as square, black: basalt/pink: andesite/blue: dacite & trachydacite) and Bajocian (Porphyritovaya Series as star, orange: basalt/green: andesite/ orange: trachyte/black: rhyolite ) times; and (c) Location of volcanic rocks according their sampling sites (see Fig. 3) with identification of both magmatic signatures and ages (P, Pliensbachian; A, Aalenian; B, Bajocian).

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Fig. 17. TiO2 v. SiO2 diagram showing the correlation between the TiO2-poor and TiO2-rich groups and the magmatic signatures as defined by trace element compositions. (See Fig. 13 for key to symbols.)

of the Bajocian rhyolite (03RU18A) could have resulted from partial melting of an E-MORB type gabbro underplated at the base of the continental crust just prior to the extrusion of the rhyolite, that is, during early Bajocian or middle to late Aalenian times. Indeed such E-MORB-type basic magma genesis occurred during both periods. During the periods of maximum production of acid magmas (Pliensbachian to Early Aalenian) basic magma underplating, via mantle melting, could also have played a role since mantle magma could have initiated or contributed to crustal melting. The trace element signatures of magmatic rocks are widely used to decipher the related geodynamic setting in locations where a certain consistency in time and space can be recognized. Analysis of the trace element signatures for the GC basic to intermediate volcanics indicates that (1) the coeval rocks display a large spectra (for example calcalkaline and E-MORB) of signatures which can be also observed for rocks from a single location (e.g. from a specific formation); and (2) some of the signatures are recurrent or/and simply disappear (as for the N-MORB signature) through the set of Pliensbachian to Bajocian rocks. Hence, there is no evidence of a consistent trend thus precluding any link between the magmatic signatures and any specific geodynamic contexts. Major and trace element magmatic signatures of the most basic rocks are related to the major and trace element compositions of the possible mantle sources, the compositions of the latter being dependent on its history prior to melting. The trace element signatures of coeval basic rocks (e.g.

Aalenian or Bajocian) vary markedly both within and between the sampling sites (Fig. 16). This would suggest that the involved mantle sources varied in composition from one location to another and according to the depth of melting. During Pliensbachian times, the only representative sample (a trachyandesite) is calc-alkaline. Such a signature is usually interpreted as being oceanic subduction-related, that is, typical of a basic magma derived from the melting of a depleted asthenospheric mantle contaminated by magma extracted from the oceanic subducted plate at depth (Wilson 1989; Best & Christiansen 2001). However, based on previous work in the region, there is no geological evidence in the Western Greater Caucasus which could corroborate such a hypothesis. Another suggestion (Chalot-Prat & Boullier 1997; Chalot-Prat & Girbacea 2000; Chalot-Prat et al. 2007 and references therein) is that the mantle source was located within the already metasomatized continental lithospheric mantle, the metasomatic history of which predated its melting in Pliensbachian times. During middle to late Aalenian times, the N-MORB signature was predominant (sites 19, 22). The asthenospheric depleted mantle (worldwide source of N-MORB; Chalot-Prat, 2005 and references therein) was involved in order to generate such a magma. Such a model, however, presupposes that eruptions occurred in a continental extensional context during which the thinning of the lithosphere was rather extreme but not enough to lead to the generation of oceanic crust. Laterally (sites 2, 24), there is evidence of E-MORB signatures. E-MOR

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

basalts often erupt at recent mid-oceanic ridges (see the Indian, Atlantic and Pacific E-MORB patterns on Fig. 15). Such a signature would derive from the melting of an asthenospheric source enriched in the most incompatible elements. In a mid-oceanic ridge context, such a source could be the overlying oceanic mantle residue metasomatized during N-MORB percolation (Piccardo et al. 2007). In the context of the Western Greater Caucasus with a thinned continental lithosphere, such a source could also be located at the asthenospherelithosphere boundary, just above the N-MORB source. To the north of the previous sites the mafic rocks have an alkaline signature (sites 10, 11), suggesting that the mantle sources were located in the metasomatized continental lithospheric mantle (cf. Chalot-Prat & Boullier 1997; Chalot-Prat & Girbacea 2000; Chalot-Prat et al. 2007 and references therein). During Bajocian times, the calc-alkaline signature of the mafic rocks was predominant (site 17). The eruptions were coeval with E-MORB to the NW (site 16), while alkaline basic eruptions occurred along strike to the east (site 12; Fig. 16). Such compositional diversity within the mafic volcanics of the same area can be understood if the different mantle sources were located at various depths within a heterogeneous continental lithospheric mantle. Indeed the E-MORB type magmas would have been derived from greater depths (asthenosphere/lithosphere boundary) than the calc-alkaline type ones, whereas these and the alkaline type magmas would have been derived from a heterogeneous metasomatized lithospheric mantle (cf. Chalot-Prat & Boullier 1997). In this context, it is of interest to note the existence of an active oceanic subduction zone located to the south (i.e. in the Lesser Caucasus), as evidenced by the presence of an ophiolitic suture zone. This subduction zone could have been the motor for plate deformation in the region which could then have induced, via adiabatic decompression processes at different mantle depths and at the same time (i.e. production of both E-MOR and calc-alkaline magmas), partial melting within the mantle and the crust.

Sr and Nd isotopic compositions Observations. Initial (indexed ‘i’) Rb –Sr and Sm–Nd isotopic data (Table 4) have been determined for some of the most primitive basic samples from the Aalenian-age Tchataltapa Fm (alkaline-03RU06; N-MORB-03RU29/30) and the Bajocian-age Atchichkh Fm (alkaline-03RU11) and Porphyritovaya Series (E-MORB-03RU15; calc-alkaline-03RU17/19/24C). Rb and Sr data were obtained using an inductive coupled plasmamass spectrometer (ICP-MS) (Table 3). Initial

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values of 143Nd/144Nd and 87Sr/86Sr were calculated at 170 Ma (i.e. eruptions dating from 189.6 to 167.7 Ma; 170Ma  boundary Aalenian – Bajocian – International Geological Time Scale Chart 2009; http://www.geosociety.org/science/ timescale/timescl.pdf). On the 1Ndi v. (87Sr/86Sr)i diagram (Fig. 18a), basalts occupy an elongate restricted field across the NW and NE quadrants with 1Ndi between þ0.3 and þ7.4 and (87Sr/86Sr)i between 0.70384 and 0.70541. A significant subhorizontal shift towards elevated 87Sr/86Sr is probably due to secondary alteration, and the observed oblique trend does not have any meaning in terms of mantle source compositions. In contrast, the two N-MORB- (03RU29/ 30) type basalts have the highest 1Ndi (6.6 and 7.4), close to the depleted mantle source of the N-MORB reference (1Ndi at 170 Ma ¼ 8; from Salters & Stracke 2004). Additionally, all of the 1Ndi are positive, suggesting that the Nd isotopic signatures are mantle dominated (i.e. that mantle sources were metasomatized by continental crust products). On a diagram illustrating the evolution of Nd isotopic compositions of mantle sources of the studied samples (Fig. 18b), all of the trends cross the evolution line of the depleted mantle (Salters & Stracke 2004) within 400 and 900 Ma, with 50% occurring in a narrow age range (c. 600 Ma + 50; Late Neoproterozoic; International Geological Time Scale Chart 2009; http://www.geosociety. org/science/timescale/timescl.pdf). The evolution of the Sr isotopic compositions (Fig. 18c) suggests a much larger range of values (400 –3800 Ma) than is credible. As in many other studies, Rb and Sr were much more mobile during alteration processes than the rare earth elements. Interpretations. The positive 1Ndi values of the basaltic samples, independent of their magmatic signature and their age, suggest that the Nd isotopic signatures are mantle dominated. The high 1Ndi values of the Aalenian N-MORB (close to those of recent MORB) provide strong evidence for a depleted asthenospheric mantle source, as already suggested by the TEN patterns. Additionally, it is suggested that the mantle sources of the basalts with slightly (to significantly) lower 1Ndi values were metasomatized. Such a conclusion is supported by the measured major and trace element compositions of the rocks (see above). Indeed, as noted in previous papers (Chalot-Prat & Boullier 1997; Chalot-Prat et al. 2007 and references therein), such initial isotopic values imply that mantle sources plot on mixing trends between depleted mantle and continental crust (or already metasomatized continental mantle) end-members. Such isotopic signatures are typical for a heterogeneous

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Table 4. Rb, Sr, Nd and Sm isotopic data of Aalenian and Bajocian whole rock basalts from the Western Greater Caucasus Rb and Sr isotopic data of Aalenian and Bajocian whole-rock basalts from the Western Greater Caucasus Sample

Tchataltapa/site 10

03RU06

Atchichkh/Site 12 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Tchataltapa/Site 19 Tchataltapa/Site 19

03RU30

Age

Trace element magmatic signature

Rb

Sr

87

Rb/86Sr

87

Sr/86Sr

(2#)

87

Sr/86Sr initial

alkaline

14.57

261

0.162

0.705801

14

0.705410

03RU11 03RU15

Middle/Late Aalenian Bajocian Bajocian

alkaline E-MORB

52.78 16.15

432 832

0.354 0.056

0.704980 0.704598

14 13

0.704124 0.704462

03RU17

Bajocian

calc-alkaline

54.85

388

0.410

0.705360

10

0.704370

03RU19

Bajocian

calc-alkaline

72.46

454

0.462

0.705857

18

0.704741

03RU24C

Bajocian

calc-alkaline

45.40

184

0.715

0.705569

11

0.703840

03RU29

Middle/Late Aalenian Middle/Late Aalenian

N-MORB

4.99

186

0.078

0.704678

10

0.704490

N-MORB

6.63

279

0.069

0.705391

17

0.705225

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Formation/site no.

Nd and Sm isotopic data of Aalenian and Bajocian whole-rock basalts from the Western Greater Caucasus Formation/ site no.

Sample

Tchataltapa/site 10

03RU06

Atchichkh/Site 12 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Porphyritovaya Series/Site 16 Tchataltapa/Site 19 Tchataltapa/Site 19

03RU30

Age

Trace element magmatic signature

Nd

Sm

147

Sm/ Nd

144

143

Nd/144Nd

(2#)

143

Nd/144Nd initial

1Nd

alkaline

13.60

3.43

0.153

0.512907

17

0.512737

6.2

03RU11 03RU15

Middle/Late Aalenian Bajocian Bajocian

alkaline E-MORB

30.55 15.43

6.09 3.96

0.121 0.156

0.512718 0.512835

13 13

0.512583 0.512662

3.2 4.7

03RU17

Bajocian

calc-alkaline

13.91

3.28

0.143

0.512653

7

0.512494

1.5

03RU19

Bajocian

calc-alkaline

25.46

6.07

0.145

0.512724

8

0.512563

2.8

03RU24C

Bajocian

calc-alkaline

18.32

4.18

0.138

0.512591

12

0.512437

0.3

03RU29

Middle/Late Aalenian Middle/Late Aalenian

N-MORB

13.16

4.38

0.202

0.513021

14

0.512796

7.4

N-MORB

14.56

4.83

0.201

0.512983

21

0.512759

6.6

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Fig. 18. (a) Initial Nd and Sr isotopic ratios of some representative basalts from the Western Greater Caucasus. Depleted Mantle with DM isotopic values after Salters & Stracke (2004). Aalenian group in red – N-MORB, triangle; alkaline, diamond; Bajocian group in blue – E-MORB, square; alkaline, diamond; calc-alkaline, circle. (b) Nd and (c) Sr model ages for mantle sources of selected basalts. Vertical dotted lines indicate minimum and maximum model ages.

trace element enriched mantle as found in the continental lithospheric mantle (Wilson 1989; ChalotPrat & Boullier 1997 and references therein; Weinstein et al. 2006 and references therein). The range of 1Ndi values can be explained by percolation, within a rather homogeneous depleted mantle (asthenosphere, as defined for the N-MORB source), of contaminant melts with similar isotopic ratios but with slightly different Nd/Sr ratios and in different proportions (see Faure 1986 for explanations). These processes would result in small-scale mantle heterogeneity. The evolution of the primitive mantle residue was thus disturbed during this period of percolation/metasomatism. The Nd model ages of the Aalenian and Bajocian basalts would support such a hypothesis since they indicate that significant modifications (partial melting and/ or metasomatism) of the involved mantle (whether lithospheric or asthenospheric) occurred at c. 600 Ma + 50 Ma (Late Neoproterozoic; i.e. during the Baikalian orogenic event, cf. in Saintot et al. 2006b). Thus, if an oceanic plate (the present-day

exhumed Lesser Caucasus ophiolitic suture) was subducting northwards below an active margin represented by the Southern Slope of the GC during Bajocian times, the coeval basalts would be the markers of the previous history of the lithospheric mantle along this plate margin. Plate deformation and shallowing of the subducting slab at the beginning of the Bajocian (as suggested by Saintot et al. 2006a, 2007) is believed to have been the main cause of partial melting of the asthenospheric and lithospheric mantles and continental crust by adiabatic decompression and at a range of depths.

Jurassic volcanism and the geodynamic evolution of the Western Greater Caucasus Spatio-temporal distribution and dynamics of eruptions As noted above, 57 different volcanic rocks (Table 2) were sampled from Jurassic successions

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in the Central Greater Caucasus and the Southern Slope areas (Fig. 3). In contrast, the Fore-Caucasus area lacks volcanics. Volcanism was active during Pliensbachian and Aalenian times in the Central Greater Caucasus area, with a break during Toarcian times. In the Southern Slope area, volcanic eruptions occurred only during the Bajocian. Volcanism appears to have migrated from north to south during Jurassic times. In both the Central Greater Caucasus and Southern Slope areas, only some zones (see above), and within these zones only some formations, were affected by volcanism. This would suggest that in the western Greater Caucasus area, volcanoes occupied only a limited area (c. hundreds of square kilometres; up to 1000 km2 for the Porphyritovaya Series (Abkhazo-Racha Zone – Southern Slope)) separated by broader areas where the predominant process was one of siliciclastic deposition. In general, the eruptions and resultant deposits were subaerial, although flow and fall emplacement in nearshore or shallow marine settings were also common. These latter environments were predominantly siliciclastic shelves, extending out into deeper waters (see below). Across the shelf, it is probably that at least some of the various eruption locations were situated on island settings, whose formation was coincident with the onset of volcanic activity. In terms of volcano dynamics, the eruptions of the rhyolitic magmas were dominantly explosive with the formation of nue´es ardentes and related deposits (ash flows/ash surges/ash falls). Some dacitic to trachydacitic flows also occurred. Mafic magma eruptions were predominantly effusive. However, in Bajocian times, explosive eruptions with emissions of large ash plumes were dominant during the second half of the eruptive period. In Aalenian (Tchataltapa Fm – Psekhako-Berezov Fm – Central Greater Caucasus) and Bajocian (Porphyritovaya Series – Abkhazo-Racha Zone – Southern Slope) times, the volcanic activity was so intense that sedimentation did not occur between the extrusive events. However, it should be noted that at the geological timescale, an eruption is very transient, and that even the relatively thick volcanic units within the Porphyritovaya Series (c. 2000 m) could have been emplaced over a period of some months or less.

Volcanic sources and Jurassic geodynamics Volcanics were either acid and/or basic, suggesting that magma sources included both mantle and continental crust. From Pliensbachian to Early Aalenian times, evidence of acid eruptions suggests that partial melting took place almost exclusively within the continental crust. However, mantle melting and crustal underplating could also have occurred

leading to the initiation of crustal melting. The rare calc-alkaline Pliensbachian mafic eruptions suggest that mantle melting took place within the lithospheric mantle. Basic to intermediate magma eruptions spanned from Middle Aalenian to Bajocian times. N- and E-MORB lava eruptions characterized the Aalenian while calc-alkaline basic eruptions typified the Bajocian. Additionally, alkaline basic volcanoes formed synchronously close to the N- and E-MORB eruptive centres in Aalenian times and to the calc-alkaline basaltic eruptive centres in Bajocian times. In addition, E-MORB lavas erupted close to the calc-alkaline basaltic lavas in the Bajocian. The variety of lavas recorded at a single locality reflects the synchronous involvment of several mantle sources and the related differences in composition. This also highlights how the mantle may be inhomogeneous in space. In terms of the geodynamic context of eruptions, the Aalenian volcanic zones (?island chains) were the sites of significant continental lithospheric thinning; this thinning was extreme but there is no evidence of any rupture of the continental crust. The contemporaneous alkaline basalts were derived from the older lithospheric mantle. The Bajocian-age calc-alkaline volcanic area formed c. 100 km to the south of the Aalenian-age one, that is, not so far from the oceanic subduction zone evidenced by the Lesser Caucasus ophiolitic suture, and possibly at the front of the active margin. Both the magmatic signatures and the Nd initial isotopic ratios of the Bajocian volcanics suggest that the genesis of the magmas was related to the partial melting of an old (600 Ma minimum) and heterogeneous lithospheric mantle at different depths and extending as far down as the boundary with the asthenosphere. This mantle most likely formed part of the overlying plate.

Early Mesozoic history of the Greater Caucasus region Analysis of the various datasets of the Western Greater Caucasus region provides evidence of a very large, shallow to deep marine basin, which was sometimes emergent. The degree of plate deformation (extension/compression) varied from one portion of crust to another, as evidenced both by the sedimentary and the volcanic record. The tectonic, magmatic and sedimentary evolution of the western part of the Western GC region in Triassic and Jurassic times in the Fore-Caucasus, Central Greater Caucasus and Southern Slope areas, as described above, serves to better constrain the history of the region (Figs 19, 20 & 21). As noted, the successions in the Western GC are generally not continuous, but rather comprise

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227

Fig. 19. Palaeogeographic maps of the Sinemurian, Pliensbachian, Toarcian, Aalenian, Bajocian and OxfordianKimmeridgian of the Western Greater Caucasus area. Ages after Ogg 2004. (For easier orientation, the main towns in the region are indicated, cf. Fig. 3.)

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Fig. 20. Possible evolution of the geodynamic setting of the GC region in Early and Middle Jurassic times (based on data from this study and Adamia et al. 1981; Gamkrelidze 1986; Dercourt et al. 2000; Nikishin et al. 2001; Gaetani et al. 2005; Okay et al. 2006; Saintot et al. 2006a, 2007).

isolated outcrops which are mainly correlated on the basis of biostratigraphic and lithostratigraphic data. Additionally, the areas are subdivided into ‘zones’ which are difficult to trace in the field; the

boundaries between different zones are generally mapped as major faults, but these faults are not always seen to crop out. Given these difficulties, and the fragmentary nature of the stratigraphic

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

229

Fig. 21. Summary of the evolution of the Greater Caucasus region (based on data from this study and Adamia et al. 1977, 1981; Gamkrelidze 1986; Adamia & Lordkipanidze 1989; Dercourt et al. 1993, 2000; Okay & S¸ahintu¨rk 1997; Nikishin et al. 1998a, b, 2001; Okay 2000; Brunet et al. 2003; Ershov et al. 2003; Robertson et al. 2004; Gaetani et al. 2005; Okay et al. 2006; Saintot et al. 2006a, b, 2007).

record in the Western GC, it is problematic at best to reconstruct the precise history of the region. However, the recognition of important lateral facies changes (e.g. in Aalenian times, see below) would suggest that the subdivision of the area into ‘zones’ may be more a facies-related feature than a real subdivision into classical zones (i.e. in a tectonic/palaeogeographic context). Our overview of the evolution of the region, commences with a review of the Triassic, followed by a more detailed presentation of the Jurassic, subdividing the

stratigraphy into three stratigraphic units (following the international subdivision into Lower, Middle and Upper Jurassic).

Triassic The Triassic record of the Western Greater Caucasus suggests that Lower and Middle Triassic depositional settings were generally continental to nearshore (with some evidence of distal shelf settings in the Tkhach Group). In the Sakhrai Group,

230

T. MCCANN ET AL.

the settings are mainly alluvial and deltaic, although there is also some suggestion that distal shelf environments were present. Additionally, there is evidence of deformation by late Anisian times, as evidenced by the presence of folding and vertical faulting (e.g. in the Acheshbok Fm, the uppermost formation of the Tkhach Group). The deformed units are overlain by a Late Anisian –Ladinian-age conglomerate (Sakhrai Group) which contains a variety of lithic fragments. These include carbonates from the underlying shelf deposits, but also metamorphic and magmatic fragments. Gaetani et al. (2005) suggested that these volcanic clasts were derived from coeval volcanic activity, although there is no evidence of synvolcanism (e.g. interbedded ash/tuff layers, volcaniclastics) within these successions. Indeed, this argument rests solely on the presence of the transported fragments, and given the deformational history of the period, as well as the predominance of transform activity, it must be assumed to be equivocal. The Late Triassic succession is characterized by the carbonatedominated Khodz Group, unconformably deposited on the subjacent deformed units in a shelf setting. In the Malaya Laba area, a Lower Triassic conglomeratic unit which includes serpentine schist pebbles has been recorded, and Gaetani et al. (2005) have suggested that these clasts were derived from rocks metamorphosed during oceanic subduction southwards of the Caucasus region. Based on this interpretation, the relative abundance of serpentine schist grains (present in the formations up into the Norian at which point they become negligible, Gaetani et al. 2005) suggests that either exhumation and/or the diapiric rise of hydrated mantle slivers or Palaeozoic meta-ophiolites, detached from the subducting Palaeo-Tethyan slab, along either a normal fault or a trench-related strike–slip fault occurred (Gaetani et al. 2005). This led Gaetani et al. (2005) to suggest that there was a major deformation phase (i.e. Eo-Cimmerian event) in Middle Triassic times. The Mesozoic evolution in the Caucasus region commenced in (?Permo-)Triassic times with a possible period of subduction. A magmatic arc, linked to this subduction event(s) is recorded in the Pontides, and the presence of ophiolites there provide further evidence for ocean crust consumption (Robertson et al. 2004). In Early and Middle Triassic times the Scythian Platform region was rifted (Tikhomirov et al. 2004), while the Greater Caucasus area (separated from the former by the Stavropol High) was, for the most part, a stable subsiding region until the late Middle Triassic phase of compressive deformation (Fig. 21). However, evidence from serpentinite pebbles in Lower Triassic conglomerates in the Fore-Caucasus area would suggest that the Peredovoy area was uplifted (i.e. related to the

Gaetani et al. (2005) Eo-Cimmerian event). However, the uplift and erosion of metamorphic rocks and the deposition of clastic wedges can also occur in continental rift settings. Additionally, it should be noted that the Devonian ophiolitic belt of the Peredovoy Range was deeply buried, and its exhumation would have necessitated a significant amount of uplift. As noted above, there is evidence for a deformational event in late Anisian times. Indeed, Gaetani et al. (2005) have suggested that this deformation episode was probably the most important one for the Cimmerian Orogeny in the NW Greater Caucasus. Folded units were observed in the Tkhach Group, whose lowermost part contains highlydeformed limestones (similar to the Anisian-age limestones displaying chevron folding, site 30 – see Fig. 8c in Gaetani et al. 2005) and in the upper part comprise conglomerates (some possibly debris flows) which rework the underlying limestones. Our observations would suggest that the inversion of the GC basin occurred prior to the Late Triassic and this may have been synchronous with the onset of compressional deformation along the subduction zone located further to the south, in the Lesser Caucasus and Turkey. As previously noted, there is evidence for the accretion of oceanic plateaux, oceanic islands and volcanic arcs along the subduction zone in the Turkish Pontides (Okay 2000; Robertson et al. 2004; Okay et al. 2006) and to the collision of the Iran continental plate (cf. Dercourt et al. 1993, 2000) which has been constrained by new data within the framework of MEBE program as not later than Carnian times (Zanchi et al. 2006). The docking of terranes along the margin resulted to the partial closure of PalaeoTethys with the concomitant inversion of the Caucasus and Transcaucasus regions. Collision resulted in crustal thickening (evident by Norian times) and the development of Neo-Tethys to the south of the Iran plate. Evidence from the Iranian ophiolites suggests that subduction was rapid. Collision-related uplift meant that there was little deposition of Rhaetianor Hettangian-age sediments in the Greater Caucasus region (where present, they tend to be continental). Far to the north in the area of the Scythian Platform, there is evidence of rifting in Late Triassic times. Here, a coeval volcanosedimentary unit is observed to be overlain by the Rhaetian-age Zurmutin Fm succession of continental sediments (c. 300 m, and which includes rhyolitic tuffs). This evidence of uplift of the continental lithosphere and the related extensive subaerial volcanism in the Scythian Platform region (Nazarevich et al. 1986; Tikhomirov et al. 2004; Chalot-Prat et al. 2007) was coeval with a significant episode of rapid and very extensive regression, possibly global, which occurred at the end of the Triassic (Hallam 2001).

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

Lower Jurassic (Hettangian-Toarcian) (Figs 19, 20 & 21) The oldest Jurassic-age sediments recorded in the Western Greater Caucasus are Sinemurian in age. The underlying Hettangian and Rhaetian– Norian are generally missing (although in the northern part of the area, some Hettangian-age sediments are present). Additionally, there are a number of unconformities recorded, particularly significant ones being present at the base of the Sinemurian, the boundary of the Sinemurian –Pliensbachian and the boundary of the Pliensbachian –Toarcian. Furthermore, the boundary between the Lower and Middle Jurassic is marked by a major unconformity. The Lower Jurassic sedimentary succession from the Fore-Caucasus region comprises the Bugunja and lower part of the Psebai formations. The former was deposited in a coastal setting, while the sediments of the Psebai Fm were laid down in a clastic-dominated shelf setting, thus forming a deepening succession. In the Central Greater Caucasus area, the sediments present (Turovoy & Bzerpia fms) are typical of a distal shelf environment (with possible turbidites present). In the Southern Slope area, the depositional setting is characterized by nearshore/coastal floodplain environments (Estasadok, Tchvejips & Illaryonov fms). At the onset of the Jurassic, the Western Greater Caucasus region was located on a passive margin since subduction has been interrupted/or disturbed by the accretion of both oceanic and continental slivers along the subduction zone (see above). Subsequent to the inversion (?Eo-Cimmerian) of the Permo-Triassic basins in the GC region, Jurassic sedimentation commenced in Sinemurian times. At this time low-energy conditions were widespread across the region, and there is no evidence of tectonic activity. We would suggest that this period was related to the postorogenic establishment of a passive margin along the southern Greater Caucasus region. The subsequent Early and Middle Jurassic evolution of the Western Greater Caucasus region corresponds to a renewed period of rift basin formation, with associated volcanism in the Greater Caucasus region. The Sinemurian –Pliensbachian-age sediments of the Western GC are represented by coastal and shelf sandstones in the northern part of the belt, deeper-marine mudstones and sandstones in the central part and by nearshore/coastal mudstones and sandstones with intercalated Pliensbachian-age rhyolitic tuffs and basalts in the south. In Toarcian times, in northern and central areas, shelf to deepmarine siliciclastics were deposited without any record of volcanic activity in the western part (although extrusives have been reported from the

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eastern GC). The Southern Slope region is characterized by nearshore/floodplain environments. In Pliensbachian times, there is evidence of normal faulting together with volcanic activity (Lordkipanidze 1980; Lordkipanidze et al. 1989). In the north (Bugunja & Turovoy fms) there is evidence for shelf sedimentation, with gradual deepening between the Northern Caucasus and the Central Greater Caucasus (Psekhako-Berezov Zone). In this latter zone, volcanic activity in the form of rhyolitic and trachyandesitic tuffs is also recorded. These were mainly subaerial in character, suggesting emergent conditions (volcanic islands?). Nearshore/coastal conditions continued in the south (e.g. Tchvejips Fm). The Toarcian was a period of basin deepening and the successions in the north (e.g. Psebai & Bzerpia fms) were deposited in deeper shelf settings. In the Southern Slope area (Krasnaya Poliana region), however, shallow marine conditions were still extant (i.e. Illaryonov Fm). The development of the GC rift basin in Pliensbachian times, following a Sinemurian period of quiescence (see above), suggests the renewal of subduction, southwards of the earlier accretionary belt (i.e. Eo-Cimmerian). This early phase of subduction had been interrupted due to the Triassic accretion of the Iranian plate along the Talesh-Alborz-Aghdarband Zone, south of the South Caspian Sea, and possible oceanic slivers along the Pontides (Okay 2000; Okay et al. 2006). Renewed subduction in Early and Middle Jurassic times was probably likely centred south of the Lesser Caucasus and the subduction-related volcanic arc, encompassing the Transcaucasus region, more specifically the ArtvinBolnisi Zone (Fig. 1; Gamkrelidze 1986; Adamia & Lordkipanidze 1989). This region is considered to form the eastern edge of the subduction-related volcanic arc of the Pontides (Ruban 2006). In the Transcaucasus area, there is evidence of shallowmarine to continental deposition as well as extensive calc-alkaline volcanism (Adamia et al. 1981; Gamkrelidze 1986; Adamia & Lordkipanidze 1989; Panov 2004). Adamia et al. (1977, 1981), Gamkrelidze (1986) and Ruban (2006) have all suggested that the Transcaucasus area was subdivided into two blocks (Northern and Southern Transcaucasian subduction-related arcs), each of which formed an archipelago of small islands within the basin. In such a scenario, the Greater Caucasus Basin belonged to a back-arc rift ribbon that extended from the Pontides (e.g. Okay & S¸ahintu¨rk 1997) possibly as far as the South Caspian Basin (Brunet et al. 2003). The basin itself was probably not a single entity, but rather formed a series of rhombshaped en echelon depocentres (Banks & Robinson 1997) which were interconnected, and thus formed a regionally-extensive area of sedimentation. Ruban (2006) suggests that the basin (i.e. his Caucasian

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Sea) was at its widest and deepest to the west, and was open to the Neo-Tethys Ocean.

Middle Jurassic (Aalenian-Callovian) (Figs 19, 20 & 21) The Middle Jurassic succession is generally well represented in the Western Greater Caucasus. A major unconformity separates this succession from the underlying Lower Jurassic. Additionally, there is evidence of a major unconformity between the Bathonian and Callovian (indeed, this interval may be entirely absent in some areas). In the ForeCaucasus region, the Middle Jurassic succession commences with the Psebai and Djangur formations which were deposited in a clastic shelf setting. The younger (and related) Kamennomostsky Fm (Bathonian–Callovian) is carbonate rich, and was probably deposited in a carbonate shelf system. In contrast, the Aalenian-age successions of the Central Greater Caucasus area range from subaerial (Tchataltapa Fm) through to coastal/floodplain (Makarov Fm) to full marine conditions, that is, below storm wave base (Naujni & Mount Indyuk fms). The observed trend varies between a shallowing succession in the Psekhako-Berezov Zone, to a deepening succession (into Bajocian times) in the adjacent Goitkh Zone. The Aalenian-age sedimentary succession in the Southern Slope area begins with nearshore/coastal floodplain sediments (Krasnopoliana Zone) which continue through to the Bajocian. In the adjacent Abkhazo-Racha Zone the Bajocian-age succession is similar, with nearshore/ continental floodplain sediments predominant. The Middle Jurassic succession across the Western Greater Caucasus is predominantly characterized by shelf sediments in the Fore-Caucasus area, while further south both subaerial and shallow-(to deeper-)marine settings are represented. This succession would appear to be the most diverse in terms of the depositional environments present and lateral variations in sedimentary facies are marked. The presence of calcareous rocks towards the top of the succession suggests the growth of a carbonate shelf/ramp in areas which were shallow enough and/or had little clastic input. The Aalenian was characterized by renewed rift activity (Lordkipanidze et al. 1989; Se´brier et al. 1997) characterized by regional east –west-oriented extension, extending even as far as the South Caspian Sea area (Brunet et al. 2003). The pronounced increase in lithic fragments within the sedimentary record possibly signalled this significant Aalenian-age rifting event (i.e. rift reactivation), which was coeval with an increase in mainly subaerial volcanic activity. The east –west-orientated extension in the GC resulted in significant strike–slip activity along the margin and the

basins which were likely pull-apart in type (Se´brier et al. 1997; Saintot et al. 2006a; this paper). The spatial distribution of the mudstones from the Tchataltapa Fm (observed at different localities) would thus, for example, not have been in continuity, but rather deposited in two distinct settings, both of them low energy, but one located in a shallow environment while the other was both deeper and distal to source. In addition, the deep-marine mudstones are associated with pillow lavas whereas the shallow depositional setting is associated with subaerial flows. Such an interpretation would certainly fit with the idea of structural highs and lows within the basin. Additionally, the abrupt facies changes observed are typical of pull apart basins (cf. Nilsen & Sylvester 1995). This example also demonstrates the degree of lateral variation present within the region, and which is not reflected in the Russian mapping scheme (i.e. the ‘zones’ alluded to above; at some future time, these will be superseded by more detailed regional mapping). As noted above, volcanic activity became more prevalent in Aalenian times, particularly in the northern part of the GC basin and the level of volcanic activity increased markedly into Bajocian times (possibly as a result of the northward consolidation of the Transcaucasus arc in Early Bajocian times, Ruban 2006). During Aalenian times, the evolution of the Western Greater Caucasus region was similar to that of an oceanic spreading back arc since N-MORB and E-MORB successions were emplaced on a continental crust and gabbros were probably underplated at the base of this continental crust. Hence, there was coeval destabilization and melting of the asthenosphere and the asthenosphere/lithosphere boundary. The two Bajocian basalt types, namely, calc-alkaline and E-MOR, were derived from the melting of the subcontinental mantle close to the Moho and close to the lithosphereasthenosphere boundary, respectively. These mantle melts at different depths would have occurred because of the shallowing of a northwards subducting oceanic plate which in turn would have induced uplift and adiabatic decompression of the overlying lithospheric plate (Fig. 20). The subaerial nature of the eruptions during the second half of the Bajocian volcanic period would suggest that uplift occurred. The precise setting of oceanic subduction at this time, however, is only constrained by the outcropping coeval ophiolitic suture in the Lesser Caucasus (Fig. 1). However, the fact that the mantle-derived volcanics in the Western Greater Caucasus region have a calc-alkaline signature does not necessarily imply that they were emplaced above a subducting oceanic plate. From the literature, it would appear that calc-alkaline basalts can erupt during active continental rifting processes either immediately

EARLY MESOZOIC EVOLUTION OF THE WESTERN GREATER CAUCASUS

following continental collisional events (ChalotPrat & Girbacea 2000; Chalot-Prat et al. 2007) or over a more extended time frame afterwards, as is the case for intraplate settings (Chalot-Prat et al. 2007 and references therein). The Bajocian-age volcanic activity in the Greater Caucasus region was localized on the southern part of the present-day belt and has been related to the expansion of the subduction-related volcanic arc from the area of the Lesser Caucasus to the southern margin of the GC back-arc basin (Saintot et al. 2006a; Fig. 20) or to the consolidation of a secondary Transcaucasus arc (Ruban 2006). However, the northernmost Bajocian unit cropping out along the northern flank of the present-day western Greater Caucasus belt (i.e. Zeitun Fm) is non-volcanic and this suggests that at this time the northern part of the Greater Caucasus Basin was distal to any volcanic activity, since it did not receive any significant influx of material from the Bajocian-age volcanic belt which had developed along the southern basin margin. The Bathonian sedimentary succession (where present) comprises siliciclastic deep-marine deposits in the basin centre and coal-bearing continental sediments on the southern basin margin. However, the Betag Fm (500 m) in the Krasnopoliana Zone shows evidence of volcanoclastic sandstones, mudstones and siltstones. A possible fluvial/continental depositional environment can be postulated for this formation. The Callovian saw the establishment of a broad carbonate depositional environment across the Greater Caucasus area. Deposition occurred in a ramp or possibly a platform setting. The depositional environment continued uninterrupted into the Oxfordian. As noted above, in the central and eastern part of the Greater Caucasus Basin, the Callovian is frequently unconformable on the Bathonian, whereas in the rest of the region, the boundary is transgressive and discordant (Gamkrelidze 1986; Zonenshain et al. 1990; Nikishin et al. 1998a, b). Those parts of the Greater Caucasus where there is no evidence of an unconformity correspond to the deepest parts (depocentres) of the Middle Jurassic Greater Caucasus Basin (Zonenshain et al. 1990). The unconformity has been interpreted as being related to a Bathonian-age compressional event which resulted in uplift of the basin margins. In the southern part of the Greater Caucasus, Bathonian folds are observed, and these are overlain by Callovian-age sediments. In both the northernmost and western parts of the Greater Caucasus, there is no evidence of an angular unconformity between the transgressive Callovian and the underlying strata (although the outcrops are very sparse and tracing the continuity of formations is problematic) (Saintot et al. 2006a).

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Therefore, it is difficult to say if there was any preCallovian deformation since much of the deformation would appear to be Cenozoic in age. If there was a pre-Callovian deformation phase, it has been so strongly overprinted by subsequent Alpine deformation that it is no longer recognizable. Indeed, in the Early –Middle Jurassic-age rocks of the northern part of the belt, where the Alpine folding and thrusting attenuates, there is no evidence of intense deformation (e.g. folds) which might be related to a Bathonian compressional event. The localized angular unconformity at the base of the Callovian rocks may thus record a ‘soft’ phase of compression of the GC basins locally affecting some units but not being directly related to the complete inversion of the Greater Caucasus Basin (cf. Zonenshain et al. 1990). Indeed, in early Middle Jurassic-age rocks there is no evidence of intense deformation which might be related to a Bathonian compressional event – again much of the observed deformation is Cenozoic. Thus, the Bathonian compressional event only resulted in inversion along the basin margins, as previously suggested by Adamia et al. (1981). However, it could also be that this unconformity is simply related to the cessation of rifting and the onset of post-rift basin development (cf. Frostick & Steel 1993; Ingersoll & Busby 1995; McCann & Saintot 2003). Brunet et al. (2003) have noted that the regional Bathonian unconformity around the South Caspian Basin may be a ‘break-up unconformity’ marking the onset of sea-floor spreading rather than the occurrence of a compressive tectonic event. Thus, the pre-Callovian unconformity remains a matter of debate. It could record either the transition between the syn- and post-rift phases in the GC Basin, or, alternatively, it may record a weak compressive event related to the accretion of crustal blocks along the active continental margin to the south (Sengo¨r 1984) and/or the expansion of the subduction-related volcanic arc from the Lesser Caucasus to the southern margin of the GC basin (Saintot et al. 2006a and see above). The Bajocian-age volcanics of the GC represent the volcanic arc which expanded from the Pontides – Lesser Caucasus to the southern margin of the GC (via Transcaucasus) – Shatsky Ridge and Crimea (where a huge calc-alkaline pile of volcanics crop out). In such a situation, the volcanic arc would have needed to be very wide (some hundreds of kms). Such volcanic arc expansion occurs when there is shallowing of the subducting oceanic slab (Fig. 20). Associated with this shallowing was the onset of compression in the back-arc region (i.e. producing a pre-Callovian unconformity; cf. Lallemand et al. 2005) as well as uplift (resulting in the formation of Bathonian continental coal-bearing successions). The return of the slab to normal

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steepness would have resulted in the subsidence of the back-arc region and the development of related accommodation space (subsequently infilled by Callovian-Late Jurassic sediments, and particularly by the thick Cretaceous –Cenozoic succession). The opening of the Black Sea (and more specifically of the Eastern Black Sea Basin) ruptured the wide Bajocian volcanic arc (Fig. 20) and remnants of it are found along the northern and southern margins of the Black Sea (in Crimea and the Pontides, respectively). This situation continued until the onset of Eocene inversion (Fig. 21). Thus, Bathonian compression may have resulted from changes in subduction dynamics, leading to or resulting from changes in either the slab dip (as outlined above), the relative convergence rate (which could also vary along strike), and the heterogeneity of the subducted material, and not as a result of collision or some related mechanism (e.g. the supposed Meso-Cimmerian orogenic phase). Additionally, subduction was continuous from the Pliensbachian (following the Eo-Cimmerian event) until at least the Late Cretaceous and later in the area to the south (i.e. along the Lesser Caucasus and Pontides).

Upper Jurassic (Oxfordian-Berriasian) (Figs 19, 20 & 21) The Upper Jurassic succession of the Western Greater Caucasus region was not studied in any detail. The succession is bounded by an unconformity at the top (intra-Berriasian) and the base (base Oxfordian), while less extensive unconformities are also present, most notably in the lower Berriasian. In the Fore-Caucasus area, the Upper Jurassic is characterized by a calcareous succession deposited in a ramp/platform setting. The only Upper Jurassic sediments present in the Greater Caucasus region are within the Abino-Gunai Zone, and comprise shelf carbonates, deposited in a shallow-marine setting. To the south, in the Southern Slope region, the sediments present are also carbonate rich, and were deposited in a similar depositional environment. Thus, there is a distinct difference between the depositional environments present in the Lower and Middle Jurassic and those present in the Upper Jurassic. The older sediments were deposited in clasticdominated settings while the Upper Jurassic sediments are carbonate rich and were deposited in shallow-marine environments. The Late Jurassic marked the beginning of a period of relative tectonic quiescence. The deepening of the basin corresponds to the thermal subsidence of the GC basin in its postrift stage (although part of the subsidence could have been triggered by a steepening of the subducting slab; see discussion above).

This paper benefited from fruitful discussions with Russian colleagues, including: S. Korsakov, P. Fokin and P. Tikhomirov. We would like to thank D. Ruban for sending us his publications. Figures were partly drawn by B. Krumbiegel. Part of this research was funded by the MEBE programme. This manuscript was greatly improved by the reviews of M. Gaetani, M. Allen and two anonymous reviewers.

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Evolution of the Late Cenozoic basins of Georgia (SW Caucasus): a review SH. ADAMIA1*, V. ALANIA2, A. CHABUKIANI3, G. CHICHUA3, O. ENUKIDZE2 & N. SADRADZE2 1

Institute of Geophysics, 1 M. Alexidze str., 0193, Tbilisi, Georgia

2

Geological Institute, 1/9 M. Alexidze str., 0193, Tbilisi, Georgia

3

National Agency for Oil and Gaz, 45 Kazbegi av., 0177, Tbilisi, Georgia *Corresponding author (e-mail: [email protected])

Abstract: The tectonic and geological evolution of Georgia and the Caucasus, on the whole, are largely determined by its position between the still converging Eurasian and Africa– Arabian lithosphere plates, within the wide zone of a continent– continent collision. The region in the Late Proterozoic– Early Cenozoic belonged to the now-vanished Tethys Ocean and its northern (Eurasian) and southern (Africa–Arabian) margins. Within this convergence zone there existed a system of island arcs, intra-arc rifts, back-arc basins characteristic of the pre-collisional stage. During syncollisional (the Oligocene–Middle Miocene) and post-collisional (the Late Miocene–Quaternary) stages, at the place of back-arc basins were formed fold and thrust belts of the Greater and Lesser Caucasus separated by the Transcaucasian intermontane lowland. Starting from the Late Miocene and as far as the end of the Pleistocene, in the central part of the region, simultaneously with formation of molassic basins and accumulation of coarse molasses there took place volcanic eruptions in subaerial conditions. According to the numerous data obtained during past decades we present a review on the lithological and structural characteristics of these collisional basins and on the coeval magmatic events.

Tectonics and geological evolution of Georgia and the Caucasus, or Black Sea –Caspian Sea region, on the whole, are largely determined by its position between the still converging Eurasian and Africa– Arabian lithosphere plates. The Phanerozoic development of this area has been considered and discussed during past decades in a great number of publications (e.g. Khain 1974; Dercourt et al. 1986, 1990; Robertson et al. 1996). According to some authors (Adamia et al. 1977, 1981; Zakariadze et al. 2007), the Black Sea –Caspian Sea region in the Late Proterozoic, Palaeozoic, Mesozoic and Early Cenozoic belonged to the now-vanished Tethys Ocean and its northern (Eurasian) and southern (Gondwanian/Africa –Arabian) margins. Within this convergence zone, a system of island arcs, intra-arc rifts and back-arc basins occurred. They are characteristic of the pre-collisional stage (Late Proterozoic–Early Cenozoic). The continent– continent collision related to syncollisional (the Oligocene –Middle Miocene) and post-collisional (the Late Miocene– Quaternary) stages of the Late Alpine tectonic cycle caused inversion of the relief. Consequently, at the place of backarc basins were formed fold and thrust belts of the Greater and Lesser Caucasus with the Transcaucasian intermontane lowland instead of

Transcaucasian rigid blocks (microcontinent, island arcs). The marine basins were replaced by hemi-closed basins of euxinic type (Paratethys) and later on (the Late Miocene, c. 9–7 Ma) by intracontinental basins with subaerial and coastal conditions of sedimentation. In this paper we present the main lithological characteristics of these upper Cenozoic basins which recorded the tectonic events due to the collision stages.

Main tectonic zones The Caucasus is divided into several large tectonic units, which differ one from another by their stability-lability degree; there are distinguished rigid (platform, sub-platform, quasi-platform) and fold-thrust units (Papanikolaou 1997). They are (from the north to the south): the Scythian (pre-Caucasus) young platform, the fold-thrust mountain belt of the Greater Caucasus including zones of the Northern Slope, Fore Range, Main Range and Southern Slope; the Transcaucasian intermountain depression superimposed mainly on the rigid platform zone (the Georgian Block); the Achara –Trialeti and Talysh fold and thrust belts; the Artvin– Bolnisi rigid sub-platform zone; the Bayburt–Karabakh mountain belt; the North

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 239–259. DOI: 10.1144/SP340.11 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Anatolian– Lesser Caucasus ophiolitic suture zone; the Lesser Caucasian part of the Taurus – Anatolian– Iranian platform and, at the extreme south of the Caucasus, the Araks intermontane depression. Neogene –Quaternary continental volcanic formations overlie the Eastern Anatolian, Armenian and South Georgian highlands. Some centres of extinct volcanoes in the Greater Caucasus (Elbrus, Chegem, Keli, Kazbegi) occur. The territory of Georgia includes only some of the mentioned zones: the Greater Caucasus and Achara –Trialeti fold-and-thrust belts, the Rioni and Kura intermontane depressions (foreland basins), the North-transcaucasian (the Georgian Block) and South-transcaucasian (the Artvin– Bolnisi Block) terranes (Papanikolaou 1997), the Javakheti and Keli –Kazbegi volcanic highlands and extinct volcanoes (Fig. 1). In this region, sedimentary, magmatic and metamorphic rocks as Late Proterozoic –Phanerozoic are the results of various palaeogeographic and geodynamic environments: ocean and small oceanic basins, intercontinental areas, active and passive continental margins (Adamia et al. 1977, 1981).

Late Cenozoic, syn- and post-collisional stages The Oligocene epoch is traditionally considered as a beginning of syncollisional stage of development of Georgia and the Caucasus, on the whole (Milanovsky & Khain 1963; Gamkrelidze 1964). The deposits of the Oligocene –Lower Miocene resulted at the most part of the territory of the Caucasus from accumulation of detrital and gypsiferous facies in hemi-closed basins of the Paratethys. They are known as the Maikop series with thickness of several kilometres in the deepest parts of the intermontane basins (Rioni, Kura). The Oligocene rocks only locally are represented by some other facies (Gamkrelidze 1964; Yılmaz et al. 2001): siliceous (opoka, spongiolite in the Dzirula massif), marine terrigenous and continental coal-bearing deposits (the central part of the Achara –Trialeti). Accumulation of predominantly shallow-sea and lagoon –lacustrine terrigenic clastics with layers of organogenic carbonate rocks (coquina) lasts almost throughout the Miocene epoch. Only at the end of the Miocene time (Tortonian), shallow-marine environment is replaced by subaerial surroundings and simultaneously clastic material becomes coarser. Marine environment of the sedimentation remains only within adjacent areas to the Black, Azov and Caspian Sea basins. Starting from the Late Miocene and as far as the end of the Pleistocene, in the central part of the region, simultaneously with formation of molassic

troughs and accumulation of coarse molasses there have occurred volcanic eruptions in subaerial conditions (e.g. Skhirtladze 1958; Gamkrelidze 1964; Milanovsky & Koronovsky 1973). Volcanic highlands and plateaus were formed in Javakheti (South Georgia); Elbrus, Chegem, Kazbegi and Keli plateaus, in the Greater Caucasus. Emplacement of small hypabyssal intrusions occurred in the Transcaucasus and Greater Caucasus (Arakelyants et al. 1968). The two main stages were distinct: 1) the Late Miocene –Early Pliocene; and 2) the Pliocene –Quaternary. Lavas predominate over volcanoclastics, especially during the second phase of eruptions (Skhirtladze 1958; Adamia et al. 1965).

Late Cenozoic molassic basins of the Transcaucasus In the Palaeogene period, the Caucasus was a part of the extensive arc-back-arc system of the Tethys (Dercourt et al. 1986, 1990). Under conditions of rapidly shrinking oceanic space, the Tethys Ocean was reduced to intercontinental basins of Mediterranean type. It is classically considered that its northern oceanic branch (North Anatolian –Lesser Caucasian) had been closed by the Late Palaeogene –Early Miocene time when the Southern Armenian Block was accreted to the Southern Transcaucasus (e.g. Knipper 1980). However some recent results indicate that the closure of the northern Tethys branch occurred since the Paleocene (Sosson et al. 2010). The collision between the Africa– Arabian and Eurasian plates caused inversion of the relief and at the place of intra-arc and back-arc basins were formed two fold-andthrust belts of the Greater and Lesser Caucasus with the Transcaucasian intermontane depression in between (Fig. 2). The marine basins of the Tethys were replaced by basins of euxinic type (the Paratethys) and later on – by continental basins with subaerial condition of sedimentation (Figs 3 & 4). The Dzirula high subdivides the Transcaucasian intermontane depression into two basins: Rioni (Black Sea) in the west and Kura (Caspian Sea) in the east.

Rioni basin The Oligocene–Lower Miocene (the Maykopian series) is represented (like within the Kura basin) mainly by alternation of gypsiferous clays with sandstones (Gamkrelidze 1964; Ananiashvili et al. 2000) that is highly specific for the Eastern Paratethys. These rocks are featured by their brown colour, poor carbonate content, scarcity of faunal fossils (except for fish scales), abundance of jarosite on bedding planes, intralayers and veinlets of

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Fig. 1. Tectonic map of Georgia, scale 1:1 000 000; A and W, location of the lithostratigraphic columns; location of cross-sections: Od, Odishi; R, Rioni; Ka, Kartli; A, Alazani; K, Kura; Kv, Kvabebi.

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columns 1–3, Fig. 4b). This time a sublatitudinal marine basin was developing within the area between the Southern Slope of the Greater Caucasus in the north and the Achara–Trialeti belt in the south. Maximal thickness of the Middle – Upper Miocene sediments achieves 1500–3500 m. In the axial parts of these subsided structures, within the intermontane depression, deep-marine carbonate clays and sandstones with marls were accumulated, while in their marginal zones shallowmarine terrigenous-carbonate rocks of limited thickness were deposited. Littoral facies show the presence of coarse clastic rocks (Gamkrelidze 1964; Eastern Paratethyan 1985; Ananiashvili & Minashvili 2004). Abundance of various marine fossils (molluscs, foraminifera) (Gamkrelidze 1964; Janelidze 1970, 1977; Popkhadze 1983; Ananiashvili & Sakhelashvili 1984; Jones & Simmons, 1997) observed from the outset of the Middle Miocene (Tarkhanian regiostage) indicates establishment of wide connections of the basin with the ocean and the replacement of euxinic environment by normal marine conditions. However, later on, the connections with the oceanic basin periodically either ceased (the Chokrakian, Karaganian stage, Middle –Sarmatian times) or were impeded (the end of the Karaganian, the onset and the end of the Konkian; the end of the Middle Sarmatian) (Maisuradze 1971; Eastern Paratethyan 1985).

Fig. 2. Correlation of the Oligocene– Quaternary scheme of the Mediterranean and Black Sea–Caspian Sea regions (Gamkrelidze 1964; Haq & Van Eisinga 1987; Yilmaz et al. 2001; Taktakishvili 2000; Remane et al. 2006).

gypsum. Very frequent are disk-shaped concretionsseptaria of spherosiderite, up to 0.5–3.0 m in diameter, located, as a rule, along the bedding planes (see Fig. 3a, columns 1– 3 and Fig. 4a). Middle Miocene: Tarkhanian, Chokrakian, Karaganian and Konkian regiostages – Late Miocene: Early and Middle Sarmatian regiostage (Fig. 3a,

Late Miocene: Late Sarmatian regiostages (Fig. 3, columns 1–3, Fig. 4c). From the Late Sarmatian time nearly all the study territory became a dry land except for small parts of the Rioni foreland. Maximal thickness of the Upper Sarmatian deposits is about 1500 m. The continuing development resulted in the further uplift of clastic source areas and subsidence of subaerial –sedimentation areas where predominantly coarse molasses with subdued sandstones and clays were deposited. According to the composition of clastic material, one can easily recognize the northern (mainly the Cretaceous –Palaeogene flysch of the Greater Caucasus) and southern and southwestern source areas (the Mesozoic –Cenozoic volcanic rocks of the Achara–Trialeti zone). Meotian regiostage (Fig. 3, columns 1–3, Fig. 4c). Marine Meotian rocks are predominantly represented by coastal shallow water deposits: conglomerates, gravelites, sandstones, sandy clays, clays and rarely, by shell-beds. They transgressively overlap the Sarmatian and older, Cenozoic and Mesozoic rocks and, in turn, are overlain by the Pontian and younger beds. In some sections, Upper Meotian deposits continuously are replaced by Lower Pontian rocks. The thickness of the

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Fig. 3. Lithostratigraphic columns of the Oligocene, Neogene and Quaternary deposits of (a) Western Georgia (Black Sea basin) and (b) Eastern Georgia (Caspian basin). Columns: 1. Central Apkhazeti; 2. Central Colkhida (well #4 Chaladidi); 3. South Colkhida (Guria); 4. Dzirula high; 5. Kartli basin north; 6. Kartli basin south; 7. Gare Kakheti (Kura foreland); 8. Alazani basin; 9. Akhaltsikhe-Javakheti: Location of the columns see Figure 1.

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Fig. 3. Continued.

Meotian and Pontian sediments varies from a few to several hundred metres or even 1000 m as in the Rioni basin. The most complete sections are in Southern Abkhazeti, bearing rich and variable assemblages of molluscs, foraminifera, ostracoda, moss fossils, fish otoliths and sponge spicules (Gamkrelidze 1964; Popkhadze 1975; Eastern Paratethyan 1985; Maisuradze 1988; Jones & Simmons 1997). Pontian regiostage (Fig. 3, columns 1–3, Fig. 4c). The Pontian of Western Georgia was regressive, as it was in the other parts of the euxinic basin. The marine basins of the consequent stages are gradually shrinking. Their individuality is manifested by mollusc fauna including not only forms common

for the Black Sea realm but also species, frequent in the Dacian and Pannonian basins (Gamkrelidze 1964; Chelidze 1973; Eastern Paratethyan 1985). Upper Pontian section is quite common in Western Georgia; it is represented by both shallow water and relatively deep water facies. The first one is build up of sandstones, sandy clays, carbonates, conglomerates and bear rich mollusc assemblages (Gamkrelidze 1964; Chelidze 1973; Eastern Paratethyan 1985). Pliocene (Fig. 3, columns 1 –3, Fig. 4d): Cimmerian regiostage. Like the Pontian regiostage, the Cimmerian one in Western Georgia has many in common with other contemporaneous formations of the Black Sea area though showing some

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Fig. 4. Palaeogeographic maps of Georgia: (a) Oligocene– Early Miocene, c. 23– 15 Ma BP; (b) Middle–Late Miocene, c. 15– 11 Ma BP; (c) Late Miocene– Early Pliocene, c. 10– 5 Ma BP; and (d) Late Pliocene–Pleistocene, c. 3– 0.01 Ma BP (without palinspastic reconstruction). Rb, Rioni basin; Kb, Kura basin (Gamkrelidze 1964; Chubinishvili 1982; Adamia et al. 1990; Go¨ru¨r et al. 2000; Zirakadze 2002).

individual features. The Cimmerian regiostage consists mainly of sandstones, claystones and conglomerates (thickness c. 200– 300 m). Mollusc assemblage includes several endemic forms

evidencing decreasing salinity in several parts of the Rioni Bay (Gamkrelidze 1964; Eastern Paratethyan 1985; Taktakishvili 1984, 2000). Kuyalnikian regiostage (terrigene clastics, thickness c. 200 m)

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Fig. 4. Continued.

is characterized by complete profiles and richer brackish mollusc assemblages. Recent investigations have shown the greater presence of Western Georgian Kuyalnikian regiostage as compared to the other parts of the euxinian basin due to the presence of transitional Cimmerian –Kuyalnikian beds, almost completely missed in other parts of

the Black Sea Area (Gamkrelidze 1964; Eastern Paratethyan 1985; Taktakishvili 1984, 2000). Pleistocene: Gurian regiostage (Fig. 3, columns 1–3, Fig. 4d). In the Black Sea area, deposits of this regiostage crop out only in Guria, although they were penetrated by boreholes in other areas

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of the Rioni basin. These are sandy clayey rocks (thickness c. 100 –500 m) with homogenous but poor mollusc fauna (Gamkrelidze 1964; Eastern Paratethyan 1985; Taktakishvili 1984, 2000). Dzirula high. Within the Dzirula salient (Fig. 3, column 4), the Oligocene –Lower Miocene is represented by a thin (200 –400 m) manganese-bearing formation of spongolitic sandstones and clays with opoka and chalcedony. The formation unconformably rests over the Cretaceous deposits (Gamkrelidze 1964). The Middle Miocene sequence, which starts with the Tharkhanian regiostage, conformably follows the Early Miocene and is represented by sandstones, clays and marls (thickness 1 –30 m). The Chokrakhian deposits unconformably overlie the older sediments, which themselves are followed by the Karaganian and Konkian regiostages having aggregate thickness 30– 20 m. They are represented by shallow-sea, often variegated sandstones, gravelites, clays, calcareous clays, sandy and oolitic limestones. Their Middle Miocene age was generally confirmed by molluscs (Gamkrelidze 1964; Ananiashvili & Sakhelashvili 1984; Eastern Paratethyan 1985) and foraminifera findings (Janelidze 1970, 1977).

Kura (Mtkvari) basin Oligocene –Early Miocene time (Fig. 3b, columns 5–8, Fig. 4a). Rocks of Oligocene –Early Miocene age in some places conformably follow sandy –clayey sediments of the Late Eocene and all over the Kura foreland are represented by uniform, mostly clayey sequence termed here ‘Maykop series’, locally thin beds and lenses of hematite are observed. Maximal thickness of the Oligocene –Lower Miocene deposits is reported within the basin of the river Kura (c. 2.5–3.0 km). In the vertical section one can observe the gradual change from more sandy varieties (the Khadum horizon of the Early Oligocene age) to more clayey rocks in the top of the section. However, near Tbilisi and further westward, the upper part of the section (the Sakaraulian and Kotsakhurian regiostages) shows considerable increase of coarse-grained, locally cross-bedded sandstones (Robinson et al. 1997). Biostratigraphy of the Maykop formation is chiefly based on foraminifera, nannoplancton and molluscs findings (Gamkrelidze 1964; Janelidze 1970, 1977; Ananiashvili et al. 1998, 2000). Middle Miocene: Tarkhanian, Chokrakian, Karaganian and Konkian regiostages (thickness 500 –700 m). Rocks of the Tarkhanian regional stage are represented mainly by argillites, sandy

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argillites and fine-grained sandstones (Fig. 3b, columns 5 –8, Fig. 4b) containing molluscs, echinids, crayfish, foraminifera, and ostracodes. In the northern and southern margins of the Kura depression these rocks, partly and at some localities fully, were washed out by the transgression of the Chokrakian Sea. Chokrakian rocks are represented by sandy-clayey sediments, however, here, the role of coarse-grained rocks, sandstones, gritstones and conglomerates is more noticeable. Locally, marls and limestones are developed. Biostratigraphy of the Chokrakian regiostage is based on fossil relicts of endemic fauna characteristic of the Eastern Paratethys (molluscs, foraminifera, nannoplancton; Ananiashvili & Minashvili 2004). The Karaganian and Konkian regiostages are represented chiefly by argillites with rare intercalations of fine-grained sandstones and marls. However, in the southern margin of the Kura basin amount of coarse-grained rocks, sandstones and gritstones increases; there appear conglomerates and sandy limestones as well. The both stages contain mollusc and microfauna fossils. Cross-bedding observed in sandstones and gritstones is characteristic of the Konkian stage in the most extreme southern part of the region (Gamkrelidze 1964; Eastern Paratethyan 1985; Ananiashvili et al. 2000). Late Miocene: Early and Middle Sarmatian regiostages (Fig. 3b, columns 5–8, Fig. 4b). The Lower and Middle Sarmatian (thickness 1000–1500 m) are represented mainly by carbonate clays and finegrained sandstones with marl intercalations and oolitic limestones. South- and eastward these rocks pass into coarse-grained sandstones and there appear layers of gritstones. In some localities, the sandstones contain a lot of marine-mollusc fossils that confidently evidence the age of the enclosing rocks. In some places, the carbonate rocks yield an abundance of microforaminifera (Gamkrelidze 1964; Eastern Paratethyan 1985; Maisuradze et al. 2004). Late Miocene: Late Sarmatian regiostages (Fig. 3b, columns 5 –8, Fig. 4c). In the northern part of the region, the Middle Sarmatian rocks grade into the Late Sarmatian (the Natskhori formation, thickness 1000– 1500 m) ones. The bottom of the latter is built up mainly of claystones and sandstones, and top is constructed by clays, sandstones and conglomerates. Clastic material of the coarse-grained rocks is represented almost entirely by the wellrounded clastic rocks of the Mesozoic flysch of the Greater Caucasus (Gamkrelidze 1964). In the central and southern parts of the Kura depression the Late Sarmatian (the Eldari formation, thickness 200–500 m) is represented by coloured clays, crossbedded sandstones, gritstones and conglomerates

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with clastic material derived from the Lesser Caucasus: volcanic, metamorphic and crystalline rocks of the Artvin –Bolnisi and more southern Loki–Karabakh zones, volcanic rocks of the Achara –Trialeti belt (Chubinishvili 1982). The Upper Sarmatian rocks are highly gypsiferous. The rocks show no marine fauna and their age is determined on the basis of their stratigraphic position between the Middle Sarmatian and Meotian regiostages and also by presence of mammal relicts (Eastern Paratethyan 1985; Chkhikvadze et al. 2000). Late Miocene: Meotian and Pontian regiostage (Fig. 3b, columns 5–8, Fig. 4c). The Meotian and Pontian sediments of the region known as the Dusheti suite in Kartli (thickness c. 2000m) and the Shiraki suite in Gare Kakheti (thickness 1500– 2000 m) at 70% consist of conglomerates, often coarse-clastic, represented by Cretaceous flysch rocks from the Greater Caucasus. Another rock types are sandstones and sandy clays interbedded with conglomerates (Gamkrelidze 1964). The amount of coarse-grained rocks increases northward. Along the southern margin of the Kura depression, the noticeable role among the clastic material in conglomerates and sandstones belongs to clastics of volcanic and plutonic rocks transported from the Achara–Trialeti belt, the Artvin –Bolnisi block and the Loki –(Baiburt) Karabakh zone (Lesser Caucasus). The Meotian and Pontian age of the Dusheti and Shiraki formations has been established on the basis of mammal (Chkhikvadze et al. 2000) and tortoises fossils (Gabashvili et al. 2000). Late Pliocene–Pleistocene (Fig. 3b, columns 5–8, Fig. 4d): Akchagylian and Apsheronian regiostages. The Akchagylian regiostage is represented here by continental and shallow marine molasses unconformably overlying the older rocks (maximal thickness c. 1500). Marine facies of the lower part of the section is made up of sandy-clayey rocks; its upper part – of sandstones and conglomerates (Gamkrelidze 1964). Fossil fauna is represented here by endemic brackish molluscs of the Caspian type, and also by ostracoda and foraminifera. Near Kvabebi (Gare Kakheti), along with marine molluscs, there were found relicts of fossil mammals (Eastern Paratethyan 1985). The continental facies of the Akchagylian and Apsheronian regiostages are represented mostly by conglomerates with insignificant sandstones and sandy clays. Locally, thin layers of volcanic ashes occur (Gamkrelidze 1964). The geological time-scale and stratigraphy of the Oligocene –Quaternary formations of the Caucasus have been elaborated in sufficient details (Fig. 2) and are based on long-term (more than a century)

investigations conducted by a great number of scientists from both the former Soviet Union and western countries (e.g. Gamkrelidze 1964; Haq & Van Eisinga 1987; Yilmaz et al. 2001; Taktakishvili 2000; Remane et al. 2006). The geochronological and stratigraphic schemes and correlations of the Mediterranean province (including the Caucasian region) widely accepted at present by majority of researchers are based on the combined interpretation of data from various fields of geology: biostratigraphy, magnetostratigraphy and palaeomagnetism, radiometric timing of minerals and rocks, geomorphic and palaeo-glaciological studies. Since mainly the upper part of the geological time scale of the eastern Caucasus (the Caspian Sea basin) has been subjected to serious revision and modification, in this paper we shall consider only this part of the geochronological scale in respect of the validity of considerable rejuvenation (by 0.6 million years, from 1.8 Ma to 1.2 Ma) of the Akchagyl/Apsheron boundary recently proposed by some authors (e.g. Mitchell & Westaway 1999). One of the most informative geological sections for dating and correlation of the Akchagyl regiostage of the Eastern Georgia is the Kvabebi section located in the Kura molasse depression (see Kv on Fig. 1). The Akchagyl regiostage is represented here by littoral marine sediments containing rich fossil fauna of molluscs, ostracoda and foraminifera (Vekua 1972; Eastern Paratethyan 1985; Chkhikvadze et al. 2000). The lower and middle parts of the section yield marine molluscs, ostracoda and foraminifera. In the upper horizons of the sections, there appear representatives of freshwater molluscs and ostracoda. The Kvabebi section is notable because of alternation of marine-fauna-containing layers with beds bearing reptiles and mammals fossil remains. The whole biostratigraphic material obtained from the Kvabebi section allows, with a good reasons, to correlate the Akchagyl regiostage of the region with the Kuyalnik regiostage of the Black Sea region and also with the Piacenzian regiostage of the Mediterranean area (Fig. 2). Magnetostratigraphic data allow to attribute the strata containing fossil remains of vertebrate and mammals to the Gauss epoch of normal polarity, whereas radiological dating of volcanic ash from the middle part of the section provide figures around 2.55 Ma (Eastern Paratethyan 1985). More detailed information on the absolute age of the deposits assigned to the Akchagyl stage in Kvabebi presented by Zubakov & Kochegura (1974).

Akhaltsikhe basin Oligocene series. The Oligocene section of the Akhaltsikhe depression is complete and well dated

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by fauna (mainly by molluscs). The Oligocene deposits of the Akhaltsikhe depression are subdivided into six suites (Yilmaz et al. 2001): (1) (2) (3) (4) (5) (6)

Poskhovi; Otskhe; Tsakhana; Vale; Pareha; and Benara suites (Fig. 3b, column 9, Fig. 4a).

The lowermost Poskhovi suite is represented by calcareous clays and sandy clays comprising interlayers and wedges of siltstones, sandstones and marls (the latter occur mostly in its upper parts). The Lower Oligocene molluscs assemblages are met at all levels of the suite (Tatishvili 1965; Kazakhashvili 1984). The Lower Oligocene age of the Poschovi suite is also confirmed by microforaminifera (Kacharava 1977) and ichthyo-fauna findings. Thickness of the Poskhovi suite is 600–700 m. The basal part of the Poskhovi suite is represented by the Karatubani sandstones. The upper part of the Karatubani sandstones (20–25 m) hosts Oligocene fish fossils, microforaminifera and Karatubani mollusc assemblages described by Kazakhashvili (1984). The mollusc assemblage of the Poskhovi suite (including the Karatubani sandstones) is very similar to that of the Khadumian horizon of the Northern Caucasus. On the whole, the Poskhovi and Khadumi mollusc fauna resembles the fauna of Ruppelian stage of the Early Oligocene age of the Western Europe. The Otskhe beds are built up of clays, siltstones and sandstones. In some localities, the rocks are calcareous. They contain brackish molluscs (Kazakhashvili 1984). The upper limit is lithologically very distinct, its lower boundary is less clear. In unclear cases the lower boundary of the Otskhe beds must be positioned below the first beds with brackish mollusc. Thickness is up to 20 –25 m. The Tsakhana suite is thick-bedded in many localities and contains mainly medium- to coarsegrained psammitic sandstones with conglomeratic layers and wedges. Thickness is up to 65 m. This lithology is termed in different literature sources as the lower sandstones of Tskurta and Tsakhana or as Lower Corbula sandstones of Tskurta– Tsakhana. The Vale (coal bearing) suite includes clays with interlayers and successions of sandstones, siltstones and lignite. It bears molluscs fossils (Gamkrelidze 1964; Kazakhashvili 1984). Thickness is up to 150 m (this suite contains industrial deposits of brown coal). These deposits have been formed under conditions of tropical to subtropical climate

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in the lagoon surrounded by swampy forests that sourced material for coal deposits. The Pareha suite is predominantly represented by thick-bedded sandstones. North of the village of Pareha they contain mollusc fossils (Kazakhashvili 1984). The thickness is 25–50 m. The Benara suite of 400–450 m thick is built up of reddish-grey, grey and greenish-grey clays (sometimes slightly calcareous) with beds and up to 4.5 m thick successions of sandstones. The sandstone succession near the village of Benara hosts fauna of terrestrial vertebrates (Gabunia 1964) who inhabited Late Oligocene swamps, humid forests and also dry savannah. Remains of this fauna have been accumulated in the lagoon that later on transformed into the brackish corbulabearing marine basin.

Structure of the basins The structure of the Oligocene –Quaternary deposits of the Transcaucasian intermontane troughs has been studied for many years as they host the most part of the oil-gas deposits of Georgia. It is accepted, that these rocks are deformed (linear folds, reverse faults, strike–slip faults). Rock deformation is especially strong within the boundary zones between the Kura and Rioni forelands and foldthrust belts of the Greater and Lesser Caucasus. Analysis of the data derived from large scale geological mapping, drilling and seismic profiling let geologists surmise presence of nappe structures in Transcaucasian basins (e.g. Chichua in Adamia et al. 1987; Adamia 1990; Banks et al. 1997). Seismic reflection profiles show that the Kura foreland fold-and-thrust belt of the eastern Caucasus is an active thin-skinned fold-and-thrust belt and it preserves the growth strata that record the tectonic and stratigraphic evolution. The structure of the Kura and Rioni foreland fold-and-thrust belts is interpreted as due to the southward thrusting of Greater Caucasus wedge and includes fault-bend folds, fault-propagation folds and duplexes (Fig. 5). Analysis of synorogenic (or growth strata) deposits in seismic section documents that evolution of deformation has been continuing during the last c. 5– 4.5 Ma together with the thrust system kinematics (Alania et al. 2008, 2009). In recent years, with our colleagues E. Szymanski and T. O’Connor (Boston College, USA), we have carried out special studies with the aim of improving knowledge of the tectonics of the Kartli basin. The Kartli basin, the eastern part of the Transcaucasian intermontane depression, occupies the middle stream of the river Mtkvari and constitutes the eastern margin of the Caspian oil-bearing

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Fig. 5. Geological profiles across the River Rioni (R), Kura (K) forelands and Alazani (A) piggyback basin. Legends: Q, Quaternary deposits; Q-P, Quarternary–Pliocene; N2, Pliocene molasse of the Rioni basin; N2ak þ ap, Akchagylian-Apsheronian molasse of the Kura basin; N1mp, Meotian-Pontian molasse; N1S, Sarmatian marine and continental molasse; N1, Middle Miocene marine molasse; N1-2, Miocene marine molasse; N, Neogene molasses; P3-N1, Oligocene–Lower Miocene Maykopian series; P, Paleocene–Eocene terrigene turbidites and volcanics (Achara– Trialeti basin), platformal limestones and marls (Rioni basin); K, Cretaceous platformal limestones; J3,Upper Jurassic reef limestones and evaporites; J, Jurassic terrigene clastics, volcanic complexes and evaporites (for location of the figures see Fig. 1).

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domain. Within the Kartli basin, the thickness of the sedimentary layer according to geophysical data achieves 10 km. This layer is composed mainly of the Jurassic, Cretaceous and Early Cenozoic predominantly shallow water formations of island-arc type and also of the Upper Cenozoic marine and continental molasse. The molasse forming upper part of the stratigraphic column of the Kartli foreland basin is overthrusted from the north by tectonic slices of the Southern Slope zone of the Greater Caucasus, while from the south by the Achara – Trialeti fold and thrust belt (Gamkrelidze 1964). Detailed structural data show that the Oligocene –Neogene deposits of the studied area are strongly deformed. Folds and faults are widespread as well as nappes and slices. Two groups of nappes are distinguished. To the south of the Kartli basin nappes are displaced to the north. In the northern part of the Kartli basin nappes are displaced towards its southern margin (Fig. 6). The south-vergent nappes are represented by sheets of the Natskhori (Late Sarmatian age) and Dusheti (Meotian–Pontian regiostages) series. They are characterized by redeposited material made of wash-out of Cretaceous, Paleocene and Eocene flysch of the Southern Slope zone of the Greater Caucasus. The north-vergent nappes are represented by the Early and Middle Sarmatian (Kaspi), Late Sarmatian (Eldari) and Meotian–Pontian (Kavtiskhevi) series, sediments of southern margin of the Kartli basin. Cretaceous and Eocene igneous rocks of Achara –Trialeti, as well as Paleocene –Eocene sandstones and argillites of the Borjomi flysch of the same zone were source zones of redeposited material. The de´collement level of the sheets is made up mainly of ductile clays of the Maykop (Oligocene –Early Miocene age) and Eldary series (Later Sarmatian age). According to the geological maps of 1:50 000 scale, in the central part of the Kartli basin there are multiple imbricated structures of the Natskhori and Dusheti series. The authors of the maps interpret this phenomenon by the south-vergent reverse faults formed at the base of the Natskhori series. The geometry of these structures was not sufficiently studied. It is not unlikely that their angle is decreasing along with deepening, faults are flattening out and transforming into overthrust structures (see profiles). Due to lack of available data, such as well data and three-dimensional (3D) seismic profiles of a good quality, this problem could not be solved at the moment. The north-vergent tectonic sheet can be observed at the right bank of the Kura river, southward of the Kaspi city, where the sheet (the Niabi nappe) made of Lower and Middle Sarmatian limestones thrusts over the Maykopian clays (Fig. 6). The bore holes drilled in the same area along the southern margin

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of the Kartli basin show that the Lower to Middle Sarmatian rocks directly overthrust Palaeogene rocks. The southernmost (lower) south-vergent nappe (so-called here the Shiomgvime nappe), includes the rocks of the Natskhori and Dusheti series. The overthrust plane of the Shiomgvime nappe is passing between the variegated Eldari and Natskhori series. Horizontal displacement of the Shiomgvime nappe is increasing eastward and near the villages Mtskheta and Dzegvi (Fig. 1, profile 3Ka) it overlies all the folds of the southern stripe of the Kartli basin, thus, the overthrust is, at least, as wide as 10 km. On the left bank of the river Mtkvari, opposite to the Dzegvi valley, the Natskhori series along the thrust form the immediate contact with the Middle Eocene igneous rocks of the Achara–Trialeti belt (Fig. 1, profile 3Ka). Nappe structure, apparently, is also characteristic of the central and northern parts of the Kartli basin; this is evidenced by the section which shows multiple repetitions of the Upper Sarmatian Natskhori series above the Meotian–Pontian Dusheti series. Tectonic sheets following the Shiomgvime nappe northward are termed respectively: the Mukhrani, Misaktsieli and Odzisi nappes (Fig. 6). Such interpretation of the structure of the Kartli basin due to the lack of available data (wells, deep seismic profiles) is to a certain extent speculative. However, there are numerous data, including large-scale geological mapping that indicates that the molasse of the Kartli basin are characterized by nappe structure. It is to be noted that some other authors, for example, Adamia et al. (1987, 1990) and Banks et al. (1997) also have pointed to the nappe structure of the region. A lot of folds were revealed in the Oligocene– Neogene molasses of the Kartli basin. In the southern contact with Achara–Trialeti belt of the Kartli basin, in between Gori and Dzegvi, some folds were mapped (Fig. 6). Several folds of the southern stripe of the Kartli basin are buried under Quaternary alluvial deposits of the Kura river. Their maximal width reaches 8–10 km. They are constructed by Oligocene–Neogene molasses of the south margin of the Kartli basin and, at present, are covered by Shiomgvime nappe consisting of molasses of the north margin of the same basin.

Neogene – Quaternary volcanic formations The Cenozoic collision between the Eurasian and Arabian continental masses is coeval with an Upper Cenozoic magmatic activity that extends from Eastern Turkey via the South Caucasus to South Russia. It includes a number of volcanoes which are active at present or were active within

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Fig. 6. Geological profiles (1–3) across the Kartli basin (for location of the profiles see Fig. 1). K, Cretaceous– Eocene rocks of the Achara– Trialeti fold–thrust mountain belt; P3, Upper Oligocene, Maycopian series; N11, Lower Miocene: Uplistsikhean, Sakaraulian and Kotsakhurian deposits; N21, Middle Miocene: Tarkhanian, Chokrakian, Karaganian and Konkian deposits; N31s1-2, Upper Miocene: Lower and Middle Sarmatian deposits; N31s3e, Upper Miocene: Upper Sarmatian Eldari formation; N31mpk, Upper Miocene: Meotian and Pontian Kavtiskhevi formation; N31s3n, Upper Miocene: Upper Sarmatian Natskhori formation; N31mpd, Upper Miocene: Meotian and Pontian Dusheti formation; Q, Quaternary rocks. Dotted lines – overthrusts.

the past million years: volcanic areas in Turkey, Armenia, Georgia and Russia. Volcanic activity in South Georgia has taken place within a broad north– south-trending belt extending from the Arabian foreland basin in the SW to the Caucasus in the NE (Fig. 1). In places, volcanic products exceed 1 km in thickness and cover a wide compositional spectrum from basalt

to high-silica rhyolite. Intense volcanic activity started in the Late Miocene time (the Sarmatian regiostage) covering the entire region (Fig. 7) and continued into the Holocene epoch. Sarmatian – Holocene calc-alkali to shoshonitic volcanic rocks spread east- and westward from Iran–Azerbaijan – Lesser Caucasus into Eastern Iran, Pontides and Anatolia. This belt is cut by a north–south-trending

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Fig. 7. (a) Distribution of the Neogene–Quaternary volcanic formations of Georgia. 1. Pleistocene –Holocene basaltic andesite, dacite and rhyolite: Khumurdo, Borjomi-Bakuriani, Keli, Kabarjina and Kazbegi formations; 2. Upper Pliocene-Pleistocene basalt: Akhalkalaki and Tsalka formations; 3. Upper Miocene–Lower Pliocene andesite, dacite and basalt: Goderdzi formation (Adamia et al. 2008). (b) Geological cross-section of the Javakheti near the Georgia-Turkey border. 1. Quaternary basalts; 2. Upper Pliocene– Pleistocene volcanic complexes; 3. Upper Miocene– Lower Pliocene – Goderdzi formation; 4. Oligocene– Lower Miocene coalbearing terrigene clastics; 5. Upper Eocene terrigene formation; 6. Middle Eocene volcanogene formation; 7. Eocene terrigene and volcanic formations; 8. Faults.

belt (chain, volcanic highlands and large extinct volcanoes). The latter belt crosses the northern part of the Arabian Shield, the Van accretionary prism, the Pontides and the Central Caucasus (Lordkipanidze et al. 1989). According to Lordkipanidze et al. (1989), the Upper Miocene to Holocene volcanic rocks are related to the transverse Van – Transcaucasian uplift belong to the calc-alkaline basalt –andesite –dacite– rhyolite series with andesites and dacites as the dominant types. The volcanic rocks of the Central Caucasus (Keli, Kazbegi) are regarded as hybrid rocks that originated from the mixing of mantle-derived tholeiitic melt and upper crustal magma. Two main stages of volcanic activity are distinct: the Late Miocene –Early Pliocene and the Pliocene–Quaternary. Lavas predominate over volcanoclastics, especially during the second phase of eruptions. According to their mineral-chemical composition, the rocks of the both stages are attributed to calk-alkaline and sub-alkaline series (Tutberidze 2004). Data on absolute age demonstrate that the first stage of eruption in Georgia

happened c. 9–5 Ma ago, while the second occurred c. 3.5–0.01 Ma ago (Tutberidze 2004).

Volcanic formations of South Georgia Post-palaeogene volcanic formations in the central part of the Artvin– Bolnisi block and Achara– Trialeti belt are represented mainly by subaerial lava flows, volcanoclastites and also by alluvial, lacustrine and talus deposits. Highlands and plateaus of Javakheti, Akhalkalaki (Fig. 1), Tsalka and Gomareti are made up of volcanic rocks which also filled ancient river-beds (Kura, Khrami, Mashavera, Borjomi–Bakuriani lava flows). There are two series reported here: the Upper Miocene – Lower Pliocene Goderdzi formation and the Upper Pliocene –Pleistocene Akhalkalaki, Tsalka, Kumurdo and Samsari formations (Figs 1 & 3b, column 9, Fig. 4c, d). The Quaternary Borjomi– Bakuriani lava flows are also attributed to the latter. Upper Miocene– Lower Pliocene, Goderdzi formation. Basaltic lavas and pyroclastic rocks

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represent lower, basal level of the Goderdzi formation. In some places this level contains economical diatomite deposits (Kisatibi). Middle part of the section is represented mainly by volcanoclastic rocks. Pyroclastic rocks in the vicinity of the Goderdzi pass contain remains of petrified subtropical wood which date the rocks as the Late Miocene– Pliocene (Skhirtladze 1958; Gamkrelidze 1964). K– Ar dating of tuffs point to their Late Miocene age (9.8 Ma) (Aslanian et al. 1984). Maximal thickness of this part of the Goderdzi formation is about 500 m. For the upper part of the Goderdzi formation laminated and/or banded andesite and dacite lavas with volcanoclastic interlayers are common. The andesite is a dominant rock unit. The total thickness of this part of the formation is about 250 –300 m. According to Aslanian et al. (1984), Lebedev et al. (2004), K –Ar age of the andesites and dacites varies from 9.4 Ma to 7.0 Ma. Upper Pliocene– Holocene, Akhalkalaki, Tsalka, Kumurdo and Samsari formations. The Akhalkalaki formation is widespread within the South Georgian highland (Javakheti), especially within the Tsalka and Akhalkalaki plateaus. Typical sections are located along the deep canyons of the Kura, Khrami and Mashavera rivers. Basaltic (doleritic) lavas are dominant rock unites. In some places, they contain lenses of fluviatile to lacustrine and alluvial deposits, also pyroclastic rocks. Andesitic basalts are subordinate, more acidic rocks are rare. Due to their low viscosity lavas could spread over large territories. They covered ancient relief and formed extensive flat plateau. The total thickness of the formation is approximately 100 –300 m. The age of the lower part of the Tsalka and Akhalkalaki formations are identified through mammalia fauna as Late Pliocene –Pleistocene. Radiometric age of the basalts supports the faunistic data (2.0 + 0.5 Ma) (Tutberidze 2004). Mitchel & Westaway (1999) claim that the start of volcanism in the Lesser Caucasus occurred around 1.5 Ma ago. This cardinal statement for the recent evolution of the Lesser Caucasus has been made on the basis of five absolute age measurements taken at eight prospects in Armenia. It is advisable that this question be more thoroughly studied, taking into consideration both the older well-known data and most recent data obtained from the Georgian part of the Lesser Caucasus. A well-studied section of the formation is located near Dmanisi (the Mashavera river gorge), where the ruins of ancient town of Dmanisi are perched over the small plateau, which occupies the watershed of the Mashavera and Pinazauri rivers. This territory is considered to be inhabited since the Early Bronze Age (La Ge´orgie 2000), however,

Palaeolithic monuments were also found during the most recent archaeological investigations. Here, in the lacustrine sediments resting on the basalt of the Mashavera river doleritic flow have been collected and identified remains of mammals. The complex of fossils identifies the age of enclosing rocks as the Eo-Pleistocene corresponding to the Apsheron regiostage of the Caspian basin regional scale, which is correlated to the upper part of the Villafranchian regiostage of the Mediterranean area. In the same sediments the remains of hominides have been found, identified by Gabunia et al. (1999) and Gabunia & Vekua (1995) as the representative of Homo ErectusErgaster. As is known, the representative of Homo Ergaster inhabited the Earth some 1.5–1.8 Ma ago. The absolute age of the basaltic lavas directly underlying the bone bearing horizon was determined by 40Ar/39Ar laser probe dating of its plagioclase phenocrysts. In total, 13 analyses of multiple plagioclase crystal samples, and five 40 Ar/39Ar laser probe analyses of single plagioclase crystals were carried out at the Geomar Tephrochronology Laboratory (USA). Both analytical series yield isochron ages ranging from 1.95 + 0.22 Ma to 2.04 + 1.37 Ma. Additional plagioclase samples from three contemporaneous basalt lava flows of the Mashavera lava flow give 40 Ar/39Ar laser probe ages ranging from 1.6 + 0.2 to 3.2 + 0.3 Ma (Djaparidze et al. 1989; Gabunia et al. 1999; La Ge´orgie 2000). The magnetostratigraphic investigations carried out during the last two decades in the Georgian part of the Caucasus (Djaparidze et al. 1989) have recognized within the Late Miocene– Quaternary sequence of volcanic rocks all the known standard palaeomagnetic chrons and subchrons: Brunhes, Matuyama (with Jaramillo, Cobb Mountain, Olduvai, Reunion 2 and Reunion 1 subchrons), and Gauss (with Kaena and Mammoth subchrons). These data indicate the place of basalts of Dmanisi as the top of the Akchagylian regiostage (c. 1.8 Ma.). The older basalts of this part of the Lesser Caucasus developed within the Akhalkalaki plateau yielded an Ar/Ar age c. 2.9 Ma (Ferring et al. 1996), 3.69 + 0.008 and 3.83 + 0.09 Ma (Camps et al. 1996). These measurements corroborate dating obtained in the laboratories of the former Soviet Union, namely 2.36 + 0.60 Ma for andesite– basalts of Trialeti (Rubinstein et al. 1972), 2.25 + 0.2 Ma for andesite of Satkhe (Maisuradze et al. 1980), 2.4 + 0.35 Ma for andesite-basalts of the volcano Emlikli (Djigauri 1991). Khumurdo formation. Quaternary andesitic lava sheets derived from numerous parasitic centres of Javakheti highland. They unconformably overlie

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the Akhalkalaki basalts and are not folded. Total thickness of the suite is about 30– 40 m (Yilmaz et al. 2001). This unit did not yield any faunistic data; however, K –Ar age of the Khumurdo formation is 1.1 + 0.1 and 1.2 + 0.1 Ma (Maisuradze et al. 1980, 1999). Samsari formation. Andesite –dacites and dacites of the Samsari and Kechuti ranges crown the outcrop of the Lower –Pliocene –Quaternary volcanic formation of the Javakheti highland (extinct volcanoes Tavkvetili, Shavnabada and some others) (Adamia et al. 1965). According to K –Ar dating, the older dacites of the Samsari range are represented by lavas of the volcano Babakhno (c. 760 Ka.). Later on (400 –170 Ka.) there have generated the platy dacitic lavas in the central part of the Javakheti highland (the Abul mountain, lake Paravani, caldera Samsari). Volcanic activity in the region came to halt, probably, in the end of the Pleistocene, about 30 Ka ago (Lebedev et al. 2004). Borjomi–Bakuriani lava flows. Quaternary volcanic rocks directly occur in the central part of the Achara-Trialeti belt (Fig. 1) in the Borjomi district (Kura gorge). Here, between the gorges of the Gujaretistskali and Borjomula rivers (right tributaries of the Kura river) there rises the plateau built up of young andesite lava flows. The Borjomi – Bakuriani lava includes particularly vitreous andesites. Maximal thickness of the flow reaches approximately 100 m. The volcanic centre of this lava flow is situated in Bakuriani (the mount Mukhera). This cone mountain has a central craterlike flat-bottomed depression. From this centre lava flowed toward the ancient valley of the Bakurianistskali and Gujaretistskali rivers and descended into the Kura valley (Gamkrelidze 1964).

Quaternary volcanoes of the Greater Caucasus The central, Georgian, part of the Greater Caucasus in the upper reaches of the Aragvi and Terek rivers (Tergy), hosts three more or less distinct groups of extinct volcanoes (in south–north direction): 1. Keli, the Upper Tetri Aragvi; 2. Kabarjina, the right bank of the Upper Terek and 3. Kazbegi, the left bank of the Upper Terek (see Figs 1 & 7). Each group includes some lava flows, whose brief characteristic is set below according to Dzotsenidze (1970), Tsagarelli (1982), Koronovsky & Demina (1999) and Tutberidze (2004). Keli group. Gudauri flow. The Gudauri –Mleta andesite-basaltic lava flow is over 12 km long. The lava has spread down the ancient valley of the Aragvi river covering thick alluvial and prolluvial

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deposits. The volcanoes Sakokhe and Sadzele of the Keli Group raise in the NE while the Patara and Didi Nepiskalo volcanic massifs in the west. The Sakokhe volcano is a polygenetic one, formed primarily in the Pleistocene time. At first, it produced andesites, later andesite-basalts, which spread over a great distance. The volcano cone has a poorly recognizable crater-like hollow. The Patara Nepiskalo volcanic massif rises in the eastern part of the Keli highland. This is a highly dissected range with a teeth-shaped crest of seven more or less isolated peaks (the ‘Mountain of Seven Brothers’) arranged in a semicircle. The volcano Nepiskalo was formed in the middle of the Pleistocene. It is a complex polygenetic volcano, which has erupted, at least, three times not counting a later (Holocene?) eruption from the place known as the ‘Flat Top’. The volcano Didi Nepiskalo has a strict conical form and is made up of rhyolite-dacite. Kabarjina group. The Kabarjina volcano’s main structure was formed in the Middle Pleistocene time. The volcano was formed by extrusions, lava eruptions and explosions. It is built up of tuffbreccia, tuffs and lavas of dacitic and andesite – dacitic composition. When flowing down, the Kabarjina lavas formed isolated masses or short tongues. Kazbegi group. The Kazbegi summit (the Mkinvartsveri) rises 5033 m above sea level. The mountain Kazbegi has two-headed cone and it represents a polygenetic volcano active throughout the Quaternary period. The youngest lavas occur at the top, and the products of the older eruptions are spread over the sides of the massif covering a distance of several km. Several flows and contact zones with slag-like brecciated lavas are seen in the outcrops. The lava has spread predominantly over the left bank of the Terek River. Farther below, it dammed the river forming a temporary lake, which inundated the forested flood plain. The forests were buried by lacustrine-alluvial deposits. Later the ‘dam’ was sawed by the Terek as deep as 10 –12 m and the lake was drained. A succession of the lake beds can be seen in the outcrop at the right bank of the river. Two radiocarbon dates (5950 + 90 years and 6290 + 90 years) are available for the wood fragments collected from the lake beds. These dates indicate that the lake beds and the Tkarsheti lava flow, apparently, could be attributed to the Middle Holocene (Tsagarelli 1982). The lava flow descended into the Terek valley along the ancient gorge of the river and over the eastern side of the Kazbegi Mountain. The lava layers alternate with coarse clastic formations, alluvium and tuff. Predominant rock type is andesite. The total thickness of the flows is a few dozen metres.

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The Mnadon lava flow also descended from the Kazbegi volcano. It moved by a distance of 15 to 18 km. The floor of the ancient Mnadon valley is armoured by this lava at an elevation of 200 – 250 m above the Terek riverbed. Some authors believe that the Mnadon lava flow outpoured in the Early Pleistocene. Further upstream, the Terek valley rapidly narrows turning into a gorge. The gorge is made up of the andesite-dacite lava flow which descended from the volcano Khorisar down the northern side of the Greater Caucasus Main Range. The Khorisar volcano is two-cone shaped. The surface of the lava flow is unmarked by glacial erosion. The age of the Khorisar lavas is post-glacial, the Holocene. The described volcanic rocks of the Central Greater Caucasus show two well distinct petrochemical series: calk-alkali and sub-alkali. The calk-alkali series is subdivided into two sub-series: low-potassium and high-potassium. The volcanic rocks are harshly predominated by the rocks of calk-alkali series (Tutberidze 2004). Fifteen bulk samples of the volcanic rocks representing the most significant centres of Quaternary volcanism of the Kazbegi volcanic area were studied in the Laboratory of Isotope Geochemistry and Geochronology of Academy of Sciences of Russia (Chernishev et al. 2000). K –Ar determination yield ages ranging from 15 to 455 Ka with their prevailing part restricted to 150 –295 Ka which corresponds to the Middle Pleistocene time. At this time, the body part of the Kabarjin and Didi Nepistskalo magmatic formations of the Keli plateau was formed, as well as some formation of the Kazbegi volcanic centre. The youngest formations of the Kazbegi volcanic centre are represented by lavas of the Arsha flow (50 + 20 Ka), which correspond to the upper part of the Late Pleistocene time. The youngest formations of the Keli complex represent products of eruption of the Sharkhokh volcano (20 + 15 Ka) and volcanic rocks of the lava flow of the western margin of the Keli plateau (15 + 15 Ka) which well correlate with the Pleistocene –Holocene boundary (Chernishev et al. 2000).

Conclusion The Oligocene – Neogene– Quaternary basins of Georgia were formed after the closure of the Tethys ocean within the Caucasian transect. As a result of topography inversion at the place of preOligocene back-arc basins of the Greater Caucasus and Achara –Trialeti (the Lesser Caucasus) there were formed fold-thrust mountain belts, while at the place of separating rigid platformal blocks (the Georgian and Artvin–Bolnisi blocks) were formed intermountain molasse troughs of the

Transcaucasus. Oligocene– Quaternary molasse usually are divided into lower, fine-grain ones (the Oligocene–Lower Miocene, Middle Miocene, Lower and Middle Sarmatian deposits) and upper, coarse-grain molasse (the Upper Sarmtian –Pleistocene deposits). The lower molasse are represented mainly by marine facies, while the upper molasse – mainly by subaerial deposits. Thickness of the lower molasse (like the upper molasse) varies in a wide range achieving within the mostly downwarped depressions of the Transcaucasus about 5000 m. Maximal thickness of the upper molasse exceeds 5000 m. The Oligocene– Pleistocene deposits are deformed as a result of tectonic compression followed by formation of linear folds and reverse faults, especially intensive deformations are observed within the boundary zones of the Kura and Rioni basins bordering fold-thrust belts of the Southern Slope zone and Achara–Trialeti. Nappe structures are common features for molasse of the Kartli basin. In Georgia, Neogene –Quaternary volcanic rocks are represented only by subaerial formations. In South Georgia (Lesser Caucasus), volcanic activity started in the Late Miocene and after lull period in the middle of the Pliocene time it recurred in the Late Pliocene time and lasted throughout the Pleistocene epoch (Holocene?). The area of volcanic eruptions was gradually decreasing and in the end of the Pleistocene epoch it was reduced to some small sites in South Georgia. Within the Greater Caucasus, volcanic eruptions occurred only during the Pleistocene and beginning of Holocene times. This work was accomplished due to support of CNRS, France; New England Research, Inc; Geology and Geophysics, Boston College, USA; and technical assistance by T. Kurdagia and N. Zviadadze.

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Cenozoic –Recent tectonics and uplift in the Greater Caucasus: a perspective from Azerbaijan JON MOSAR1*, TALAT KANGARLI2, MARTIN BOCHUD1, ULRICH A. GLASMACHER3, ANNICK RAST1, MARIE-FRANCOISE BRUNET4 & MARC SOSSON5 1

De´partement de Ge´osciences – Sciences de la Terre, Universite´ de Fribourg, Chemin du Muse´e 6, CH-1700 Fribourg, Switzerland

2

Geology Institute of Azerbaijan, National Academy of Sciences, H. Javid av. 29A, Baku AZ1143, Azerbaijan

3

Institute of Earth Sciences, Im Neuenheimer Feld 234, 69120 Heidelberg, Germany

4

UPMC Univ Paris06-CNRS-INSU, UMR Institut des Sciences de la Terre Paris Case 129, 4 place Jussieu, 75005 Paris, France 5

CNRS -GeoSciences Azur, UMR 6526, 250 Rue Albert Einstein - Sophia Antipolis, 06560 Valbonne, France *Corresponding author (e-mail: [email protected]) Abstract: The Greater Caucasus is Europe’s highest mountain belt and results from the inversion of the Greater Caucasus back-arc-type basin due to the collision of Arabia and Eurasia. The orogenic processes that led to the present mountain chain started in the Early Cenozoic, accelerated during the Plio-Pleistocene, and are still active as shown from present GPS studies and earthquake distribution. The Greater Caucasus is a doubly verging fold-and-thrust belt, with a pro- and a retro wedge actively propagating into the foreland sedimentary basin of the Kura to the south and the Terek to the north, respectively. Based on tectonic geomorphology – active and abandoned thrust fronts – the mountain range can be subdivided into several zones with different uplift amounts and rates with very heterogeneous strain partitioning. The central part of the mountain range – defined by the Main Caucasus Thrust to the south and backthrusts to the north – forms a triangular-shape zone showing the highest uplift and fastest rates, and is due to thrusting over a steep tectonic ramp system at depth. The meridional orogenic in front of the Greater Caucasus in Azerbaijan lies at the foothills of the Lesser Caucasus, to the south of the Kura foreland basin.

The Caucasus orogen lies at Europe’s cross-road with Asia and Arabia, and is one of the world’s outstanding mountain ranges (Fig. 1). It is Europe’s highest mountain range with Mount Elbrus culminating at 5642 m a.s.l. in the western Greater Caucasus. It consists of the Greater Caucasus (GC), intramontane basins (Kura-Kartli-Rioni; c. 200 m elevation) and the Lesser Caucasus. North of the Greater Caucasus the deep sedimentary Terek and Kuban foreland basin (.6000 m thick; up to 1600 m elevation) form the transition to the Scythian platform. NNW of Mount Elbrus, the Stavropol ‘high’ forms a basement uplift, and in the east the northern slope is formed by the Dagestan foreland fold-and-thrust belt. The southern Greater Caucasus foreland, SW of Tbilisi is one of the world’s earliest sites of human society with 1.8 Ma old hominoid remains of Dmanisi (Georgia) (Lordkipanidze et al. 2007). The Lesser Caucasus with lower topography (c. 3000 m), is a zone of important volcanic

and seismic activity. In the east and west, the Caucasus topography is bound by two very deep sedimentary basins, the South Caspian Sea and the Black Sea, hosting some of the world’s largest oil and gas provinces. The Caucasus orogen is caused by the north directed movement of the Arabian plate squeezing a Jurassic to Early Palaeogene subduction related volcanic arc (Lesser Caucasus) as well as Jurassic to Pliocene marine sedimentary rocks and sediments (northern Lesser Caucasus, substratum of KuraKartli Basins and Greater Caucasus Basin) towards the Scythian plate (Gamkrelidze 1986; Nikishin et al. 2001; Stampfli et al. 2001; Popov et al. 2004; Hafkenscheid et al. 2006; Kazmin & Tikhonova 2006; Sosson et al. 2010). Recent plate tectonic models and GPS-based convergence rates (Gamkrelidze & Kuloshvili 1998; Vernant et al. 2004; Reilinger et al. 2006; Kadirov et al. 2008) suggest a moderate anticlockwise rotational

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 261–280. DOI: 10.1144/SP340.12 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. General map of the larger Caucasus area with topography and earthquake distribution. Not all existing thrusts are shown only a selection of the major thrusts of the Greater Caucasus relevant for the discussion. Transparent red indicates the pre-Mesozoic core. Yellow highlighted area is the Adjara-Trialet fold-and-thrust-belt. Black triangles are mountain summits of the Greater Caucasus, black stars correspond to localities cited in text.

component of convergence and a complex plate boundary with vertical and horizontal strain partitioning (Jackson 1992). Recent convergence rates of up to 14 mm a21, strong earthquakes, landslides, active volcanoes, and extreme subsidence and surface uplift rates are indicative for the dynamics of the continent– continent collision. From east to west, the morphological shape and the structural features are strongly influenced by the rotational convergence of the Arabian plate and westward escape of the Anatolian Plate causing distinct tectonic regimes in the Caucasus. The Lesser Caucasus area is dominated at present by a strike–slip regime, whereas the Greater Caucasus is dominated by thrust tectonics with a main NNE –SSW direction of movement. The dominant movement is top to the south in the main range and the southern slopes. Top-to-the-north motion is observed in the areas in the north and in Dagestan. Hereafter we will present different aspects of Cenozoic and recent tectonics, and tectonic geomorphology, especially based on detailed structural studies carried out over several years in the eastern Greater Caucasus in Azerbaijan. We shall discuss their relevance for understanding the thrust kinematics and the links between tectonics,

topography, seismicity and uplift in the Greater Caucasus.

Regional tectonics and geodynamics The geodynamics of the Greater Caucasus orogen correspond to an intercontinental collision zone inverting a deep Mesozoic –Cenozoic basin (Fig. 2) that is not located above a subduction regime, but bordered east and west by super deep sedimentary basins that have their origin in the Mesozoic and are filled with Cenozoic– Quaternary sediments. To the north and south of the Greater Caucasus are the foreland basins of the Terek-Kuban and the Kura-Kakheti-Kartli-Rioni, respectively (Ershov et al. 1999; Mikhailov et al. 1999; Ulminshek 2001; Daukeev et al. 2002; Ershov et al. 2003); to the east and west are the Caspian Sea and the Black Sea, respectively (Shikalibeily & Grigoriants 1980; Berberian 1983; Ismail-Zade et al. 1987; Narimanov 1992; Abrams & Narimanov 1997; Mangino & Priestley 1998; Nikishin et al. 1998; Allen et al. 2002; Brunet et al. 2003; Nikishin et al. 2003). The Lesser Caucasus is situated above an old, possibly detached subduction slab (Hafkenscheid et al. 2006). An incipient subduction

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Fig. 2. General crustal-scale cross-section through the eastern Greater Caucasus and the Lesser Caucasus (location see Fig. 1). The Lesser Caucasus is associated with a northward subduction and possibly a detached slab [based on tomography after Hafkenscheid et al. (2006)]. No subduction is seen under the Greater Caucasus. The Greater Caucasus is a doubly verging orogenic wedge with the dominant thrusting towards the pro-wedge to the South. A retro-foreland fold-and-thrust belt develops to the north in Dagestan (Russia) see also Figure 1. Three different types of crust have been distinguished according to their geodynamic belonging: to the South a crust intruded and associated with the Jurassic– Cretaceous suprasubduction arc volcanism in the Lesser Caucasus, in the centre the thinned and rifted and intruded southern part of the supra-subduction backarc basin, and to the north the northern part of this extended backarc rift system with the important Mesozoic sedimentary series of the Greater Caucasus Basin. Some major faults such as the Main Caucasus Thrust (MCT) are highlighted. The structure and the position of the thrusts at depth remains speculative, but indicate underthrusting of the terranes to the south of the Greater Caucasus and strong imbrication over a ramp system in the Greater Caucasus.

is believed to have occurred at the northern edges of the Black Sea, whereas in the east the subduction process was already initiated in Pliocene times, when the South Caspian Basin started subducting to the north under the eastern termination of the GC and the Apsheron Ridge (Allen et al. 2002; Knapp et al. 2004). The detailed link of the incipient subduction to the structures such as the Main Caucasus Thrust (MCT) in the Greater Caucasus remains to be investigated. The depth of the Moho changes from about 40 km beneath the Kura basin to more than 50 km beneath the eastern Greater Caucasus and rises to 40 km again under the northern foreland basin (Brunet et al. 2003; Ershov et al. 2003). The Greater Caucasus is a doubly verging mountain-belt (Fig. 2) with two external fold-andthrust belts (FTB) and a complex nascent axial zone (Sholpo 1993; Khain 1997). The main tectonic underthrusting appears to be towards the north – similarly as the subduction sense in the Lesser Caucasus – creating an overall asymmetry of the mountain range. The southward propagating meridional foreland FTB together with a large part of the axial zone form the pro-wedge (front) of the orogen (Khain 1975; Adamia et al. 1977, 1981;

Gamkrelidze 1986; Philip et al. 1989; Gamkrelidze 1997; Gamkrelidze & Shengelia 2005). The KuraKartli and also the Rioni foreland basins are dissected by, and incorporated into, the outward propagating foreland FTB to the south of the main range. Deep seated southward migration of the orogenic front led to the inversion of the Pliocene to Late Pleistocene sediments, and the transport of the Alasani basin (Figs 1 & 3a) as a piggy back basin towards the south. Unlike in the western Greater Caucasus, a broad north-directed foreland FTB develops in the NE, in Dagestan and is part of the retro-wedge of the orogen (Kopp & Shcherba 1985; Dotduyev 1986; Zonenshain et al. 1990; Sobornov 1994, 1996; Djavadova & Mamula 1999). The Terek basin since the early Pliocene has subsided more than 4000 m, and recently has exhibited pitted gravels of Early Pliocene age at 1600 m elevation. This northern FTB, similarly to the southern FTB, propagates into the Cenozoic –Quaternary series of the Terek Basin in the Dagestan area (Fig. 2). Whereas the axial zone of the Greater Caucasus comprises Jurassic sedimentary rocks (Azerbaijan), a pre-Mesozoic basement (Georgia, Russia), and Pliocene intrusions, both external fold-and-thrust

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Fig. 3.

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Fig. 3. (Continued) (a) Simplified tectonic model of the GC linking topography and major geomorphic features with the major tectonic thrusts. The map shows the active fault strings of the MCT (in terms of seismicity) and the possible propagation outward to lower structural levels in the western GC in the Racha area. (b) Topography v. tectonics: map showing the main areas of uplift and tectonic activity. Interpretative model which shows the link between thrusting and topography and seismic activity. The model proposes that due to motion over a ramp system at depth one can achieve fast uplift in narrow zone bound to forward thrust and back-thrust generated over the tectonic ramp.

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belts consist mainly of Cretaceous and Cenozoic sedimentary rocks (Khain 1997). The Greater Caucasus basin has developed in a back-arc setting to the southerly subduction-related volcanic arc of the Lesser Caucasus. Intrusive rocks are frequently found up into the Early Cenozoic, but mainly affect the southern parts of the basin (Sosson et al. 2010). Volcanoclastic series derived from the Lesser Caucasus volcanic arc are now found imbricated and folded in the southern foothills of the Greater Caucasus where they form distinct tectonosedimentary units (Kangarli 1982, 2005). In situ intrusives remain however rare and are associated with igneous activity on the margins to the south of the Greater Caucasus Basin (Mengel et al. 1987; Mustafayev 2001; Chalot-Prat et al. 2007). Pliocene to Quaternary igneous activity is observed in the central part of the mountain range, in the border areas between Georgia and Russia (Tchechenia). The most outstanding examples are Mount Elbrus with 5642 m a.s.l. and further East Mount Kazbek (5047 m a.s.l.). These intrusions are mainly late-collisional, subalkaline granitoids that roughly range between 4.5 and 1.5 Ma (Hess et al. 1993; Gazis et al. 1995; Nosova et al. 2005;

Lebedev & Bubnov 2006), and culminate with Quaternary volcanism reaching into the Holocene (Lebedev 2005; Chernyshev et al. 2006). Several successive tectonic events are documented in the Greater Caucasus sedimentary record. Precambrian and Palaeozoic (pre-Hercynian and Hercynian) tectonic phases are recorded in the preAlpine basement or Palaeozoic core (for discussion and references see Gamkrelidze & Shengelia 2005; Kazmin 2006; Saintot et al. 2006a, b; Somin et al. 2006) and are followed by palaeotectonic events related to the Tethyan oceans (Palaeo- and Neotethys) (Nikishin et al. 1997; Barrier et al. 2008). These palaeotectonic events included extensional structures recorded throughout the Mesozoic cover of the Greater Caucasus Basin (Dotduyev 1986), but also unconformities thought to result from compressive phases such as the ‘Eo-Cimmerian’ (Triassic) and the ‘Mid-Cimmerian’ (Callovian – Bajocian) which is well documented in northern Azerbaijan (Fig. 4). The link of the latter unconformity to possible orogenic events remains speculative and debated. The geometry of the Greater Caucasus sedimentary basin is of passive margin type with numerous

Fig. 4. Cross-section through the Sahdag mountain in north-eastern Azerbaijan. The section highlights the structural style at the northern edge of the mountain range. Thrusting is mainly to the north and of Plio-Pleistocene age. The Sahdag is built by a tectonic klippe that duplicates the Jurassic–Cretaceous series. The tectonic contact between the klippe and the underlying series is possibly a synsedimentary fault associated with the collapse of the carbonate platform during Cretaceous. To the north the Oxfordian overlies discordantly older folded series of the Middle to Lower Jurassic. This unconformity reflects the Mid-Cimmerian event. On the summit of the Sahdag Sarmatian rocks of marine origin document a vertical uplift of some 3700–4000 m. Plio-Pleistocene detrital series including conglomerates are uplifted and folded (seen further east along strike the same structure). Photos: (a) View towards the west on the Sahdag Klippe which is on of the zones of fastest uplift in the whole area. Rocks range from Cretaceous to Jurassic in age. (b) View to the SW from Sahgdu¨zu¨ (2500 m a.s.l.) to Bazardu¨zu¨ range which is the highest range in Azerbaijan culminating at more than 4000 m a.s.l. Rocks are mainly of Lower Jurassic (Aalenian) age.

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tilted blocks. The central part of the orogen – where the oldest series outcrop, and topography is the highest – represents a distal basin between a platform domain to the north and a distant deeper domain with a structural high (tilted block) to the south. The foreland basins associated with the orogenic evolution are filled with Cenozoic and Quaternary sediments. In the south they build on top of the former distal, stretched continental margin (Greater Caucasus Basin), in the north they build on a flexural foreland underlain by a carbonate platform (Ershov et al. 1999). During the growth of the orogen since the Early Cenozoic the thrust front is propagating out into its own foreland basins. Especially the southern basins develop into a succession of piggyback foreland basins, subsequently and progressively abandoned (relic thrust fronts) as the orogenic front migrates southward. A typical example of an abandoned basin is the Cenozoic–Quaternary Alasani Basin (Philip et al. 1989) (Figs 1 & 3a). Distinct tectonic zones, from north to south, are separated by major thrusts (Dotduyev 1986). They correspond to the original palaeogeographic setup and build upon inherited, pre-existing structures (Egan et al. 2009). Lateral correlations and differences can be made between the western region in Crimea (Saintot et al. 1998; Saintot & Angelier 2000; Saintot et al. 2006a), through Georgia (Gamkrelidze & Rubinstein 1974; Gamkrelidze & Gamkrelidze 1977; Banks et al. 1997) to the Caspian Sea (Kangarli 1982, 2005; Allen et al. 2003; Egan et al. 2009). The Adjara-Trialet FTB in Georgia located to the south of the southern limit of the Greater Caucasus in Georgia (Banks et al. 1997; Gudjabidze 2003) is of particular interest since thrusting is top to the North, opposite the direction in the GC (Gamkrelidze & Kuloshvili 1998). One of the major structural features in the GC is the Main Caucasus Thrust (MCT) (Dotduyev 1986). This large thrust can be observed along strike of the mountain belt over a distance of more than 1000 km (Figs 5 & 3). Displacement on this major thrust fault is top to the South, possibly in excess of 30 km in some places. In the west in Russia and Georgia, the MCT separates the Palaeozoic metamorphic core of the mountain range from the Jurassic cover series to the South. Further east in Georgia, Dagestan (Russia) and Azerbaijan it is found in the core of the orogen, separating rocks of different Jurassic ages. The definition of the MCT used here is according to Dotduyev (1986). Some recent papers (Kadirov et al. 2008) mistakenly label the thrust separating the Alasani Basin from the terrains in higher topographic elevations to the north as MCT. This latter thrust is believed to be a relic thrust front of early Cenozoic age. In eastern Azerbaijan, east of Mount Bazardu¨zu¨ (the highest summit in Azerbaijan, Fig. 1), we lose the trace of

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the MCT and fieldwork has shown that is relayed by a string of fault-related folds.

Active tectonics, convergence and uplift Earthquakes and active faults The Greater and Lesser Caucasus are seismically active zones linked to the rapid and non-uniform plate convergence between Arabia and Eurasia (Philip et al. 1989; Jackson 1992; Priestley et al. 1994; Triep et al. 1995; Jackson et al. 2002; Allen et al. 2004, 2006) (Fig. 1). The Lesser Caucasus and the adjoining Anatolian Plateau show a predominance of strike–slip focal mechanisms associated with a system of vertical faults. In the Greater Caucasus, on the contrary, convergence is accommodated predominantly by reverse focal mechanisms associated to thrusting with a general north–south to NE–SW compression (Koc¸yigit et al. 2001; Barazangi et al. 2006; Copley & Jackson 2006; Tan & Taymaz 2006), see also discussion in (Jackson 1992; Allen et al. 2004). Slip vectors based on earthquake focal mechanisms show a general top-to-the-south thrusting. Strike – slip mechanisms exist but are rare. Present seismicity is unevenly distributed across the GC (Figs 1 & 3). A zone with a higher seismic activity is observed on the south slope of the Greater Caucasus west of Tbilisi (Georgia) in the Racha area (Triep et al. 1995). Studies of focal mechanisms and focal depths show that this seismicity is linked to several active fault strings in the subsurface of the Gagra-Dzhava zone (Triep et al. 1995; Gamkrelidze & Kuloshvili 1998). They show south directed slip vectors. These faults are located to the south and are structurally in the footwall of the MCT. To the west this fault system links to the MCT precisely where this latter shows an important bend, and is stepping back (to the north) into the mountain range (Figs 1 & 3a, b). We suggest that the MCT is developing a new splay, and that the higher seismicity in this region is due this propagation of the MCT to the SW and to a lower structural level (Fig. 3a). To the east this fault system may be correlated with the thrust fault at the front of the Alasani Basin. It is relevant to notice that elsewhere in the Greater Caucasus the largest earthquakes known (earthquakes . magnitude 6, both historical and measured) are all located in the vicinity of the MCT. We interpret this to show the importance of the MCT to the present day in the deformation processes, since large earthquakes occur along large faults accommodating important displacement. The MCT appears to be a major thrust in the development of the Greater Caucasus. Seismicity is extending into the Middle and South Caspian Sea (Kovachev et al. 2006). In

268 J. MOSAR ET AL. Fig. 5. Faults v. topography: map showing majors faults (see also Fig. 1) limiting zones of changing topography; MCT is highlighted. The zone of highest topography is mostly in the green colour. Notice position of relic mountain/thrust front to the North of Alasani Basin. Karamarian anticline with water and wind gaps is at the northern edge of the Kura plain in Azerbaijan (see location in Fig. 1).

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the Apsheron zone focal mechanisms show NNE– SSW oriented thrusting (Jackson et al. 2002) and seismic activity may be linked with an extension/termination of the Greater Caucasus towards the East and/or with young north-directed subduction of the South Caspian Basin to the North under the Apsheron (Allen et al. 2002; Knapp et al. 2004). The seismicity further south as well as in the Gobustan desert area shows a westward component of motion relative to Eurasia, suggesting underthrusting towards the west (Jackson et al. 2002). Some seismic activity is also seen in the central part of the eastern Greater Caucasus, as well as in the Kura basin. On the northern slopes, the Dagestan FTB and the recent faults in the Terek Basin show a higher concentration of earthquakes pointing to active thrust tectonics in this area. Studies on palaeoseismology remain rare but confirm the existence of inherited faults and the possible 2000 year recurrence of high magnitude events (Rogozin et al. 2002; Rogozin & Ovsyuchenko 2005).

Convergence and uplift Studies based on GPS technologies in the larger Caucasus area, including Turkey, Arabia and Iran have confirmed the global picture of convergence across the Caucasus (McClusky et al. 2000; Nilforoushan et al. 2003; Reilinger et al. 2006). The average convergence of Arabia and Eurasia of 18 –23 mm a21 is transformed into a deformation of 14 mm a21 with a north–south direction across the Greater Caucasus, mainly the southern part (Vernant et al. 2004). Detailed studies in Azerbaijan (Kadirov et al. 2008) confirm a rotational convergence between Arabian and European plates. Shortening is distributed between the Northern Kura Basin and the outermost thrusts of the Dagestan FTB. Present-day slip rates decrease from 10 + 1 mm a21 in eastern Azerbaijan to 4 + 1 mm a21 in western Azerbaijan (Kadirov et al. 2008). A similar study in Georgia shows the opposite thrust directions between south-dipping thrusts in the Adjara-Trialet FTB and the Greater Caucasus front, which shows relative motion of 6.9 + 1.1 mm a21 to the SW on north-dipping thrusts (Gamkrelidze & Kuloshvili 1998). These studies also show a marked change (decrease) in velocities across the MCT. Indeed north of the MCT velocities are almost 0, indicating no longitudinal displacement. All deformation appears to be taken up in uplift. Across the more or less north –south oriented West Caspian Fault – at the transition Kura basin to Gobustan area near the eastern shores of the South Caspian Sea – a recent study indicates a

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dextral strike –slip motion and calculates a differential movement of 11 + 1 mm a21 (Kadirov et al. 2008).

Uplift/subsidence and tectonics v. topography Subsidence studies The Greater Caucasus basin was initiated by Mesozoic back-arc extension related to the subduction of the Tethys Ocean to the south (Brunet et al. 2003; Barrier et al. 2008). A large part of the subsidence occurred during this rift-related crustal extension, prior to the subsequent Cenozoic mountain building and the subsidence observed both north and south of the Greater Caucasus. Rapid subsidence occurred there in a foreland basin setting, in different phases during the Cenozoic. The North Caucasus foredeep (Kuban and Terek basins mainly) can be described in three major periods: Early Jurassic to Late Cretaceous related to initial rifting then cooling followed by a Late Cretaceous to Middle Eocene phase of alternating subsidence and uplift as a far field effect of the suturing of the Tethys Ocean to the south (Mikhailov et al. 1999). The period of Late Eocene to present relates to the development of a foreland basin coeval with shortening and uplift in the adjacent Greater Caucasus range. In the Black Sea and South Caspian basins (Brunet et al. 2003), the much more rapid Pliocene–Quaternary phase of sedimentary infill as well as subduction related subsidence in the north of South Caspian (Egan et al. 2009) occurred simultaneously with the asymmetrical subsidence of Caucasus-related molasse basins to the north and south (Ershov et al. 1999, 2003). A crustal/lithospheric – scale model suggests that crustal thickening and removal of lithospheric roots are responsible for supporting the Caucasus Mountains. Subsidence is explained by loading of the lithospheric roots (Ershov et al. 1999). The Eocene– Early Oligocene phase of subsidence is associated with cessation of subduction in the southern areas of the Lesser Caucasus, while the rapid Middle Miocene to Present subsidence is linked to the final closure and inversion of the Greater Caucasus Basin.

Uplift-exhumation The Caucasus offers numerous geomorphic field lab conditions to apply classic methods such as terrace deposits studies or ‘young’ marine sediment age/ altitude studies, palaeogeography and palaeoclimate/environment, and more modern ones such as fission track studies on apatite or zircon.

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Geomorphologic research (Mitchell & Westaway 1999) and minor thermochronological investigations show a very young heterogeneous rock and surface uplift in the Greater Caucasus. A study on pre-alpine basement rocks shows ages based on apatite fission track studies younger than 68 Ma in agreement with the general idea of an uplift that initiated in the Cenozoic (Kral & Gurbanov 1996). Patterns of age distributions also indicate uplift of 7–4 Ma in some areas, suggesting a strong influence of recent tectonics in the process of exhumation/ uplift. Studies on cooling history of recent granites in the western central Greater Caucasus show ages between 2.5 and 1.2 Ma and suggest uplift rates of 4 mm a21 (Hess et al. 1993), but thermal modelling difficulties linked to the close vicinity of hot intrusives may significantly modify modelling results leading to overestimates of the uplift rates. Compilation of uplift amount and uplift rates (Philip et al. 1989; Ershov et al. 2005), confirm the very young uplift (Fig. 6a, b) but show uplift rates in excess of 12 mm a21, which in light of field evidence (see Mitchell & Westaway 1999 for discussion) seem very fast. However, they consistently show that the fastest and highest uplift is in the centre of the range. A more recent study on apatite fission tracks on the Early Miocene Maikop series of the western Greater Caucasus shows the detrital provenance of the clasts and suggests a minimum Early Oligocene age for subaerial uplift of the mountain range (Vincent et al. 2007). Work in progress by our group on apatite fission track data in Azerbaijan, showed similar inherited ages of 12 –56 Ma for samples taken in conglomerates of Pliocene age north of Mount Sahdag (Fig. 4). Ages of 21.8 Ma from samples taken in Aalenian sandstone in the central part of the range south of Mount Sahdag (Fig. 4) confirm the Cenozoic uplift. Detailed studies on geomorphology, including the young terraces along the Black Sea and Caspian Sea (Brod 1962; Krasnov et al. 1974), or the Kura Basin (Shirinov 1973; 1975) exist but are all in Russian and difficult to access (Budagov 1969, 1973; Shcherbakova 1973; Shirinov 1975). One of the most prominent geomorphologic archives directly relating to uplift are the marine and river terraces. Up to 14 levels of terraces are recognized and reach into the valleys of the Greater Caucasus. A terrace at 475 m a.s.l. is given as latest Pliocene in age (Fig. 6c), and is

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found along the western shore of the Caspian Sea (in Azerbaijan) and in the Kura basin along the foothill of the Greater and Lesser Caucasus. It reflects the location of the Pleistocene marine coastline corresponding to the Late Akchagyl transgression of the Caspian Sea (Krasnov et al. 1974; Popov et al. 2004). Its age is given as between 2.5 and 1.8 Ma possibly as young as 1.2 Ma (see Mitchell & Westaway 1999 for discussion and references). This is the same period that the oldest hominids in the Caucasus area were living in the Dmanisi (Georgia) area, near the ocean shores of the eastern Paratethys (Lordkipanidze et al. 2007). Many other terraces can be seen up to altitudes of 3000 m into the mountain valleys; their potential to help date uplift or show intramontane erosional/depositional processes remains to be investigated. Evidence from river incision of several hundred of metres since the last glaciations suggests uplift rates of 10 mm a21 (Rastvorova & Shcherbakova 1967). Similar deep incisions are also observed in Plio-Pleistocene sediment of the Samur river in northern Azerbaijan, however detailed studies need to confirm the fast uplift rates. The development of palaeo-rivers such as the Kura (Djavadova & Mamula 1999), Volga, Terek and Samur rivers, and provenance studies have been given great attention in recent years due to oil exploration studies (Reynolds et al. 1998; Guliyev et al. 2003; Morton et al. 2003). They transport the sediments from their source area to their final sinks. All this development, younger than 10 Ma, possibly even 5 Ma, led to the formation of very high quality reservoirs making the circum Caucasus basins unique oil provinces. The river systems and their deltas also underline the rapid lateral changes between areas with high and low topography or below and above sea-level. The longlived river systems have also left a complex system of terraces, both in the valleys cutting into the mountain belt (mainly parallel and perpendicular to the main structural trend, (Lukina 1981)). The existence of several levels and ages of river terraces document continued incision possibly due to uplift of the mountain range (Shirinov 1973). This is also shown by the important cliffs of Quaternary material cannibalized on the northern slopes of the eastern Greater Caucasus near Quba (Figs 1 & 4) (Kangarli 1982). Connecting these ‘events’ to the different levels and ages of terraces along the

Fig. 6. (Continued) Maps of uplift rates, vertical movement, and Lower Pleistocene coastlines. (a) Uplift rates: map shows that highest rates of 10– 12 mm a21 are found in the central part of the western Greater Caucasus (Philip et al. 1989). High subsidence is occurring in the Kura Basin. (b) Total vertical movement since Sarmatian (10–13 Ma Miocene). Strong subsidence is seen north and south of the eastern termination of the Greater Caucasus in Azerbaijan (Ershov et al. 2005). (c) Lower Pleistocene coastline – map of coastline (Krasnov et al. 1974) which separates the area of the former marine domain and zone above sea-level with Pleistocene deposits, and area with unknown amount of palaeotopography. Notice the position of Dmanisi (Georgia) where hominid remnants dated back to 1.8 Ma have been discovered.

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Caspian Sea (Shcherbakova 1973; Mamedov 1997) opens the prospect of quantifying uplift v. subsidence and basin infill. One of the best data sets to directly quantify uplift are marine sediments such as those of Sarmatian age now found at altitudes of 3700– 4000 m a.s.l. such as in the Sahdag area of Azerbaijan (Budagov 1963) (Fig. 4). Akchagylian marine clays are found in the Budur (localiser) syncline at altitudes of 2000 m a.s.l. in the Azerbaijan Greater Caucasus (Kangarli 1982). Combined information from different datasets suggest uplift rates between 0.33 and 1.00 mm a21 for the last 10 Ma in eastern Greater Caucasus (Mitchell & Westaway 1999) but detailed rates for the different periods remain uncertain.

Palaeogeography During the Cenozoic the Greater Caucasus formed a vast island within the landlocked Paratethys Sea. Detailed palaeogeographic reconstructions indicate that the Greater Caucasus is emerging above sea level as lowland only after the mid Middle Miocene (13 –14 Ma) to grow into a mountain range after Late Miocene (11– 10 Ma) to Middle– Late Pliocene 3.4–1.8 Ma) (Popov et al. 2004). Palaeoclimate studies using spores and mostly pollen from Sarmatian (12–8 Ma), mostly marine sediments, in Georgia north of Tbilisi show the existence of pine forest at moderate to high altitudes and reflecting important climatic changes linked to mountainous topography and orogenic evolution (Shatilova 1992; Shatilova et al. 2008; Kokolashvili & Shatilova 2009). Prior to this Mid Miocene period, palaeogeographic reconstructions only indicate lowlands or islands or locally in the western Caucasus area mountainous areas. This is at odds with the fact that tectonic convergence and thrusting started already in Late Eocene times and most likely generated considerable topography. This is corroborated by the fact that important quantities of sediments are deposited in peripheral basins starting in Eocene –Oligocene, and that are most likely sourced in the nascent Greater Caucasus. An early Oligocene uplift is also confirmed by provenance studies and fission track investigations in the Maikop series (Early Miocene) of the western Greater Caucasus (Vincent et al. 2007).

Tectonic geomorphology Only a few combined studies on tectonics and geomorphology (Burbank & Anderson 2001; Keller & Pinter 2002; Delcaillau 2004) are available in the Greater Caucasus such as an investigation of the Alasani basin on the southern slope of the Greater Caucasus in Georgia (Philip et al. 1989; Triep et al. 1995).

This latter basin, runs parallel to the southern slope of the Greater Caucasus (Figs 1 & 5) and is filled with Pliocene–Quaternary sediments. This basin is bound to the north by one or several thrusts separating it from the steeper slopes rising into the high mountain range (Philip et al. 1989). To the east in Azerbaijan it can be shown that these faults are thrusting Cretaceous rocks over Oligocene series in the Basqual area (north of the town of Agsu; Fig. 7) (Kangarli 1982). Its front to the south is bound by a thrust-related fold which leads to a sharp topographic drop down to the Kura plain. The Alasani basin is thus an intramontane, piggyback-type basin with a relic thrust front that was active in the Miocene along its northern edge and a frontal thrust which is active in PlioceneQuaternary just south of the basin. In Georgia the Gagra-Dzhava zone located in the southern foothills of the main range is playing a similar role (Gamkrelidze & Gamkrelidze 1977; Dotduyev 1986; Gamkrelidze 1991). Its structure is similar to the Alasani Basin, but in addition shows somewhat farther travelled outliers of the main range, as is the case in the Basqual nappe structure in Azerbaijan (the eastern equivalent of the Alasani basin). Further to the south the Quaternary sediments of the northern part of the Kura–Kartli basin are folded and thrusted. Very prominent in the morphology the Karamarian Quaternary Anticline in Azerbaijan (Girdimanchai River – city of Agsu; (Figs 1 & 7) is a large doubly plunging anticline showing well developed water gaps formed by tributaries of the river Girdimanchai (Fig. 7) (Shirinov 1975). A large asymmetric south-verging anticline is developed over a blind thrust with splays. The water gaps were cut during thrust-related folding and the fast changes in tectonics (leading to local uplift that was not matched by erosion) caused the change to wind gaps for some valleys. In western Azerbaijan the active thrust front extends all the way into the foothills of the Lesser Caucasus, folding alluvial fans and forming gentle topographic highs and tilted terraces. Vertical faults with a strike–slip motion, are clearly reflected in the morphology as seen in examples from the Lesser Caucasus in Turkey and Georgia, as also discussed from earthquake focal mechanisms (Rebaı¨ et al. 1993; Koc¸yigit et al. 2001). The importance of strike –slip tectonics in the Greater Caucasus has been suggested and discussed in many instances, however without conclusive or convincing evidence. Remote sensing studies of lineaments (rivers, and mountains crest, as well as faults (Cloetingh et al. 2007)) and field investigations on faults with strike –slip movements show a general pattern of NNE –SSW oriented young anti-Caucasian faults cross-cutting all the major fold-and-thrust structures. The general

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Fig. 7. Pictures of the geomorphology of the Karamarian anticline (top image; Landsat 7 image) in the northern Kura basin of Azerbaijan, near the city of Agsu. The doubly plunging anticline is cut by the river Girdimanchai to form water gaps due to river incision into the growing structure. Detailed topography along profile A-B (top right) shows the asymmetry of the Karamarian anticline, suggesting a buried thrust with top to the south movement. Dashed black line shows topographic limit of the southern steep slopes of the Basqual nappe unit (equivalent to the Alasani basin) to the north. Lower picture shows lateral view of the Karamarian anticline looking towards the south.

palaeostress orientation resulting from fault kinematic analyses shows a north–south to NNE– SSW oriented compression, with east– west extension and vertical intermediate stress axis confirming that they are recent faults that have suffered no rotation. However, no large-scale vertical faults or fault system crosscutting the whole mountain range could be observed.

Topography and thrusts Based on a digital elevation model at 90 m we have extracted a contour map (Fig. 5) and superposed the main thrusts bounding the large changes in topography. Many thrusts coincide with changes in topography. In many other instances of topographic changes no thrusts are known, opening the

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possibility to denominate promising locations for thrusts, which remain to be ground-truthed. The most striking feature is the MCT which forms the southern limit of the zone of highest topography throughout the Greater Caucasus. The relationship topography – thrust trace is even more clearly shown on two topographic profiles through the eastern Greater Caucasus (Fig. 8). Especially the relic thrust fronts of the Alasani basin and the folds and thrusts of the Karamarian-Kakheti zone stand out from the flat surface of the Kura basin. The morphological changes associated with the present active thrust front at the foothills of the Lesser Caucasus are too small to show on the profiles. A second set of profiles showing the changes in slope emphasize the position of the main thrusts! The combined use of topography and slope helps to determine the position of the major thrusts which can be refined by the use of more detailed geological maps and satellite imagery (work in progress on a new tectonic map of the Greater Caucasus). This method can be used as a prospective tool to determine active (in combination with seismicity), but also fossil thrusts or orogenic fronts. The major topographic changes and thrusts can be followed throughout the whole Greater Caucasus.

Discussion and conclusion New field evidence on thrust geometry, tectonic geomorphology, and tectonics combined with literature data on Cenozoic tectonics in Azerbaijan led us to investigate the relationship between tectonics, topography and uplift in the eastern Greater Caucasus and correlate and expand our findings along strike to the whole Greater Caucasus. † The Greater Caucasus is an intracontinental doubly verging orogen resulting from the inversion of the Mesozoic, rift-related Greater Caucasus basin. The convergence between Arabia and Eurasia led to the closure of this basin and generated underthrusting and stacking of tectonic units. No subduction zone appears to exist under the Greater Caucasus, though incipient/young subductions are described for the western and eastern terminations of the Greater Caucasus into the Black Sea and Caspian Sea, respectively. During the orogenic development important foreland basins developed on both sides of the mountain range. Their development is contemporaneous with the infill of super deep sedimentary sinks in the Black Sea and the South Caspian Basin. † The Greater Caucasus is a fast growing orogen that started to build to its present topography since Late Eocene. Detailed timing of uplift above sea level remains elusive, but there is

clear evidence, including age and position of relic mountain fronts, that the orogen was growing in Oligocene and probably in Late Eocene already. The climax of growth was during the Late Cenozoic, starting in the Middle Miocene and accelerating in Plio-Pleistocene. Information about uplift remains very unevenly distributed. Total uplift since Sarmatian is in excess of 3700 m as documented in the Eastern Greater Caucasus in Azerbaijan. Uplift rates in excess of 10 –12 mm a21 in the central part of the mountain range and based on river incision since the last glaciations, remain questionable and need to be confirmed especially when confronted with data showing more moderate uplift rates of 0.33 to 1.00 mm a21 over the last 10 Ma. † The Main Caucasus Thrust is a discrete major thrust stretching across the whole Greater Caucasus from the shore of the Black Sea in the west to Azerbaijan in the east. In Azerbaijan and towards the shore of the Caspian Sea the MCT splits into several fault strings and affects a more diffuse zone, where both top-to-the-north and top-to-the-south thrusting is observed. The zone of highest topography, which correlates with the zone of fasted uplift, is bound to the south by the MCT which shows important top-to-the-south movement. To the north this zone is bound by south dipping thrusts with top-to-the-north movement. Detailed cross sections from northern Azerbaijan and insight from the Dagestan FTB suggest that this northverging backthrusting is linked to a thrust ramp system at depth. Thus the zone of highest uplift forms a triangular shaped domain limited to the south by the MCT and to the north by backthrusts. This triangular zone of fast uplift and its associated major topographic/tectonic features can be correlated across the whole Greater Caucasus. Uplift over such a more or less steep ramp system may also explain why we observe such important uplift (a minimum of 3700 m since Sarmatian) in a rather restricted area. † Major displacement has been accommodated along the MCT as suggested from the location of the strongest earthquakes known is the area (.6 Mg). The 1991 earthquake of high magnitude in Georgia in the Racha area, but also by historical earthquakes in Azerbaijan such as in 1667 in the Shamakhi area occur on thrusts located to the south of the MCT. These quakes area in similar structural position on a thrust system structurally beneath and to the south of the MCT and we interpret them as outward (to the south) propagating splays of the MCT. † Progressive southward migration of the thrust front during the Cenozoic formed successive active thrusts that were subsequently abandoned

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Fig. 8. Sections across the Greater Caucasus showing the average evolution of topography along two transects as well as the changes in slope along the same transects (position of cross-sections see Fig. 5).

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to form relic thrust fronts. The progression is registered in the geomorphology by clear changes in topography, deviation of rivers and water gaps. This is clearly documented in the thrust faults north, than south of the AlasaniBasqual (Basqual is a lateral structural equivalent of Alasani basin in Azerbaijan) structure, followed to the south by the Kartli FTP and the structures such as the Karamarian anticline. The active orogenic front zone, both to the north in Dagestan and to the south in the foothills of the Lesser Caucasus, can be deduced from morphological, structural and seismicity criteria. † Unlike the western Greater Caucasus, the eastern Greater Caucasus is seismically very active. This high seismicity correlates with higher present convergence rates in this area. We suggest that this reflects the continued growth of the foreland basins outboard of the main range into the Quaternary to recent sediment filled foredeeps. Simultaneously, as suggested by earthquakes in the central parts of the orogen, we observe continued growth and uplift in the central part. These zones correlate with the areas of high topographic gradients such as the MCT or other areas west of Tbilisi on the southern slope of the Georgian Greater Caucasus. The difference between east and west is due to different factors such as the east to west decreasing plate convergence rate. But, also the inherited structure possibly plays a major role. Indeed, the basement units present in the west are not observed to the east and this may be interpreted as a difference in the initial palaeotectonic setting that was subsequently differentially inverted during compression. In addition the observed recent magmatic activity in the west may suggest a difference in the crustal or whole lithosphere structure and strength contributing to a different evolution east and west. † Strain partitioning across, and parallel to, the Greater Caucasus is very heterogeneous, both vertically and horizontally. Present day plate convergence as observed from GPS data is oblique with a larger convergence in the East. Overall this convergence converted into a 6– 14 mm a21 horizontal deformation across the Lesser Caucasus and the Greater Caucasus. An important drop in velocities across the MCT in the Greater Caucasus, indicates, in agreement with what we suggest in this paper, that there is no horizontal displacement, but dominantly uplift! The MCT is a main boundary north of which we observe the development of triangular zone of uplift related to an important tectonic ramp system at depth. West of Tbilisi (Georgia) in Racha area seismicity indicates partitioning of strain between the MCT and the thrusts beneath

the Gagra-Dzhava zone. As seen from GPS data and tectonic geomorphology the meridional FTB and the thrust front is clearly propagating, but the present activity is the Dagestan FTB and foreland basin. An example of vertical strain partitioning between surface and deeper structures is suggested at the eastern transition of the GC into the South Caspian basin. Here in the Gobustan area GPS data suggest a dextral strike–slip motion along a major fault, the West Caspian Fault, whereas earthquake focal mechanisms suggest a thrust motion to the east which may be linked with westward subduction connected with subduction of the South Caspian Basin. Clear links between geomorphology, seismicity and tectonics can be observed in the Greater Caucasus. With the exception of the Racha area in Georgia, mainly the eastern Greater Caucasus appears to be active at present. Strain partitioning across the mountain belt is very heterogeneous. In the absence of enough detailed data precise timing and assessment of rates of uplift and deformation and the partitioning of strain into uplift and horizontal deformation, as well as the partitioning between thrusting and strike–slip remain difficult to assess. We would like to thank especially the MEBE Programme and its sponsors for supporting this research. We are further indebted to the Swiss national science foundation for their financial support (SCOPES project IB7320110973) as well as INTAS (project 06-1000017-8930) and the University of Fribourg. We also would like to thank all the people in Azerbaijan and Georgia that were involved in this project and that helped in one way or another. We thank the reviewers for their constructive suggestions.

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Late Palaeozoic – Early Cenozoic tectonic development of the Eastern Pontides (Artvin area), Turkey: stages of closure of Tethys along the southern margin of Eurasia ¨ MER1* & ALASTAIR H. F. ROBERTSON2 TI˙MUR USTAO ¨ niversitesi, 34850-Avcılar, I˙stanbul, Turkey Jeoloji Bo¨lu¨mu¨, ˙Istanbul U

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School of GeoSciences, Earth and Planetary Science Group, University of Edinburgh, West Mains Road, Edinburgh, EH9 3JW, UK *Corresponding author (e-mail: [email protected]) Abstract: Early Carboniferous – Eocene units exposed in the Arvin area document the development of the southerly, active continental margin of Eurasia. The oldest rocks exposed in the area are Early Carboniferous granites that regionally intrude schists and gneisses. The continental terrane rifted along the entire length of the Pontides (.1000 km east–west) during the Early– Middle Jurassic. Subsidence of the rift basin in the Artvin area was accompanied by terrigenous debris flows, turbidites and deep-sea radiolarian muds, and was associated with local extrusion of chemically ‘enriched’ basalts. Swarms of subduction-influenced basic, intermediate, to locally silicic dykes, intruded high-grade metamorphic basement within the rift. A basement horst within the rift was covered by condensed pink ammonite-bearing pelagic facies. Large volumes of subduction-influenced basalts erupted during the later stages of extensional basin development (Mid-Jurassic), associated with volcaniclastic sedimentation. The Artvin Basin is interpreted as a supra-subduction rift associated with incipient arc magmatism. The basin was stratigraphically inverted in response to Late Middle Jurassic ‘Neo-Cimmerian’ deformation. It was then partially eroded and covered by Upper Jurassic continental, to shallow-marine sediments, together with localized eruption of ‘enriched’ (non-subduction-influenced) basalts. The margin collapsed during the Late Jurassic–Early Cretaceous, initiating deposition of pelagic carbonates and mixed terrigenous, biogenic and volcaniclastic gravity flows. Subduction during the Late Cretaceous then constructed the east Pontide magmatic arc and a thick volcaniclastic fore-arc apron to the south. Supra-subduction-type ophiolites and accretionary melange formed within Neotethys to the south during the Late Cretaceous and were emplaced regionally northwards onto the leading edge of the Pontide active continental margin during the latest Cretaceous. Continental collision during the Mid-Eocene telescoped the distal part of the active margin which was emplaced northwards onto the east Pontide continental basement. The geological evolution of Artvin area correlates with the Pontides further west and with the southern and northern Transcaucasus to the east. Our favoured tectonic model involves long-lived, episodic, northward subduction of Tethys. Finally, there is no evidence of ‘Palaeotethyan’ ophiolites in the eastern Pontides region.

The geology of northeastern Turkey sheds light on the tectonic development of the southern continental margin of Eurasia and Tethys to the south (Robinson et al. 1995). Here, we report evidence from the Artvin area, in the eastern Pontides, adjacent to the Caucasus of Georgia, Armenia and the Russian Federation to the east (Fig. 1). The eastern Pontides is bounded by the Black Sea to the north and the ˙Izmir–Ankara –Erzincan– Sevan suture zone to the south. The eastern Pontides can usefully be compared with the central and western Pontides (Yılmaz et al. 1997) and the Caucasus to the east (Nikishin et al. 1998, 2001). A near consensus exists that the Early to Mid-Mesozoic rocks of the Pontides, the Crimea and the Caucasus relate to the genesis of one, or several, back-arc basins

behind a northward-dipping Tethyan subduction zone (e.g. Adamia et al. 1977, 1981, 1995, 2001). In this paper, we will discuss and interpret the Artvin area in the eastern Pontides, named after the regional city (Fig. 2). The Artvin area is dominated by the eastern Pontides magmatic arc of Late Cretaceous age that forms mountains up to 4000 m high (Taner & Zaninetti 1978; Akıncı 1984; Bektas¸ et al. 1999; Yılmaz et al. 2003). Older units are exposed beneath the arc and also within regional-scale thrust sheets further south that were finally emplaced northwards during the Mid-Eocene closure of Tethys. The Upper Palaeozoic basement of the eastern Pontides is constructed of several different types of magmatic and metamorphic rock. In some places Upper Carboniferous

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 281–327. DOI: 10.1144/SP340.13 0305-8719/10/$15.00 # The Geological Society of London 2010.

282 ¨ MER & A. H. F. ROBERTSON T. USTAO Fig. 1. Tectonic map of the Pontides, Caucasus and adjacent regions. The study area in the far NE of Turkey is marked by a red box. Several additional units that are mentioned in the paper are too small to show at the scale of this map: CUZ, Chortchana–Utslevi unit; DS, Dzirula Salient; LS, Loki Salient (all in the Caucasus); KTB, Kocaeli Triassic Basin (in NW Turkey).

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Fig. 2. Outline geological map of the eastern Pontides in Turkey showing the main tectonic units in the region. The large box shows the main area discussed in this paper. Modified from the 1:500 000 geological map of Turkey (MTA 2002).

to Permian, mainly clastic sedimentary rocks, unconformably overlie the basement, followed by Lower Jurassic – Paleocene terrigenous, calcareous and volcanogenic lithologies. We will also discuss successive stages in the development of the eastern Pontides from the Late Palaeozoic to Early Cenozoic. New and existing literature will be integrated and synthesized to test alternative tectonic models. The time scale of Gradstein et al. (2004) is used throughout.

Previous work Mapping of the Artvin area was carried out by the Mineral Research and Exploration Institute (MTA) during the 1960s (Baydar et al. 1969, 1977) and more recently (Konak & Hakyemez 1996, 2001). The latter authors divided the region into four main tectonic zones bounded by major tectonic contacts (Fig. 3). The first is the East Pontide Autochthon in the north, which is structurally overlain

by elongated NE–SW-trending allochthonous units that we term the Lower Slice Complex (c. Artvin-Yusufeli Zone of Konak & Hakyemez 1996, 2001). This includes metamorphic basement, swarms of sheeted dykes and Jurassic sedimentary and volcanic units. This unit is structurally overlain by the Upper Slice Complex (c. Olur-Tortum Zone of Konak & Hakyemez 1996, 2001). This includes small exposures of basement lithologies, Middle Jurassic volcanic rocks and Upper Cretaceous –Lower Cenozoic volcanic and sedimentary successions. Structurally above comes the ˙IzmirAnkara-Erzincan suture zone (c. Erzurum-Kars suture zone of Konak & Hakyemez 1996, 2001), which is made up of ophiolitic rocks, minor metamorphic rocks (e.g. blueschist) and latest Cretaceous – Paleocene sedimentary rocks. S¸engo¨r et al. (1980) and Yılmaz & S¸engo¨r (1985) reported the existence in the Artvin area of an intact pre-Upper Jurassic ophiolite and used this to support a regional model of mid-ocean ridge spreading

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Fig. 3. Regional geological map of the area studied during this project. Modified from the 1:500 000 geological map of Turkey (MTA 2002). The area mapped during this work is marked as a box (see Fig. 4).

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followed by southward subduction of a Palaeotethyan ocean. They described the Artvin area as a NE–SW-trending regional-scale antiform (exposed on either side of the C ¸ oruh River). They reported an intact ophiolitic pseudostratigraphy, with serpentine at the base, passing upwards into gabbro, microgabbro and diabase, then basalt, although sheeted dykes were not reported. Overlying sedimentary units were described as sandstone-shale ‘distal turbidites’, pelagic shale and also manganiferous ribbon chert. The ‘Artvin ophiolite’ was correlated with apparently similar units in the central Pontides, namely the ‘Ku¨re and C ¸ angaldag˘ ophiolites’, as parts of an originally wide Palaeotethyan ocean (Yılmaz & S¸engo¨r 1985). Adamia et al. (1995, 2001) presented a regional cross-section and correlations with the Caucasus, as exposed in Georgia to the east. They identified radiolarians and ammonites of Early Jurassic age above a metamorphic basement in the Artvin area. These authors envisaged the eastern Pontides as part of a Late Palaeozoic– Early Jurassic back-arc basin related to northward-dipping subduction. They also favoured correlation with a Ku¨re ophiolite in the central Pontides. Dokuz et al. (2006) correlated the east Pontide crystalline basement in the Artvin area (Karadag˘ unit) with the metamorphic Pulur Massif further west (Tanyolu 1988). According to this work, the metamorphic basement in the Artvin area includes schists, gneisses and metabasites, with medium to strong schistocity, local ptygmatic and migmatitic fabrics, and occasional mylonites. The basement is cut by NE –SW-trending lenticular intrusions (c. 30 km2), ranging from relatively mafic plagioclase-cumulates, to quartz-diorites. Both the metamorphic basement and the mafic intrusive rocks were seen as being cut by two NE –SWtrending felsic intrusions of tonalite, trondhjemite, subordinate quartz-diorite and granodiorite. Sub-parallel felsic dykes crosscut relatively mafic intrusive rocks; also, felsic veins locally intrude older mafic intrusive rocks. A single large granitic body was reported to be cut by gabbroic intrusive rocks. Based on chemical analysis (36 samples) Dokuz et al. (2006) inferred that an earlier, relatively basic gabbro-diorite intrusion could have been created by 15–20% melting of lherzolite within a mantle wedge located above a subduction zone. Laterstage more felsic tonalites and trondhjemites were derived from similar source rocks, but not from the mantle wedge directly, or by fractionation of mafic magma. Dokuz et al. (2006) pointed out that the felsitic intrusions show a low-K, calc-alkaline trend towards trondhjemite, as in some arc-related complexes. They envisaged the deep levels of a continental margin arc, similar to the Cretaceous

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Sierra Nevada of Western USA, as a suitable tectonic setting. Dokuz & Tanyolu (2006) also gave a general description of the sedimentary cover of the metamorphic basement. A lower unit c. 1000 m thick of pillowed/massive basalt and basaltic andesite was described as being interbedded with thin-bedded silts and shales (c. Kelkit Formation of Bergougnan 1975; ‘Hamurkesen Formation’), followed conformably by an upper unit of shallow-water sedimentary rocks with interbedded basic volcanic rocks. Petrographic study by Dokuz & Tanyolu (2006) of Lower –Middle Jurassic sandstones suggested that feldspathic wackes (arkosic arenites) predominate; these plot in the continental block field (with basement uplift), to transitional arc field on petrographic discrimination diagrams. Major element, trace element and rare earth elements (REE) analysis of volcaniclastic sandstones suggested a mainly magmatic arc source, whereas more quartz- and shale-rich samples plotted in the passive margin field on tectonic discrimination diagrams. In addition, chemical data suggested a single cycle of clastic sediment weathering and deposition. The main source of the Lower Jurassic clastic sediments was inferred to be Upper Palaeozoic granites and a no-longer exposed carapace of arc-type volcanic rocks. Our work has resulted in some significantly different field observations and interpretations, as discussed below.

Tectono-stratigraphy and structure Part of the Artvin area was mapped at 25 000 scale (Tortum G47 a1, a2, a3, a4, b1, b4, Artvin F47 c4, d3 sheets) during 1997– 2002 by the first author (Ustao¨mer 1998), as shown in Figures 4 and 5. Both authors carried out additional fieldwork during 2001. In this paper, we use an informal tectono-stratigraphy, as summarized in Figure 6. The East Pontide Autochthon is overlain, first by the Lower Slice Complex (c. Artvin-Yusufeli Zone), which itself consists of several internally deformed units (Ustao¨mer 1998). Above comes the Upper Slice Complex (c. Olur-Tortum Zone), which comprises up to five thrust sheets. Upper Cretaceous ophiolites occur at still higher structural levels. The tectonic contacts between the units of the Lower Slice Complex are relatively steeply inclined compressional shear zones, in contrast to mainly low-angle thrusts that characterize the Upper Slice Complex. The Lower Slice Complex experienced compression-related deformation prior to erosion and the accumulation of an Upper Jurassic sedimentary cover. Lower-Middle Jurassic successions overlie basement units, known as the Narlık Granite and the Karadag˘ Metamorphics (Fig. 6)

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Fig. 4. (a) Geological map of the area studied in detail during this work, based on 1:25 000 mapping by the first author (1997–2002). Note: Additional, higher thrust sheets occur within the Upper Slice Complex to the south of the mapped area (see Fig. 3). The southerly part of the area was mapped by Baydar et al. (1969), Konak & Hakyemez (1996) and Dokuz et al. 2006); (b) Key to the units.

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Fig. 5. Representative cross-sections of the map area. (a) NW–SE section, A-A’; (b) NE– SW section, B-B’ (see Fig. 4 for locations).

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Fig. 6. Summary of the tectono-stratigraphy of the Artvin region as used in this paper. See text for explanation.

and are deformed into large-scale (hundreds of metres) north-facing recumbent, to isoclinal folds (Fig. 5b). However, similar folding does not affect the unconformably overlying Upper Jurassic – Lower Cretaceous successions within the Upper Slice Complex, or the East Pontide Autochthon to the north. The Upper Slice Complex experienced northward-directed thrusting and folding that

mainly resulted from Mid-Eocene suturing of Tethys to the south.

Hercynian basement and intrusive rocks Basement units are exposed in four tectonic positions: within the East Pontide Autochthon in the north, within two units of the Lower Slice

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Complex, and within the lower part of the Upper Slice Complex (Fig. 6).

East Pontide Autochthon Within the autochthonous foreland in the north, the basement is mainly composed of unmetamorphosed granitic rocks. Separated zircons have been dated by the laser ablation-SF ICP-MS method as Early Carboniferous (Visean) in age (Ustao¨mer et al. 2010). An unmetamorphosed granite (Narlık Granite) in the north of the area studied intrudes a low-grade assemblage of meta-conglomerates, meta-sandstones and phyllites. The granite is petrographically similar to that of the Pontide autochthon regionally, although the geochemistry differs (see below).

Demirkent Intrusive Complex Crystalline basement rocks, known as the Demirkent Intrusive Complex are exposed within the

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structurally overlying Lower Slice Complex (Fig. 6). They occur as volumetrically minor screens between swarms of basic dykes that dominate the Demirkent Intrusive Complex (Fig. 7). In general, the basement host rocks are restricted to less than several-metre-wide screens between dykes. Detailed study of the field relations reveals that the rocks within the screens developed in three main stages (Fig. 8). The first stage involved metamorphism and deformation of mafic protoliths to form banded amphibolite (S1). The second stage was related to the intrusion of granitic (i.e. metatonalite) veins (mostly ,10 cm across); these were later foliated (S2). Granitic segregations have yielded Early Carboniferous (Serpukhovian) ages (Ustao¨mer et al. 2010). The S1 and S2 fabrics are sub-parallel in most outcrops, but locally show an angular discordance of up to 208 (as seen in road sections near Demirkent village; Fig 4). Swarms of basaltic, to locally felsic, dykes were intruded during a third stage of development. Associated pegmatitic granite intrusions contain muscovite

Fig. 7. Summary of the lithologies exposed in the East Pontide Autochthon and the overlying Lower Slice Complex (see Fig. 6).

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Fig. 8. Sketch showing an outcrop of a screen of crystalline basement rocks between a swarm of younger sheeted intrusions within the Demirkent Intrusive Complex. The figure shows banded amphibolite cut, in turn by small-scale granitic intrusions (meta-tonalite) and then by basaltic dykes. The meta-tonalite intrusion is Early Carboniferous in age. Near Demirkent village (Fig. 4), on the Artvin-Yusufeli main road.

that yielded a Bajocian age using the Ar/Ar method (unpublished data). Long intervals (tens of metres) of the Demirkent Intrusive Complex comprise up to 100% of different intrusive lithologies, as seen in a road section near Demirkent village (Fig. 9). The basic intrusive rocks range from individual dykes several centimetres wide, to larger intrusive bodies, up to tens of metres across, all with well-developed chilled margins. The intrusions strike at c. 0108 – 0308,

generally parallel to the strike of the bounding tectonic contacts of the Demirkent Intrusive Complex (Fig. 4). The dykes are generally subvertical, or steeply inclined to the NW or SE. The foliation in the basement host rocks is locally offset across the dykes in a manner that suggests that the intrusion was accompanied by extensional faulting. Individual intrusions range from gabbro to plagiogranite (Fig. 9). The largest basic intrusion is gabbroic, with local pegmatitic segregations. The

Fig. 9. Details of intrusive relations within the Demirkent Intrusive Complex (see Fig. 6). The amphibolitic host rocks are successively cut by dyke-like intrusions of gabbro, diabase, basalt, porphyritic andesite and granite. Near Demirkent village, on the Artvin-Yusufeli main road. See text for explanation and Figure 4 for location.

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gabbros are intruded by sub-parallel diabase and basaltic dykes, and also by later-stage tonalite and rare dacite dykes. In one well-exposed interval near Demirkent early micrograbbo dykes (,5 m wide) are intruded by later diabase dykes (,70 cm wide); these are bisected by narrow (,30 cm) tonalitic dykes.

Karadag˘ Metamorphics Crystalline basement rocks, known as the Karadag˘ Metamorphics occur further south, within a structurally higher part of the Lower Slice complex (Fig. 6). Banded gneiss and quartz-mica schist are cut by occasional basaltic dykes and small granitic intrusions (,30 m wide). The host rocks exhibit ductile to brittle fabrics, transitional to mylonites. Many exposures are strongly folded and exhibit small ptygmatic folds.

Metamorphic and intrusive rocks of the Upper Slice Complex To the east of our main study area (near Olur; Fig 3) the lower structural levels of the Upper Slice Complex have been mapped as including low-grade pelitic and psammitic schists (c. Kıs¸la Metamorphics of Konak & Hakyemez 1996, 2001). In addition, a tectonically bound sliver of high-grade metamorphic rocks is cut by swarms of basic, to felsic, dykes (c. Gu¨vendik Dike Complex of Konak & Hakyemez 1996, 2001). The crystalline metamorphic rocks of the East Pontide Autochthon and the Lower and Upper Slice Complexes can be restored as different parts of the regional basement of the East Pontide Autochthon.

Geochemistry of the East Pontide Autochthon Geochemical analysis of mainly granitic and basic volcanic rocks (see Tables 1 & 2) was carried out to help determine tectonic setting. The analysis of major and trace elements was mainly carried out by X-ray fluorescence (XRF), using the method of Fitton et al. (1998). In addition, a small number of samples were analysed by ICP MS at Minnesota University, USA. Upper Palaeozoic granitic rocks from the East Pontide autochthon (five samples) exhibit patterns that are suggestive of a non-subduction influenced, intrusive setting on ocean ridge granite (ORG)-normalised spider diagrams (Fig. 10a) and chondrite-normalized spider diagrams (Fig. 10b) (i.e. they lack a negative Nb anomaly). Occasional, cross-cutting dykes (three samples) that were analysed by ICP MS exhibit Hf and Zr depletion but no relative Nb depletion (Fig. 10a). The marked

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Table 1. Representative major, trace and rare earth element (REE) analysis of granitic rocks from the East Pontide Autochthon. 1 – 3, C¸amlıkaya granite; 4 – 6, Narlık granite. 1 – 3 were analysed by XRF and 4 – 6 by ICP-MS Sample locations (sample number: 1:25 000 topographic map sheetgrid reference) CN 27: Tortum G46 d4, 73.98-90.98 CN 28: Tortum G46 d4, 73.96-90.98 CN 31: Tortum G46 d4, 76.76-92.30 NAR 390: Artvin F47 c4, 32.56-42.78 NAR 600: Artvin F47 c4, 32.38-42.63 NAR 640: Artvin G47 b1, 32.61-42.45

Eu depletion (Fig. 10b) is suggestive of plagioclase fractionation. The Narlık Granite (11 samples) plots in the field of Volcanic Arc Granite on the Y þ Nb v. Rb diagram (Fig. 11a), and on the Volcanic Arc granite plus the syncollisional granite field on the Nb v. Y diagram (Fig. 11b). Table 2. Representative major and trace element, by analysis XRF, of basaltic rocks from the East Pontide Autochthon and the overlying allochthonous units. 1 – 8, Diabase dykes from the Demirkent Intrusive Complex; 9 – 10, Host amphibolites from the Demirkent Intrusive Complex; 11– 12, Pillow basalt from the Zeylinlik lava unit; 13 – 15, Basalt from the Irmakyanı Formation; 16 – 17, Early Jurassic basalt from the East Pontide Autochthon; 18 – 21, Basalts from the Kınalıc¸am Formation. See Figure 6 for tectono-stratigraphy Sample locations (sample number: 1:25 000 topographic map sheet grid reference) TU97.26: Tortum G47 a3, 23.81-21.41 PO.01.34: Tortum G47 a3, 27.52-27.34 PO.01.35: Tortum G47 a3, 27.66-27.82 TU97.35b: Tortum G47 a2, 29.75-29.84 TU97.36: Tortum G47 a2, 29.70-29.83 TU97.37: Tortum G47 a2, 29.68-29.84 TU97.38: Tortum G47 a2, 29.66-29.80 TU97.39: Tortum G47 a2, 29.60-29.81 TU97.49: Tortum G47 a2, 30.35-30.64 TU97.51: Tortum G47 a2, 30.18-30.73 ZEYP 4: Artvin F47 c4, 38.22-54.34 PO.01.01: Artvin F47 c4, 38.34-54.40 TU97.22: Tortum G47 a4, 18.54-21.33 TU97.23: Tortum G47 a4, 19.58-21.62 TU97.24: Tortum G47 a4, 20.34-22.05 PO.01.66: Tortum G46 d2, 85.60-01.12 PO.01.67: Tortum G46 d2, 86.72-01.33 PO.01.42: Tortum G47 b1, 36.58-34.15 PO.01.43: Tortum G47 b1, 36.34-33.67 PO.01.44: Tortum G47 b1, 36.27-33.70 PO.01.45: Tortum G47 b1, 36.24-34.01

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Fig. 10. Geochemistry of rocks from crystalline basement units. (a) Ocean-ridge granite (ORG)-normalized spider diagram of Upper Palaeozoic granite from the East Pontide Autochthon and cross-cutting granitic dykes (near Ispir); (b) Chondrite-normalized REE spider diagram of the same granitic basement and granitic dykes (,5 m wide). See text for explanation.

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lower parts of the Lower Slice Complex (ArtvinYusufeli Zone) on either side of the Demirkent Intrusive Complex (Figs 4, 6 & 7).

East Pontide Autochthon

Fig. 11. Tectonic discrimination of Narlık granite on, (a) Y þ Nb v. Rb and, (b) Y v. Nb diagrams (Pearce et al. 1984). Syn-COLG, Syn-collisional granite; WPG, Within-plate granite; VAG, Volcanic arc granite; ORG, Ocean ridge granite. The samples plot in the VAG and in the VAG þ syn-collisional granite field. See text for explanation.

In addition, 34 samples of amphibolitic hostrocks and cross-cutting dykes were analysed from the Demirkent Intrusive Complex. Only a small number of these samples are of basaltic composition. On multi-element spider diagrams the basiccomposition amphibolites show Nb-depletion relative to Ce and a near flat trend from Zr to Y relative to MORB (mid-ocean ridge basalt) (Fig. 12a). In contrast, the dykes show LIL (large-ion lithophile) element enrichment, with Nb depletion relative to La and a negative slope between Zr and Ti (Fig. 12b).

Lower –Middle Jurassic sequences Lower– Middle Jurassic sequences are widely exposed in the East Pontide Autochthon. Sequences of this age span are also exposed in the upper and

The Upper Palaeozoic granitic basement is unconformably overlain by a mixed terrigenous, calcareous and volcanogenic succession, as well exposed c. 25 km west of Yusufeli, near ˙Ispir (Fig. 3; MTA 1998). The succession (c. 500 m thick) can be correlated with the regional Kelkit Formation (‘Hamurkesen Formation’) (Fig. 13), begins with lenticular conglomerates, overlain by coarse-grained sandstones and dark calcareous siltstones and mudstones. The carbonate component includes ooids, oolitically coated grains, shell fragments, reworked oolitic limestone, echinoderm plates, microbial carbonate and coral fragments. The non-calcareous component is mainly large quartz grains, quartzite lithoclasts, altered siliceous volcanics, rhyolite (with partially resorbed quartz phenocrysts) and recrystallized siliceous volcanic rocks. There are then grey bioclastic limestones, lenticular terrigenous sandstones with mudstone partings, muddy sandstones, silty carbonates and nodular limestones. The succession passes upwards into wellstratified, coarse, pebbly volcaniclastic sandstones, alternating with red bioclastic sandstones and red muddy micritic limestones. The limestones include benthic foraminifera (Involutina, Trocholina sp., Lenticulina sp., Viladina sp. Lingulina and Lagenidae), of inferred Early Jurassic age (MTA 1998). Associated volcaniclastic sandstones locally contain black glassy basalt, feldspar microphenocrysts and rare feldspathic basalt grains. Massive lavas (.100 m thick) above are locally columnar jointed (Fig. 13) and intercalated with red volcaniclastic siltstones and mudstones. The volcanic rocks are cut by several small (,10 cm-wide) neptunian dykes, indicative of crustal extension.

Lower Slice Complex Lower-Mid Jurassic basinal succession. Within the lower, southerly part of the Lower Slice Complex (Fig. 6), the Narlık Granite basement and its host meta-sedimentary rocks are unconformably overlain by a thick sedimentary and volcanogenic succession, c. 1.5 km thick, that we term the Irmakyanı Formation (c. Zeytinlik Subzone of Konak & Hakyemez 1996, 2001). The type area is near Irmakyanı Village (Fig. 4). We introduce this new formation name because the sequence differs significantly in facies from the Lower– Middle Jurassic Kelkit Formation (‘Hamurkesen Formation’), as exposed both in the

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Fig. 12. MORB-normalized spider plots. (a) Basic amphibolite screens within the Demirkent Intrusive Complex; (b) Basic dykes within the Demirkent Intrusive Complex. See text for explanation. Normalizing values are from Pearce (1983).

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Fig. 13. Measured log of the Lower-Middle Jurassic part of the succession unconformably overlying Hercynian granitic basement in the East Pontide Autochthon (near Ispir). See Figure 3 for location.

East Pontide Autochthon (see above) and also further west in the Pontides (e.g. Koc¸yig˘it & Altıner 2002). During this work we measured numerous sedimentary logs of different outcrops (e.g. Fig. 14) to produce a composite succession (Fig. 7b). The Irmakyanı Formation is divided into a mainly terrigenous lower part and a mainly

volcanogenic upper part. The lower, mainly terrigenous part of the succession is well exposed in the Cevizlik-Narlık area (Fig. 3), whereas the upper volcanogenic unit is well exposed near Irmakyanı (Fig. 4). In addition, intervening siliceous biogenic sediments are seen along the main road beside the C¸oruh River (e.g. near C¸ag˘layan, Fig. 4).

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Fig. 14. Measured logs of Lower-Middle Jurassic sequences within the Lower Slice Complex (Irmakyanı Formation), (a) lower part; (b) mid-part; (c) upper part. See Figure 4 for locations. (a– c) can be correlated along strike as a single intact stratigraphic succession (see Fig. 7b).

The Irmakyanı Formation is commonly steeply dipping, cleaved, folded on a hundreds-of-metres scale, and exhibits several phases of crenulation cleavage and kink banding. Most outcrop-scale folds are upright, or overturned to the NW. Matrix-supported debris flows with clasts of gneiss, quartz-mica schist and granitic rocks characterize the lowermost c. 300 m of the sequence above the granitic basement (Fig. 7b). Individual lenticular depositional units, up to tens of metres thick, exhibit channelized bases. Near the basal unconformity, clasts are ,30 cm in size and mainly well rounded but become more angular upwards. The succession grades into sandstone turbidites and shales, estimated as several hundred metres thick, after taking account of folding and faulting (e.g. near C¸ag˘layan; Fig. 14a). Thick-bedded sandstone turbidites predominate low in the succession; these are mainly

well cemented, medium- to coarse-grained, and exhibit partial Bouma sequences (Bouma 1962). The sandstones are mainly well-sorted, sub-mature sublitharenites and subarkoses, including fresh plagioclase, perthite (exsolved feldspar), monocrystalline and polycrystalline quartz, muscovite and minor opaque grains, set in a fine-grained matrix of clay minerals and quartz. Near Zeytinlik, a fault-bound volcanic interval is exposed within the overall succession of siliciclastic turbidites (Fig. 15). This starts with pillow lava, overlain by massive lava and then by further pillow lava. These extrusive rocks are overlain by volcaniclastic sediments, black shale and siltstones. Massive lava follows with rare interlayers of siliciclastic sandstones and black shale (Fig. 15). The upper part of this interval includes basic bodies that are interpreted as sills (see below).

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Fig. 15. Measured log of Lower Jurassic sequence in the higher part of the Lower Slice Complex. Basic lavas (Zeytinlik Lava) are interbedded with siliciclastic sedimentary rocks of the Irmakyanı Formation (see Fig. 7b).

The sandstone turbidite succession as a whole passes upwards into an interval of cleaved, dark grey to black, locally reddish brown shales, several tens of metres thick. The shales are interbedded with subordinate, thin-bedded (c. 5 cm) siltstone turbidites, mainly composed of fine quartz and sericite (Yag˘cılar-Tarakc¸ılar; Fig. 14b). Associated

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intervals of non-bioturbated dark, organic-rich shales are locally interbedded with siltstone turbidites. The shales pass upwards into c. 5 m of reddish (chemically oxidized), to greenish (chemically reduced) siliceous mudstones and ribbon cherts, rich in radiolarians of Early? Jurassic age (Adamia et al. 1995). In one sample, K. Tekin (pers. comm. 2002) identified Hexasaturnalis tetraspinus (Yao) (Late Toarcian-Middle Bathonian) and Archaeohagiastrum sp. cf. T. longipes BAUMGARTNER (Late Toarcian– Early Callovian), suggesting a Late Toarcian –Middle Bathonian age. In another sample, he identified Archaeohagiastrum longipes BAUMGARTNER (Late Toarcian –Early Callovian), Triactoma jakobsae CARTER (late Middle Toarcian –Late Bajocian), Hexasaturnalis sp., Hagiastrum sp. cf. H. egregium RUST (late Middle Toarcian –early Late Aalenian), suggesting a Late Toarcian –early Late Aalenian age. In addition, a Bajocian ammonite has been reported from a nearby loose block (Adamia et al. 1995). The succession continues upwards into mainly volcaniclastic sandstone turbidites, siltstones and shales, several hundred metres thick (e.g. in the Batmankaya area; Fig. 14c). Individual beds (,5 cm thick) can be traced laterally for hundreds of metres with little change in thickness. Coarser interbeds predominate higher in the succession. The sandstones in the upper part of the succession are mainly volcaniclastic litharenites with irregularly shaped basalt grains, ranging from vesicular volcanic glass of basic to silicic composition, to flow-banded plagioclase-phyric basalt. Other components are pyroxene crystals (reworked phenocrysts), variably altered volcanic lithoclasts and glassy basalt grains that occasionally retain a bubble-wall eruptive texture. There is then a gradation to predominantly volcaniclastic conglomerates that form laterally continuous depositional units, up to 5 m thick. Angular, to sub-angular, clasts (15 cm) are set in a volcanogenic matrix. Some individual clasts exhibit chilled margins. Massive basalt (c. 100 m thick) appears above this with occasional thin pillow lava flows. This is followed by volcaniclastic breccias and volcaniclastic conglomerates, in beds up to 5 m thick, with several thin flows of pillow lavas. In a second area, near Madenler (Fig. 3), c. 8 km to the NE of C¸ag˘layan, the base of the succession comprises c. 300 m of terrigenous, locally graded, sandstones and siltstones. Volcaniclastic sediments higher in the succession include possible psepherite. Typical volcaniclastic sandstones contain poorly sorted, sub-rounded, to sub-angular clasts of basalt, andesite and silicic igneous rocks. The interbedded volcaniclastic sandstones are mainly composed of basalt, dolerite, silicic extrusive rocks,

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altered plagioclase, chlorite, hyaloclastite and large quartz crystals (phenocrysts). Lava and tuff (c. 60 m) pass upwards into bioclastic limestone/tuff alternations (c. 20 m). The Irmakyanı Formation as a whole is cut by isolated dykes, ranging from basalt-dolerite, to intermediate-silicic compositions. Stratiform basic igneous sheets (,1 m thick) also occur locally, especially within the finer-grained facies low in the succession (Fig. 15). These igneous rocks are believed to be bedding-parallel sills, although an origin as interbedded thin lava flows cannot be ruled out because no clear chilled margins, top or bottom, were observed. Lower-Middle Jurassic platform succession. Further south within the Lower Slice Complex (Fig. 6), the Karadag˘ metamorphic unit is unconformably overlain by a contrasting Lower to Middle Jurassic unit named the Kınalıc¸am Formation. The succession (Fig. 7d) differs strongly from the coeval basinal formation (Irmakyanı Formation), discussed above, and also differs from similar-aged successions elsewhere in the Pontides (e.g. Koc¸yig˘it & Altıner, 2002; see below). The type succession is near Kınalıc¸am village (Fig. 4). This unit is generally equivalent to the Gu¨nyayla Sub-zone of Konak & Hakyemez (1996, 2001). Similar to the basinal succession (Irmakyanı Formation), the platform-related succession (Kınalıc¸am Formation) is divisible into a lower, mainly terrigenous to calcareous part, and an upper, mainly volcanogenic part. Unlike the basinal succession, however, it is generally uncleaved and only locally folded. The Karadag˘ Metamorphics (Fig. 7d) are unconformably overlain by conglomerates that infill depressions in the unconformity surface. Individual hollows are up to tens of metres wide and several metres deep. Clasts infilling the hollows are mainly well-rounded cobbles and pebbles of quartz, micaschist and granite. The coarse clastic sediments grade upwards, over several metres, into siltstones, mudstones and impure fossiliferous limestones, tens of metres thick. The sandstones are mainly medium bedded, laterally continuous over tens of metres, and exhibit planar and parallel lamination. En echelon lenses of limestone are up to several metres thick and tens of metres in lateral extent. Fossiliferous pink, to red, micritic limestones contain orthocones (e.g. Orthoceras sp.), crinoids (e.g. Pentacrinus sp.), bivalves and brachiopods (e.g. rhychonellids). Attempts to date this assemblage have so far been inconclusive. However, an Early Jurassic (Pliensbachian– Toarcian) age is probable because the pink pelagic limestone, known as Ammonitico Rosso, is similar to facies exposed further east in the

¨ zer 1984) and throughBayburt-Gu¨mu¨s¸hane area (O out the Pontides generally (Koc¸yig˘it & Altıner 2002). Above the interval containing Ammonitico Rosso, the sequence grades transitionally (over several metres) into dark shales, siltstones and finegrained sandstones, several hundred metres thick, as exposed on the upper southern and eastern slopes of Karadag˘ (Fig. 4). Sandstone beds, up to tens of centimetres thick, exhibit parallel lamination, and show sharp upper and lower contacts within finer grained facies. Interbedded mudstones are calcareous, bioturbated and commonly contain bioclastic debris, including brachiopods (e.g. rhynchonellids), crinoids and rare ammonites. In a second fault-bounded outcrop c. 25 km further NE (near Sakut; Fig. 4) the Kınalıc¸am Formation comprises shales interbedded with relatively fine-grained, well-sorted, normal-graded sandstones that we interpret as turbidites. Litharenites predominate with angular to sub-angular grains of quartz, plagioclase and muscovite. Flute marks (16 measurements) in this outcrop indicate NNE to SSW sediment transport. Trace fossils locally occur on the soles of sandstone beds. The mainly terrigenous sequences (e.g. Karadag˘ and Uc¸ar) are directly overlain by basic extrusive rocks (Kınalıc¸am Volcanics; Fig. 7d). For example, on the flanks of Karadag˘ (Fig. 4), a mainly pillow basalt interval, up to 200 m thick, can be traced c. 20 km to the NE before thinning and disappearing, while another pillow lava unit (c. 4 km long  250 m thick) appears further NE (near Sakut). In general, individual lava flows are up to several tens of metres thick, with small volumes of associated volcaniclastic sediments and interpillow hyaloclastite. In all exposures (e.g. on the slopes of Karadag˘), the pillow lavas pass upwards into laterally variable, coarse volcaniclastic sediments, ranging from thick-bedded volcaniclastic breccias with angular extrusive rock clasts (up to tens of centimetres in size to thinner-bedded, finergrained volcaniclastic facies rich in hyaloclastite. Occasional pink dacitic dykes (,3 m wide) cut the succession in several areas. In addition, there are occasional dome-shaped silicic intrusions, up to 500 m across. These post-date the weak deformation fabric of the Kınalıc¸am Formation, where present.

Geochemistry of Jurassic igneous rocks Basalts from the Lower-Middle Jurassic intervals of the East Pontide Autochthon and from the Lower Slice Complex (Irmakyanı and Kınalıc¸am Formations) were plotted on MORB-normalized spider diagrams. The basalts from the autochthon show relative enrichment in LILE (large ion

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Fig. 16. MORB-normalized spider plots of Lower Jurassic basalts from the East Pontide Autochthon (near I˙spir), showing an ‘enriched’ pattern. See text for explanation.

lithophile elements) v. HFS (high field strength elements) with no evidence of a subduction influence (Fig. 16). The basalts from the lower, terrigenous part of the Irmakyanı Formation (Zeytinlik Lava; Fig. 17a) show moderately ‘enriched’ patterns of high field strength (HFSE) elements relative to MORB. In contrast, the basalts from higher in the succession (Irmakyanı Lava; Figs 7b & 17b) show near-MORB patterns, with a marked negative Nb anomaly i.e. indicative of a subduction influence. Basalts from the upper, volcanogenic part of the Kınalıc¸am Formation (Fig. 17c) also show a subduction influence.

Upper Slice Complex Middle Jurassic? magmatic arc-type unit A regionally important magmatic unit, up to c. 1000 m thick, known as the Olur Magmatics, is extensively exposed in the upper part of the Upper Slice Complex (Thrust Sheet 3; Fig. 6). Thrust sheets 1 and 2 of the Upper Slice Complex (Fig. 18a, b) are relatively thin and discontinuous such that in places the Jurassic rocks directly overlie in the Lower Slice Complex. For example, an assemblage of basic, intermediate and silicic

extrusive rocks, and also pyroclastic and epiclastic sediments was mapped near ˙Is¸han (Fig. 4). A thicker sequence, exposed further east (Oltuc¸ayı unit of Yılmaz 1985; Bozkus¸ & Yılmaz 1983), includes a silicic intrusion, several kilometres wide (Olur Dacite of Bozkus¸ & Yılmaz 1983), bounded by swarms of silicic dykes. Further SW, a large sill pinches and swells laterally, reaching a maximum thickness of 50 m in the Olur area. The age of this magmatic unit is currently constrained only by a Late Oxfordian age of fossils in the overlying cover sequence (see below). However, a Middle Jurassic age is generally assumed (e.g. Konak & Hakyemez 1996, 2001). Where studied in detail (e.g. near I˙s¸han and Uzundere; Figs 3 & 4), the Jurassic magmatic unit, c. 500 m thick, is represented at the base by massive, basic- to intermediate-composition lavas that are up to tens of metres thick and continuous laterally for hundreds of metres. The lavas are overlain by massive, to weakly stratified, lava breccias, in unit up to tens of metres thick. Amygdaloidal pillow lavas occur at three main levels in the lava breccia, together with small volumes of interflow hyaloclastite. A middle flow, exposed near Tas¸lıko¨y, exhibits a massive interior of plagioclasephyric lava. The unit as a whole is cut by silicic

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Fig. 17. MORB-normalized spider plots of Lower-Middle Jurassic basalt from the Lower Slice Complex. (a) Basalts from the lower, siliciclastic part of the Irmakyanı Formation (Zeytinlik Lava). (b) Basalts from the upper part of the sequence (Irmakyanı Lava). Samples in (a) are ‘enriched’. (c) Basalts from the Kınalıc¸am Formation, further south (see Fig. 7d). These show a near-MORB trend but with a strong negative Nb anomaly, indicative of a subduction influence. See text for explanation.

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Fig. 18. Local cross-sections of the area studied in detail during this work, near the Tortum River (Fig. 4), showing the relationships between The Lower Slice Complex and the Upper Slice Complex. (a) The Lower Slice Complex, overlain by the Upper Slice Complex, which includes correlatives of Thrust sheets 1 and 3. (b) The Lower Slice Complex, overlain by the Upper Slice Complex made up of Thrust sheets 1, 2 and 3. See text for explanation.

sheets, up to 500 m wide in outcrop and dykes, up to c. 1 m thick. The Jurassic magmatic unit is also exposed east of the Tortum River as a NE –SWtrending anticlinal structure, c. 35 km long by c. 15 km wide. In addition, correlative magmatic rocks occur in the uppermost structural unit of the Olur-Tortum Zone (see below). The basalts from the Jurassic Olur magmatic unit (Figs 19a & bi– bii) are highly ‘depleted’ with very low concentrations of Ti, Y and Zr and U-shaped patterns, similar to some high-magnesian andesites (i.e. boninites) (e.g. Flower & Dilek 2003). On the ORG-normalized discrimination diagram the granitic intrusions in the Olur magmatic unit plot in the volcanic arc field (not shown).

Upper Jurassic – Upper Cretaceous sequences During Late Jurassic to Early Cenozoic time three contrasting areas of deposition and magmatic

activity existed, namely the East Pontide Autochthon; within two units of the Lower Slice Complex, and with in all of the units of the Upper Slice Complex (Fig. 6).

East Pontide Autochthon Within the East Pontide Autochthon the Lower – Middle Jurassic succession (Fig. 7a) passes upwards without a perceptible break into 150– 300 m of bioclastic limestones (Berdiga Formation; MTA 1998). These carbonate rocks contain bivalve fragments, benthic foraminifera, echinoderm plates and coral fragments, mixed with detrital grains. The clastic grains include unstrained quartz phenocrysts, ferromagnesian minerals, altered plagioclase, biotite, muscovite, altered extrusive igneous rocks and quartzite. Foraminifera include Trocholina alpina, T. elongata, Hedbergella sp., Ticinella sp., Orbitolina sp., Textularia sp., Grobigerinelloidees sp., Miliolinidae, Lagenidae, Valvulinidae and

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Fig. 19. Field relations and geochemistry of the pre-Upper Jurassic Olur magmatic unit (Upper Slice Complex). (a) Sketch of part of the succession that was sampled for geochemical anlaysis (from near Tas¸lıkoy; Fig. 3). (bi) MORB-normalized spidergram. (bii) Chondrite-normalized REE plot. The MORB-normalized spider plots (bi) of the pillow basalts show a strongly depleted, U-shaped pattern, characteristic of boninite series volcanic rocks. See text for explanation.

Ophtalmidae of Late Jurassic –Cenomanian age (MTA 1998). The overlying succession (c. 400 m thick C ¸ atak Formation; MTA 1998; Fig. 7a) begins with shales, followed by reddish pelagic limestones with planktic foraminifera including Globotruncana lapparenti, G. linneiana, G. Marginotruncana, Globigerinella sp., Globegerinollides sp., Gumbelina and Hedbergella sp., of Turonian–Santonian age (MTA 1998). The succession continues with felsic lavas and pyroclastic rocks (100 –500 m thick) of Santonian age. Above come basic volcanics and volcaniclastic sediments (c. 100 –600 m thick) of Campanian –Maastrichtian age. A several kilometre thick succession of Santonian–Campanian age is also exposed c. 100 km east of the area studied during this work (Hopa-Borc¸ka area; Fig. 3) (Konak & Hakyemez 1996). This is dominated by volcaniclastic gravity flows with subordinate basaltic, to andesitic, lavas. The succession continues upwards into Late Campanian–Early Maastrichtian turbiditic sediments, Middle Maastrichtian –Paleocene argillaceous limestones and calciturbidites, and ends with

Eocene intermediate-silicic volcanic rocks and volcaniclastic sediments, cut by granitic intrusives (Konak & Hakyemez 1996; MTA 1998).

Upper Jurassic– Upper Cretaceous of the Lower Slice Complex Upper Jurassic successions are exposed above the Jurassic basinal succession in the north (Irmakyanı Formation; Fig. 20a) and above the platform-related succession further south (Kınalıc¸am Formation; Fig. 20b). A Late Jurassic age is inferred from the position of these sediments between Lower-Middle Jurassic sediments unconformably below and calcareous sediments above that are dated as Kimmeridgian (Konak & Hakyemez 1996, 2001). Cover of the Jurassic basinal succession. Where ¨ g˘dem (Fig. 4), the basinal sucwell exposed, near O cession (Irmakyanı Formation) is unconformably overlain by a deltaic unit of sandstones, conglomerates, mudstones with thin layers of coal (Fig. 20a). Interbedded conglomerates and sandstones (50 – 60 m thick) are overlain by cross-bedded sandstones

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Fig. 20. Field relations. (a) Upper Jurassic cover of the Lower-Middle Jurassic Irmakyanı Formation showing an unconformable relationship; (b) Upper Jurassic cover of the Lower-Middle Jurassic Kınalıc¸am Formation showing an unconformity (see Fig. 7b, d).

with lenses of bioclastic limestone. Pebbly sandstones just above the unconformity contain mixtures of plutonic, metamorphic, volcanic and neritic detritus. Many grains are well rounded, but poorly sorted. Lithoclasts are mainly fresh basalt, altered basalt, hyaloclastite, altered diabase, quartzose sandstone, silicic lava (with plagioclase phenocrysts), rhyolite (with quartz phenocrysts) and quartzite (with fused quartz grains). Shell fragments, echinoderm plates and microbial carbonate are also present, together with large quartz grains. The succession above is an overall finingupward sequence (c. 1000 m thick) of coarse,

red channelized conglomerates, sandstones and laminated siltstones, interbedded with lava flows (individually ,10 m thick). The conglomerates are clast supported (clasts ,30 cm in size), but without observable clast imbrication. Some conglomerate beds show reverse grading. Higher parts of the succession include thin micritic limestones, caliche nodules, lava flows and cross-cutting dykes. Limestone clasts include Early Jurassic fossils (Konak & Hakyemez 1996). Towards the top of the succession, medium- to thick-bedded limestones contain coral and other fossils of Barremian– Aptian age (Konak & Hakyemez 1996).

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In a second area, c. 25 km to the NE (near Madenler; Fig. 3) the contact with the Jurassic basinal succession (Irmakyanı Formation) is faulted. Elsewhere; Konak & Hakyemez (1996, 2001) report a correlative volcanogenic sequence that is unconformably overlain by red conglomerate, with lenses of Oxfordian aged coralline limestones. This conglomerate is thinner and generally finer grained than the one in our study area. This succession (Fig. 7b) continues upwards into dark-coloured shallow-water limestones of Valanginian age. Unconformably above come Upper Cretaceous (Late Santonian–Early Campanian) volcanogenic facies. Late Campanian–Middle Paleocene shallow-water limestones occur transgressively above this and finally a siltstone, marl, sandstone and redeposited limestone sequence of Late Paleocene age (Konak & Hakyemez 1996, 2001). In a third area, c. 10 km to the NE (near Zeytinlik village; Fig. 3), the basinal succession (Irmakyanı Formation) is unconformably overlain by red, nonmarine conglomerates with local intercalations of coralline limestone of Kimmeridgian age. These sediments are overlain by thick Upper Cretaceous volcanogenic, volcaniclastic and pyroclastic sediments, indicating that a major non-sequence exists. Cover of the Jurassic platformal succession. The sequence, as exposed near Kınalıc¸am village (Fig. 4), begins with grey mudstones and occasional lenticular sandstones with scattered polished volcanic clasts, up to 120 cm long (Fig. 20b). Conglomerate-sandstone alternations (c. 30 m) follow, then siltstone and sandstone (c. 70 m). There is next a sharp transition to well-bedded, matrix-supported bioclastic conglomerates (i.e. debris flow deposits), in units up to 3 m thick. Clasts, up to tens of centimetres in size, are set in a white chalky micritic matrix. Higher levels include internally disrupted slump horizons (up to 4.5 m thick). Facies equivalents are exposed on strike c. 10 km to the NW, on the northwestern flanks of Karadag˘ (Fig. 4). There, siltstones and well-sorted, lenticular, cross-bedded, yellow sandstones are interbedded with local intercalations of coal, including seams up to 1.3 m thick. Poorly exposed shallow-water limestones, red sandstone and conglomerate occur above this. The succession is cut by andesitic dykes and sills, each of which is up to several metres thick. (e.g. at Muzrat Yayla; Fig. 4). Elsewhere, non-marine conglomerates are reported to be interbedded with lenticular shallowwater coral-bearing carbonates of Late JurassicEarly Cretaceous age (Konak & Hakyemez 1996). The sequence passes upwards, above an angular discordance, into Upper Cretaceous shallow-marine clastic sediments with lenses of reef limestone.

Above come shallow-water carbonates of Coniacian (?) –Santonian age. There is then a sequence of Upper Cretaceous volcanogenic sediments, culminating in hemipelagic carbonates with sandstone and pyroclastic interbeds towards the top (Konak & Hakyemez 2001).

Upper Jurassic– Paleocene units: Upper Slice Complex Sedimentary and volcanic rocks of Late Jurassic – Paleocene are exposed in up to five thrust sheets that constitute the Upper Slice Complex (Figs 6, 21 & 22). Thrust Sheet 1. The lowest unit is exposed in the hanging wall of Thrust Sheet 1 and forms a c. 2 kilometre-wide outcrop on both sides of the Tortum River (Fig. 18a, b). A generalized log is shown in Figure 21a. Above a basal thrust, the succession locally begins with sandstones, crinoidal limestones and breccia-conglomerate (near Morkaya; Fig. 18a). Above come poorly exposed soft-weathering volcaniclastic sandstones, siltstones and mudstones with pelagic carbonate intercalations. A succession exposed 6 km on strike to the NE (Fig. 18b) begins with medium, to thick-bedded bioclastic limestones with shell fragments, microbial carbonate, echinoderm plates and large foraminifera (miliolinids). This is followed by silicic lavas, tens of metres thick (Fig. 21a), then quartzose sandstones with well-rounded quartz grains and rubbly-bedded debris flows set in a matrix of pink pelagic carbonates and volcaniclastic sandstones. Pelagic limestones containing Globotruncana sp. are interbedded with volcaniclastic sandstones, c. 100 m thick, terminated upwards by a thrust (Fig. 21a). Thrust Sheet 2. The overall succession (c. Konakyeri Formation of Akdeniz et al. 1994) begins with lithic sandstones, shales and marls and passes upwards into nodular silty bioclastic limestones with interbeds of reworked, matrix-supported debris flows (Fig. 21b). Upwards, the debris flows increase in relative abundances and thickness, whereas limestones gradually disappear. The uppermost debris flows (c. 80 m), individually up to c. 20 m thick, contain limestone blocks up to 2 m across, interbedded with micaceous sandstone. The limestones contain Foraminifera of Late JurassicBerriasian age (MTA 1998). There is then an upward transition to brecciated and sheared, rubbly limestone debris flows, forming repeated units up to 70 m thick, each traceable laterally hundreds of metres, with clasts up to 10 cm in size. Reddish limestone conglomerates crop out above this. A local angular discordance (possible

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Fig. 21. Generalized logs of Upper Jurassic–Lower Cretaceous sequences of the thrust sheets making up the Upper Slice Complex. See text for explanation and literature sources.

unconformity) is followed by a fining-upward sequence of graded debris-flow conglomerates, with mainly basaltic clasts. Above, come siltstones, with nodules of red caliche and normal-graded conglomerates with scoured bases. Sandstones are commonly cross-laminated, with rip-up clasts, and are interbedded with caliche-rich mudstones, which then pass into red siltstone/mudstone/fossiliferous limestone alterations. Above come massive greengrey calcareous mudstones and sandstones with conglomeratic limestone. Brownish, medium- to thick-bedded sandstones follow, rich in carbonate or quartzitic grains. Angular to sub-rounded, basement-derived granitic clasts then appear within fining-upward breccia intercalations (Fig. 21b). Higher in the overall succession (Fig. 21b) there is a switch to mainly volcanogenic turbidites and debris flows (c. 110 m thick). The sandstones are mainly well bedded and graded, with quartz, plagioclase, biotite, quartzite and large muscovite grains, set in siliceous mudstone. Above come pale, white, fissile tuff, grey muddy limestone (6 cm), volcanogenic debris flows (2 m), pebbly limestone with pink detrital clasts, and then wellbedded pelagic limestones. A calcareous interval

above this includes shell fragments, echinoderms, microbial carbonate and pyrite framboids. The clastic components are mainly chlorite, muscovite, biotite and rare silicic volcanic and muscoviteschist. The succession continues upwards through alternations of pink pelagic limestones, pale fissile tuff, tuffaceous conglomerates, volcaniclastic sandstones and mega-breccia (38 m thick), entirely composed of limestone clasts. Thrust Sheet 3. Further south, near Tortum Lake, an unconformable relationship (Figs 21c & 22a) is exposed between the Olur magmatic unit and its Upper Jurassic sedimentary cover. The Olur magmatic unit is unconformably overlain by lava breccias and conglomerates and then by well-bedded volcaniclastic sandstones and siltstones, followed in turn by thin rhyolite flows. Higher levels of the sequence, well exposed south of ˙Is¸han (Fig. 4) begin with columnar-jointed andesite, followed by .300 m of volcaniclastic sandstone turbidites, volcanogenic mudstones and debris flows (individually up to 9 m thick) (Fig. 21c). The sandstones typically contain variable mixtures of sub-rounded quartz grains, basic volcanics, exsolved feldspar,

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Fig. 22. Field relations. (a) Mid-Jurassic Olur magmatic unit (Thrust Sheet 3), unconformably overlain by its Upper Jurassic– Lower Cretaceous sedimentary cover, and then by Thrust Sheet 4. (b) Detail of volcaniclastic facies in Thrust Sheet 3. (c) Lavas and volcaniclastics cut by a gently inclined dacitic intrusion in Thrust Sheet 3.

perthite, altered plagioclase, silicic volcanics, peletal micrite, pisoliths, echinoderm plates and shell fragments. The succession then changes abruptly upwards to quartzitic sandstone turbidites with common plant debris and occasional granitic clasts, together with limestone and basaltic pillow lava intercalations (c. 20 m). These lavas plot in the within-plate basalt field on the Zr v. Zr/Y diagram (Fig. 23a). They show a humped pattern, typical of the within plate basalts, on the MORBnormalized spidergram (Fig. 23b). Sandstones above the basalt contain well-rounded plutonic quartz grains, plagioclase crystals, neritic limestone, vesicular basalt, glassy basalt, quartzite and granite. Further east, near Olur (Fig. 3), according to Yılmaz (1985), the Upper Jurassic succession, begins with Tithonian-Berriasian Calpionellid limestones (Yes¸ilbag˘lar Limestones c. 850 m thick; Bozkus¸ 1992). Biostratigraphical studies of sections exposed along the western banks of the Olur River (c. 2 km SE of Olur) demonstrate the presence of four Calpionellid biozones ranging from earliest Upper Tithonian to Late Berriasian in age, although pelagic limestones may continue into the Hauterivian (Tunc¸ 1992). In this area, we observed laterally continuous, repeated units of well-bedded white pelagic limestones of Tithonian– Berriasian age with occasional sandstone turbidite beds, up to several metres thick. The turbidites include clasts

of coral and redeposited limestone, together with abundant bioclastic debris. Some of the limestone clasts contain echinoderm plates, calcite-replaced sponge spicules and both thick-and thin-walled (Posidonia-type) bivalve shell fragments, microbial carbonate pellets, Calpionellids and calcite-replaced radiolarians. The limestones are abruptly overlain by alternations of dark green sandstone turbidites, siltstones, claystones, marl and pelagic limestones, all of Hauterivian–Campanian age. The succession includes 30– 40 m of black fissile calcareous mudstones, as also exposed near Tortum Lake (e.g. near C ¸ ag˘layan; Fig. 4). The section (Fig. 21c) continues with a prominent interval (c. 100 m thick) of volcaniclastic turbidites, mudstones and debris flows, with clasts up to 1 m in size. A Hauterivian –Campanian age is indicated by the presence of planktic foraminifera, including Globotrunca (e.g. G. stuartifomis; G. subspinusa; G. linneina) (c. Karmaso¨r Formation; Bozkus¸ 1992). Sandstone-limestone alternations follow, with a prominent horizon of black chert (c. 1 m thick), black mudstones, pelagic limestones and sandstone turbidites. Similar facies are also well exposed along the road bordering Tortum Lake. Unconformably above come alternations of sandstone, sandy and muddy limestones of Turonian– Late Campanian age (Karagu¨ney Formation;

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Fig. 23. Geochemistry of the Late Jurassic pillow lava in Thrust Sheet 3. (a) Zr v. Zr/Y diagram (Pearce & Norry 1979). (b) MORB-normalized spidergram. Normalizing values are from Pearce (1983). See text for explanation.

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Bozkus¸ 1992), and then tuffaceous sediments of inferred Late Campanian–Early Maastrichtian age (Killis Tepe Formation). The highest part of the Late Cretaceous succession comprises pelagic limestones, fine-grained calciturbidites, marl and mudstone (Akbayır Formation; Bozkus¸ 1992), dated as Upper Campanian-Maastrichtian using planktic foraminifera, including Globotruncana fornicata, G. bulloides and G. ventricosa (Bozkus¸ 1992). An Upper Paleocene (Thanetian) sequence follows, including shallow-water limestones, limestone conglomerate, redeposited carbonates, marls and siltstones (Konak & Hakyemez 1996). Elsewhere, Bozkus¸ (1992) reports an unconformity, followed by conglomerate, tuffaceous sediments, and finally Lower-Middle Eocene (YpresianLutetian) clastic facies. However, the Lower Cenozoic facies were not studied during this work. Thrust sheets 4 and 5. Several additional, structurally higher, thrust sheets are reported in the literature to the south of our main study area (Fig. 21d). Thrust sheet 4 (Aksu subzone of Konak & Hakyemez 1996) begins with Upper Jurassic sandstones, limestones and marl, overlain by pelagic carbonates with ammonites of Tithonian– Aptian age. Above, there are pelagic limestones, sandstones, siltstones and marl of Aptian –Maastrichtian age (Bozkus¸ & Yılmaz 1993; Konak & Hakyemez 1996). These sediments are then overlain, with a faulted contact, by Upper Cretaceous pillow lavas and massive lavas, cut by dykes. Non-marine, coarse clastic sediments of Lower –Middle Eocene age occur unconformably above this, cut by dacitic intrusive rocks. Thrust sheet 5 (C ¸ ardaklı subzone of Konak & Hakyemez 1996) begins with basaltic-andesitic volcanics and pyroclastic sediments of possibly Middle Jurassic age, unconformably overlain by Upper Jurassic –Lower Cretaceous mixed carbonate, clastic and volcaniclastic sequence, similar to the succession in the underlying thrust sheet. Above this, poorly exposed volcanogenic turbidites pass into Upper Cretaceous hemipelagic carbonates. Middle Eocene shallow-marine clastic sediments occur unconformably above this. Thrust sheets 4 and 5 are likely to be local imbricates of essentially the same regional unit.

I˙zmir-Ankara-Erzincan suture zone In general, ophiolitic rocks, ophiolitic melange and occasional blueschists are described as being overlain by latest Cretaceous (Campanian) clastic sediments that then pass upwards into deeper-water facies. This was followed by a regression and deposition of Early-Middle Eocene shallow-water clastics. The entire assemblage is reported to

be unconformably overlain by Upper Eocene (Bartonian –Priabonian) shallow-water clastic sediments (Bozkus¸ & Yılmaz 1993; Konak & Hakyemez 2001). Additional ophiolitic rocks are exposed c. 50 km further east (near S¸enkaya; Fig. 2) and large bodies of ophiolitic peridotite also occur further west between Erzurum and Erzincan (Fig. 3). In the west (near the Bayburt-Erzurum road) the oldest sediments overlying emplaced ophiolitic rocks are reported to be Maastrichtian neritic limestones (Kapıkaya Formation; Okay et al. 1997). More commonly the oldest cover sediments (Sipiko¨r Formation) are dated as Paleocene –Eocene (Rice et al. 2006, 2009). Ophiolites and related units apparently were initially emplaced northwards during latest Cretaceous time (Okay et al. 1997), although there is little reported evidence of this in the vicinity of our main study area. This was followed by pervasive northward emplacement, together with continental margin units during Middle Eocene time. Hakyemez & Konak (2001) comment that much of the MidEocene (Lutetian) is absent from the cover of Upper Slice Complex (Olur-Tortum zone). In contrast, only a small unconformity seems to exist between the Late Paleocene (Thanetian) and the Early Eocene (Late Ypresian) within the cover of Lower Slice Complex (Artvin-Yusufeli Zone). Two basins apparently co-existed during this time, separated by a NE–SW-trending, subaerially exposed high (Hakyemez & Konak 2001).

Regional comparisons Comparable units are exposed the length of the Pontides (.1000 km), in the Crimea to the north, and in the Caucasus (Figs 1 & 2).

Eastern Pontides The closest comparisons can be made with the more westerly parts of the eastern Pontides (Bergougnan 1975, 1987; Korkmaz et al. 1995; Yılmaz et al. 1997). The metamorphic basement in the Arvin area (Fig. 6) can be correlated with other outcrops in the eastern Pontides (Fig. 24), notably the Pulur Massif (Okay 1996; Okay et al. 1997; Topuz et al. 2004, 2007). These units include medium, to high-grade, feldspathic gneiss and metabasites. The high-grade basement of the Pulur Massif (Cenvi Group) is overthrust by lower-grade metamorphic rocks (Hasso Group; Okay 1984), dated as Permian (Topuz et al. 2004). Lower-grade rocks were thrust northwards over high-grade basement prior to the Jurassic (Okay 1996). Radiometric dating of the Pulur Massif has yielded Early

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Fig. 24. Tectono-stratigraphic diagram showing the sedimentary successions of the East Pontide Autochthon and the overlying thrust sheets in the well-documented Gu¨mu¨shane-Bayburt area to the west of the Artvin area discussed in this paper. See text for sources.

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Carboniferous ages (Topuz et al. 2004). Small cross-cutting bimodal intrusive bodies are locally dated as Early Jurassic in the Pulur Massif (Topuz et al. 2004). In addition, the Ag˘vanis Massif in the eastern Pontides (Fig. 2) comprises greenschist facies meta-basalts, meta-carbonates and metacherts, cut by tonalitic intrusions (Okay 1984). Large outcrops of Upper Palaeozoic basement rocks are also widely exposed further west, including the Tokat Massif and the Devrekani metamorphic unit in the central Pontides (Figs 1 & 2; e.g. Yılmaz et al. 1997). The Lower-Middle Jurassic sedimentary successions in the Artvin area (Fig. 2) can be compared with successions of the same age further west in ¨ zer the Gu¨mu¨s¸hane, Bayburt and Erzincan areas (O 1984). The sequences exposed in the East Pontide Autochthon in the north and in several thrust sheets towards the ˙Izmir-Ankara-Erzincan suture zone further south are particularly informative (Fig. 24). Additional relatively autochthonous units are exposed near the Black Sea coast, south of Trabzon (Fig. 2). In the far north, within the East Pontide Autochton, Jurassic rocks are exposed in several fault-bound inliers of the Cretaceous eastern Pontide arc (S¸en 2007). The lowest exposed horizons in one unit (I˙kisu; Fig. 2) are thick pillow lavas, overlain by tuffs and massive basalts, followed by porphyritic andesites, all of inferred Lower-Middle Jurassic age. The succession grades up into Upper Jurassic shallow-water carbonates. Two other exposures in the same area are mainly massive lava, interbedded with tuffaceous sediment and overlain by shallow-water carbonates. An additional Upper Jurassic–Lower Cretaceous sequence shows an upward gradation to calciturbidites, breccia with basaltic-dacitic blocks (20 m), limestone and siliciclastic sediments (S¸en 2007). Palaeontological evidence indicates foundering of a regional carbonate platform during Albian-Aptian time. The volcanic rocks are mostly basaltic andesites, minor basalts, andesites and dacites. MORB-normalized, immobile element patterns (including REE) indicate a subduction influence. Sr and Nd isotopic ratios of two samples from this area and a further four from further south (see below) also indicate a subduction influence. S¸en (2007) concluded that the Jurassic basalts of the eastern Pontides were influenced by contemporaneous northward subduction. Sequences further south are exposed in a more southerly part of the East Pontide Autochthon, and in three overlying thrust sheets (Fig. 24). These become structurally higher southwards towards the ˙Izmir-Ankara-Erzincan suture zone. The East Pontide Autochthon (Cebre Relative Autochthon of Okay & Leven 1996) exposes an

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important mixed sandstone, siltstone, shale, limestone sequence c. 1000 m thick (C¸atalc¸es¸me Formation), but with no exposed base (Okay & Leven 1996). Fine-grained interbeds contain foraminifera of Late Carboniferous (Late Kasimovian–Early Gzelian age), to possibly Early Permian age (Hardis¸i Formation) (Okay & Leven 1996). Plant fossils are interpreted to have a Eurasian affinity. The East Pontide Autochthon is also exposed near Gu¨mu¨shane (Korkmaz & Baki 1984; Yılmaz & Kandemir 2006; Fig. 2). The Hercynian basement there (Gu¨mu¨shane-Bayburt Metamorphics) shows evidence of extensional faulting, including neptunian dykes filled with Lower Jurassic sediments. Overlying, laterally variable siliciclastic conglomerates, sandstones and organic-rich sediments (coal) pass upwards into mixed volcaniclasticsiliciclastic turbidites, up to .2000 m thick (S¸enko¨y Formation). The turbidites include tuffaceous sediments, pyroclastic rocks, pillow basalt and andesite. Overlying pink Ammonitico Rosso contains ammonites, belemnites, crinoids, brachiopods, bivalves, gastropods and large foraminifera, of inferred Early Jurassic age (Yılmaz & Kandemir 2006). Overlying shallow-marine carbonates of Callovian?–Cenomanian age (Berdiga Formation; c. 300 –500 m thick) begin with conglomerates and breccias that were eroded from an underlying shallow-water carbonate sequence, and also from an underlying Lower-Middle Jurassic volcanicsiliciclastic succession. This evidence implies the existence of an unconformity, comparable with the one that was identified in the Artvin area (Fig. 6) Neptunian dykes that again occur locally are infilled with Upper Cretaceous sediment. Overlying calciturbidites with chert nodules contain benthic foraminifera, ostracods and algae. Upper Cretaceous pelagic carbonates and marls grade upwards into Campanian-Maastrichtian siliciclastic sandstone turbidites, interbedded with hemipelagic marls (Kermutdere Formation). Sandstone-filled neptunian dykes again characterise this interval. The sandstones were mainly derived from the granitic basement of the eastern Pontide volcanic arc and its underlying Mesozoic sequence. Unconformably above come Eocene shallow-marine volcaniclastic and Nummulitic calcareous sediments (Alibaba Formation). A comparable succession is exposed further ¨ zer 1984), where crystalsoutheast, near Bayburt (O line basement (Gu¨mu¨shane-Bayburt Metamorphics) is unconformably overlain by c. 600 m of alternating marl, sandstone, red Ammonitico Rosso, basalt, tuff, agglomerate and lava, collectively dated as Lower Pliensbachian –Aalenian. The succession passes upwards into Bajocian-Oxfordian? dolomitic and siliceous platform carbonates, in turn overlain by medium-bedded limestones

(calciturbidites?), thin- to medium-bedded sandstones and marls of Early Cretaceous age. The succession is then unconformably overlain by polymict conglomerates, followed by sandstones and tuffaceous sediments of Early-Middle Eocene (Ypresian¨ zer 1984). Lutetian age) (O The first of the overlying thrust sheets, exposed c. 50 km south of Gu¨mu¨s¸hane (‘Hamurkesen Thrust Sheet’; Fig. 24) begins with another outcrop of Upper Carboniferous sediments (C ¸ atalc¸es¸me Formation), which is unconformably overlain by c. 1000 m of volcaniclastic sediments of Pliensbachian-Kimmeridgian age (Kelkit or ‘Hamurkesen’ Formation). These are, in turn, overlain by oolitic carbonates and siliceous biomicrites of Early Cretaceous (Berriasian) age (Korkmaz & Baki 1984; Okay et al. 1997). The thrust sheet above this (As¸utka Thrust Sheet; Fig. 24) begins with metamorphic basement (Pulur metamorphic complex), unconformably overlying by a Lower Jurassic succession (Tanyolu 1988). This begins with basal conglomerate and quartzitic, to feldspathic, sandstone (c. 100 m thick), tentatively dated as Early Lias using Foraminifera (e.g. Trocolina multispire, Involutina sp., Endothyranella sp.). A mixed volcanic-sedimentary succession (c. 1500 m thick) follows, made up of sandstone, siltstone, sparitic limestone, tuff, agglomerate, basalt-diabase, andesite and dacite. This is overlain by late Upper Jurassic– Berriasian sediments, beginning with oolitic limestones, and passing into siliceous micritic limestone, dated by foraminifera and calpionellids. Upper Cretaceous ophiolitic melange occurs above this, in turn unconformably overlain by Upper Paleocene –Lower Eocene clastic sediments. Finally, the highest thrust sheet (I˙malıdag˘ and C¸imendag˘ Thrust Sheets; Fig. 24) is located closest to the I˙zmir-AnkaraErzincan suture zone and is characterized by relatively distal Lower Cretaceous pelagic carbonates.

Central and western Pontides We now summarize the successions in the central and western Pontides to provide a more regional perspective. Koc¸yig˘it & Altıner (2002) focused on northwestern Anatolia (notably the Bayırko¨y Formation) and inferred a similar overall tectonostratigraphic development based on a correlation of exposures in 14 areas extending over .1000 km east –west. Specific information on the central Pontides is given by Bergougnan (1975), Okay and S¸ahintu¨rk (1997); Rojay (1995), Rojay & Altıner (1997) and Yılmaz (2006). In general, basal conglomerates rest with an angular unconformity, or non-conformity, on crystalline basement, where exposed. Unsorted, to poorly sorted, polymict fan-type conglomerates

LATE PALAEOZOIC–EARLY CENOZOIC TECTONIC DEVELOPMENT

contain pebbles, to cobbles, of granite, quartz, Permian–Triassic crystalline limestone, schist, phyllite, spilitic basalt, serpentinite, meta-ultramafic rocks, chert and radiolarite, set in a litharenite matrix. Overlying sandstones are thick bedded or massive, lenticular, locally cross-bedded and show evidence of pebble imbrication, graded bedding, growth faults and plant debris. Stratigraphically higher, in different areas, there are medium- to thick-bedded shallow-marine sandstone, thin-bedded, or laminated yellow siltstone, shale, marl, ‘turbiditic’ sandstone, red mudstone, and marl with volcanogenic intercalations. Ammonitico Rosso includes ammonites, brachiopods, bivalves, gastropods, belemnites, crinoids and foraminifera, dated as Late Hettangian– Pliensbachian in age. Successions generally continue upwards into continental, to shallow-marine limestones, 100 –700 m thick (e.g. Bilecik area). Higher levels of the succession are associated with volcanogenic units that are generally more abundant in the east. These include alternations of volcanogenic sandstone, tuffs, volcanic breccia, pillow basalt, andesite, lava, shale, mudstone, sandstone turbidites and pelagic carbonate. Wood is locally abundant, and diabase sills are rarely present. Volcanogenic facies are locally interbedded with Ammonitico Rosso, indicating a syn- or prePliensbachian, to Callovian age. In several areas, the volcanogenic rocks are overlain by variably thick (40–500 m) successions of gypsum, shale, marl, cross-bedded sandstone and coal. The coal is locally dated as Bajocian– Bathonian, but so far only in the eastern Pontides (e.g. Bayburt and Kelkit; Fig. 2). The mainly terrigenous sediments pass upwards into oolitic limestone in all areas, except Artvin, where there is conglomerate (c. 1000 m thick), as described earlier in the paper. The oolitic limestones mark the base of a thick carbonate platform succession (100 –1500 m) of Bajocian-Barremian age. A very thin condensed sequence (1–2 m thick) includes ammonite-rich stromatolitic and pelletal, condensed packstone facies. These sediments initially accumulated in an open-marine, relatively shallow-water setting. This was followed by contrasting platformal and basinal sequences in different areas. The platformal units (1000–1100 m thick), of Kimmeridgian –Aptian age, exhibit both upward-fining (transgressive), and upward-coarsening (regressive) successions in different areas. A prominent palaeo-high was developed near Ankara (Aktas¸-I˙neko¨y palaeo-high). Elsewhere, basinal successions (up to c. 1.6 km thick) comprise thinning- and fining-upwards sequences of sandstone and shale turbidites, breccias, marl and porcellaneous micritic limestones. Thick- and thin-bedded carbonate turbidites (up to

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1000 m thick), of Oxfordian–Berriasian age, form thinning-and fining-upward sequences of mixed shallow-water terrigenous and carbonate sediments. The most distal successions include large amounts of pelagic carbonate and relatively thin-bedded turbidites. Both the proximal and more distal sequences, of Tithonian –Early Valanginian age, are characterized by unsorted breccias and chaotic megabreccias, up to 1000 m thick. Megablocks, up to tens of metres in size, occur especially in proximal settings. The blocks are mainly shallow-water platform carbonates, whereas the host sediments include volcaniclastic and terrigenous sediments. The successions in different areas are characterised by stratigraphic gaps that are well dated in the west. In many areas Upper Jurassic facies (Bathonian or Callovian) rest conformably on Lower Jurassic (Pliensbachian) facies. Also, the Early Cretaceous (Valanginian– Aptian) interval is missing in many areas. Koc¸yig˘it & Altıner (2002) explain the stratigraphic breaks, major facies changes and redeposition events as the result of repeated phases of normal faulting related to the development of an extensional rift basin during Hettangian –Aptian time. Continental fans were deposited by debris flows and braided rivers during initial rifting. Various Ammonitico Rosso-type facies accumulated on submerged fault-controlled highs. Associated volcanism was extension related in this interpretation. Extension also influenced the development of the platforms and basins during the Middle-Late Jurassic (Bathonian–Callovian), followed by ‘post-rift subsidence’ (Kimmeridgian– Aptian), according to Koc¸yig˘it & Altıner (2002).

Ku¨re Complex and related units: Central Pontides We next summarize the regional occurrences of ‘ophiolitic’ units to allow comparison with the Artvin area. Yılmaz & S¸engo¨r (1985) compared an ‘Artvin ophiolite’ in the eastern Pontides with a ‘Ku¨re ophiolite’ in the Central Pontides (Fig. 1). In the Central Pontides, ophiolitic rocks form part of a c. 20 km thick imbricate thrust stack, dominated by siliciclastic turbidites (c. Akgo¨l flysch) that was termed the Ku¨re Complex (after the town of Ku¨re; Fig. 2) by Ustao¨mer & Robertson (1994). Partial ophiolitic sequences, up to 400 m thick, occur within dismembered thrust sheets and blocks. The most complete sequences within individual thrust sheets comprise serpentinized peridotite at the base, overlain by layered cumulate gabbro, isotropic gabbro (cut by diabase dykes), sheeted dykes, then pillowed or massive lavas and fragmental

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extrusive rocks. The extrusive rocks are associated with economic massive sulphide ore bodies near Ku¨re town (Fig. 1). Locally, the extrusive sequence is depositionally overlain by dark, terrigenous and organic-rich shales (e.g. near Ku¨re opencast). These sediments pass depositionally upwards into sandstones turbidites, terrigenous siltstones and shales (up to many hundreds of metres thick). Immobile element major- and trace element geochemistry has demonstrated the presence of MOR-type, to subduction-influenced ophiolitic basalts (Ustao¨mer & Robertson 1999). A subduction-influenced setting for the Ku¨re ophiolite is also supported by the chemical compositions of pyroxene phenocrysts and chrome spinel grains within ultramafic rocks (Ustao¨mer & Robertson 1994, 1999). The sandstone turbidites were mainly derived from plutonic igneous and metamorphic sources. The lower contacts of the ophiolitic rocks are commonly intensely sheared black shale in which C/S fabrics, small-scale thrust duplexes and asymmetrical folds mainly indicate northward tectonic transport (Ustao¨mer & Robertson 1994). The dismembered Ku¨re ophiolite and its sedimentary cover remain poorly dated. Upper Palaeozoic palynomorphs within black shales were probably reworked from coal-bearing sequences of the Palaeozoic of Istanbul (Istanbul Fragment) or the Scythian Platform to the north (Fig. 1). The sedimentary cover of the Ku¨re ophiolite has been inferred to be of Middle– Late Triassic age, based on the occurrence of Hallstatt-type Ammonitico Rosso (Anisian-Norian) in the north (Kozur et al. 2000). However, our recent work (unpublished data) shows that the dated Triassic facies occur in detached blocks and debris flows, and so their relationship (if any) to the typical siliciclastic Ku¨re turbidites is unclear. Near the Black Sea coast the Ku¨re Complex includes localised small blocks of Upper Carboniferous to Middle Permian carbonate rocks that are possibly equivalents of ‘olistoliths’ that occur within turbiditic sedimentary rocks in the Crimea to the north (Taurian series) (Kozur et al. 2000). Near the Black Sea coast the Ku¨re Complex is unconformably overlain by an Upper Cretaceous volcano-sedimentary succession (Karadag˘ Unit; Tu¨ysu¨z 1990). The likely age of the Ku¨re ophiolite is, therefore, Triassic, and its conformable sedimentary cover, Early-Middle Jurassic. In the south, the Ku¨re Complex is in tectonic contact with a crystalline basement unit known as the Devrekani Metamorphic Unit, from which Hercynian K– Ar radiometric ages have been obtained (Ustao¨mer & Robertson 1997; Nzegge et al. 2006). The Complex is composed of highgrade gneisses and amphibolites, overlain by metamorphosed carbonates. It is also intruded by Upper Jurassic granites and unconformably overlain by

Upper Jurassic–Lower Cretaceous cover sediments (Tu¨ysu¨z 1990). The Devrekani Metamorphic Unit is bounded to the south by high-angle, to low-angle, tectonic contacts with a thick volcanogenic unit, known as the C ¸ angaldag˘ Complex (Fig. 2); this comprises thrust-imbricated greenschist facies meta-igneous rocks, c. 10 km-thick. A basement of sheeted dykes and basic extrusive rocks, interpreted as the upper levels of an ophiolite, is depositionally overlain by thick basic, to evolved, volcanic rocks and volcaniclastic sedimentary rocks (Ustao¨mer & Robertson 1997). Granites intrude the C ¸ algaldag˘ Complex in the north. The inferred ophiolitic basement includes basaltic rocks of tholeiitic, calc-alkaline and boninitic, compositions, consistent with an island arc, or possibly a fore-arc setting. The thick overlying extrusive sequence has a subduction-influenced composition and is interpreted as part of an oceanic volcanic arc (Ustao¨mer & Robertson 1999). The highest levels of the volcaniclastic sequence are black phyllites, deepwater limestones and debris flows. These are interpreted to record the collapse of a volcanic arc, possibly related to the opening of the Ku¨re basin to the north (Ustao¨mer & Robertson 1997). The C¸algaldag˘ Complex is assumed to be Triassic in age (pre-Late Jurassic), although it is not specifically dated. The Ku¨re Complex clearly includes the remnants of a complete ophiolite of pre-Late Jurassic age. This ophiolite is subduction influenced rather than of mid-ocean ridge, as suggested by S¸engo¨r & Yılmaz (1985); also, there is no evidence of associated ‘Palaeotethyan’ oceanic units (e.g. radiolarian cherts). Ustao¨mer & Robertson (1993) inferred that the Ku¨re ophiolite formed in back-arc basin behind a northward-dipping subduction zone, a view since supported by many authors (e.g. Adamia et al. 2001; Stampfli & Borel 2002; Kazmin & Tikhonova 2006). The C¸angaldag˘ Complex is interpreted as a forearc-type ophiolite, overlain by an oceanic magmatic arc; this was not an ‘ordered’ Palaeotethyan ophiolite, as suggested by S¸engo¨r & Yılmaz (1985).

Crimea The northern extension of the Pontides split off to form the Crimea as a result of opening of the western and eastern Black Sea basins during Late Cretaceous –Early Cenozoic time (e.g. Go¨ru¨r 1988; Okay et al. 1994). To the south of the Scythian platform (Fig. 2) the Triassic to Mid-Jurassic time interval is dominated by Middle Triassic siliciclastic turbidites, including quartzo-feldspathic sandstones and shales, known as the Taurian Series (e.g. Mazarovich & Mileev 1989). These sediments are petrographically similar to the Ku¨re

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turbidites. However, in contrast, the Triassic turbidites in the Crimea are unconformably overlain, by Lower Jurassic (Sinemurian-Pliensbachian) arkosic sediments, beginning with a basal conglomerate (Voznesensky et al. 2001). The Norian –Early Jurassic interval contains intermediate-silicic composition volcanics and tuffs. Toarcian –Middle Jurassic facies include cross-bedded sandstone and coal, with conglomerate, sandstone and shale stratigraphically above. There is then a calc-alkaline volcanogenic sequence including basic, to silicic volcanics, volcaniclastic sediments and tuff of Late Bajocian age. Bathonian-Callovian sediments, including limestone, accumulated in a tectonically active basin (Permyakov & Maidanovich 1984; Mazarovich & Mileev 1989). Many authors (e.g. Adamia et al. 2001; Kazmin & Tikhonova 2006) envisage the Taurian Series as part of a Triassic marginal basin that opened between the Scythian platform to the north and a continental fragment (i.e. ‘Rhodope-Pontide fragment’) to the south. However, the basement of any such basin is unknown because no oceanic crust or ophiolites are exposed in the Crimea.

Caucasus From south to north the main units of the Caucasus are the Lesser Caucasus, the Transcauscus (northern and southern parts), the Great Caucasus (Southern Slope Basin and Main Range), the Fore-range and the Bechasin Zone (Khain 1975; Adamia et al. 1977, 1981, 1995, 2003; Zakariadze et al. 1983, 1988; Bogdanovski et al. 1992; Gamkrelidze 1997; Saintot et al. 2006; Fig. 1). In general, the Artvin area can be correlated with the northern part of the Lesser Caucasus and the southern part of the Transcaucasus (Adamia et al. 2001; Yılmaz et al. 2001). The ophiolitic rocks of Sevan–Akeran part of the Lesser Caucasus (Zakaraidze et al. 1983) and related units are similar to those of the I˙zmirAnkara-Erzincan suture zone further west (Fig. 1). In Armenia, the ophiolites (Sevan –Akera and Vedi ophiolites) have recently been shown to be of Lower-Middle Jurassic age based on 40Ar/39Ar dating of amphibole in gabbro (Rolland et al. 2005, 2007), somewhat older than the Late Jurassic ages previously reported by Zakariadze et al. (1983, 1988) and Bogdanovski et al. (1992). Also, Upper Triassic radiolarians were identified in associated ophiolitic melange (Knipper 1980). These ophiolites are probably older than those of the ˙Izmir-Ankara-Erzincan suture zone in the eastern and central Pontides that are believed to be Late Cretaceous (Rice et al. 2006, 2009). The southern Transcaucasus has a frontal part known as the Baiburt-Karabakh unit, which is

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broadly equivalent to the Lower and Upper Slice Complexes of this study (Fig. 6). The Artvin-Bolnisi unit to the north is equivalent to the Eastern Pontide Autochthon. The Baiburt-Karabakh unit in the south includes imbricated and isoclinally folded Upper Palaeozoic? –Lower Mesozoic black slate-chertdiabase, slivers of plagiogranites and plagiogranite gneisses (locally dated as Hercynian), blocks of gabbro-diabase, pillow basalt, radiolarite, serpentinitic melange, mafic and ultramafic cumulates, and Jurassic–Lower Cretaceous volcanic-turbiditic sequences. All of these units are unconformably overlain by an Upper Cretaceous arc-type volcanicsedimentary sequence, similar to the eastern Pontides (Adamia et al. 1977, 1981, 1995). The Artvin-Bolnisi unit in the north includes a basement of granitic rocks, schists, phyllite, migmatites and gabbro amphibolites. This is intruded by an older assemblage of quartz diorites, plagiogranites and tonalites, dated as 310–330 Ma by the K –Ar and Rb–Sr methods; younger microcline granites are also reported. The basement is overlain by Late Palaeozoic? Mesozoic and Early Cenozoic shallow-marine sediments and subduction-related volcanics. Early Cenozoic sequences exhibit hiatuses and are strongly faulted (Adamia et al. 2001). High-grade metamorphic and ophiolitic rocks also occur further north, notably in the Dzirula massif and associated units (Knipper & Khain 1980). Migmatites in the Dzirula massif have been dated as Ordovician by the U –Pb method, and crosscutting granitic rocks have yielded Variscan ages using the K– Ar method. Quartz diorites in the comparable Loki massif were dated at 370 Ma by the U –Pb method. The metamorphic basement has in places been assigned a Neoproterozoic Pan-African age, suggesting that some units originally rifted from Gondwana (Zakariadze et al. 2007). The metamorphic basement units of the Dzirula, Loki and Kharami massifs, specifically are overlain by Upper Carboniferous-Lower Permian ‘molasse’, as seen locally in the eastern Pontides. In addition, volcanic rocks erupted subaerially during the Late Triassic (Svanidz et al. 2000). Yılmaz & S¸engo¨r (1985) compared the Ku¨re ophiolite with the Dzirula massif (ChorchanaUtslevi zone; Fig. 1), from which various members of a metamorphosed ophiolite were reported, including serpentinite, gabbro, gabbro-amphibolite, amphibolite and diabase, together with possible remnants of a magmatic arc. However, the internal structure and stratigraphy of the Dzirula massif remain poorly documented and the sedimentary cover appears to be significantly older (Late Palaeozoic) than that of the Ku¨re Complex (Late Jurassic). The northern part of the Great Caucasus, known as the Main Range Zone (Fig. 1), comprises

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Precambrian–Palaeozoic metamorphic rocks intruded by granitic rocks of at least partially Carboniferous age, overlain by coal-bearing Upper Carboniferous clastic sedimentary rocks. The southern part of the Great Caucasus, known as the Southern Slope Basin, is a thick sequence of Devonian-Triassic, mainly terrigenous, basinal clastic sediments with subordinate limestone and silicic volcanogenic rocks in the lower part. This sequence, known as the Dizi Series, is widely correlated with the Taurian Series in the Crimea (e.g. Adamia et al. 2001). The Dizi Series culminates in an unconformity, above which there is a several kilometres thick sequence of relatively fine-grained meta-clastic sediments of Early Jurassic age, including MOR-type meta-basaltic intercalations. Above come Middle Jurassic (Bajocian) silicic lavas, Upper Jurassic limestones, and then a Paleocene–Eocene cover. The southern slope basin was intruded by granitic rocks prior to the Late Jurassic, comparable with the Central Pontides and also the Istranca zone (Fig. 1) in the western Black Sea region. In the north, the Fore-range Zone is widely interpreted as a south-facing passive margin (e.g. Adamia et al. 2001), together with over-riding oceanic units, accretionary melange and magmatic arc rocks that were tectonically emplaced then covered by Upper Carboniferous clastic sediments. The Caucasus are generally considered to reflect the evolution of the southerly, active continental margin of Eurasia during Late Palaeozoic –Early Mesozoic time, associated with the genesis and emplacement of one, or several subduction-related marginal basins (Adamia et al. 1977, 1981, 1987, 2001; Knipper & Khain 1980). Recently, Kazmin & Tikhonova (2006) have suggested that the Transcaucasus and a more northerly, small crustal fragment, the Shatsky Rise, rifted from Eurasia associated with opening of the Tauric (Ku¨re) basin during the Early Jurassic (Toarcian) to form an inferred marginal oceanic basin. In this interpretation, the Transcaucasus and Shatsky Rise microcontinents collided with the Eurasian continent during the Late Bathonian, while Neotethys remained open to the south. However, the Ku¨re ophiolite probably predates the Early Jurassic, and the Jurassic Artvin Basin is more likely to be laterally equivalent to the inferred Jurassic marginal basin in the Transcaucasus.

Discussion of depositional and tectonic settings There are several similarities and differences between the Artvin area of the eastern Pontides and the elsewhere in the circum-Black Sea region,

including more westerly parts of the Pontides, the Crimea and the Caucasus. Similarities: First, the Late Palaeozoic metamorphic basement of the Artvin area (Fig. 2; area in box) is similar to areas further west (e.g. Pulur Massif; Gu¨mu¨shane-Bayburt Metamorphics; Fig. 2) that include small crosscutting bimodal intrusions of Early Jurassic age. The granite-gneiss basement units in the Artvin area are interpreted as windows into a once extensive Upper Palaeozoic continental margin arc, dominated by arc-type granitic rocks. At least locally, the arc was uplifted by Late Carboniferous time, shedding clastic detritus, as seen in the Gu¨mu¨s¸hane area. Second, the basal conglomerates of Early Jurassic age, including debris flows and alluvial fan-type facies, are similar throughout the Pontides. Third, many areas, including Arvin, have Ammonitico Rosso of Early Jurassic (Hettangian –Pliensbachian) age, where dated. Fourth, in most areas the basal sediments pass into siliciclastic sandstones, with variable volcanogenic intercalations. Fifth, the Upper Jurassic–Lower Cretaceous time interval is characterised by a similar range of platform, slope and basinal facies throughout the Pontides. On the other hand, there are a number of differences between the Artvin area and the more westerly Pontide areas. First, only in the Artvin area is the metamorphic basement known to be cut by swarms of basic, to felsitic, dykes, suggesting a larger amount of crustal stretching and related magmatism than elsewhere. Second, Lower-Middle Jurassic successions, especially in the central and western Pontides, are mainly shallow-marine, to non-marine. Turbidites have been mentioned (Koc¸yig˘it & Altıner 2002), but these could instead be graded storm deposits because the associated facies appear to be shallow marine. Fourth, the preUpper Jurassic basinal succession of the Lower Slice Complex in the Artvin areas (Irmakyanı Formation; Figs 6 & 7) includes radiolarian mudstones and radiolarian cherts that indicate the existence of a deep-water, open-marine basin, whereas facies further west in the Pontides appear to be of more shallow-water type. Fifth, despite localized hiatuses, the Upper Jurassic–Lower Cretaceous sequences in the western and central Pontides appear to be conformable, whereas in the Arvin area the base of the Upper Jurassic succession (Oxfordian) is a strong angular unconformity. Sixth, the Upper Jurassic (Oxfordian–Callovian) succession in most areas is shallow marine, whereas in the Arvin area the cover successions in the Lower Slice Complex (Irmakyanı and Kınalıc¸am Formations) begin with thick non-marine conglomerates (up to 1000 m thick) that were eroded from locally underlying units, including metamorphic basement.

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The sequence in the Arvin area is generally similar to the Taurian Series of the Crimea, in that siliciclastic turbidites are overlain by Jurassic volcanogenic facies. However, there is no definite evidence in the Arvin areas of Triassic lithologies, unlike the Crimea. The eastern Pontide autochthon, including the Artvin area, can be closely compared with the northern part of the Transcaucasus (Adjara-Trialetzi zone), whereas the overlying slice complexes are similar to the southern Transcaucasus. On the other hand, the Artvin rocks are dissimilar to the ophiolitic rocks of the Northern Transcaucasus (Dzirula massif).

Development of the East Pontide active continental margin We now draw on all of the information and comparisons presented above to outline several important stages in the development of a long-lived active continental margin.

Carboniferous The oldest rocks in the Artvin area are gneisses, amphibolites and schists, exposed beneath various cover units. The basement rocks of the Karadag˘ metamorphic massif (Fig. 6) were intruded by granitic rocks and later foliated together with country rocks. The high-grade metamorphic host rocks of the Demirkent Intrusive Complex (Fig. 6) are locally cut by plagiogranite segregations, dated as Early Carboniferous. These Early Carboniferous granitic basement rocks can be generally correlated with similar metamorphic rocks cut by granitic intrusive rocks elsewhere in the Pontides, the Transcaucasus and the Greater Caucasus. All of the granitic rocks are interpreted to relate to northward subduction of Tethys and the development of an Andean-type active margin along the southern margin of Eurasia (Adamia et al. 1977, 2001; Ustao¨mer & Robertson 1993; Stampfli & Borel 2002; Okay et al. 2006; Kazmin & Tikhonova 2006). However, the granitic rocks of the East Pontide Autochthon studied during this work do not show the expected subduction influence, showing that further work on the Hercynian basement is still needed.

Triassic The oldest dated, Early Jurassic deep-sea sediments of the Arvin basin are underlain by several hundred metres of undated siliciclastic sediments. A Late Triassic age for these cannot be excluded. However, there is no confirmed record of Triassic rocks beneath the comparable Jurassic successions

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elsewhere in the Pontides. The Arvin successions that are floored by continental basement are dissimilar to the Triassic Karakaya Complex, as seen further south throughout the Pontides (e.g. Okay et al. 1996). The Karakaya Complex is widely interpreted as a subduction complex related to northward subduction of Palaeotethys (Pickett & Robertson 1996, 2004; Okay 2000). Throughout areas stretching from the western Pontides to the Caucasus, latest Triassic time was characterized by compressional deformation related to an ‘Eo-Cimmerian’ orogenic event (Nikishin et al. 1998, 2001; Kazmin & Tihkonova 2006). In the western Pontides, oceanic-derived units of the Karakaya Complex were emplaced northwards onto the Sakarya continent, a unit that has been correlated with the Upper Palaeozoic basement of the Pontides (Okay 2000; Okay et al. 2006). In Iran, a ‘Cimmerian’ (Gondwana-derived) microcontinent is inferred to have rifted from Gondwana in the Permian, drifted northwards and eventually collided with the Eurasian active margin during latest Triassic time (e.g. Rolland et al. 2007). The eastern Pontides may also have experienced Late Triassic emplacement of oceanic or continental margin units but, if so, they were removed (tectonically or by erosion) before the Early Jurassic, or are simply not exposed.

Jurassic We interpret the Jurassic units of the Artvin area as a small, incipient, rifted marginal basin that formed above a subduction zone during Early-Mid Jurassic time (Fig. 25a). The sheeted dykes of the Demirkent Intrusive Complex are interpreted as intrusions near the axis of the supra-subduction zone rift. The dykes are assumed to have fed overlying lava flows, as exposed in the upper parts of the sedimentary successions to the north (Irmakyanı Lava) and south (Kınalıc¸am Lava). During early-stage extension, the crystalline basement rifted creating a subsiding continental, to shallow-marine, basin in which terrigenous sediments accumulated as debris flows, deltaic deposits and then turbidites. Interbedded relatively enriched basalts (Zeytinlik Lava) probably represent lowdegree melts during early-stage extension. In response to continuing extension, organic-rich muds and terrigenous silts and sand turbidites accumulated on the floor of a subsiding deep-water basin. Radiolarian deposition is indicative of nutrient upwelling and relatively high plankton productivity in a relatively deep-water open-marine basin. The metamorphic rocks of the Karadag˘ Metamorphics and their cover of relatively condensed sediments, including Ammonitico Rosso are inferred to have formed a horst (i.e. submerged

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Fig. 25. Alternative regional tectonic models for the tectonic development of the Artvin Basin. (a) Northward subduction triggered development of a supra-subduction zone rift during the Early Jurassic, followed by construction of an incipient continental margin arc, including local boninite-type volcanism during the Middle Jurassic. (b) During the late Middle Jurassic a continental fragment or volcanic edifice possibly collided with the active continental margin triggering compressional deformation and uplift of the Artvin rift basin. (c) After consumption of a hypothetical crustal edifice, subduction resumed during Late Jurassic time, resulting in slab rollback and extensional collapse of the continental margin. (d) With continuing subduction, the eastern Pontide magmatic arc developed and shed volcaniclastic sediments into a forearc basin. Late collision-related events are shown in Fig. 26e, f.

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platform) between grabens to the north and south. In the south, a magmatic arc began to be constructed during the Middle Jurassic (Olur magmatic arc unit), as indicated by the presence of calc-alkaline, intermediate- silicic volcanic rocks. In addition, the local presence of high-magnesian andesites (boninites) is suggestive of a fore-arc setting, based on modern comparisons, for example, with the Mariana-Bonin arc (e.g. Stern & Bloomer 1992). The horst was later covered by volcaniclastic sediments and subduction-related basalts (Kınalıc¸am and Irmakyanı lavas).

Late Jurassic (pre-Oxfordian) Many areas of the southern margin of Eurasia, including the eastern Pontides (Okay et al. 1997; Yılmaz et al. 1997) and the Transcaucasus (Nikishin et al. 1998, 2001; Kazmin & Tichonova 2006) show evidence of Late Jurassic (pre-Oxfordian) ‘Neo-Cimmerian’ compression. We have found good evidence of an important unconformity affecting several of the units in the Artvin area, notably within the Lower Slice Complex and Thrust Sheet 3 of the Upper Slice Complex (Fig. 6). By contrast, the Early-Middle Jurassic, to Eocene, succession of the East Pontide Autochthon to the north is conformable (Fig. 7a). Hiatuses are reported in successions of this age in the Pontides further west (e.g. in the Bayburt-Gu¨mu¨s¸hane area) but without evidence of an Upper Jurassic angular unconformity. In the Artvin area, underlying units were tilted by up to 908, and local recumbent folds are sealed by unconformably overlying clastic sediments. We infer that previously stretched, magmatically active crust of the Artvin basin was compressed, resulting in stratigraphic inversion of extensional faults, and that this was sufficient to cause uplift, tilting and emergence (Fig. 25b). Such inversion could have been restricted to rheologically weak grabens and not have affected the East Pontide Autochthon tens of kilometres to the north, perhaps because this crust was thicker and less easily deformed, or simply further from the locus of compression.

Oxfordian – Berriasian During the Late Jurassic, mixed carbonate, siliciclastic and volcaniclastic sediments accumulated on the East Pontide Autochthon in the north, accompanied by minor volcanism (Fig. 7a). Within the Lower Slice Complex, deltaic facies including coal accumulated initially, followed by thick shallow-marine, deltaic and alluvial sediments (i.e. above the Irmakyanı Formation in the north). These sediments, of Oxfordian –Callovian age, record renewed subsidence, perhaps re-activating

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older faults within the metamorphic basement (Figs 25c & 26a). In contrast, thinner, non-marine to shallow-marine sediments accumulated further south, above the former, Jurassic basement horst and its platformal cover (Kınalıc¸am Formation; Fig. 7d). Shallow-water carbonates of Late Jurassic age also accumulated within the Upper Slice Complex (e.g. Thrust Sheet 3), associated with sparse basaltic volcanism of relatively enriched, non-subduction-influenced type. The Late Jurassic –Early Cretaceous sequence of Thrust Sheet 2 in the Upper Slice Complex documents spectacular collapse, as indicated by redeposited units including multiple debris flows, detached limestone blocks and calciturbidites (Fig. 26b). Pelagic carbonates then accumulated on a deeperwater slope, represented by Thrust Sheet 3 (Fig. 26c). Background pelagic sedimentation was interrupted by gravity deposition of mixed terrigenous, bioclastic and volcaniclastic sediment derived from the north. Sediments accumulated on a relatively fertile deep-water slope, or base of slope setting, where calcareous and siliceous plankton flourished, as reflected in calpionellid and radiolarian abundance and local diagenetic chert formation. Shelf- and upper-slope-derived carbonate was redeposited basinwards as calciturbidites and debris flows, together with terrigenous and basementderived clastic material. Local intraformational angular discordances reflect tiling of the slope.

Late Cretaceous During the Late Cretaceous, the East Pontide Autochthon submerged and was covered by pelagic carbonates. Volcaniclastic sediments, derived from the eastern Pontide arc that had begun to develop further north, followed this. After a hiatus, similar volcanogenic units developed during the Late Cretaceous within each of the tectonic units further south. Thick volcaniclastic sequences, mainly turbidites and debris flows, constructed a thick slope-apron, interbedded with pelagic carbonates, as exposed in Thrust Sheets 3 –5 (Figs 25d & 26d). Local occurrences of black shales could reflect a Late Cretaceous oceanic anoxic event. During the Upper Cretaceous, supra-subductiontype ophiolites are assumed to have been created within the Tethyan ocean to the south, as well exposed in the Erzincan area further west (Rice et al. 2006, 2009). During the latest Cretaceous ophiolites, ophiolitic melange and related blueschists apparently were emplaced northwards onto the distal edge of the Pontide margin, eroded and then transgressed by terrigenous sediments, locally by Maastrichtian time. (Because there is little evidence in our study area, this inferred regional stage is, however, omitted from Fig. 26.)

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Fig. 26. Reconstructions of the Arvin area is a fore-arc setting during Late Jurassic–Late Cretaceous time. (a) Extension-related faulting and volcanism. (b) Extensional collapse. (c) Subsidence and pelagic deposition. (d) E Pontide arc magmatism. (e) Upper Cretaceous (pre- or syn-Maastrichtian) ophiolite and melange emplacement. (f) Telescoping of the entire active margin during Mid-Eocene final closure of Neotethys.

Paleocene –Eocene Clastic sediments accumulated on emplaced ophiolitic and related rocks during the Late Paleocene – Early Eocene. A Lower Cenozoic basin exposed

further SW in the Erzincan area is interpreted as a fore-arc basin that then emerged, signifying the onset of continental collision (Rice et al. 2006, 2009). Any remaining Tethyan oceanic crust to the south subducted followed by collision of the

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Eurasian and Iranian/Tauride continental units. This resulted in regional-scale imbrication and northward thrusting of the east Pontide active margin, extending into the southern Transcaucasus (Fig. 26e). Emplacement is likely to have been oblique (i.e. diachronous), which would explain the high-angle, sheared nature of many tectonic contacts, especially within the Lower Slice Complex. In addition, the suture was modified by later strike– slip, as indicated by fault planes with mainly left-lateral slickenlines throughout the area.

Modern and ancient comparisons The early stages of rifting of the Artvin basin can be compared with the Late Miocene –Recent opening of the SW Pacific Woodlark Basin (Taylor et al. 1999a, b; Robertson et al. 2001). In this case, a Late Miocene forearc rifted during the Pliocene, associated with episodic westward propagation of a spreading axis. Rifting was initially non-volcanic but became magmatically active directly ahead of the propagating spreading tip. Westward rift propagation could similarly explain the westward decrease in the volume of Lower-Middle Jurassic volcanic rocks in the Pontides, although it is not known if the Artvin rift passed eastwards into a spreading axis similar to the Woodlark Basin. Interestingly, Adamia et al. (2001) interpreted the southern and northern Transcaucasus as separate units with a marginal oceanic basin between. The tectonic setting is also generally similar to the Late Miocene–Recent development of the Tyrrhenian Sea, as revealed by land-based study and ocean drilling (Kastens et al. 1988). An older arc-related continental basement, exposed in Sardinia to the NW, rifted and subsided, associated with coarse terrigenous sedimentation (Robertson et al. 1990) and the eruption of ‘enriched’-type volcanics (similar to the Zeytinlik Lava) as marginal seamounts (Beccaluva et al. 1990). The subsiding crust was later covered by terrigenous and hemipelagic sediments, associated with chemically variable basaltic volcanism, ranging from MORB to subduction influenced (Beccaluva et al. 1990). Also, in common with the Artvin Basin, rifting preceded the onset of arc magmatism on the southeastern margin of the basin, as represented by the Aeolian Islands and the adjacent Marsili Basin. The more northerly Vavilov Basin is underlain by exhumed continental crust and sub-continental mantle lithosphere, including serpentinized peridotite (Kastens et al. 1988). By contrast, the incipient Artvin marginal basin is underlain by dyke-intruded metamorphic basement, suggesting a lesser amount of crustal extension.

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The Artvin marginal basin also shows similarities with several tectonically emplaced marginal basins. In northern Greece, northward subduction is inferred to have lead to rifting of continental crust of the Serbo-Macedonian zone to the north to open a marginal basin (Be´bien et al. 1987). Swarms of basaltic dykes cut continental basement rocks (De Wet et al. 1989). However, in contrast to the Artvin marginal basin, extension was much greater there allowing oceanic crust to form within the marginal basin, as now represented by Guevgueli and other ophiolite (e.g. Be´bien et al. 1987; see Robertson 2002). Elsewhere in the eastern Pontides (Erzincan area) ophiolitic rocks of inferred Late Mesozoic age include swarms of diabase dykes cutting host rocks (e.g. marble). These dykes have been interpreted to relate to the opening of an Upper Cretaceous marginal basin along the southern margin of Eurasia (Rice et al. 2006; 2009). Closure of this supra-subduction zone marginal basin led to the northward emplacement of ophiolitic rocks on the distal Pontide margin.

Alternative tectonic models There are four main tectonic models to explain the Lower Jurassic–Upper Cretaceous tectonic setting of the Pontides, and the Artvin area, specifically. In the first model (Fig. 27a), oceanic crust to the north of a Pontide –Transcaucasus crustal unit (i.e. Palaeotethys) subducted southwards during Early-Middle Jurassic, resulting in progressive rifting of the Northern Neotethys (Go¨ru¨r et al. 1983; Bektas¸ et al. 1987; Koc¸yig˘it & Altıner 2002). Final continental break-up occurred during the Late Jurassic, followed by Early Cretaceous passive margin subsidence (Yılmaz et al. 1997). The main problems with this interpretation are that the supposed conjugate margin of the Pontides as a whole is in the Taurides, which shows evidence of much earlier rifting, during Early-Middle Triassic (e.g. Go¨ncu¨og˘lu et al. 2004; Robertson et al. 2004; Mackintosh & Robertson 2009). This was followed during the Jurassic by passive margin subsidence, without further volcanism (Collins & Robertson 1998; Andrew & Robertson 2002; Go¨ncu¨og˘lu et al. 2004, 2006; Robertson et al. 2004; Okay et al. 2006). The Tauride carbonate platform shows little evidence of fault-controlled sedimentation during Lower-Middle Jurassic, as would be expected if it formed the hanging wall of an extensional detachment. Also, deep-water radiolarian sedimentation already existed within the ˙Izmir-Ankara-Erzincan ocean (northern Neotethys) by Late Triassic time (Bragin & Tekin 1996; Go¨ncu¨og˘lu et al. 2006).

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Fig. 27. Alternative plate tectonic reconstruction of the Pontide continental margin. (a) Southward subduction of either a wide ‘Palaeotethys’ or a smaller Ku¨re-Tavric marginal basin (both to the north) leading to the opening a back-arc rift to the south. (b) Northward subduction of Palaeotethys opening a back-arc rift to the north. (c) Formation as a strike–slip controlled basin. b, is preferred (see text for explanation).

In a related model (not shown in Fig. 27) the Lower-Middle Jurassic sedimentary and magmatic rocks of the Pontides –Transcaucasus formed as an extensional rift, or incipient marginal basin, again above a southward-dipping subduction zone, in this case related to southward closure of a Triassic marginal oceanic basin to the north (Ku¨re-Tauric basin). The Ku¨re marginal basin then closed related to southward subduction during Late Middle Jurassic time (Bajocian?). Transmitted stress led to gentle uplift of Pontides causing erosional hiatuses. Compression was more intense in the east, where the rift basin was widest deepest and most magmatically active. Compression resulted in stratigraphic

inversion, uplift and erosion of pre-existing fault blocks. After collision the Pontides began to subside and were covered by shallow-marine carbonates during the Kimmeridgian. Northward subduction was then initiated beneath the amalgamated southern margin of Eurasia during the Late Cretaceous. The main problems with this model are the lack of evidence that a sufficiently wide Ku¨re-Tauric oceanic basin existed to the north in order to fuel Mid-Jurassic arc volcanism. Also, it is surprising that the pre-latest Jurassic compression is not documented within the Pontide autochthon that was located closest to the closing marginal basin in this interpretation.

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In a contrasting model (Fig. 27b), Tethys subducted northwards during Triassic to Late Cretaceous time (Robertson & Dixon 1984; Dercourt et al. 1986, 1993; Adamia et al. 1995, 1997; Golonka 2004; Robertson et al. 2004; Kazmin & Tikhonova 2006). The Mid-Jurassic calc-alkaline magmatism (Olur magmatic arc unit) developed on the oceanward side of an incipient marginal basin that became deeper to the SE (S¸en 2007). Trench roll-forward, or collision of a topographic edifice (seamount, or microcontinent) then caused the pre-latest Jurassic uplift and erosion in the Artvin area (Fig. 25b). Slab roll-back during the Late Jurassic then resulted in extension and collapse of the east Pontide margin, associated with localized intra-plate-type volcanism during latest Jurassic time. The main problems with this interpretation are the absence of any known Jurassic –Lower Cretaceous accretionary prism bordering the Pontides to the south, or in the Lesser Caucasus. Also, the Upper Jurassic basalts of the Artvin region include non-subduction-influenced, within-plate-type basalts, unexpected above a subduction zone. However, a Jurassic accretionary prism might have been removed related to Upper Cretaceous subduction, or Early Cenozoic continental collision. Also, the pre-existing inferred Early-Middle Jurassic subduction zone was possibly disrupted, or detached, during the pre-latest Jurassic inferred regional compression event; this could have allowed nonsubduction influenced within-plate-type volcanics to be sourced from depth. In another contrasting tectonic model (Banks & Robinson 1997; Fig. 27c,d) the Artvin area of the eastern Pontides experienced transtension during the Late Triassic– Early Jurassic, opening a pullapart basin, correlated with the Ku¨re ophiolitic basin and a supposed Early Jurassic ophiolitic basin in the Arvin area. The authors propose that the southern margin of the East European Platform including the North Dobrogea was separated by the Polish Trough, Moesian Platform and the Strandzha by an important left-lateral strike –slip fault zone (Peceneaga-Camena Fault Zone). To the east, the Eurasian margin changed orientation from NW–SE to more east –west, resulting in a switch to transtension in the Pontides. This triggered the development of pull-apart basins in the Pontides including development of supposed Artvin, Ku¨re and C¸angaldag˘ ophiolitic basins. It is, indeed, likely that strike –slip played an important role in the genesis and deformation of basins in the Pontides. However, there are several problems with this particular model. There is no oceanic crust in the Artvin Basin, which is dissimilar to both the Ku¨re supra-subduction zone-type oceanic crust and the C ¸ angaldag˘ arc-type crust. The stratigraphy of

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the Artvin Basin is generally similar to that of the Pontides as a whole suggesting a common tectonic setting, although the Artvin Basin was deeper and more magmatically active. Also, there is no confirmed evidence of the required through-going transform fault in the Pontides (although one might exist), and oblique opening of the Early –Middle Jurassic Artvin Basin is by no means excluded. The available evidence presented in this paper is insufficient to prove any specific tectonic model, although we favour Model 2 (Fig. 27b). In this scenario, the Artvin marginal basin reflects stages in the progressive northward subduction of Tethys beneath Eurasia. The Early Jurassic rifting relates to re-establishment of the subduction zone after the regional emplacement of the Karakaya accretionary complex onto the Eurasian continent during latest Triassic time (seen to the west of our study area). The Middle Jurassic subduction-related magmatism reflects early-stage arc magmatism bordering an incipient marginal basin. Radiometrically dated Jurassic plagiogranites cutting ophiolitic ultramafic rocks in the Ankara Melange of Central Anatolia (Dilek & Thy 2006) could then record northward subduction within Neotethys. Also, Jurassic supra-subduction zone ophiolites are known in the Lesser Caucasus (Rolland et al. 2007). The Late Middle Jurassic (‘Eo-Cimmerian’) event might, therefore, record the collision of a topographic edifice (i.e. seamount or microcontinent) with the Eurasian active margin. Subduction then stepped back resulting in extensional collapse of the entire Pontide continental margin during Late Jurassic-Early Cretaceous time. Northward subduction resulted in Cretaceous accretion of HP/LT rocks, as exposed in the central Pontides (Okay et al. 2006). Continuing northward subduction during the Late Cretaceous constructed the eastern Pontide magmatic arc, followed by regional Late Cretaceous ophiolite emplacement (documented to the west of our study area). Continental collision ensued during Middle Eocene time and the distal edge of the Eurasian active continental margin was detached, imbricated and thrust northwards over the east Pontide platform.

Conclusions (1)

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The Artvin area of far-northeastern Turkey documents stages in the development of an active continental margin and closure of Tethys from Late Palaeozoic to Early Cenozoic time. Metamorphic basement in the Artvin area is believed to reflect Upper Palaeozoic active continental margin-type magmatism throughout the Pontides and Caucasus. However, Upper Palaeozoic magmatic

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rocks studied in the Arvin area lack a chemical subduction influence, suggesting that not all of the granitic rocks simply represent parts of a typical continental margin arc. Metamorphic basement in the Arvin area is locally intruded by swarms of subductioninfluenced basic dykes and associated minor intrusions that are indicative of crustal extension in a supra-subduction zone setting during Early–Middle Jurassic time. Rifting of continental basement created a fault-bounded basin that subsided and infilled, initially by coarse basement-derived terrigenous sediments, followed by deep-sea radiolarian muds, and within plate-type basalts. These lithologies were covered by subduction-influenced volcanics as extension and subsidence continued. Ammonitico Rosso accumulated on a horst (submerged platform) during Early Jurassic time and was later covered by subduction-influenced volcanic rocks. The Artvin area Jurassic rifted basin was uplifted, tilted and eroded associated with a pre-Late Jurassic (‘Neo-Cimmerian’) regional compressional event. This resulted in an important angular unconformity between the underlying Lower-Middle Jurassic units and Upper Jurassic sedimentary and volcanic rocks above. Relatively thin, mixed terrigenous, calcareous and volcanogenic successions continued to form on the East Pontide Autochthon to the north during Late Jurassic –Early Cretaceous time. By contrast, further south, extensional collapse took place and large volumes of volcaniclastic, terrigenous and shallow-water biogenic sediments accumulated on an unstable continental slope to the south. Sparse basalts are of relatively ‘enriched’, non-subduction-influenced type, suggesting a break in steady-state subduction, with possible extrusion through an asthenospheric slab window. Oversteepening of the slope resulting in sliding of exotic blocks, slumping and multiple debris flows. Pelagic carbonates accumulated on the upper slope and were redeposited to the lower-slope basin plain as calciturbidites. The calc-alkaline eastern Pontide arc was constructed during the Late Cretaceous, shedding huge volumes of volcaniclastic sediments over the platform and slope to form a fore-arc apron to the south. Cretaceous northward subduction of Tethys generated supra-subduction zone-type ophiolites and accretionary melange that, based on regional evidence, were first emplaced

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northwards over the edge of the active continental margin Late Cretaceous time. The active margin (fore-arc basin), together with ophiolitic rocks, was thrust dissected and emplaced northwards, onto the east Pontide foreland during Mid-Eocene regional collision of the Pontide (Eurasian) and Gondwana-related continental units. The tectono-stratigraphy of the Artvin area can be correlated with better-known westerly parts of the eastern Pontides (e.g. Gu¨mu¨s¸hane-Bayburt area) and with the southern and northern units of the Transcaucasus to the east. There is no evidence in the Artvin area of any Palaeotethyan (i.e. Triassic or older) ophiolite. Gabbroic rocks, basic dykes and basic extrusives occur in different units of different age; also, ultramafic rocks are absent from the Artvin area. The Artvin area documents the incipient development of a rifted marginal basin above a subduction zone during the LowerMiddle Jurassic. However, subduction was interrupted during the Late Jurassic, before any large continental margin arc or oceanic marginal basin could develop. We favour a tectonic model in which the Artvin Basin developed in response to a stage in northward subduction of Tethys beneath the Eurasian continent. EarlyMiddle Jurassic subduction was possibly interrupted by the collision and subduction of a continental fragment or seamount. Subduction stepped back, collapsing the margin during Late Jurassic–Early Cretaceous time, followed by construction of the eastern Pontide magmatic arc in the Late Cretaceous.

The first author thanks I˙stanbul University Research Fund (Project No: 1510/28082000) and TUBI˙TAK (Project No: YDABCAG-370) for financial support during this work. The second author thanks the Carnegie Trust for the Scottish Universities for financial support to carry out fieldwork. We thank Dr D. James for assistance with chemical analysis that was carried out by X-ray fluorescence at the Grant Institute, University of Edinburgh, UK. Professor D. Whitney, Minnesota University, USA kindly provided some additional chemical analysis. We ¨ zbey, O. C also thank Z. O ¸ ılgın and T. Go¨ru¨m for help with preparing the figures. The manuscript benefited from review by two anonymous reviewers.

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LATE PALAEOZOIC–EARLY CENOZOIC TECTONIC DEVELOPMENT Ustao¨mer, T. 1998. Pre-Jurassic crustal extension, basin formation and associated magmatism at a consuming margin, Yusufeli-Artvin, eastern Pontides, NE Turkey. Third International Turkish Geology Symposium. Abstracts, 251. Ustao¨mer, T. & Robertson, A. H. F. 1993. Late Palaeozoic-Early Mesozoic marginal basins along the active southern continental margin of Eurasia: evidence from the Central Pontides (Turkey) and adjacent regions. Geological Journal, 28, 219–238. Ustao¨mer, T. & Robertson, A. H. F. 1994. Late Palaeozoic marginal basin and subduction-accretion: evidence from the Palaeotethyan Ku¨re Complex, Central Pontides, N Turkey. Journal of the Geological Society, London, 151, 291–306. Ustao¨mer, T. & Robertson, A. H. F. 1997. Tectonic– sedimentary evolution of the north Tethyan margin in the Central Pontides of northern Turkey. In: Robinson, A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. American Association of Petroleum Geologists Memoir, 68, 255– 290. Ustao¨mer, T. & Robertson, A. H. F. 1999. Geochemical models used to test alternative plate tectonic models for pre-Upper Jurassic (Palaeotethyan) units in the Central Pontides, N Turkey. Geological Journal, 34, 25–53. Ustao¨mer, T., Robertson, A. H. F., Gerdes, A. B. & Ustao¨mer, P. A. 2010. LA ICP MS and ion probe U–Pb dating of igneous and metasedimentary units in the NE Pontides, NE Turkey: evidence of PeriGondwananan terrane accretion, Late Palaeozoic magmatism/metamorphism and Early Mesozoic extension along the S Eurasian margin. Geophysical Research Abstracts, 12, EGU2010-5958. Voznesensky, A. I., Volozh, Yu. A., Antipov, M. P., Lipatova, V. V. & Shlezinger, A. Ye. 2001. Mesozoic palaeogeography and sedimentary basins of the Scythian-Turanian plate (Caucasus– Caspian region). In: Derman, A. S., Toksoy, F. & Yılmaz, E. (eds) Proceedings of the 2nd International Symposium on the Petroleum Geology and Hydrocarbon Potential of the Black Sea Area. S¸ile– Istanbul, Turkey, 22– 24th September 1996. Turkish Association of Petroleum Geologist, Special Publication, 4, Ankara, Turkey, 1–17.

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Subductions, obduction and collision in the Lesser Caucasus (Armenia, Azerbaijan, Georgia), new insights ¨ LLER2, T. DANELIAN3, R. MELKONYAN4, M. SOSSON1*, Y. ROLLAND1, C. MU 5 6 S. KEKELIA , S. ADAMIA , V. BABAZADEH7, T. KANGARLI8, A. AVAGYAN4, G. GALOYAN4 & J. MOSAR9 1

Universite´ de Nice Sophia Antipolis, CNRS, OCA-UMR Ge´osciences Azur, 250 rue A. Einstein, Valbonne, France 2

6 bis rue Haute 92500 Rueil Malmaison, France

3

Universite´ de Lille, Laboratoire Ge´osyste`mes, FRE 3298, 59655 Villeneuve d’Ascq cedex, France

4

Institute of Geological Sciences, Armenian Academy of Sciences, 24 Baghramian ave, Yerevan, Armenia

5

Institute of Geology, 1 M. Alexidze Street, 0193, Tbilisi, Georgia

6

Institute of Geophysics, 1 M. Alexidze Street, 0193, Tbilisi, Georgia 7

University of Baku, Baku, Azerbaijan

8

Geological Institute of Azerbaijan, Geotectonics and Geodynamics, H.Javid Av., 29A, Baku AZ1143 Baku Azerbaijan 9

University of Friburg, Earth Sciences, Friburg, Switzerland *Corresponding author (e-mail: [email protected])

Abstract: In the Lesser Caucasus three main domains are distinguished from SW to NE: (1) the autochthonous South Armenian Block (SAB), a Gondwana-derived terrane; (2) the ophiolitic Sevan– Akera suture zone; and (3) the Eurasian plate. Based on our field work, new stratigraphical, petrological, geochemical and geochronological data combined with previous data we present new insights on the subduction, obduction and collision processes recorded in the Lesser Caucasus. Two subductions are clearly identified, one related to the Neotethys subduction beneath the Eurasian margin and one intra-oceanic (SSZ) responsible for the opening of a back-arc basin which corresponds to the ophiolites of the Lesser Caucasus. The obduction occurred during the Late Coniacian to Santonian and is responsible for the widespread ophiolitic nappe outcrop in front of the suture zone. Following the subduction of oceanic lithosphere remnants under Eurasia, the collision of the SAB with Eurasia started during the Paleocene, producing 1) folding of ophiolites, arc and Upper Cretaceous formations (Transcaucasus massif to Karabakh); 2) thrusting toward SW; and 3) a foreland basin in front of the belt. Upper– Middle Eocene series unconformably cover the three domains. From Eocene to Miocene as a result of the Arabian plate collision with the SAB to the South, southward propagation of shortening featured by folding and thrusting occurred all along the belt. These deformations are sealed by a thick sequence of unconformable Miocene to Quaternary clastic and volcanic rocks of debated origin.

The Mesozoic ophiolites of the Lesser Caucasus belong to the Tethyan ophiolitic suture zone (Knipper 1975; Adamia et al. 1981; Zakariadze et al. 1983; Knipper et al. 1986) (Fig. 1). Consequently the Lesser Caucasus mountain belt recorded Mesozoic and Cenozoic geodynamic evolutions related to the closure of the Northern Neotethys ocean (Sengo¨r & Yilmaz 1981; Adamia et al. 1981; Dercourt et al. 1986; Ricou 1994; Nikishin

et al. 1998; Yilmaz et al. 2000; Stampfi et al. 2001; Robertson 2002; Golonka 2004). In this context, the NW –SE trending Sevan– Akera suture zone extends on nearly 400 km, north of the Sevan Lake from Amassia to the Nagorno-Karabagh region. To the SE of this region it is covered by Quaternary deposits along the Araks valley (Fig. 2). The outcrops of this arched zone reach a maximum width of 25 km and are characterized

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 329–352. DOI: 10.1144/SP340.14 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Structural sketch map of the Tauride–Anatolides, Caucausus and Iranian belts. Location of Figure 2 is indicated.

by many discontinuous massifs, exposing ultrabasites, gabbros, basalts and pelagic sedimentary rocks (Knipper 1975; Gevorkyan & Gevorkyan 2003; Aghamalyan 2004; Melikyan 2004) (Fig. 2). The location of this ophiolitic zone between Eurasia and the Daralagez continental terrane or South Armenian Block (SAB; a continental terrane of Gondwanian origin, Knipper & Khain 1980; Monin & Zonenshain 1987) suggests either an allochthonous unit or a real suture zone within the Lesser Caucasus belt. According to Aghamalyan (1996, 2004), the Sevan–Akera ophiolitic zone

corresponds to a suture zone where Eurasian and a microplate plates have joined together. However, other outcrops of ophiolites (the Vedi and Zangezur ophiolites) are known southwestward of the Sevan– Akera zone (Fig. 2). Consequently the ophiolites of the Lesser Caucasus could correspond to several sutures zones (Aslanyan & Satian 1977) or only one reflecting a complex west-Pacific type evolution of the Neotethys Ocean east of the Erzincan suture zone (Eastern Anatolia, Turkey) and west of the Iran ophiolitic ones (Knipper 1975; Adamia et al. 1981; Lordkipanidze et al. 1988; Yilmaz et al. 2000).

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Fig. 2. Structural map of the Lesser Caucasus, modified from Nalivkin (1976) and including our observations. A, Variscan Khrami massif and its Carboniferous volcano-sedimentary cover (Georgia); B, Variscan Loki Massif (Georgia); C, Zayemchay Valley (Azerbaijan); D, Gochkarchay Valley (Azerbaijan); E, Yasamal Valley (Azerbaijan); F, Kurakchay Valley (Azerbaijan), all these locations are discussed in the text.

In the Lesser Caucasus the apparent southwestward obduction of oceanic lithosphere on the South Armenian Block (Knipper 1975; Knipper & Khain 1980; Monin & Zonenshain 1987) occurred

during the Late Cretaceous (Sokolov 1977; Knipper & Khain 1980). It was classically admitted that the obduction was directly related to the collision between the SAB and Eurasia. As a result

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the Eurasian plate overthrusted the ophiolites southwards (Adamia et al. 1977, 1981). As for many examples around the world, in such structural setting, occurrences of ophiolites feature the obduction and collision stages of geodynamic evolution of mountain belts. Where the suture zone units have been preserved from metamorphism due to the collision, they may give significant information on the initial stages as the geodynamic setting of their formation and the subduction. This belt was deformed by superimposed tectonics which makes it difficult to decipher each structure and consequently the geodynamic processes from which they resulted. Accordingly, the main debated subjects on the Lesser Caucasus are: 1) the occurrence or not of three main ophiolitic

zones corresponding for some authors to three main sutures zones; 2) the modalities and timing of collision stages between the SAB and Eurasia, including the verging of continental subduction and age of its triggering; and 3) the origin of the widespread magmatic activity that occurred during Eocene time and then from the Oligocene – Miocene to the Quaternary. New investigations were deemed necessary to solve these questions all related to the geodynamic evolution of the Neotethys Ocean and its closure. According to new field surveys (mapping and acquisition of structural data) and new analytic data (Ar/Ar dating, geochemistry, micropalaeontological dating: Tables 1 and 2) obtained thanks to the MEBE programme support, we discuss

Table 1. Upper Cretaceous Nannofossils dating with (WGS84) GPS locations Location

Vedi Valley

Sample number

Latitude

Longitude

Elevation m

AR S 16

N 39896551

E 44892178

1392

Avazan Valley North Sevan Lake

AR 14 05 AR 17 05 AR 21 05 AR 22 05 AR 94 05 AR 100 05

N 40832581 N 40832581 N 40832921 N 40832921 N 40832461 N 40832461

E 45870898 E 45870898 E 45872668 E 45872668 E 45869376 E 45869376

2291 2291 2459 2459 2242 2242

Vedi Valley

AR 63 05 AR 64 05 AR 65 05 AR 67 05 AR 72 05 AR 73 05

N 39896291 N 39896291 N 39896291 N 39896291 N 39893349 N 39893349

E 44894482 E 44894482 E 44894482 E 44894482 E 45801987 E 45801987

1385 1385 1385 1385 1915 1915

AR 20 05

N 40832529

E 45871949

2371

Avazan valley NE Sevan Lake

Dali Valley North Sevan Lake

AR 39 05

AR 01 05

N 40848414

N 40851028

E 45842111

E 45837118

Age

Corollithion exiguum, Eiffellithus turriseiffeli, Eprolithus floralis, Manivitella Late Cenomanian pemmatoidea, Podorhabdus albianus, Prediscosphaera cretacea, Watznaueria barnesae, W. biporta

Santonian

Watznaueria barnesae, Micula staurophora Eiffellithus turriseffeli, Eiffellithus eximius Eprolithus floralis, Prediscosphaera cretacea Zygodiscus diplogrammus, Marthasterites furcatus Tranolithus orionatus, Reinhardtites anthophorus Quadrum gartneri, Lucianorhabdus cayeuxii

Late Santonian

Watznaueria barnesae, Reinhardtites anthophorus Micula staurophora, Lucianorhabdus cayeuxii Eiffellithus eximius, Prediscosphaera cretacea Lithastrinus grillii, Marthasterites furcatus Quadrum gartneri, Broinsonia parca expansa

2118

2021

Early Campanian

Artanish Valley North Sevan Lake AR 04 05

N 40851456

E 45837415

2092

Nannofossil assemblages

Late Campanian

Gartnerago obliquum, Eiffellithus eximius Eiffellithus turriseiffeli, Micula staurophora Prediscosphaera cretacea, Broinsonia parca Arkhangelskiella cymbiformis, Broinsonia parca, Eiffellithus eximius, Quadrum gothicum, Q. trifidum, Reinhardtites anthophorus

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Table 2. Paleocene–Eocene Nannofossils dating with (WGS84) GPS locations Location

Artanish valley North Sevan Lake

Sample number

Latitude

Longitude

Elevation m

AR 03 05

N 40851343

E 45837745

2048

AR 07 05 AR 08 05

N 40851812 N 40852080

E 45837207 E 45836939

2236 2308

AR 54 05

N 39893931

E 44886781

1127

AR 101 05 N 40800623 AR 107 05 N 39898047

E 44890870 E 44888994

1564 1460

Khosrov valley

here some key points allowing us to clarify the subduction, obduction and collision processes all along the Sevan –Akera suture zone. Moreover we present a new interpretation for the structure of the Lesser Caucasus belt along a NE–SW crosssection.

General structural setting As presented on the structural sketch map (Fig. 2) and the NE– SW trending cross-section (Fig. 3) summarizing previous data (Nalivkin 1976) and our observations, three main lithostructural domains can be distinguished in the Lesser Caucasus, from NE to SW: 1) Eurasia; 2) the Sevan–Akera suture zone; and 3) the South Armenian Block (SAB). This structural framework was previously proposed by Milanovski (1968). The main difference between the Milanovski’s section and ours is the occurrence of ophiolitic nappes (the Vedi ophiolites) in front of the Sevan–Akera ophiolitic zone (Fig. 3). Satian et al. (1979, 2005) consider the Vedi ophiolites as remnants of a small palaeo-ocean. Knipper & Sokolov (1976) and Sokolov (1977) provided evidence for an allochthon position of the ophiolitic unit on the base of structural data. Others differences concern the occurrence of post-Eocene reverse and thrust faults which accommodated shortening in the basement of the SAB and the Sevan–Akera ophiolitic zone (Figs 2 & 3).

Age

Nanofossil assemblages

Coccolithus pelagicus, Sphenolithus radians Sphenolithus moriformis, Dictyococcites dictyodus Late Eocene, Reticulofenestra umbilica, NP 18 Cyclococcolithus formosus Cyclicargolithus floridanus, Chiasmolithus oamaruensis Late –Middle Coccolithus pelagicus, Sphenolithus radians Eocene Sphenolithus moriformis, to Late Dictyococcites dictyodus Eocene Reticulofenestra umbilica, Cyclococcolithus formosus Cyclicargolithus floridanus Coccolithus pelagicus, Late Ericsonia subpertusa Paleocene Chiasmolithus danicus, NP 5 Fasciculithus tympaniformis Cruciplacolithus tenuis Coccolithus pelagicus, Early Ericsonia subpertusa Paleocene Cruciplacolithus tenuis, NP 3 Zygodiscus sigmoides Chiasmolithus danicus

Evidence for subduction processes Subduction along the Eurasian margin. As documented by previous studies performed on the Lesser Caucasus the subduction of the Tethys is evidenced by a thick and mainly calcalkaline volcanogenic and volcanoclastic series dated as Bajocian to Santonian (e.g. Adamia et al. 1981 for a review). At this period of time the northern Lesser Caucasus was characterized by an island arc domain called the Somkheto-Karabakh Island Arc (Knipper 1975; Adamia et al. 1977, 1987; Ricou et al. 1986). The basement formations are quite similar to those known all along the Eurasian margin. Indeed they are exposed in three Variscan metamorphic massifs in Georgia (Transcaucasian massif): the Khrami (A on Fig. 2), Loki (B on Fig. 2) and Dzirula salients (Kazakhashvili 1950; Adamia 1968; Kekelia & Khutsishvili 1980; Adamia et al. 1983, 1987; Adamia 1984; Shengelia et al. 1989; Abezadze et al. 2002). In the Khrami massif the Carboniferous volcanoclastic series unconformably overlie the Variscan metamorphic basement (Adamia 1968, 1984; Adamia et al. 1983). This volcanogenic-sedimentary formation is transgressively overlain by the Liassic, Upper Jurassic– Lower Cretaceous and Cenomanian sediments and volcanogenic series. Upper Palaeozoic metamorphic rocks are also described in the Azerbaijan part of the Northeast Lesser Caucasus flank (Nalivkin

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Fig. 3. Cross-section of the Lesser Caucasus. Location is shown on Figure 2.

1976) (see mark C on Fig. 2). We studied the Mesozoic volcanogenic series in four valleys near Ganca (Azerbaijan) (Fig. 2) from west to east: Zayemchay (C on Fig. 2), Gochkarchay (D on Fig. 2), Yasamal (E on Fig. 2), Kurakchay (F on Fig. 2). These observations allow us to present a synthetic stratigraphic log of the Eurasian margin of the northeastern part of Lesser Caucasus (Fig. 4). Bajocian –Bathonian formations (+2400 m) present a basal conglomerate on the Variscan basement (see mark A on Fig. 2) and upwards volcanoclastic series, sandstones, marl with Ammonites and Cancellophycus, basalts, volcanoclastic turbidites and andesite on top. This series indicates a major magmatic activity during the Middle Jurassic (Maghakyan et al. 1985; Lordkipanidze et al. 1988). The Oxfordian stage is composed of a transgressive sequence with conglomerate, siltstones, mudstones and reef limestones (Fig. 4). The Kimmeridgian features massive intrusions (granodiorites, gabbros, quartz-diorites, plagiogranites) (Melkonian 1976; Lordkipanidze et al. 1988). The gabbroic Dachkasan intrusion is responsible for the metamorphism of limestones into white marbles (Fig. 4). The intrusions caused many ore-deposits: 1) Middle and Upper Jurassic epigenetic copper and barite-polymetallic deposits; 2) Middle Jurassic gold-bearing porphyry copper deposits; 3) Upper Jurassic iron skarn deposits; and 4) Lower Cretaceous porphyry copper deposits (Kekelia et al. 2004; Melkonian & Akopyan 2006; Babazadeh et al. 2007). Coniacian formations unconformably overlie (unconformity U2, mark E on Fig. 2) the Upper Jurassic intrusions and Lower Cretaceous sedimentary rocks that are very much

reduced in thickness in that part of the belt (Fig. 4). The Coniacian stage consists from bottom to top of a red conglomerate, sandstones and reef limestones indicating a shallow water environment at that time. A similar palaeoenvironment is also featured by the Santonian formations. They are characterized by a transgressive sequence (conglomerate, sandstones and limestones) including prismatic basalt flows. In some places, the series contains on its top pillowed and massive basalt lava flows. Consequently the magmatic arc was still active during this period of time along this part of the Eurasian margin. The end of magmatic activity occurred in the Campanian to Maastrichtian stages. The formations are made of thin layered pelagic limestones with some carbonates turbidites. The compiled lithostratigraphic log (Fig. 4) characterizes an arc-type evolution of the active Eurasian margin from Bajocian to Late Jurassic all along the Lesser Caucasus northern flank. In this part of the Eurasian margin some ages of volcanogenic series (Late Cretaceous) are younger than in the northwestern part (Transcaucasian massif, Georgia). The southeastward younging of magmatic ages may result from migration of the magmatic arc during the Late Cretaceous. On this basis we suggest that an oceanic plate subducted northward (if we exclude any major rotation since this time) beneath the Eurasian plate. At least it is well evidenced from Bajocian to Santonian (100 Ma), with younger ages present to the SE. This oceanic plate can be correlated to the northern Neotethys part or to one marginal basin related to an intra-oceanic subduction. In

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Fig. 4. Synthetic lithostratigraphic log of the Eurasian margin in the Lesser Caucasus (Armenia and Azerbaijan).

order to test these hypotheses we present below new evidences for an intra-oceanic subduction in the Neotethys domain during these times.

Evidence for an intra-oceanic subduction According to petrological and geochemical data from volcanites, Lordkipanidze et al. (1988)

suggested a Middle Jurassic Marianas-type subduction in the Lesser Caucasus region. New evidence for an intra-oceanic subduction is described in detail in a companion paper (Rolland et al. 2010), consequently we point out and summarize here the main results. The northwestern part of the Sevan–Akera suture zone (SW of the Stepanavan town, Fig. 2)

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is characterized by tectonic units outcropping within a nappe anticline in which the core is made of blueschist slices (Aghamalyan 1981, 2004; Galoyan et al. 2007; Galoyan 2008; Rolland et al. 2007). The blueschist units protoliths correspond to unroofed sediments along a subduction zone. The HP-LT metamorphism was dated 90–80 Ma by K –Ar methods (Aghamalyan 1998). Further, it was found to be a two-stage evolution using Ar/Ar dating technique (on phengites), with HP-LT stage at 94 –90 Ma followed by a retrograde Green Schist/epidote amphibolite metamorphism during exhumation in the Late Cretaceous (71–74 Ma, Ar/Ar on phengites, Rolland et al. 2007, 2010). The blueschist unit is overthrusted by unmetamorphosed ophiolites, Late Jurassic in age (Danelian et al. 2007). In that area and in the other ophiolitic zones of the Lesser Caucasus (Sevan –Akera, Vedi), this geodynamical setting is also well documented by the magmatic evolution of the ophiolites. Petrologically and geochemically (major, trace, REE and isotopic geochemistry) these ophiolites are characterized by a slight island arc signature (Rolland et al. 2009). Such geochemical features are typical of oceanic crust formed in a back-arc setting (Galoyan 2008; Galoyan et al. 2009). Above the ophiolites, an Upper Cretaceous calc-alkaline volcanic series is evidenced, lying on a sedimentary unconformity marked by Albian conglomerates and limestones (Aghamalyan 1998; Galoyan et al. 2007). The geochemical data, including isotopic Sr –Nd compositions, obtained on these lavas indicate a volcanic arc-type series on top of the ophiolites (Galoyan et al. 2007, 2009; Galoyan 2008; Rolland et al. 2009). Moreover, the structural position of the Stepanavan blueschists unit in the core of a nappe anticline overthrusted by ophiolites and arc series suggests an intra-oceanic subduction context in the stage preceding collision of the SAB with the Eurasian margin. According to these structural, petrological and geochemical results one can suggest that this part of the Neotethys Ocean was, at least from Middle Jurassic to Campanian in age, featured by an intraoceanic subduction to its south and a subduction below the Eurasian continental plate to the north. These new results are significant constrains to the geodynamic processes evolution of the Lesser Caucasus at these periods of time.

Obduction modalities As previously described by other works, several ophiolitic series are found in North and Central Armenia (Fig. 2). They have slightly different structures as a result of their lithologic and structural

settings, but are thought to be remnants of a single obducted sequence (see Rolland et al. 2010). The ophiolitic series of the North Sevan Lake, Amassia and Stepanavan (Fig. 2) (Sevan–Akera ophiolites) (Knipper 1975; Adamia et al. 1977; Adamia et al. 1980; Maghakyan et al. 1985; Adamia et al. 1987) are found along the main suture joining the SAB and the Eurasian margin, while the Vedi ophiolites are an allochthonous nappe thrusted over the SAB (Knipper & Sokolov 1977; Sokolov 1977; Adamia et al. 1981). The Zangezur ophiolites located between the Sevan–Akera suture zone and the Vedi ophiolites correspond to some slivers of ophiolites unroofed along thrust faults during collision (Fig. 2) (Aslanyan & Satian 1977, 1982; Knipper & Khain 1980). One of the main characteristics of these ophiolitic rocks is the very low grade and even, in most cases, the absence of any metamorphism related to subduction/collision stages. The main parts of the ophiolitic units are very well preserved excepted in some shear zones, especially at the base of the obducted sequence, in which some low greenschist facies metamorphic crystallisations occur. The ophiolites record a low metamorphic grade imprint related to heterogeneous hydrothermal alteration dating from oceanic stages (Rolland et al. 2010). Even in the Stepananvan area, the main ophiolitic body is unmetamorphosed. There, its rests on a tectonic me´lange mainly composed of palaeo-accretionary prism blueschist facies rocks (Rolland et al. 2007). The obduction event is well documented in the Vedi area (Southeast of Yerevan, Armenia) (Figs 2 & 5), where the frontal obduction contact is exposed (Figs 6 & 7). In that part of the Lesser Caucasus, ophiolitic rocks are thrusted over platform series of the SAB (Sokolov 1977). In order to describe the obduction modalities it is important to identify at first the nature of the autochthon. Is this autochthon a part of the Eurasian margin or does it correspond to an exotic terrane originating from a southerly position in the Neotethys domain?

The SAB autochthon In the Lesser Caucasus, the SAB (also called Daralagez continental terrane) is mainly exposed in Armenia and corresponds of a Proterozoic metamorphic basement, an incomplete Palaeozoic sedimentary succession, Triassic and some Jurassic formations in Nakhijevan province (Azerbaijan), Upper Cretaceous sedimentary series, Paleocene to Oligocene detrital and volcanogenic rocks, Miocene volcanogenic rocks and Plio-Quaternary volcanites (Nalivkin 1976) (Figs 3 & 5). The Proterozoic basement is made of metamorphic rocks characterized by gneisses, micaschists

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Fig. 5. Structural map of the Vedi area (Armenia).

and leucogranite intrusions. It is well exposed NE of Yerevan, in the Dzarkuniatz massif (Aghamalyan 1998) (Fig. 2). The Palaeozoic rocks of the SAB in Armenia are mainly localized in the southwestern part of the country (Figs 2 & 5). From bottom to top, the following succession of sedimentary rocks has been distinguished (Paffenholtz 1959; Karyakin 1989).

(1)

(2) (3)

Upper Devonian series (Frasnian –Famennian) composed of at least 1000 m thick detrital rocks including quartzites, sandstones and argillites. 500– 700 m thick Carboniferous reef limestones. A Permian platform type section composed of 500– 1500 m thick black limestones and

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Fig. 6. Geological cross-section of the Vedi area (Armenia).

Fig. 7. Photographs of the obduction contact in Vedi area (a) and on the formations overlying the ophiolites (b). Unconformities U1 is related to the obduction event and U2 is related to the first collision stage.

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marls with occurrences of bitumen. As the Upper Devonian formations they could represent a source rock for hydrocarbon. They disconformably overlie the Devonian and Carboniferous formations. Some syn-sedimentary normal faults occur in the sequence. The contact of the sedimentary cover with the basement is unknown in these areas (Figs 5 & 6). The metasedimentary rocks exposed in the Dzarkuniatz massif could belong to the Proterozoic or to the Upper Devonian-Permian formations. Normal faults are suspected in the Devonian-Carboniferous formations as deduced by the geometry of the bedding within the Cenozoic structures (fold), (Fig. 5 and cross-section Fig. 6). Lower Triassic rocks are conformably overlying on the Upper Permian limestones (Fig. 5). The series are composed of thin layered conodont rich limestones (Grigorian 1990), which evolve upward to detrital series composed of sandstones, black siltstones and mudstones of Late Triassic age. The thickness of the Triassic formations ranges from 700 to 900 m. In Armenia, rocks from the Jurassic period are unknown in the southern part of the SAB. However, tens of kilometres southeastward in Nakhijevan (Azerbaijan) and in Iran (south to Araks valley) (Fig. 5), 500 m thick Lower and Middle Jurassic are described on top of the Upper Triassic section (Lordkipanidze et al. 1988). Palaeomagnetic data from Middle Jurassic alkaline basalts and from Aalenian to Bajocian sediments indicate a palaeo-latitude of 21.58N + 3.78 during this period of time (Bazhenov et al. 1996). The sedimentary sequence is characteristic of a platform containing Ammonites, Briachiopodes and Pelecypodes (Bazhenov et al. 1996). Cenomanian and Turonian carbonates unconformably overlie the previous formations (Fig. 6). They are composed of thick reef limestones with some thin marl intercalations (Eghoyan 1955; Paffenholtz 1959; Sokolov 1977; Hakobyan 1978). In Vedi Area, the reef limestones are overlain upward by a Upper Cenomanian flysch (new dating by nannofossils, sample AR S 16, Table 1) thicker southward than northward (Sokolov 1977) and by a Upper Coniacian –Santonian olistostrome (Fig. 8) (Eghoyan 1955; Rengarten 1959; Hakobyan 1976, and ages we obtained from nannofossils, samples AR 63 05, AR 64 05, AR 65 05 and AR 67 05, Table 1) made of ophiolitic blocks contained in a muddy matrix (Fig. 4). According to the stratigraphic and lithological data of the SAB, in comparison with the Eurasian characteristics (Fig. 4) the SAB clearly differs from the Eurasian margin. This is also evidenced by tectonic phases recorded in the SAB and the Eurasian margin basement.

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The SAB metamorphic basement is Proterozoic (Aghamalyan 1969, 1978, 1998; Belov & Sokolov 1973), the Eurasian margin is mainly Variscan (Adamia et al. 1983; Adamia 1984, Zakariadze et al. 1998). The Palaeozoic series of the SAB mainly consist of platform-type carbonates whereas the sedimentary successions overlying the Eurasian crystalline basement are metamorphic schists dating back to the Visean and older. Moreover we do not find any sign of rifting all along the Eurasian margin series which could explain a drift to the south of the SAB after the Variscan orogeny. The palaeomagnetic data indicate a southerly palaeo-latitude of the SAB during the Early and Middle Jurassic (Barzhenov et al. 1996). All these data argue for a Gondwanian origin of the SAB located during Middle Jurassic, 2000 km south of its present location (Barzhenov et al. 1996).

Ophiolitic unit This ophiolitic sequence outcrops within a folded klippe (Figs 2, 5 & 8). According to the palaeontological data (Radiolarians), the Vedi ophiolites are mainly Middle Jurassic (Middle to Late Bajocian) (Danelian et al. 2008, 2010) and Late Jurassic (Danelian et al. 2010) and probably Early Cretaceous in age (Belov et al. 1991). The geochemistry undertaken on the oceanic crust indicates an N-MORB origin (Galoyan et al. 2007, 2009; Galoyan 2008; Rolland et al. 2009, 2010) and some basalts located on top of the previous ones are signs of an alkaline component (Tsameryan et al. 1988; Galoyan 2008). The geochemical compositions and the fact that peridotites, gabbros and plagiogranites were on the ocean-floor together (the radiolarites cover all these rocks) (Galoyan et al. 2008) indicate a slow-spreading oceanic crust, put in place in a back-arc domain, and overlain by a hot-spot series (see Rolland et al. 2010 for a synthesis). According to previous works, the whole ophiolitic sequence was interpreted as an ophiolitic melange (Sokolov 1977; Zakariadze et al. 1983). In contrast, we distinguish ophiolitic units comprising of preserved slices of oceanic crust thrusted over the Upper Coniacian –Santonian olistostrome of the SAB reworking the ophiolites (Figs 6 & 8). On top of the ophiolites, we observed reef-type sediments, OIB lavas and arc-type volcanic rocks (andesites lava flows), which unconformably cover the pillow-lavas (Fig. 8). These intra-oceanic arctype volcanic series have been evidenced from geochemical data (Galoyan et al. 2007, 2009; Galoyan 2008). An Upper Coniacian to Santonian transgressive series disconformably overlies the ophiolitic Vedi unit. This transgressive series is characterized

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Fig. 8. Geological cross-sections of the Vedi ophiolites, locations are shown on Figure 5.

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upwards by breccias and conglomerates, which rework ophiolitic rocks and the Cenomanian– Turonian autochthonous limestones (Fig. 7). The conglomeratic formation grades laterally into reef limestones composed of Upper Coniacian Hippurites fossils. Green mudstones, siltstones and thin turbiditic layered reef limestones with some pink levels overlie the reef limestones (Fig. 7). Late Coniacian –Santonian ages of the green mudstones were obtained by nannofossils (samples AR 72 05, AR 73 05, Table 1). One can note that in the Vedi area the Campanian and Maastrichtian are rare or missing. This indicates erosion or non-deposition due to the obduction uplift starting during the Late Coniacian and ongoing during Campanian to Maastrichtian times. Moreover, the HP metamorphism of the Stepanavan Blueschist unit is dated at 94 –90 Ma and its exhumation at 74– 71 Ma (Rolland et al. 2007). This exhumation is ascribed to the underthrusting of the SAB continental lithosphere still attached to the subducted slab below the oceanic crust section during the first stage of the obduction process. In this case, unroofing the Blueschist unit within the subduction channel could be the result of driven forces conducted by the low density of the continental crust. In summary the back-arc basin is thrusted over the SAB during Late Coniacian –Early Santonian time, as evidenced by 1) the ophiolitic olistostrome

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age; and 2) the occurrence at the same time of reef limestones on top of the ophiolites which suggest a shallow marine environment indicating uplift of the oceanic crust along the obduction thrust. This obduction stage was still active during Campanian and Maastrichtian explaining the absence or the pauce of sedimentation in the Vedi area south of Lesser Caucasus.

Collision The collision stages in the Lesser Caucasus are not well known. This is mainly due to superimposed tectonics which occurred in this region from Late Cretaceous to present. The first aim of our studies was to clearly identify the timing of collision triggering or more exactly that of the continental subduction of the SAB. Major unconformities and structures were dated allowing the reconstruction of the tectonic evolution of this collision event.

Beginning of the collision between SAB and Eurasia Along the Sevan –Akera suture zone North of Sevan Lake, late-Middle to Upper Eocene nummulitic series unconformably overlie the ophiolite units, the Eurasian margin and the SAB (Unconformity U2) (Figs 2, 9 & 10). According to these stratigraphic

Fig. 9. Structural sketch map of the northern Sevan Lake area (Sevan– Akera suture zone), (Armenia and Azerbaijan). Modified from Nalivkin (1976), and including our observations.

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Fig. 10. Geological cross-section of the northern Sevan lake area (Armenia and Azerbaijan). Photographs represent the Nummulites of the Nummulitic sandstone of late-Middle Eocene age and the unconformity U2 of the late-Middle Eocene on the Campanian pelagic limestones.

data the onset of the collision between the SAB and Eurasia occurred before the late-Middle Eocene. In detail the unconformable sedimentary deposits are characterized upward by conglomerates reworking Upper Cretaceous pelagic limestones, ophiolites and Lower Eocene andesites. Sandstones and nummulitic limestones overlie the conglomerate. The sedimentary succession of few tens of metres thickness is interbeded with andesitic lava flows.

Upwards, the series is characterized by thick volcanic layers made of andesitic and trachytic lavas (Fig. 10). The late-Middle to Late Eocene age (NP 18) of the base of this formation was obtained by nannofossiles (samples AR-03-05, AR-07-05 and AR-08-05, Table 2). In the western part of North Sevan area this volcano-sedimentary series unconformably overlies thick (maximum 2000 m thick) Lower to Middle Eocene calcalkaline volcanogenic

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series mainly made of andesite, trachyte, rhyolites and sandstones (Lordkipanidze et al. 1988). This is still well expressed on the Eurasian margin in Armenia and Azerbaijan where the lateMiddle Eocene formations unconformably cover the Middle Jurassic to Upper Cretaceous formations (Fig. 2). Consequently the late-Middle Eocene deposition occurred after a period of erosion. The sedimentary products of this erosion event have to be found southwestward.

Syn-collisional basin The Paleocene is not well expressed in the area of the Sevan –Akera suture zone. Danian sandstones have been described (Abovyan 1961; Sokolov 1977) but they are not clearly exposed all along the suture zone. In this part of the belt the Paleocene time is characterized by uplift, erosion and folding. Cross-sections and maps from the suture zone in North Sevan area (Figs 3, 9 & 10) illustrate well Paleocene folds unconformably covered by the lateMiddle Eocene nummulitic sandstones and limestones. There, the Campanian pelagic limestones that overlie the ophiolites unit are folded (Fig. 10). Nevertheless to the south in the Vedi area the Paleocene detrital series disconformably overlie the obducted ophiolites and the SAB. These formations are well developed and characterized by thick molassic deposits (more than 1000 m in thickness) (Figs 5 & 6). The detrital sequences are made, from bottom to top, of conglomerates, sandstones, greywackes, siltstones, marls and some algae limestones interlayers. The detrital series mainly reworked the ophiolites and the SAB autochthon. The Paleocene age of this formation was obtained by nannofossils (samples AR-54-05, AR-101-05, AR-107-05, Table 2) and are in agreement with the previous dating (Nalivkin 1976). The Paleocene basin to the south has to be considered as a foreland basin in front of a collisional belt located to the north (Figs 2, 5 & 6).

Structures and deformation related to the SAB collision In Vedi area, N1308E trending folds deform the Palaeozoic to Santonian formations of the SAB, the ophiolites and the foreland basin (Fig. 5). Paleocene to Oligocene molasses covering the ophiolites are deformed by ramp anticlines (Fig. 6). As a result of the southwestward folding propagation, Eocene and Oligocene series are deposited with onlap relationships over the Palaeozoic –Mesozoic anticline cores (Fig. 6). The anticline cores are made of Palaeozoic formations from the SAB. For example, to the SE part of Vedi area the Oligocene formations unconformably cover the Permian.

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The corresponding tectonic phase is due to the southwestward propagation of the collisional deformation. These types of structures, featured by folds and ramps are related to de´collement levels within the Devonian series and the ophiolites (serpentinites) (Fig. 6). In contrast, the Miocene volcanic rocks seal the structures related to this deformation phase (Unconformity U3), as they appear to be monoclinal (Fig. 5). As a result the folding event can be bracketed between the Oligocene and Miocene periods. More precision on the age of these Neogene series is needed to accurately constrain the age range of the NE–SW shortening event responsible to folding. The disconformity (U4) of the Plio-Quaternary volcanites on the Miocene volcanic and sedimentary rocks is well exposed in this area (Figs 5, 6 & 8). The structural interpretation on Figure 6 takes into account the occurrence of reactivated normal faults in the Palaeozoic series and a crustal-scale active fault (the Garni Fault: a currently dextral strike –slip fault, Rebai et al. 1993; Philip et al. 2001) precisely located at the anticline axis of the main structure (Figs 5 & 6). This interpretation is in favour of inherited structures within the SAB basement which have been reactivated during the Cenozoic, and still currently active faults (Philip et al. 2001; Karakhanian et al. 2004; Avagyan et al. 2005). In the north area of the Sevan Lake two main unconformities have been observed. The first one (U1) corresponds to the deposition of the Upper Cretaceous pelagic series over the ophiolitic sequence (Fig. 10). We dated these formations by nannofossil assemblages from Late-Coniacian to Late Campanian (Avazan and Artanish valleys, Table 1). The Maestrichtian would be present in some parts (Nalivkin 1976). The second one (U2), described above corresponds to the deposition of the late-Middle to Late Eocene series over deformed ophiolites and Late Cretaceous formations (Figs 9 & 10). N50 –808E trending folds deform the Campanian pelagic limestones and are related to a shortening event that occurred before the deposition of Nummulitic sediments. Moreover, N120– 1308E trending reverse faults and thrusts cross-cut the Upper Eocene series all along the North Sevan Lake from Artanish to Zod (Figs 9, 10 & 11). These structures also occur all along the suture zone from the Transcaucasus massif to the Karabakh (Fig. 2) and in front of the belt from Transcaucasus domain to the southeastward end of the Lesser Caucasus in Armenia and Nakhijevan.

Timing of the collision stages According to the stratigraphical and structural data the collision between SAB and Eurasia started

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during the Paleocene as evidenced by the formations of a foreland basin in the southeastern part of the belt and by the folding and uplift of the Sevan–Akera suture zone. Following this uplift and erosion phase it is very probable that during Early Eocene the north flank of the suture zone and the Eurasian margin subsided, resulting in deposition of Lower Eocene detrital rocks and magmatism. Nevertheless uplift to the South, along the suture zone, and subsidence to the north along the southern Eurasian margin must have been coeval, resulting in a huge erosional surface on which late-Middle Eocene shallow marine sediments were deposited. Some normal faults in the Upper Eocene volcanites were reactivated as post Eocene thrust faults are present in the North Sevan Lake area (e.g. Fig. 10, cross-section B). During the Paleocene to late-Middle Eocene in the southern (Armenian) part, a flexural molassic basin covered the obducted ophiolite and related structures, and was progressively deformed along thrusts and de´collement faults developing a fold-and-thrust belt type until the Miocene (Fig. 6). Some of the normal faults within the SAB Palaeozoic formations and the basement have been reactivated as reverse faults (Fig. 6). We presume that these faults were normal faults responsible for titled blocks of the north SAB passive margin. The Miocene epoch corresponds to a drastic transition in the deformation style of the belt. As clearly evidenced by the Vedi area cross-section, folding is stopped and a huge erosion surface marks the transition (Figs 2, 3, 5 & 6). However, strain field evolved, and the shortening direction changed from NE–SW to NNW –SSE (Avagyan et al. 2005). This shift in strain direction produced reactivation of the main N1308E trending previous structures as left-lateral strike –slip faults with a reverse component. Since then deformation has remained with similar features, which resulted in the opening of NW–SE elongated volcanic clusters in the main shortening direction, and in a general uplift of the area (Karakhanian et al. 2004; Avagyan et al. 2005, 2010).

Geodynamic evolution and discussion According to the new results presented above and with the help of numerous previous data mainly acquired during the Soviet period we can propose a new geodynamic model of the Lesser Caucasus evolution since the Late Jurassic (Fig. 12).

Middle Jurassic to Coniacian According to palaeomagnetic data (Bazhenov et al. 1996) the SAB was located 2000 km south of its present position during the Middle Jurassic. Moreover from Middle Jurassic to Late Coniacian,

two main subduction zones are responsible for: (1) the huge magmatic activity on the Eurasian active margin (Somkheto-Karabakh Island arc); (2) the calcalkaline magmatism on top of the oceanic lithosphere which have been obducted since Late Coniacian; and (3) occurrence of Blueschist unit (94– 90 Ma). From the overall geometry, we can suspect a subduction towards the NW beneath the Eurasian margin, but the preceding stages of intra-oceanic subduction may be featured by different orientations. Based on stretching lineations, mineral HP-LT mineral lineations and senses of shear in the Stepanavan Blueschist unit, Rolland et al. (2007) suggest an east-verging intra-oceanic subduction zone. However this direction could have been rotated during collision. Whatever in this hypothesis the ophiolites of the Lesser Caucasus belongs to a marginal basin opened in back-arc position within the Neotethys Ocean (lherzolite ophiolite type [LOT] ophiolites, low spreading rate, Galoyan 2008; Galoyan et al. 2009; Rolland et al. 2009). All these results and the geochemical characteristics of the oceanic basalts evidence a SSZ type for the intra-oceanic subduction zone (Rolland et al. 2009). On the Eurasian margin we have observed a preConiacian erosional surface and disconformity of Coniacian and Santonian arc-type series characterized by a basal reddish conglomerate reworking the island arc series and their Upper Jurassic intrusions. Consequently during the Early Cretaceous an active erosion event took place, which resulted in the unroofing of plutons of the magmatic arc (Azerbaijan, mark E on Fig. 2). This erosion event is the result of significant uplift and denudation during the Early Cretaceous. The reasons for such a change in the Eurasian active margin strain field could be the subduction of the spreading ridge of the back-arc basin (see Fig. 12) or the entrance in the subduction zone of an oceanic plateau or an intra-oceanic ridge, which is suspected from the geochemical analysis of alkaline basalt on the ophiolites (Galoyan et al. 2007; Galoyan 2008; Rolland et al. 2009).

Coniacian to Paleocene From the Coniacian to the Paleocene the intraoceanic subduction (SSZ) evolved to a continental subduction of the SAB beneath the intra-oceanic arc and the marginal basin (Fig. 12). This event is supported by HP-LT metamorphism at 94 –90 Ma of oceanic formations within an accretionary prism identified in the Stepanavan area (Fig. 2) (Rolland et al. 2007; Galoyan 2008). This step conducts around 5 Ma later to the obduction of the ophiolites over the SAB. This is well evidenced by: 1) the ophiolites uplift evidenced by the Upper

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Fig. 11. Photographs of the post Eocene thrust faults in the northern Sevan Lake area; location of photo A shown on Figures 9 and 10, location of photo B shown on Figure 9.

Coniacian –Santonian reef limestone deposits on top of them in Vedi Area; 2) the coeval olistostrome formation beneath the ophiolitic thrust (Vedi and Sevan areas, Figs 8 & 9) and overlying Cenomanian– Turonian flysch of the SAB.

The Campanian–Maastrichtian formations are absent in the belt’s frontal part, while they are present as pelagic sediments on the ophiolites of the Sevan–Akera zone. This lateral variation is ascribed to the southern frontal part’s uplift

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Fig. 12. Geodynamic model of the evolution of the Lesser Caucasus from Late Jurassic to present (see explanations in the text).

produced by the southward obduction thrust propagation and by a remnant oceanic crust covered by pelagic limestones to the north. The southern slope of this remnant oceanic crust (on the back of

the obduction unit) was characterized by a northward dipping evidenced by north verging slumps within the Campanian pelagic deposits (Fig. 10 cross-section A). During the same time one notes

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the exhumation of the Blueschist unit along the previous intra-oceanic subduction zone (Rolland et al. 2007).

Paleocene to Lower Miocene The occurrence of a foreland basin in front of the belt and the folding and erosion of the Sevan–Akera ophiolitic zone suggest the entrance of the SAB in the subduction zone below the Eurasian margin since the Paleocene, pulled by the dense eclogitized oceanic slab to which it is still attached. Due to the low density of the SAB continental crust entering in the subduction contact under the Eurasian margin, the continental subduction produces: 1) the uplift, shortening, folding and erosion of the obducted unit on the SAB (The Sevan–Akera zone); 2) the thrusting and folding of the Eurasian margin formations; and 3) the southwestward transport of the erosional products to the flexural basin in the belt front. The Lower to Middle Eocene magmatism on the Sevan–Akera suture zone (Lordkipanidze et al. 1988) could correspond to the first stage of a slab retreat triggered by the continental subduction and the break off of the continental lithospheric subducted mantle. Thermo-mechanical physical modelling has evidenced this mechanism (Boutelier et al. 2003, 2004). But more studies on the geochemical analysis of this magmatism must to be performed in order to test this model for the Lesser Caucasus region. One prominent result of this study is that the time of oceanic closure, indicated by the late-Middle Eocene unconformity on the SAB, the suture zone and the Eurasian margin clearly pre-dates this epoch. Therefore, the huge magmatic activity occurring from Late Eocene to Miocene (Lordkipanidze et al. 1988; Karapetian et al. 2001) has to be explained in this collisional context. We suggest that the slab retreat and break off processes lead to an asthenospheric upwelling below the suture zone (Fig. 12). This mechanism could explain the location of magmatism all along the Lesser Caucasus. However, some geochemical studies would be necessary to test this assumption.

Upper Miocene to present The continental subduction seems to be stopped but deformations still occur. The Arabian plate collision with the SAB to the South (in the Bitlis region, Oberha¨nsli et al. 2010 for a review), is responsible for the strain field’s reorganisation (Philip et al. 2001; Avagyan et al. 2005 and references herein). This produces the reactivation of the thrusts as strike– slip faults and the opening of volcanic clusters (Avagyan et al. 2005). Due

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to the collision, one may explain the widespread magmatic activity by partial melting of overthickened continental crust (Mitchell & Westaway 1999; Karapetian et al. 2001). Another possibility could be a slab break off process beneath Eastern Anatolia (Faccenna et al. 2006; Lei & Zhao 2007). However, this second hypothesis does not seem appropriate to explain all the magmatic activity known in the Lesser Caucasus and Greater Caucasus regions, which are relatively far from Eastern Anatolia. Another solution could be that of asthenospheric upwelling (if this one has actually existed) producing significant weakening of the SAB continental lithospheric mantle, which begins its delamination. This could explain the magmatism and uplift of the Lesser Caucasus since Pliocene time. The question if whether asthenosphere upwelling is related or totally unrelated to the lithosphere-scale plate tectonic processes remains open.

Conclusions The results coming out from studies undertaken in the Lesser Caucasus region allow to clarify the main questions on its tectonic and geodynamic evolutions. (1) The Sevan–Akera suture zone represents the only suture of the belt. The other ophiolitic units (Vedi and Zangezur, Fig. 2) correspond to ophiolitic nappes due to the obduction of oceanic crust over the SAB or to tectonic slivers reworked by the Upper Oligocene– Lower Miocene thrusts and reverse faults. The collision stage has significantly deformed the initial obduction contact by folding and thrusting and then more recently by strike– slip faulting. Consequently all the identified ophiolitic units in the Lesser Caucasus have been remobilized and outcrop today thanks to these collisional structures, as in the cores of anticlines (Sevan Lake, Stepanavan areas) (Figs 2 & 9) or in synclinal klippes (Vedi area) (Fig. 6) or as slivers along thrust faults (Zangezur and Amassia areas) (Fig. 2). In our interpretation of the structure of the Lesser Caucasus the ophiolites are continuous from Central to North Armenia, deep below the Sevan Lake (Fig. 3). The consequence of this structure is that the SAB is underthrusted below this huge ophiolite sliver, and may continue also below the Sevan Lake. This point is important in terms of natural resources because of the hydrocarbon potential of source rocks in the SAB Palaeozoic section. Indeed these source rocks could be capped by the allochthonous ophiolites, which could

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represent an overscreen under the Sevan Lake and the neighbouring regions. The onset of collision or the continental subduction of the SAB below the Eurasian margin is dated as Paleocene. This process occurred around 20 Ma later than the obduction (Late Coniacian –Santonian, 88 –83 Ma) of the marginal basin over the SAB. It led to the uplift of the Sevan–Akera suture zone, its folding, erosion and to the transfer of detrital

materials in a flexural basin in front of the belt, above the obduction structures. From late-Middle Eocene to Miocene all the belt is deformed by thrusts and reverse faults in the internal part (Transcaucaus to Karabakh, Fig. 2) and by N1308E trending fold and thrust belt at the chain front. One can note an evolution from thin-skin to thick-skin tectonics in front mainly due to the reactivation of the previous normal faults in the SAB

Fig. 13. Campanian, Maestrichrian and Ypresian palaeotectonic maps of the Lesser Caucasus region and its neighbouring areas (modified from Barrier & Vrielynck 2008).

LESSER CAUCASUS GEODYNAMIC

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basement at depth (Vedi and Southeast Armenia area, Figs 2 & 3). The widespread Eocene to recent magmatic activity is the result of various, still enigmatic processes. The Eocene magmatic activity occurred after the onset of the collision. Consequently it cannot be directly related to the oceanic subduction processes except for the Adjara-Trialet massif in Georgia considered as a back-arc basin due to a oceanic subduction of a small remnant oceanic lithosphere NW of the Sevan– Akera Suture zone (Adamia et al. 1981; Barrier & Vrielynck 2008). We suggest the possibility of a slab retreat and a break off allowing the rising of the asthenospheric mantle, which could have heated the overthickened continental crust (Fig. 12). The Miocene to recent magmatic activity seems to be related to 1) a possible asthenosphere upwelling beneath the suture zone after the break off and, 2) strain field reorganization at these times due to the Arabian plate collision with the SAB producing opening of volcanic clusters along main inherited faults (Karakhanian et al. 2004; Avagyan et al. 2005, 2010).

The Lesser Caucasus is a key area to understand lateral connections with the Taurides –Anatolides to the west and the Iranian regions to the East. Indeed, as demonstrated by Sto¨cklin (1974) and developed by Zanchi et al. (2009) the Eo-Cimmerian 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). However, a thick comparable Shemshak-type basin is unknown in the Lesser Caucasus. In this belt the contact between the Upper Triassic detrital series and the Lower Triassic carbonates formations is not well exposed. A complete sequence has been described in the Nakhijevan part of the Lesser Caucasus (Barzenov et al. 1995). Nevertheless, no evidence of Eo-Cimmerian orogeny has been found in the Lesser Caucasus. These features conduct to expect in the study area a major stratigraphic and structural limit. Consequently, the oriental border of the continental blocks of Taurides– Anatolides and South Armenian Block of Gondwanian origin need to be clarified. This block collided with Eurasia during Paleocene Eocene times. According to the MEBE palaeotectonic map (Fig. 13) (Barrier & Vrielynck 2008) the passage between Lesser Caucasus and NW Iran (NW Alborz and Tabriz-Saveh) would be then a syntax similar to those delimiting India colliding with Asia, but it remains to be demonstrated. One way to explain the problems is the existence of NE –SW orientated Araks strike –slip fault,

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suggested as an oriental border of Lesser Caucasus (Fig. 1). But if this fault has a recent expression, it is not sure that it was present at the beginning of the collision between the South Armenian Block and Eurasia. New interpretations allow some points to be studied in more detail, particularly the origin of the widespread collisional magmatism and the connections with the Iranian belt, which remain to this day a subject of debate. This work was financed by INTAS and since 2003 by the Middle East Basins Evolution (MEBE) programme. Many thanks to the MEBE programme coordinators E. Barrier and M. Gaetani, for their support and encouragement. Fieldworks were greatly facilitated by the support of the Institute of Geological Sciences (Armenia), the Geological Institute of Azerbaijan and the Institutes of Geophysics and Geology (Georgia). We also thank the support of the French Embassy at Yerevan (Armenia) for the MAE PhD grant of G. Galoyan. The reviews by A. Tremblay and A. Zanchi are highly acknowledged.

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Melikyan, L. S. 2004. Some geological and structural features of inner structure of the Amassiya– Sevan ophiolitic zone (Republic of Armenia). Izvestia NAS of Armenia, Nauki o Zemle, 3, 32– 38 (in Russian). Melkonian, R. L. 1976. Petrology, mineralogy and geochemistry of the complexes of Alaverdi ore region. In: Petrology and Geochemistry of the Intrusive Complexes of Some Ore Regions. Ana of Armc SSR, Armenian SSR (in Russian). Melkonyan, R. L. & Akopyan, M. S. 2006. Some aspects of the interrelations between magmatism and ore formation: evidence from oxygen isotope data on ore– magmatic systems, Armenia. Petrology, 14, 413– 420. Milanovski, E. E. 1968. Neotectonics of the Caucasus. Nedra, Moscow (in Russian). Mitchell, J. & Westaway, R. 1999. Chronology of Neogene and Quaternary uplift and magmatism in the Caucasus; constraints from K–Ar dating of volcanism in Armenia. Tectonophysics, 304, 157– 186. Monin, A. S. & Zonenshain, L. P. (eds) 1987. History of the Ocean Tethys. Moscow Institute of Oceanology (in Russian). Nalivkin, P. V. (ed.) 1976. Geological Map of Caucasus at 1/500 000 scale. Ministry of Geology, USSR, Cief. Nikishin, A. M., Cloething, S., Brunet, M.-F., Stephenson, R. A., Bolotov, S. N. & Ershov, A. V. 1998. Scythian Platform, Caucasus and Black Sea Region: Mesozoic–Cenozoic tectonic history and dynamics. In: Crasquin-Soleau, S. & Barrier, E. (eds) Peri-Tethys Memoir 3: Stratigraphy and Evolution of Peri-Tethyan Platforms. Me´moire du Muse´um d’Histoire Naturelle, Paris, 177, 163– 176. Oberha¨nsli, R., Candan, O., Bousquet, R., Rimmele, G., Okay, A. & Goff, J. 2010. Alpine high pressure evolution of the eastern Bitlis complex, SE Turkey, GSL. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 461–483. Paffenholtz, K. N. 1959. Geological History of the Caucasus. Izdatelstvo NAS Arm. SSR (in Russian). Philip, H., Avagyan, A., Karakhanian, A., Ritz, J.-F. & Rebaı¨, S. 2001. Estimating slip rates and recurrence intervals for strong earthquakes along an intracontinental fault: example of the Pambak-Sevan-Sunik fault (Armenia). Tectonophysics, 343, 205–232. Rebaı¨, S., Philip, H., Dorbath, L., Borissoff, B., Haessler, H. & Cisternas, A. 1993. Active tectonics in the Lesser Caucasus: coexistence of compressive and extensional structures. Tectonics, 12, 1089– 1114. Rengarten, V. P. 1959. Stratigraphy of the Cretaceous Deposits of the Lesser Caucasus. Regional Stratigraphy of USSR. Izdatelstvo of AS of USSR, Moscow (in Russian). Ricou, L. E. 1994. Tethys reconstructed: plates, continental fragments and their boundaries since 260 Ma from Central America to South-eastern Asia. Geodinamica Acta, 7, 169– 218. Ricou, L. E., Dercourt, J., Geyssant, J., Grandjacquet, C., Lepvrier, C. & Biju-Duval, B. 1986. Geological constraints on the Alpine evolution of the Mediterranean Tethys. Tectonophysics, 123, 83–122.

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The Armenian Ophiolite: insights for Jurassic back-arc formation, Lower Cretaceous hot spot magmatism and Upper Cretaceous obduction over the South Armenian Block Y. ROLLAND1*, G. GALOYAN1,2, M. SOSSON1, R. MELKONYAN2 & A. AVAGYAN2 1

Universite´ de Nice-Sophia Antipolis, OCA, UMR Ge´oAzur, CNRS, Parc Valrose, 06108 Nice cedex 2, France 2

Institute of Geological Sciences, National Academy of Sciences of Armenia, 24a Baghramian avenue, Yerevan, 375019, Armenia *Corresponding author (e-mail: [email protected])

Abstract: Similar geological, petrological, geochemical and age features are found in various Armenian ophiolitic massifs (Sevan, Stepanavan and Vedi). These data argue for the presence of a single large ophiolite unit obducted on the South Armenian Block (SAB). Lherzolite Ophiolite type rock assemblages evidence a Lower–Middle Jurassic slow-spreading rate. The lavas and gabbros have a hybrid geochemical composition intermediate between arc and Mid Ocean Ridge Basalt (MORB) signatures which suggest they were probably formed in a back-arc basin. This oceanic sequence is overlain by pillowed alkaline lavas emplaced in marine conditions. Their geochemical composition is similar to plateau-lavas. Finally, this thickened oceanic crust is overlain by Upper Cretaceous calc-alkaline lavas likely formed in a supra-subduction zone environment. The age of the ophiolite is constrained by 40Ar/39Ar dating experiments provided a magmatic crystallization age of 178.7+2.6 Ma, and further evidence of greenschist facies crystallization during hydrothermal alteration until c. 155 Ma. Thus, top-to-the-south obduction likely initiated along the margin of the back-arc domain, directly south of the Vedi oceanic crust, and was transported as a whole on the SAB in the Coniacian times (88–87 Ma). Final closure of the basin is Late Cretaceous in age (73–71 Ma) as dated by metamorphic rocks.

The history of central and northern Neotethys can be inferred from the study of oceanic crust domains obducted in the Armenian Lesser Caucasus. It is important to depict the geodynamic evolution of oceanic domains that were formed in the NeoTethyan domain, as they provide key time and palaeogeographic data in the Middle East Basin Evolution (e.g. Sengo¨r & Yilmaz 1981; Tirrul et al. 1983; Ricou et al. 1985; Dercourt et al. 1986; Ricou 1994; Okay & Tu¨ysu¨z 1999; Stampfli & Borel 2001; Barrier & Vrielynck 2008). Furthermore, they provide constraints on the timing of oceanic closure and obduction, by the study of metamorphic rocks associated to the ophiolites. Finally, their geometry is also important to infer the preservation potential of oil resources that could be contained underneath. During the Mesozoic, the Southern margin of the Eurasian continent has been featured by closure of the Palaeo-Tethys and opening Neo-Tethys Ocean (Fig. 1). Later on, subductions, obductions and micro-plate accretions, ranging mostly from the Cretaceous to the Eocene and finally continent– continent collision have occurred between Eurasia

and Arabia. The study of Armenian ophiolites allows unravelling part of this complex history. Previous geological, petrological and geochemical works undertaken on those date back to late 1970s and 1980s, and have never been undertaken at the scale of the Armenian ophiolites. This work is particularly difficult due to the very large number of tectonic and volcanic events that have occurred after ophiolite obduction. The polyphased tectonic history of the Lesser Caucasus region includes arccontinent accretion and subduction-exhumation in or above accretionary prisms followed by continent –continent collision (e.g. Okay & Tu¨ysu¨z 1999; Rolland et al. 2007). The tectonic events have dissected the ophiolites (Avagyan et al. 2005), which were further overlain by very thick (.1 km) sequences of volcanic rocks. In this paper, we propose a synthesis of the research undertaken on Armenian ophiolites, based on recently published papers of individual ophiolite zones and new geological, petrological, geochemical and 40Ar/39Ar geochronological data obtained on these different zones. Three ophiolites have been studied, located in NW Armenia

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 353–382. DOI: 10.1144/SP340.15 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Black

Sea LC LC Albroz CCB i i

i

i i i

i

i

LC: Lesser Caucasus

Albroz CCB: Albroz CarboniferousCimmerian belt

Fig. 1. Tectonic map of the Middle East –Caucasus area, with main blocks and suture zones, after Avagyan et al. (2005), modified.

(Stepanavan), north Armenia (Sevan) and central Armenia (Vedi; Fig. 2). In the three zones, we show the presence of three superposed levels of lavas corresponding to three distinct environments, from bottom to top: (1) back-arc; (2) ‘Ocean Island Basalt’ (‘OIB’)-like; and (3) arc. Moreover, we demonstrate that these ophiolite windows should correlate with each other and be part of a

unique obducted nappe above the South Armenian Block (or SAB, Knipper & Khain 1980; Zakariadze et al. 1983). We propose that tectonic transport of this nappe onto the SAB occurred directly after the OIB ‘plateau’ event, in the Early Upper Cretaceous, and shortly preceded the final oceanic closure along the Eurasian margin in the latest Cretaceous (c. 73 –71 Ma).

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Fig. 2. Sketch geological map of Armenia, with location of the studied areas: 1, Stepanavan area; 2, Sevan area; 3, Vedi area. 4 is the location of Zangezur ophiolites, located along a NNW–SSE striking fault.

Geological setting The ophiolites are located in the northern part of the Lesser Caucasus region (Fig. 1). They are situated in three geographic zones (Figs 2–5). (1)

(2)

The Sevan– Akera zone at the northern rim of the SAB and at the southern edge of the European active continental margin (Knipper 1975; Knipper & Khain 1980; Adamia et al. 1980). In the present paper, we present detailed mapping of the Stepanavan (NW Armenia, Fig. 3) and of the Sevan (north Armenia, Fig. 4) ophiolites. The Vedi zone (Fig. 5), disposed in a more southerly position, above the SAB (Knipper

(3)

& Sokolov 1977; Knipper & Khain 1980; Zakariadze et al. 1983), or within a suture zone eventually correlating with Central Iran or Alborz ophiolites (Sokolov 1977; Adamia et al. 1981). The Zangezur zone situated along the Zangezur fault (Aslanyan & Satian 1977, 1982), between the two domains, interpreted as an ophiolite suture by Knipper & Khain (1980) and Adamia et al. (1981).

A companion paper written on the geology of the Sevan ophiolite has already put up in detail the lithologies and radiometric age of this ophiolite (Galoyan et al. 2009). Main features are summarized below. The lithological assemblages found

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Fig. 3. Sketch geological map of the Stepanavan ophiolite (NW Armenia).

agree with a Lherzolite Ophiolite Type (LOT; Nicolas 1989); these include the following. (i) A high level of fractional crystallization in the series, with cumulate olivine gabbros and two pyroxene gabbros overlain and intruded by generally amphibole-bearing gabbros, and frequent differentiated melts (diorites to plagiogranites). These most differentiated melts are

generally emplaced in ductile extensive shear zones cross-cutting the gabbros. This complete differentiation series suggests low partial melting levels and long-lived cooling as is proposed in LOT settings (Lagabrielle et al. 1984; Lagabrielle & Cannat 1990). Absolute radiometric datings indicate oceanic crust emplacement in the Middle Jurassic, constrained at 165– 160 Ma by zircon U –Pb age of one

ARMENIAN OPHIOLITES

357

Fig. 4. Sketch geological map of the Sevan ophiolite (north Armenia), after Galoyan et al. (2009), modified.

tonalite (160 +4 Ma; Zakariadze et al. 1990) and by 40Ar/39Ar amphibole age on gabbro (165.3 +1.7 Ma; Galoyan et al. 2009). (ii) Rare pillow lavas, with compositions ranging from tholeiitic basalts to andesites. The feeding dyke swarm is reduced, as rare dolerite dykes have been found crosscutting the series. The calc-alkaline affinity is also evidenced by Nb –Ta negative anomalies, which agree with some interaction with slab-derived component. These support a slow spreading rate in a back-arc setting. (iii) Peridotites are frequent and often exhumed at sea-floor level. They are generally serpentinized, and witness further hydrothermal alteration when exhumed at sea-floor level (‘listwenites’). The lherzolitic nature of the mantle-derived ultramafic rocks is then difficult to assess. The previous petrographical investigations on the serpentinized ultramafics suggest that the protoliths was mantle-derived with various compositions ranging from lherzolites to harzburgites and dunites (e.g. Melikyan et al. 1967; Harutyunyan 1967; Palandjyan 1971; Abovyan 1981; Ghazaryan 1987; Zakariadze et al. 1990). Undeformed ultramafics have a magmatic cumulative origin, shown by the poikilitic texture of olivine inclusions within large enstatite crystals (up to 10– 15 mm; Palandjyan 1971). We observe

similar textures, together with layers, contained in magmatic pods cross-cutting the serpentinites in the Stepanavan area (Galoyan et al. 2007). These latter serpentinites are strongly deformed and altered, thus it was difficult to unravel their origin. However, the ductile character of deformation is in agreement with a mantle origin for these rocks. (iv) Radiolarites are found as interlayers or as unconformably overlying the various above lithologies. The fact that they overlie gabbros, plagiogranites and serpentinites shows that these rocks were uplifted and exhumed by normal faults. Radiolarite datings undertaken in the different ophiolites all agree with oceanic accretion in the Middle– Upper Jurassic (Danelian et al. 2007, 2008). The ophiolitic sequences are weakly deformed with anchizonal metamorphism. Only some outcrops show evidence of small shear zones ascribed to the ophiolite obduction in the Coniacian (Zakariadze et al. 1983). High pressure (HP) metamorphism is described in the Stepanavan region (Figs 2 & 3), where blueschists (Aghamalyan 1981, 1998) outcrop in small km2 size tectonic windows below the ophiolite. Timing of metamorphism from radiometric 40Ar/39Ar phengite datings indicates HP metamorphic peak at c. 95 Ma, and medium

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Fig. 5. Sketch geological map of the Vedi ophiolite (Central Armenia).

pressure-medium temperature (MP-MT) retrogression at 73–71 Ma (Rolland et al. 2007). The ophiolite series are locally overlain from bottom to top by (1) alkaline lavas, which have age ranges from 114 to 95 Ma (Baghdasaryan et al. 1988; Satian & Sarkisyan 2006) recently confirmed by 40Ar/39Ar dating of 117.3 +0.9 Ma (Rolland et al. 2009); and (2) Upper Cretaceous andesites and detrital series (Dali valley; Stepanavan; Galoyan et al. 2007). The alkaline lavas are alternatively interpreted as (1) intra-continental

rifting (Satian et al. 2005) in the Vedi area; and (2) plume-derived magmatism above the oceanic crust before the obduction (Galoyan et al. 2007). The calc-alkaline series should be related to intra-oceanic arc emplacement above this oceanic crust sequence and implies the presence of a subduction zone between the ophiolite and the SAB, featured by the Stepanavan blueschists (Rolland et al. 2007). These two magmatic sequences closely predate the ophiolite obduction onto the SAB during the Coniacian (Sokolov 1977).

ARMENIAN OPHIOLITES

Analytical methods Mineral compositions were determined by electron probe microanalysis (EPMA). The analyses are presented in Figure 6. They were carried out using a Cameca Camebax SX100 electron microprobe at 15 kV and 1 nA beam current, at the Blaise Pascal University (Clermont-Ferrand, France). Natural samples were used as standards. Thirty-seven samples of magmatic rocks from the Sevan, Stepanavan and Vedi ophiolites have been analysed for major, trace and Rare Earth elements (REE) (Table 1). Samples were analysed at the C.R.P.G. (Nancy, France). Analytical procedures and analyses of standards can be found on the following website (http://www.crpg.cnrsnancy.fr/SARM). Amphiboles were separated from the gabbro sample AR-05-110, from the Vedi ophiolite. Geochronology of amphiboles was undertaken by laser 40 Ar/39Ar dating. The results are presented in Table 2. Amphibole crystals were separated under a binocular microscope. The samples were then irradiated in the nuclear reactor at McMaster University in Hamilton (Canada), in position 5c, along with Hb3gr hornblende neutron fluence monitor, for which an age of 1072 Ma is adopted (Turner et al. 1971). The total neutron flux density during irradiation was 9.0  1018 neuton cm22. The estimated error bar on the corresponding 40Ar*/39ArK ratio is +0.2% (1s) in the volume where the samples were set. Three amphibole grains (c. 500 mm in diameter) were chosen for analysis on a laser UV spectrometer of Nice (Ge´osciences Azur). Analyses were undertaken by step heating with a 50 W CO2 Synrad 48-5 continuous laser beam. Measurement of isotopic ratios was done with a VG3600 mass spectrometer, equipped with a Daly detector system; see detailed procedures in Jourdan et al. (2004). The typical blank values for extraction and purification of the laser system are in the range 4.2 –8.75, 1.2–3.9 and 2– 6 cc STP for masses 40, 39 and 36, respectively. The massdiscrimination was monitored by regularly analysing air pipette volume. Decay constants are those of Steiger & Ja¨ger (1977). Uncertainties on apparent ages in Table 2 are given at the 1s level and do not include the error on the 40Ar*/39Ark ratio of the monitor. Uncertainties on plateau ages in Figure 7 are given at the 2s level and do not include the error on the age of the monitor.

Results Field relationships Synthetic logs are drawn on Figure 8, showing the lithological associations and the structural

359

relationships in each zone. In Stepanavan (Fig. 8a, b), ophiolite sections exhibit abundant serpentinites, cross-cut by normal fault and shear zones in which gabbronorites, gabbros and plagiogranites are intrusive and deformed. Laterally, thick layers of pillow basalts are observed which interlayer with radiolarites. At the top of the ophiolite section, a thin layer of alkaline lava flows is found. Above, these lavas are uncomformably overlain by Upper Cretaceous conglomerates and limestones, and calc-alkaline pillow basalts or graywackes. The ophiolite sequence is thrust over a blueschist facies metamorphic sole, which outcrops in two km2 sized tectonic windows. In the Sevan area, sections are extremely variable laterally (Fig. 8c–e). Pillow lavas are rare, and serpentinites are frequently found at sea-floor level. Intense hydrothermal alteration (‘listwenites’) has transformed the uppermost part of exhumed serpentinites. Rare dolerites are observed. Large intrusive pods of amphibole-bearing gabbros and plagiogranites are also exhumed and overlain by radiolarites. Normal faults are observed, and are interpreted as the cause of such lateral variations, by vertical uplifting of footwall sections, and local infilling of axial rift valleys, in agreement with the LOT ophiolite model (e.g. Lagabrielle et al. 1984; Lagabrielle & Cannat 1990). Locally, thick sequences of alkaline pillow lavas are observed. The ophiolite is locally eroded, and uncomformably overlain by conglomerates and soils, overlain by an Upper Cretaceous section of reef-limestones with graywackes interlayers. In the Vedi area, the ophiolite section is much thinner (Fig. 8f–h). The basal tectonic contact is exposed, exhibiting top-to-the-south sense of shear. At the base, the ophiolite rests on a serpentinite layer. The ophiolite is intensely sheared above the basal contact with boudins of tholeiitic basalts (Fig. 8h). Laterally, the ophiolite consists mainly of gabbros (Fig. 8g) or serpentinites, which suggests similar lithological features as in the Sevan area. However, the different parts of the ophiolite are dismembered and displaced from each other as a result of obduction deformation. Above the ophiolites, layers of radiolarites are found below a very thick section of alkaline pillow lavas (Fig. 8h). These alkaline lavas are amphibole-bearing. Pillows are larger (metre-scale) than the ophiolite ones (several decimetre scale), and interlayer with thin pink limestones. At the front of the obduction, an olistolith formation with conglomerates and slided blocks in a muddy matrix is present (Fig. 8f). The age of the olistostrom is Coniacian –Santonian (see Sosson et al. 2010), it connects laterally to Lower Coniacian series below the ophiolite, and with Santonian reef limestones above the obduction. Therefore, the obduction age can be bracketed

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Y. ROLLAND ET AL. 0.10

Sevan 0.09

G150 olivine gabbro AR-05-86 olivine gabbro AR-03-10 hornblende gabbro AR-03-02 diabase

0.08 0.07

Stepanavan AR-04-02 wehrlite AR-04-03 websterite AR-04-04 websterite AR-04-51 websterite AR-04-53 anorthosite AR-04-45D hornblende gabbro AR-04-31 basaltic andesite

Ti

0.06 0.05 0.04

Alkaline

Tholeiitic

0.03

Vedi

0.02

AR-05-112A hornblende gabbro AR-05-113 hornblende gabbro AR-05-103 gabbro-diabase

0.01 0.00 0.5

0.6

0.7

0.8

0.9

1.0

Na + Ca Fig. 6. Chemical compositions of studied clinopyroxenes plotted in the Ti v. (Na þ Ca) diagram of Letterier et al. (1982). Note that a majority of data plot in the Alkaline compositional field, and a minority is in the Tholeiitic part.

to the Coniacian –Santonian (88 –83 Ma). Laterally, the upper part of the ophiolite is made of kilometrescale slided blocks, mainly comprised of alkaline pillow basalts and calc-schists. These blocks slide on a greenish mudstone rock, probably originated from the alteration of the ophiolite itself. The Upper coniacian uncomformity is variably marked by conglomerates, marls and reef limestones. As emphasized in Galoyan et al. (2009) in the Sevan area, and by Galoyan et al. (2007) in the Stepanavan area, the lithologies found in all the exposed Armenian ophiolites are in good agreement with the hypothesis of a slow expansion rate spreading centre, as described for the western Alps ophiolites (LOT; Nicolas & Jackson 1972; Nicolas 1989). These LOT features include the following, as already stressed by Galoyan et al. (2009). (1) (2)

(3)

A high degree of fractional crystallization shown by coarse amphibole-bearing gabbros and widespread plagiogranites. Both plutonic rock types are often strained in ductile-to brittle conditions of amphibolite to lower greenschist metamorphic conditions and often exhumed at sea-floor level, as for serpentinites. These tectonic features are in agreement with a slow-spreading system in which the overall morphology is dominated by normal faults. The scarcity of basalt lava flows and the restricted dolerite dyke swarm are also in

agreement with a slow spreading rate system, featured by low partial melting levels. The similar lithological and age features found in the several Armenian ophiolites suggest that they were part of the same oceanic crust section. This has to be confirmed by the comparison of geochemical data from each zone. The presence of three magmatic series: ophiolite sensu stricto (tholeiitic), ‘OIB’ (alkaline) and arc (calc-alkaline) in the same structural position (from bottom to top, respectively) has been evidenced in the three zones.

Petrography and mineral chemistry The field and microscopic analyses of the Armenian ophiolite magmatic rocks have evidenced a continuous magmatic succession from ultramafic cumulates (wehrlites, websterites) to gabbros and plagiogranites, exhumed at sea-floor level and overlain by pillow-basalts. Ophiolite plutonic rocks. Wehrlites are found in the Stepanavan ophiolite and have a poikilitic texture showing numerous clinopyroxene crystals with diopside composition (Wo45-47En48-50Fs2-4), included in large olivine Fo87-88 (.60 –65%) porphyric grains. Spinel is less than 5%, and is of magnetite composition. Gabbros are the most abundant rocks in the crustal complex. Their petrography evolves from cumulate-banded olivine gabbros in their lower

ARMENIAN OPHIOLITES

part towards more leucocratic plagioclase-rich gabbros in the upper part (Abovyan 1981; Ghazaryan 1987, 1994). The Cumulative banded olivine gabbros and websterites are described only in Sevan and Stepanavan areas, while more leucocratic gabbros are widespread. Olivine gabbros are fresh, massive, and fine- to medium-grained (0.5–1 mm and 1.2–2 mm correspondingly). They have cumulate, ophitic textures and consist of plagioclase (c. 60– 65%; An68-74, An80-89), olivine (c. 5–10%; Fo72-76) and clinopyroxene (c. 25– 35%). Clinopyroxene is of augite (Wo39-44En45-48Fs11-13) and diopside (Wo45En44 Fs11) types. Some enstatitic orthopyroxenes (Wo2 En75Fs23) are also found rimming olivine porphyrocrysts. Websterites have a granular texture with large 2–8 mm porphyrocrysts of orthopyroxene (30 – 70%), clinopyroxene (70–30%) and olivine grains (0–35%; Fig. 9a). Orthopyroxenes are enstatite-rich (Wo1-5En59-84Fs11-37) and clinopyroxenes are augites (Wo35-42En36-40Fs15-19) and olivine is relatively rich in forsterite (Fo84-88). Gabbronorites have a gabbroic texture, with plagioclase (10–60%, 1– 3 mm), clino- and ortho-pyroxene. Plagioclase is of bytownite type (An80-85), while orthopyroxenes (1–4 mm) are enstatites (Wo2-5 En59-61Fs34-37), and clinopyroxenes are augites (Wo35-42En36-40Fs15-19). Spinel is less than 5% and is of magnetite composition. Mesocratic to leucocratic gabbros of the upper section are massive, fine- to medium-grained and have gabbroic (or gabbro-ophitic), subautomorphe to xenomorphe granular texture (0.5 –4 mm), with plagioclase (c. 40– 65%; An50-75, An72-93), clinopyroxene (8– 45%; augite) and hornblende (0–40%) and lack olivine. Accessory minerals (1–10%) are apatite, titanomagnetite, ilmenite and rarely quartz. The hornblende-rich gabbros have coarse granular textures (Fig. 9b), with c. 50 –65% euhedral to subhedral plagioclase (An54-58) and (c. 35– 50%) anhedral to subhedral amphibole (Fig. 10a –d). Brown Ti-rich euhedral hornblende is presumed to be a primary mineral as it appears in the centre of phenocrysts; while a Ti-poor subhedral to xenomorphic green magnesio-hornblende (Leake et al. 1997) is thought to be a secondary phase related to hydrothermal alteration (Fig. 10d). The augite (Wo40-42 En39-47Fs11-14), diopside (Wo45-48 En40-44Fs8-15), and enstatite (Wo2En57Fs41) relicts (5–10%) are found in the magnesio-hornblendes crystals replacing the pyroxenes. However, it is not related to shear zones and fractures, and should also be a late magmatic mineral. In leucocratic gabbros, the clinopyroxene (augite Wo40-41 En33-35Fs18-19) content does not exceed 25%. Normal zoning is observed in plagioclase (from An85 to An60), which are frequently altered.

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Clinopyroxenes have alkaline to slightly tholeiitic compositions (0.8 , Na þ Ca , 0.9; Leterrier et al. 1982; Fig. 6). Pegmatitic gabbros crosscutting the plutonic sequence (Vedi zone) are composed of plagioclase and hornblende, and are mainly altered in chlorite, carbonate, sericite, albite, quartz, actinolite, and so on. Diorites (well-known in Sevan and Vedi zone) which occur as small, apparently intrusive bodies within the gabbro units (Palandjyan 1971; Abovyan 1981; Ghazaryan 1987, 1994), have a porphyritic to subhedral granular (1– 4 mm) texture and have relatively similar hornblende contents (5– 30%) as in the gabbros. Plagioclase (c. 65 – 70%) is albite-rich (An34-38) and accessory minerals (quartz, opaque oxides) are rare. Amphibole grains are magnesio-hornblendes in which actinolite and sometimes epidote aggregates are present. Laterally and towards upper sections (Sevan zone) diorites grade into quartz-diorite (quartz 5–10%). Plagiogranites (Fig. 10e, f) appear to be the most differentiated component of the gabbro-dioritic intrusives. They form diffused segregations or discontinuous networks of veins with local coarse pegmatic, or hypidiomorphic to xenomorph granular (0.5–4 mm) texture within and around gabbrodiorite intrusives. They are formed by 40–65% plagioclase (An15-30), 25–45% quartz, minor biotite (,5%), ortho-amphibole (,5%; Stepanavan), K-feldspar (0 –10%, microcline; Vedi zone; Fig. 9c) and accessory phases (titanomagnetite, hematite, sphene and apatite). Amphiboles are rarely preserved, commonly replaced by chloriteand epidote-group minerals. Ophiolite volcanic and subvolcanic rocks. Diabases are present in several locations (Sevan and Stepanavan areas) as isolated dikes, crosscutting the layered gabbros. They are generally altered (chlorite, epidote, carbonates) and have the subdoleritic texture composed of plagioclase (60– 65%; An65-75) and two clinopyroxenes (augite Wo41-44 En44-47Fs11-13 and diopside Wo45En37Fs18). The volcanic rocks of the studied Armenian ophiolites are present as pillowed and massive lava flows and pillowed breccias. In general they are relatively altered due to hydrothermalism but still relict igneous textures are preserved. The basalts and basaltic andesites are vacuolar (1– 5 mm, vesicles are filled with carbonate-calcite, chlorite and quartz) and largely aphyric (intersertal, spilitic, microdoleritic and variolitic, up to 1.5– 2 mm in diameter), composed mainly of albitized plagioclase and/or plagioclase-clinopyroxene microlites, Ti-magnetite and hematite microlites, in a devitrified (calcite þ chlorite) groundmass (Fig. 9d). Alkaline lavas. The alkaline basalts are found in several locations of the three ophiolite zones.

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Table 1. Representative whole-rock analyses of samples from the Sevan, Stepanavan and Vedi areas. Major Groups no.

Sevan ophiolitic series Flaser gabbro

Sample SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total Mg# Rb Sr Y Zr Nb Ba Hf Ta Pb Th U V Cr Co Ni Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu* (La/Sm)N (La/Yb)N

Olivine gabbro

AR-03-25 AR-05-86 45.36 13.32 14.91 0.14 7.44 10.89 3.15 0.22 2.55 0.31 1.46 99.8 52.1 2.67 231.7 61.81 175.9 4.51 36.77 4.25 0.35 – 0.08 0.11 440.7 35.43 26.81 77.98 18.16 58.57 14.89 34.75 5.057 24.09 7.38 3.11 9.32 1.60 10.47 2.22 6.37 0.94 6.28 0.95 1.15 1.27 1.60

48.09 16.72 5.94 0.11 10.51 14.1 1.65 – 0.27 0.02 2.8 100.2 79.3 0.48 102 6.41 5.28 0.08 4.1 0.22 0.01 – – – 138 802 40.2 188 102 28.3 0.33 0.94 0.18 1.23 0.58 0.34 0.87 0.16 1.14 0.24 0.65 0.10 0.65 0.09 1.45 0.36 0.34

Olivine gabbro

Gabbro

G150

AR-03-39

AR-03-24

AR-03-10

48.39 15.63 6.44 0.12 10.48 16.65 1.16 – 0.29 0.04 0.65 99.8 77.9 – 101 7.24 4.85 – – 0.21 – – – – 190 412 41.7 131 111 29.6 0.36 1.02 0.19 1.28 0.61 0.32 0.98 0.18 1.24 0.26 0.74 0.11 0.72 0.11 1.28 0.37 0.34

49.49 14.11 11.59 0.18 6.79 9.38 3.52 0.29 1.32 0.14 2.92 99.7 56.1 3.05 189.5 30.05 75.20 1.55 120.8 2.12 0.11 – 0.29 0.09 319.90 94.50 34.77 32.35 52.59 86.18 3.39 9.57 1.57 8.27 2.93 1.07 4.01 0.71 4.89 1.03 3.02 0.46 3.03 0.48 0.95 0.73 0.76

50.60 7.2 7.77 0.17 15.29 17.65 0.48 – 0.20 0.06 0.79 100 81.1 – 58.2 7.84 5.24 – 14.14 0.20 – – – – 195.9 810.3 41.88 136.2 142.3 47.98 0.28 0.99 0.22 1.40 0.67 0.28 1.07 0.20 1.36 0.29 0.83 0.13 0.86 0.13 1.01 0.26 0.22

50.68 18.17 9.09 0.16 6.75 9.26 3.21 0.15 0.36 0.05 1.21 99.1 61.6 0.84 303.7 11.93 22.89 0.50 34.36 0.81 0.04 – 0.14 0.07 222.2 104.3 33.71 29.23 44.81 71.99 1.60 4.51 0.74 3.80 1.34 0.48 1.70 0.30 1.92 0.41 1.22 0.19 1.32 0.21 0.97 0.75 0.82

Their structural position is as large massive pillowlavas flows on top of the ophiolite section or as dykes of diabase cross-cutting it (Stepanavan area), but their relationships with the ophiolite

Gabbro-norite Hornblende gabbro

Diorite

Diorite

AR-04-218 AR-03-23 55.09 13.45 8.51 0.15 9.84 8.76 2.59 0.17 0.24 0.04 1.87 100.7 71.6 1.42 207.4 7.82 21.08 0.32 34.9 0.76 0.03 – 0.07 0.04 158.1 562.7 39.6 148.1 15.63 73.35 1.44 3.73 0.57 2.87 0.94 0.35 1.13 0.20 1.31 0.28 0.82 0.13 0.92 0.15 1.03 0.97 1.06

57.41 14.10 8.84 0.14 2.24 4.92 6.36 0.12 0.87 0.16 4.33 99.5 35.6 1.81 212.5 25.22 75.07 1.83 55.59 2.27 0.14 2.62 0.75 0.25 122.6 136.9 15.76 10.09 21.02 50.45 4.62 11.73 1.82 9.05 2.91 0.97 3.68 0.69 4.60 0.97 2.95 0.47 3.25 0.52 0.91 1.00 0.96

(sensu stricto) pillow lavas remain often unclear. The first group of alkaline rocks displays large vacuoles (0.5– 3 mm), filled with carbonates and rarely chlorites, and have both phyric and aphyric

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363

oxides are in wt%, and trace elements and REE in ppm Sevan alkaline series Plagiogranite

Diabase

Trachy-andesite

Basaltic trachyandesite

Andesite

Basanite Trachybasalt Trachyandesite

AR-03-19

AR-03-02

G154

AR-03-17

AR-03-34

G142

AR-05-80

AR-03-33

74.91 12.32 3.58 0.03 0.42 3.05 4.20 0.31 0.21 0.04 0.57 99.6 20.2 2.22 145.2 27.65 71.64 2.15 65.83 2.48 0.10 1.32 1.09 0.62 50.8 1464 5.52 37.19 6.91 9.98 5.38 12.66 1.78 8.36 2.68 0.67 3.50 0.64 4.33 0.94 2.94 0.46 3.30 0.53 0.67 1.26 1.10

46.02 16.29 8.39 0.13 7.73 10.68 3.53 0.37 1.26 0.17 4.61 99.2 66.6 13.97 630.9 23.93 127.4 2.95 285.3 2.91 0.24 1.51 1.01 0.29 179.3 277.0 38.26 54.43 58.20 66.89 7.07 18.32 2.71 12.68 3.53 1.36 3.95 0.67 4.21 0.84 2.45 0.36 2.39 0.37 1.12 1.26 2.00

53.70 14.09 11.35 0.15 4.52 3.64 6.07 – 1.36 0.12 4.85 99.9 46.2 – 28.76 27.82 81.43 1.77 16.09 2.36 0.14 1.61 0.55 0.35 334.9 – 29.48 8.95 102.8 67.20 4.25 11.14 1.67 8.76 2.93 1.14 4.04 0.71 4.78 1.01 2.99 0.46 3.16 0.51 1.01 0.91 0.91

54.27 15.16 12.36 0.19 3.74 4.38 6.63 – 1.33 0.15 1.78 100 39.5 – 49.82 29.88 85.60 1.44 19.01 2.38 0.12 – 0.37 0.27 321.8 – 24.33 5.14 15.61 80.0 3.91 10.93 1.81 9.35 3.26 1.15 4.18 0.75 4.99 1.06 3.11 0.47 3.22 0.51 0.95 0.75 0.82

55.48 14.13 12.45 0.13 4.07 5.49 3.96 0.61 1.17 0.11 2.18 99.8 41.6 5.05 102.5 27.49 54.39 1.69 16.65 1.63 0.12 – 0.33 0.14 347.4 251.7 27.09 16.09 5.23 19.68 2.69 7.17 1.15 6.17 2.32 0.79 3.36 0.62 4.34 0.95 2.82 0.44 2.98 0.47 0.86 0.73 0.61

40.63 14.40 11.70 0.29 4.15 11.40 3.64 1.42 2.06 0.48 10.01 100.2 43.4 31.89 147.0 25.58 153.4 40.63 166.7 3.57 2.99 4.98 4.13 1.43 257.4 33.86 38.26 34.29 61.86 100.1 32.41 64.63 7.64 29.89 6.02 1.97 5.58 0.82 4.69 0.90 2.48 0.35 2.33 0.36 1.04 3.39 9.39

43.80 17.58 9.46 0.11 6.70 4.95 4.09 2.24 1.68 0.44 8.89 99.9 60.5 44.83 330.8 17.37 131.4 17.95 299.0 2.86 1.19 4.63 4.06 0.94 271.7 21.68 31.44 28.76 54.48 91.55 29.46 59.03 6.92 27.11 5.21 1.63 4.45 0.62 3.42 0.61 1.64 0.22 1.44 0.22 1.03 3.56 13.83

51.57 14.34 6.06 0.11 0.77 9.78 6.34 0.56 1.98 1.08 6.65 99.3 21.7 7.93 341.5 52.91 411.2 49.22 168.7 9.03 3.80 4.45 5.15 4.61 286.1 162.0 24.06 13.83 18.34 92.74 48.54 107.1 13.38 56.55 12.98 4.14 12.43 1.84 10.27 1.87 4.80 0.64 4.05 0.60 1.0 2.35 8.09 (Continued)

(Fig. 9e) intersertal textures, with plagioclase megacrysts (c. 5%; 0.5–2 mm), microliths and opaque minerals (3–10%), surrounded by a calcite-chlorite mesostase. The second group (Vedi and Stepanavan

zones) have doleritic (Fig. 9f ) to ophitic (gabbroophitic in central parts of lavas flows) textures, and are mainly composed of plagioclase (c. 40– 55%; 1– 3 mm), clinopyroxene (10–30%; 1– 4 mm),

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Table 1. Continued Groups no.

Sample SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total Mg# Rb Sr Y Zr Nb Ba Hf Ta Pb Th U V Cr Co Ni Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu* (La/Sm) N (La/Yb)N

Stepanavan ophiolitic series Websterite

Hornblende gabbro

Hornblende gabbro

Plagiogranite

Basaltic trachy-andesite

Basaltic trachy-andesite

AR-04-03

AR-04-16

AR-04-45D

AR-04-44

AR-04-20

AR-04-30

53.24 1.03 6.02 0.15 23.18 16.52 0.12 – 0.05 0.03 0.54 100.9 90.0 – 11.79 1.05 – – 3.71 – – – – – 135.1 2804 55.82 361.3 340.3 25.68 – 0.15 0.02 0.15 0.08 0.03 0.13 0.03 0.18 0.04 0.12 0.02 0.12 0.02 0.85 0.0 0.0

47.30 14.39 12.90 0.21 9.11 10.14 2.93 0.19 1.18 0.07 1.76 100.2 60.4 1.12 125.4 20.91 42.74 1.01 32.71 1.21 0.08 – 0.18 0.05 324.7 236.4 51.46 78.0 – 60.28 2.40 6.37 1.08 5.76 2.09 0.96 2.98 0.53 3.53 0.74 2.15 0.32 2.10 0.32 1.17 0.72 0.77

53.77 14.00 8.92 0.15 7.81 6.98 3.34 2.42 0.16 0.05 2.44 100.1 65.6 30.17 213.4 5.79 21.26 2.14 228.1 0.63 0.21 3.61 1.27 0.43 94.47 324.2 31.5 101.7 189.9 60.33 3.06 6.28 0.65 2.30 0.49 0.19 0.55 0.10 0.81 0.19 0.64 0.12 0.91 0.17 1.12 3.94 2.27

75.35 12.20 2.71 0.03 0.77 2.05 5.03 – 0.11 0.02 1.07 99.3 38.1 0.58 91.04 1.23 6.48 0.35 20.58 0.15 – – 0.02 0.01 30.74 421.5 7.71 24.86 189.5 23.13 2.42 3.90 0.42 1.58 0.29 0.35 0.24 0.04 0.22 0.05 0.15 0.03 0.21 0.04 4.05 5.25 7.92

51.53 14.69 14.81 0.23 4.15 4.86 5.74 0.18 1.62 0.13 1.89 99.8 37.7 1.4 61.01 35.91 86.0 1.62 19.58 2.54 0.13 2.29 0.43 0.12 459.8 99.23 38.85 22.76 64.58 130.6 4.23 11.08 1.88 9.88 3.44 1.28 4.69 0.87 5.93 1.28 3.79 0.58 3.89 0.61 0.98 0.77 0.73

48.55 13.29 8.67 0.15 6.86 10.49 4.74 0.24 1.08 0.11 6.01 100.2 63.1 7.94 95.7 26.52 68.43 1.9 21.73 1.83 0.15 1.93 0.19 0.09 305.4 316.7 42.98 109.1 132.8 81.13 2.53 7.37 1.31 7.04 2.55 0.99 3.49 0.65 4.34 0.93 2.75 0.42 2.84 0.44 1.02 0.63 0.60

amphibole (c. 25%; 1–3 mm) and accessory Ti-magnetite (.5– 10%), apatite (c. 3%; prismatic, acicular, 0.5 –1.5 mm) and rarely biotite. Apatites are frequent. Vitreous interstices are filled by

carbonates or carbonate-chlorite assemblages. The tabular plagioclase laths show a transitional zoning with bytownite to labrador (An72-60) or labrador to andesine (An55-32) compositions. Thin

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Stepanavan alkaline series

365

Stepanavan calk-alkaline series

Basaltic trachy-andesite

Basaltic trachy-andesite

Diabase

Olivine basalt

Basaltic trachy-andesite

Basaltic trachy-andesite

AR-06-02

AR-03-53

AR-04-05

AR-04-32

AR-04-40A

AR-04-31

45.37 14.27 13.52 0.32 6.22 4.09 1.53 5.37 3.12 1.31 5.11 100.2 49.8 45.3 157 40.6 268 52.8 608 5.97 3.97 7.04 5.39 1.34 172 – 26.3 6.5 20.4 177 50.3 96.1 12.1 53.3 11.2 3.87 10.5 1.50 8.11 1.49 3.75 0.51 3.18 0.45 1.09 2.83 10.68

48.54 15.01 12.65 0.27 4.25 5.33 3.93 2.69 2.64 1.08 3.16 99.5 42.3 33.17 322.6 44.41 294.4 57.95 578.3 6.51 4.20 2.54 5.98 1.46 94.35 25.77 16.75 – 9.48 137.1 50.59 107.0 12.97 53.32 11.26 4.08 10.41 1.52 8.69 1.57 4.16 0.57 3.64 0.56 1.15 2.83 9.38

50.19 13.91 13.73 0.24 3.27 5.85 5.11 0.42 3.39 0.67 2.94 99.7 34.0 7.62 198.8 51.24 373.5 42.33 156.6 8.01 3.24 2.24 4.65 1.20 201.6 – 31.93 – 14.94 152.7 40.02 85.12 10.85 45.27 10.35 3.39 10.3 1.63 9.50 1.80 4.93 0.70 4.51 0.69 1.0 2.43 5.99

49.15 18.53 10.19 0.16 5.25 8.25 4.36 0.52 0.86 0.14 3.12 100.5 52.7 9.61 520.3 16.0 44.4 2.14 133.9 1.25 0.17 7.22 0.72 0.19 241.5 21.05 29.51 15.55 12.6 150.9 4.93 11.46 1.74 8.35 2.36 0.94 2.60 0.44 2.80 0.57 1.62 0.24 1.60 0.25 1.15 1.31 2.07

49.79 15.80 8.82 0.15 3.54 9.12 3.54 1.24 1.07 0.20 7.27 100.6 48.4 18.12 303.8 24.4 99.18 2.29 239.1 2.69 0.18 3.42 1.46 0.67 279.1 31.91 29.05 22.38 188.1 86.53 7.87 18.51 2.68 12.58 3.47 1.13 3.89 0.65 4.09 0.85 2.46 0.38 2.54 0.39 0.94 1.43 2.09

52.20 17.05 9.28 0.16 3.59 6.56 4.61 1.00 0.94 0.18 5.36 100.9 44.5 18.45 282.3 24.19 95.3 3.32 213.6 2.61 0.26 5.66 1.67 0.58 263.7 73.83 27.36 18.95 170.9 100.0 8.69 18.05 2.53 11.65 3.15 1.04 3.57 0.61 3.96 0.83 2.46 0.38 2.54 0.40 0.95 1.74 2.31 (Continued)

rims of pure albite (Ab – 98%) are also present. The clinopyroxenes are generally chloritized, but its relics are still recognizable and belong to diopside (Wo49En35Fs16). The amphibole is a

kaersutite (Leake et al. 1997), with zonation from kaersutite to ferro-kaersutite from core to rim, respectively. Some samples show abundant calcitefilled veins and pockets.

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Table 1. Continued Groups no.

Vedi ophiolitic series Hornblende gabbro

Sample SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total Mg# Rb Sr Y Zr Nb Ba Hf Ta Pb Th U V Cr Co Ni Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu* (La/Sm)N (La/Yb)N

Diorite

Plagiogranite

Basalt

Vedi alkaline series Basaltic andesite

Basalt

Trachybasalt

Basaltic trachyandesite

Trachydacite

AR-05-113 AR-05-110 AR-05-111 AR-05-114 AR-05-106 AR-05-78 AR-05-104 AR-05-102 AR-04-75 45.09 21.24 4.32 0.07 7.88 14.29 2.12 0.17 0.16 – 5.08 100.4 79.7 2.41 492.7 3.70 4.30 0.08 131.6 0.16 – – – – 89.5 785.2 30.98 130.7 92.4 24.06 0.231 0.71 0.14 0.78 0.36 0.24 0.53 0.10 0.64 0.14 0.38 0.06 0.35 0.06 1.64 0.40 0.44

58.57 16.03 6.66 0.11 5.18 7.08 3.94 0.47 0.33 0.04 2.08 100.5 62.7 4.12 254 10.1 42.3 0.49 57.5 1.3 0.04 1.1 0.42 0.12 185 123 22.8 42.4 18.3 50.7 2.09 4.84 0.72 3.74 1.20 0.42 1.46 0.25 1.65 0.34 1.01 0.16 1.14 0.18 0.97 1.10 1.24

70.45 14.69 4.44 0.08 1.06 3.87 4.47 0.2 0.43 0.09 1.04 100.8 34 1.11 161 13.1 86.5 0.59 33.5 2.35 0.05 1.12 0.48 0.13 44.5 9.1 6.6 6.4 – 39 2.91 6.15 0.81 4.48 1.45 0.69 1.81 0.32 2.11 0.46 1.35 0.21 1.52 0.25 1.31 1.26 1.29

47.5 16.17 8.76 0.15 8.46 8.57 3.99 0.72 0.93 0.09 4.78 100.1 67.6 3.51 134.6 20.78 54.9 0.69 141.7 1.48 0.07 – 0.15 0.06 211.3 416.7 42.44 195.5 14.05 65.52 1.96 6.25 1.09 5.9 2.15 0.9 3.0 0.53 3.48 0.74 2.08 0.32 2.15 0.33 1.08 0.57 0.62

48.3 15.03 10.14 0.16 5.97 8.01 3.96 0.18 1.2 0.12 7.14 100.2 56.0 4.14 110.1 26.91 73.98 2.46 16.2 2.04 0.20 – 0.22 0.10 238.3 324.7 49.08 122.6 81.78 94.28 3.07 8.51 1.47 7.97 2.88 1.12 4.02 0.70 4.67 0.97 2.78 0.42 2.83 0.45 1.0 0.67 0.73

44.58 12.52 9.36 0.12 2.63 15.53 3.83 – 2.35 0.33 9.17 100.4 37.8 0.63 153.5 24.94 160.8 23.22 1097 3.91 1.76 1.53 2.085 0.643 240.7 50.2 28.9 26.28 28.41 106.8 18.44 39.11 5.02 21.85 5.5 1.93 5.72 0.85 4.85 0.88 2.30 0.31 1.95 0.29 1.05 2.11 6.39

44.64 15.41 11.99 0.14 4.85 7.85 4.24 0.96 3.67 0.85 5.04 99.6 46.6 10.48 926 36.26 318.8 67.52 444.9 6.81 4.88 1.25 4.60 1.18 219.8 4.22 31.4 21.81 50.82 146.8 49.35 109.7 14.15 59.18 12.52 4.15 10.96 1.48 7.84 1.32 3.26 0.42 2.49 0.35 1.08 2.48 13.39

50.39 16.2 7.76 0.13 5.05 7.16 3.24 1.96 2.36 0.64 5.07 99.9 58.4 27.1 643 22.4 260 43.3 659 6.03 3.29 6.71 8.39 1.8 135 136 64.3 126 42.2 134 50.7 91.8 9.83 42.9 8.47 2.67 7.36 0.99 5.01 0.81 1.92 0.25 1.46 0.22 1.03 3.77 23.44

59.61 17.48 7.64 0.12 1.11 1.89 6.37 2.4 0.72 0.25 2.14 99.7 23.16 64.89 260.4 58.18 680.5 82.24 422 14.76 5.99 4.30 12.28 2.78 5.26 66.9 4.98 – 9.93 167.9 74.8 142.6 15.84 59.79 12.64 3.68 11.69 1.877 10.97 2.08 5.77 0.86 5.84 0.88 0.93 3.73 8.64

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Table 2. Step

40

367

Ar/39Ar dating results of single grain AR-05-110 gabbro amphiboles from the Vedi ophiolite

Laser power (mW)

Contamin. atmosph. (%)

AR-05-110(A) (K103) 1 500 99.437 2 620 73.967 3 690 134.382 4 850 93.450 5 1050 38.568 6 1150 19.355 7 1200 24.516 8 1300 17.488 9 2500 17.374 AR-05-110(B) (K119) 1 500 82.734 2 585 83.526 3 670 29.902 4 760 17.905 5 806 15.376 6 860 20.078 7 1800 13.653 AR05 110(C) (K132) 1 484 81.313 2 520 70.539 3 600 30.721 4 660 22.277 5 710 32.287 6 785 25.007 7 862 33.799 8 2000 18.441

39

Ar (%)

37

ArCa/39ArK

40

Ar*/39ArK

Age (Ma +1s)

1.03 0.58 0.54 0.91 11.19 39.90 16.76 5.37 23.73

16.751 5.709 7.676 17.798 34.521 40.983 44.907 35.012 37.852

0.237 3.169 2 0.599 3.440 3.920 3.601 4.116 4.009

9.792+116.946 126.879+159.432 2+ 2 24.696+125.797 137.322+10.256 155.682+5.710 143.488+6.732 163.105+17.008 159.066+4.931

1.70 1.36 9.58 27.30 23.98 7.09 28.99

13.380 18.310 50.630 48.774 43.418 42.179 40.579

30.304 15.052 30.333 29.942 30.465 29.367 30.589

178.902+22.943 91.078+30.863 179.068+4.364 176.868+2.342 179.810+2.604 173.631+5.472 180.506+2.321

1.51 1.35 8.14 17.99 12.26 18.02 11.28 29.44

12.110 11.216 45.281 38.256 46.619 52.601 62.023 34.698

6.266 3.010 4.824 4.168 3.678 4.278 5.175 4.251

243.436+50.894 121.065+52.838 190.263+10.923 165.539+4.834 146.840+8.284 169.727+10.053 203.359+6.919 168.686+2.481

A few dacitic dyke-like bodies crop out among the basaltic pillow flows in the Vedi valley. As in the pillow basalts, plagioclase is the main mineral phase and Fe-oxides (c. 5 –10%) are the accessory minerals (Fig. 9g). Some plagioclase unzoned phenocrysts (1–2 mm, oligoclase-andesine?) are distributed in the fine-grained (,0.2–0.5 mm) hyalopilitic or micro-cryptocrystalline devitrified groundmass made of albitic plagioclase, opaque microlites, and carbonate-quartz-chlorite aggregates. There are also thin hydrothermal veins made of calcite. Calc-alkaline lavas of Stepanavan zone. These consist of large pillow-lavas of basaltic and basaltic andesitic compositions with microcryptocrystalline (Fig. 9h) to intersertal textures of large andesine-oligoclase plagioclase phenocrysts (2–7 mm) and microliths, and minor augite (Wo36-38En42-43Fs13-15) clinopyroxenes. These lavas overlie Upper Cretaceous limestones, uncomformably lying on the ophiolite s.s.

Major– trace – REE geochemistry The geochemical analyses of the ophiolitic rocks from the Sevan ophiolite are of relatively alkaline composition in comparison to MORB. Major element data of pillow- lavas, diabase and gabbros show that they have predominantly basalt to trachybasalt compositions. Major elements. Major element analysis of plutonic rocks ranges from gabbros to granites ( plagiogranites) with intermediate dioritic compositions (Fig. 11a). These magmatic rocks appear to define a large trend (Le Maitre et al. 1989). Similarly, in the alkali v. iron and magnesium (AFM) diagram (Fig. 11b) most rocks lie close to the limit between the tholeiitic and calc-alkaline fields. † Overall, the rocks of the ophiolitic suite are enriched in MgO and more depleted in TiO2, K2O and P2O5 relative to the alkaline suite (Figs 11 & 12; Table 1). The volcanic rocks from the different studied areas plot in the

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Fig. 7. 40Ar/39Ar age spectra and isochrons undertaken on single amphibole (hornblende) grains of a gabbro sample (AR-05-110) from Vedi ophiolite.

ARMENIAN OPHIOLITES

Fig. 8. Representative geological logs of the Stepanavan, Sevan and Vedi ophiolites. 369

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Fig. 9.

Y. ROLLAND ET AL.

ARMENIAN OPHIOLITES

same compositional range (from basalts to andesites and trachyandesites), and are relatively richer in Na2O than the plutonic rocks of the same series. † The studied alkaline lavas from different zones plot in the same range, varying compositionally from basanite-trachybasalt to basaltic trachyandesite and trachyandesite, and are clearly in the calc-alkaline/alkaline domain of the AFM diagram (Fig. 11a, b). One of the most significant features of the alkaline lavas is their higher TiO2, K2O and P2O5 contents. † The arc-type calc-alkaline lavas, having trachybasalt and basaltic trachyandesite compositions in TAS diagram (Fig. 11), plot essentially in a transitional position between ophiolitic and alkaline domains in Harker’s diagrams (Fig. 12), except lower TiO2 and higher Al2O3, depend on the abundance of plagioclase in such rocks. Regarding now the spread of compositional variations in major elements within series, it appears that SiO2, Al2O3, MgO, correlate relatively well with variations in TiO2, an element considered as more immobile during alteration processes (e.g. Staudigel et al. 1996), (Fig. 12). Other elements, and particularily Na have relatively scattered compositions, even in individual magmatic suites, which could be ascribed to alteration or relatively complex magmatic processes. This is also supported by thin section observations and previous studies of the Armenian ophiolites (e.g. Palandjyan 1971; Abovyan 1981; Ghazaryan 1994), which indicate that the whole magmatic sequence has been affected by oceanic low-temperature hydrothermal alteration events. These processes induced modification of the most lithophile elements, as revealed by the increase of LOI (loss on ignition) in whole-rock chemistry (Table 1). Trace elements. High field-strength elements (HFSE) are not mobilized during alteration and

371

their contents reflect, without ambiguity, those of their parental magma (Staudigel et al. 1996). Trace elements contents confirm the presence of three clearly distinct magmatic suites, as defined in previous section. † Basalts and gabbros of the ophiolite suite show strong enrichments in LILE (Large Ion Lithophile Elements: Ba, Rb, K and Th) are close and up to ten times MORB values and bear negative anomalies in Nb-Ta and Ti (Fig. 13a, b), which is generally indicative of volcanic island arc environments (Taylor & McLennan 1985; Plank & Langmuir 1998). † Overall, the concentrations of each element in the alkaline basalts exceed those in the basalts from ophiolitic series (Fig. 13c). Moreover, alkaline series basalts are characterized by high abundances of LILE, high field strength elements (Nb, Ta, Zr and Ti), and light rare-earth elements (LREE). † The calc-alkaline suite rocks show strong depletions in Nb and Ta, relative to Th and La, and slight Ti negative anomaly (Fig. 13d). They globally show slightly stronger enrichments in LREE and LILE relative to the ophiolite suite rocks. These differences in normalized element patterns support that these basalts are not petrogenetically related and most likely derived from melts formed in different tectonic settings: (1) N-type MORB (and/or Back-arc basin type); (2) Ocean-island and/or within-plate alkali basalts; and (3) volcanic island arc. REE geochemistry. In the chondrite-normalized rare earth element (REE) diagrams (Fig. 13), analysed ophiolite basalts and gabbros have flat and parallel REE spectra in chondrite-normalized plots ((La/ Yb)N ¼ 0.6–0.9), showing some slight depletions in LREE and a slight enrichment in MREE (Fig. 13e, f ). No extensive Eu anomalies were

Fig. 9. (Continued) Microphotographs of representative magmatic rock types from different ophiolite complex. Plutonic and volcanic ophiolite series: (a) subautomorph granular texture of a cumulate banded websterite (sample AR-04-36, Stepanavan area, Cheqnagh valley); (b) coarse-grained hornblende gabbro with normally zoned plagioclases (sample AR-05-110, Vedi area, massif of Qarakert); (c) xenomorph granular texture of a microcline (Mc) bearing plagioclase rich leucogranite (sample AR-05-109, in the same massif ); (d) aphyric, intersertal (spilitic) and variolitic basalt composed of mainly albitized plagioclase, Ti-magnetite and hematite microlites, in a devitrified groundmass (sample AR-05-106, Vedi area, Khosrov valley). Alkaline series: (e) aphyric, intersertal basalt, totally devoid of phenocrysts, and composed of carbonatized plagioclase microlites and opaque minerals (c. 5%) in a chlorite-carbonate groundmass (sample AR-05-80, Sevan area, Tsapatagh valley); (f ) doleritic texture in a trachybasalt composed of plagioclase, chloritized clinopyroxene, kaersutite (Krs), Ti-magnetite and apatite (sample AR-05-104, Vedi area); (g) phyric trachydacite with a hyalopilitic to cryptocrystalline texture (sample AR-04-75, Vedi valley). Calc-alkaline series: (h) olivine-bearing, plagioclase phyric (15– 40%) basalt with a microcrystalline (plagioclase, quartz, opaque minerals) texture from pillow lavas suite (sample AR-04-32, Stepanavan area, Herher valley), in which the olivine phenocrysts are entirely pseudomorphosed to quartz and rims of iron oxides. From a–c under crossed nichols, and d –h under parallel nichols.

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Fig. 10. Microphotographs of plutonic rocks from the Vedi ophiolite. (a) Gabbro (sample AR-05-112A) composed of euhedral to subhedral plagioclase (Pl) with normal zoning with green euhedral amphibole (Amph), containing clinopyroxene inclusions (Cpx). (b) In the same sample, subhedral plagioclase coated by anhedral amphibole. (c) In gabbro sample AR-05-110, detailed back-scattered image showing zoning in large amphibole crystals, from magnesio-hornblende (1), to edenite (2) and pargasite (3) toward the rim. (d) Subhedral amphibole crystallized over a previous anhedral amphibole core (in sample AR-05-110), with some interstitial quartz (Q). (e) Plagiogranite (sample AR-05-111) formed by subhedral plagioclase, de anhedral quartz and little proportion of chloritized þ epidotized (Ep) amphibole. (f ) Leucocratic plagiogranite (sample AR-05-109), showing anhedral granular structure of plagioclase, K-feldspar (with microcline twining, Mcc) and quartz. Observations are under parallel (//) or crossed () nicols all at the same scale.

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(a)

373

(b) Gabbros Diorites Plagiogranite Ophiolitic lavas Alkaline lavas

Gabbros Plagiogranite Ophiolitic lavas Alkaline lavas Arc-type lavas

Gabbros Diorites Plagiogranite Ophiolitic lavas Alkaline lavas Trachydacite

Fig. 11. Plots of magmatic rocks (ophiolite, ‘OIB’ and arc series) in the (a) (Na2O þ K2O) v. SiO2 (Le Maitre et al. 1989) and (b) AFM (Irvine & Baragar 1971) diagrams.

Gabbros Diorites Plagiogranite Ophiolitic lavas Alkaline lavas

Gabbros Diorites Plagiogranite Ophiolitic lavas Alkaline lavas Trachydacite

Gabbros Plagiogranite Ophiolitic lavas Alkaline lavas Arc-type lavas

Fig. 12. Harker variation diagrams showing the compositions of the three (ophiolitic, alkaline and calc-alkaline) series.

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(a)

(e)

(b)

(f)

(c)

(g)

(d)

(h)

Fig. 13. Trace and REE plots of the three studied magmatic suites. The multi-element spiderdiagrams are normalized to the N-MORB values of Sun & McDonough (1989), and REE plots are normalized to the Chondrite values of Evensen et al. (1978). Patterns for the studied magmatic rocks: ophiolitic volcanic (a, e) and plutonic (b, f ) series; OIB type alkaline series (c, g), and arc type calk-alkaline series (d, h).

observed (Eu/Eu* ¼ 0.95–1.15), which show that almost no plagioclase fractionation has occurred. Thus, plagioclase likely remained stagnant and was enriched in the final liquid. The concentration

of REE varies from 8 to 30 times chondrite compositions in volcanic rocks, and from 1 to 15 times in the gabbros. Only one gabbro sample (sample AR-03-25) shows an extreme 60 times-chondrite

ARMENIAN OPHIOLITES

REE concentration, which may be explained by fluid alteration as it displays a flaser –structure. These features are interpreted as a result of extreme crystal fractionation involving plagioclase, clinopyroxene, orthopyroxene and, to a lesser extent, olivine accumulation (Pallister & Knight 1981). The websterite and gabbronorite have the lowest concentrations of REE (0.1– 0.9 and 1–5 times chondrite respectively) with patterns characterized by depletion in LREE (Fig. 13f). One hornblende gabbro (sample AR-04-45D from Stepanavan ophiolite) is characterized by LREE enrichment ((La/Yb)N ¼ 2.27) and some depletion in MREE (a convex downward pattern) with smaller positive Eu anomalies (Eu/Eu* ¼ 1.12). The REE patterns of the diorites (6-20 times chondrite) and plagiogranites are parallel to those of the gabbros, with smaller enrichment in LREE ((La/Yb)N ¼ 1.1). While the plagiogranite from Stepanavan (sample AR-04-44) is characterized by more depletion in the middle to heavy REE compared to other plagiogranites, and strongly positive Eu anomalies (Eu/Eu* ¼ 4.05) due to its cumulative nature that ascribed to high plagioclase contents. In contrast, chondrite-normalized REE patterns of alkaline lavas (Fig. 13g) show huge LREE enrichments and HREE depletion ((La/Yb)N ¼ 6–14), being characteristic of intraplate continental basalts, as compared to ophiolite lavas. Meanwhile, no extensive Eu anomalies were observed (Eu/ Eu* ¼ 0.95–1.15). The pattern of a trachydacites (sample AR-04-75) is parallel to those of the basanite-trachyandesite series having the highest overall REE concentration. Chondrite-normalized REE patterns of calcalkaline lavas are strongly parallel and form a narrow domain (Fig. 13h). They have similar HREE contents as volcanics of previous series with significantly more depleted LREE contents than alkaline series rocks ((La/Yb)N ¼ 2.1–2.3). These differences of trace elements behaviour between the three studied series further support that these basalts are petrogenetically unrelated and, most likely derived from melts formed in different tectonic settings. 40

Ar/39Ar dating

Three analyses have been done on amphibole single grains from one gabbro sample (AR-05-110) from the Vedi ophiolite, which is described in the ‘Petrography and mineral chemistry’ section. These are presented in Table 2 and Figure 14. In the first dating (k103), a plateau age is defined by the four last steps at 154.7 +6.9 Ma (2s) comprising 86% of 39Ar. Further, using all the steps, an

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isochron age of 154.4+8.1 Ma (MSWD ¼ 1.4) is obtained. The initial 40Ar/36Ar ratio is of 264+7 shows a slight shift from the air value which shall be ascribed to the large error bars in the low temperature steps. In the second experiment (k119) a wellconstrained plateau of 178.7+2.6 Ma (2s) is obtained, with 97% of 39Ar. The average 37ArCa/ 39 ArK ratio is high as for the latter experiment, c. 40 in low temperature steps, decreasing steadily to c. 30 in high-temperature steps. An isochron is obtained using the five steps used in the estimate of the plateau age, with an initial 40Ar/36Ar ratio close to the atmospheric value. Even with all of the steps, including the lower temperature ones, except step 2 (which has a large error and represents only 1.4% of degased 39Ar) we calculate a similar within-error isochron age of 177.6+2.6 Ma. The third experiment (k132) provides a more disturbed Ar spectra, which does not provide any plateau due to one high-temperature step, featured by a higher age. However, we obtain a weighted average age of 172+6 Ma, using steps 4–8, or a pseudo-plateau age of 167.3+6.6 when excluding step 7. The calculated isochron age including all steps and that obtained only with the HT steps provide a similar within error age of about 160 Ma (see Fig. 7). The initial 40Ar/36Ar ratio varies from 346+7 with all steps to 318+3 with the high temperature ones, which is slightly lower than the air value, and might be ascribed to some disturbance of the Ar system. Age variations within samples are ascribed to the contribution of finely inter-fingered mineral components within the dated amphiboles, following the works done by Villa et al. (2000). In the three datings, 37ArCa/39ArK ratios are close the obtained electron micro probe (EMP) values of the amphiboles, with some slight variations. The clearest case is the experiment K119, which also shows the flatter age pattern. In this sample, the slight decrease from higher 37ArCa/39ArK in low temperature steps towards lower 37ArCa/39ArK in higher temperature steps is ascribed to core-rim variations in amphibole. Indeed, the petrographic analyses show the presence of several mineral generation (see Fig. 10c, d) and EMP analyses show higher Ca/K ratios in the rim (actinolite) v. the in the (hornblende) core of minerals. Therefore, it is likely that the amphibole rim contributes more to the low temperature steps, while the core contributes more to the high temperature steps, as featured in Villa et al. (2000) and Rolland et al. (2006). Such process of amphibole recrystallization may explain the large range in obtained ages within the same sample. Plateau ages are always similar (within error) to isochron ages, with initial 40Ar/36Ar ratios very close to the air value. Thus, all the ages

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obtained here have a geological meaning. However, the slight disturbances observed in initial 40Ar/36Ar ratio of the most disturbed Ar spectra (K132) is a good argument to explain part of the age spread by some fluid-rock interaction process. The flatter Ar spectrum (k119) also provided the older age (178.7+2.6 Ma). It is therefore probable that this age is very close to the initial magmatic crystallization age. This age is also in agreement with other geological data, comprising palaeontological age of the pelagic limestones (Sokolov 1977) and radiolarian interbedded in the pillow lavas (Danelian et al. 2008) all attributed to the Middle Jurassic period. Therefore, the perturbed (k 132) and c. 155 Ma (k103) amphiboles agree for a later or long-lasting alteration process, related to the crystallization of greenschist to epidote amphibolite minerals within very heterogeneous and discrete zones. In conclusion this is the first Ar-dating undertaken on the plutonic part of the Vedi ophiolite. The initial crystallization age of the gabbro is at the limit between the Lower and Middle Jurassic (Toarcian-Aalenian), and predates shortly the deposition of radiolarites and limestones. This age is also slightly older than similar Ar-dates obtained on one gabbro from the Sevan ophiolite (Galoyan et al. 2009), and other evidence for the formation of the Sevan ophiolite in the Middle Jurassic (Zakariadze et al. 1990).

Discussion Ophiolites of the Lesser Caucasus region of Armenia are generally separated into three distinct zones: (1) the Sevan–Akera zone in the North (Knipper 1975; Adamia et al. 1980); (2) the Zangezur zone in the SE (Aslanyan & Satian 1977; Knipper & Khain 1980; Adamia et al. 1981) of Armenia, respectively; and (3) the Vedi zone to the south (Knipper & Sokolov 1977; Zakariadze et al. 1983). Due to the importance of Cenozoic volcanism that covered most of the surface of Armenia (Fig. 2), it is still difficult to conclude if whether the different ophiolites correlate with each other, or if they represent various suture zones delimitating several continental micro-blocks. As emphasized in the following discussion, we will show further that these ophiolites show some similarities and differences in their age, structure, lithological successions and geochemical features; but these features remain compatible with an origin from a sole oceanic domain. This domain opened in the Lower– Middle Jurassic and has registered several phases of magmatic emplacement, evidenced in each of the different investigated geographic zones. These

correlations provide insights into the evolution of the Tethyan domain, and in particular allow us to propose a geodynamical model for the obduction of the ophiolite over the Armenian block. In the following discussion, we will evaluate the following points. † Petrographically and geochemically, the Armenian ophiolites are similar to island-arc tholeiites. Such geochemical features are typical for oceanic crust, formed on a back-arc setting with the melting of a shallow asthenospheric source contaminated by slab-derived fluids (Tarney et al. 1981; Saunders & Tarney 1984). Such a hypothesis has already been proposed for ophiolitic gabbros from Turkey (Kocak et al. 2005), but has to be further evaluated considering isotopic compositions and partial melting rates constraints. † Lower Cretaceous Alkaline lavas of variable thicknesses overlain this ophiolitic sequence. Their origin has to be considered. (i) Do they also derive from the same ophiolitic series? (ii) Did they formed in an island-arc setting or in an oceanic plateau environment? The occurrence of alkaline magmatism prior to obduction in the the Late Lower Cretaceous may be of significant importance for the obduction model of the ophiolite crustal sequence. † Finally, the calc-alkaline lavas are Upper Cretaceous in age. Then, these volcanic arc-related series may be formed during closure of the Neo-Tethys Ocean. Their geochemical feature will be considered to evaluate this hypothesis.

Significance of Armenian ophiolites: MOR or back-arc setting? We could not find any outcrops of the Zangezur ophiolites. These might be very thin stretched rocks, which we believe mark the base of a major thrust in SE Armenia. Petrographically and geochemically, the Armenian ophiolite rocks from the different studied localities (Stepanavan, Sevan and Vedi) share the same lithological and petrological features. They have geochemical features intermediate between MOR basalts and island arc tholeiites. However, as shown by field relationships in the ‘Geological setting’ section, it is clear that these ophiolites were emplaced at oceanic spreading centres. The association of exhumed serpentinites, gabbros and plagiogranites at sea-floor levels, overlain by radiolarites is the result of intra-oceanic tectonics (Lagabrielle et al. 1984), typical of LOT ophiolites (Nicolas & Jackson 1972; Nicolas 1989; Lagabrielle & Cannat 1990). In this context, the paucity of sea-floor lava spreading explains that deep sections of the oceanic crust are exhumed by

ARMENIAN OPHIOLITES

extensional faults. Faults and shear zones have guided magmatic infiltration, which explains the intense hydrothermal activity observed in the serpentinites (‘listwenites’) and the pervasive alteration of lavas and diabase dykes. In addition, Armenian ophiolitic series are shown to be tholeiites with slight calc-alkaline character, ranging from basalts to basaltic andesites and basaltic trachyandesites. Spider diagrams show clear Nb–Ta negative anomalies (Fig. 13a, b), LILE enrichments and flat to slightly LREE-enriched spectra. These observations do not support a geochemical ‘normal’ ophiolitic crust and are more probably in agreement with typical volcanic arc settings, in which enrichments in LILE, LREE result from slab fluids/ melts contamination (Pearce et al. 1984). For these reasons it appears most likely that the Armenian ophiolites were emplaced in a back-arc setting with contamination provided by slab fluids. We did not find any evidence of the associated Upper Jurassic volcanic arc system, which we suggest might have been subducted or eroded during the obduction. The very strong (greenschist to epidote amphibolite facies) and long-lasting (.20 Ma) hydrothermal imprint which is evidenced in the 40Ar/39Ar dating experiments (Fig. 7) is also in agreement with a LOT environment. In such context, the time for magmatic crystallization and hydrothermalism is longer due to slow-spreading rate. In the present case, a time of .23 Ma for greenschist facies hydrothermalism may suggest accretion rates ,1 cm a21. The time of formation of the Vedi (178.7 +2.6 Ma, this work) and Sevan (165.3+1.7 Ma; Galoyan et al. 2009) ophiolites, constrained by 40Ar/39Ar on gabbro amphibole, show that the age of oceanic crust is older in the southern leading edge of the obducted sequence than on its northern side. This age difference is also seen in the age of radiolarians that are younger in Sevan and Stepanavan than in Vedi (Danelian et al. 2007, 2010). Such difference is in agreement with the Vedi ophiolite being very close to the rim of the former back-arc basin, the obduction may then have been triggered along the margin of the back-arc domain, directly south of the Vedi oceanic crust.

Origin of alkaline lavas Mineral chemistry and geochemistry of the alkaline volcanic series of Sevan, Stepanavan and Vedi ophiolites is similar to that of OIBs. As shown in the Mineral Chemistry section, pyroxenes bear an alkaline composition. The alkaline lava samples show strong enrichments in incompatible elements (up to 100 times chondrite values). In the Vedi area, Satian et al. (2005) already pointed out the alkaline character of the lava series, which they

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interpreted as intra-continental rifting. However, these lavas were emplaced above, and formed after the ophiolites. An age of 117.3+0.9 Ma was recently obtained with the 40Ar/39Ar dating method undertaken on single-grain amphibole by laser step-heating (Rolland et al. 2009). Moreover, they are interstratified and overlain by reef limestone, which suggests a shallow marine environment just after emplacement. Thus, we interpret these series as resulting from a plume event that occurred in an intra-oceanic setting. This plume was sufficiently large to overlain the various ophiolite zones present in Armenia, over a surface of .5000 km2, thus this plume event might have formed a plateau, which by itself may explain the obduction (Rolland et al. 2009). Such alkaline magmatism is widely documented in the Middle-East region, along the Arabian and Indian platforms, in relationship with the formation of the Neo-Tethys ocean (Lapierre et al. 2004). Similar Cretaceous alkaline series are found above the Iranian ophiolite (Ghazi & Hassanipak 1999), and in Turkey (Tu¨ysu¨z et al. 1995; Tankut et al. 1998). However, it is still difficult to relate these alkaline events due to their geographical and temporal distance, the paucity of radio-chronological and Sr, Nd, Pb isotopic data.

Reconstruction of the ‘ophiolite’ history From all the available geological data, we can propose the following model for the evolution of the Armenian Ophiolite (Fig. 14). † The SAB is of Gondwanian origin according to lithological associations found in central and SE Armenia (Knipper & Khain 1980; Kazmin et al. 1987; Aghamalyan 2004). Therefore, it is likely that the Sevan oceanic basin opened as a response to the north-dipping subduction of Neo-Tethys (Fig. 14). Emplacement of the ophiolite occurred in the Early to Middle Jurassic (Galoyan et al. 2009) in an intra-oceanic back-arc setting between the Armenian block and the active Eurasian margin. The older age of the Vedi ophiolite, with respect to that of Sevan implies that it should be at the rim of the back-arc system. † Emplacement of an Oceanic Plateau above the back-arc oceanic crust during the late Lower Cretaceous (114 –102 Ma). † The calc-alkaline lavas unconformably overlain the ophiolite and related alkaline series (Galoyan et al. 2007). These lavas have similar geochemical features as volcanic arc series. Their emplacement is bracketed in the Upper Cretaceous, as for the high pressure metamorphism constrained in the Stepanavan area (Meliksetyan et al. 1984 and references therein),

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Y. ROLLAND ET AL. Intra-oceanic subduction

Eurasian Active Margin

Arc volcanism

S South Armenian Block +

+ +

(1)

N Vedi

+

Sevan/ Stepanavan

+

+

+

?

?

+

+

+

+ +

+ + Back-arc extension and decompressional melting

Lower–Upper Jurassic Slab fluid metasomatism

Eurasian Active Margin Intra-oceanic subduction

S

Arc volcanism (reduced)

South Armenian Block

N Oceanic plateau

+

+ +

(2)

+

+

+

+

+

?

Sevan/ Stepanavan

Vedi

+

+

+ +

+

Hot spot Alkaline magmatism Crustal thickening

+

Slab retreat

Late Lower Cretaceous (Albian-Aptian, 120–115 Ma)

Eurasian Active Margin S

South Armenian Block

Volcanic arc subduction

N

Erosion

Vedi

+

+ +

(3)

+

? Subduction of the arc series (Stepanavan Blueschists) c. 95–90 Ma

Slab retreat

Arc volcanism

Frontal flysch sequence

Obduction

Erosion

Vedi

Erosion

South Armenian + Block+

Eurasian Margin

Shallow to pelagic carbonated sedimentation +grawackes

+ +

N

Sevan/ Stepanav an

+

+ +

+ +

+ +

+

+ +

Turonian - Lower Coniacian 95–88 Ma

Frontal molasse sequence

Slab retreat

Obduction Stopped

S

Upper Coniacian Uncomformity

Erosion

Erosion-slumps

South Armenian Block +

Blocking of the subduction at c. 73–71 Ma

Eurasian Margin N

Vedi

+

+ + +

+

Santonien 83 Ma

Fig. 14.

+ +

+ +

S

(5)

+

+ +

+

Cenomanien Upper Cretaceous c. 102–95 Ma

(4)

Sevan/ Stepanavan

+

anavan

+

+

+

Sevan/S tep

+

+

+ +

+ +

ARMENIAN OPHIOLITES

constrained at c. 95 –90 Ma (Rolland et al. 2007). Therefore this magmatic event can be related to the N-dipping subduction of the NeoTethys ocean below the Sevan–Akera back-arc prior to the obduction of the Armenian Ophiolites onto the SAB. † Then, the SAB enters the subduction zone in the Cenomanian (102 –95 Ma), which triggers a ‘collision’ with the thickened plateau. During this process, the volcanic arc has probably been subducted below the Oceanic plateau and metamorphosed in the blueschist facies (Rolland et al. 2007). The large variety of lithologies comprising metabasites, marls and conglomerates in a pelitic matrix, within the Stepanavan blueschists, is in agreement with such a scenario. † The obduction of the ‘ophiolite’ section over the SAB is further constrained by the Lower Coniacian frontal flysch sequences, found below and in front of the Vedi obducted sequence. The calc-alkaline series found above the Stepanavan ophiolite show that a volcanic arc was active during this time above the obducted sequence. † The end of the obduction is constrained by Upper Coniacian fauna in sediments unconformably overlying the ophiolite. The subduction below the Eurasian margin may stop at 73– 71 Ma, as shown by 40Ar/39Ar age of MT-LP metamorphism in the Stepanavan blueschists and the general tectonic uplifting of the region, witnessed by erosion and absence of sedimentary record during the Late Cretaceous-Palaeocene (Rolland et al. 2007). This 73– 71 Ma event is thus interpreted as the insight of ‘collision’.

Conclusions: geodynamic significance of Armenian ophiolites (1)

(2)

The Armenian ophiolites show evidence for the obduction of a single oceanic crust sequence above the SAB. Similar geological, petrological, geochemical and age features are found in the studied Armenian ophiolitic massifs (Sevan, Stepanavan and Vedi). The age of the ophiolite is constrained by 40 Ar/39Ar dating experiments undertaken on gabbro amphibole in the Vedi area, which

(3)

(4)

(5)

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provided a magmatic crystallization age of 178.7+2.6 Ma, and further evidence of greenschist facies crystallization during hydrothermal alteration until c. 155 Ma. As compared to the Sevan ophiolite, where oceanic crust formation is dated at 165.3+1.7 Ma with the same method, the oceanic crust sequence in the Vedi area is significantly older, suggesting that the initial lateral oceanic age relationships are preserved. The oceanic crust sensu stricto corresponds to a Lherzolite Ophiolite Type (LOT), formed in the Early-Middle Jurassic by slow-spreading accretion. The long-time span of alteration (.20 Ma) recorded in 40Ar/39Ar ages of dated amphiboles is also suggestive of such a slow spreading environment. In addition, the hybrid arc-MORB geochemical signature of the ophiolite rocks strongly suggests they formed in a back-arc basin by melting of an asthenosphere source contaminated by subducted slab-derived products. Alkaline volcanic series with OIB-type geochemical features are found above the ophiolite sequence in each of the studied areas. Hot-spot related magmatism may have led to the formation of Oceanic island(s) or even Oceanic plateau(s), with significant crustal thickening dated to the Albian –Cenomanian (114 –95 Ma) like in many other parts of the world. These alkaline series are also locally overlain by calc-alkaline volcanic series, which were likely formed in a supra-subduction zone environment. Further evidence of this subduction is provided by blueschists series dated at 95–90 Ma. Therefore Plateau formation and volcanic arc formation shortly pre-dated the obduction, which occurred in the Coniacian –Santonian (88–83 Ma). The obduction was followed by final collision of the SAB with the Eurasian margin at c. 73 –71 Ma.

This work was supported by the Middle East Basins Evolution project jointly supported by a consortium including oil companies and the CNRS. Many thanks to the MEBE programme coordinators Eric Barrier and Maurizio Gaetani for their support and encouragements. Special thanks to Marie-Franc¸oise Brunet for her

Fig. 14. (Continued) Geodynamic reconstruction of the ophiolite formation from the Lower Jurassic to the Upper Cretaceous periods. (1) Formation of the ophiolite in a back-arc setting between the SAB and the Active Eurasian margin; (2) Formation of OIB-type series above a hot-spot in the late Early Cretaceous; (3) Obduction of the ophiolite initiated directly after hot-spot magmatism in the early Late Cretaceous leading (4) to its emplacement above the SAB from the Early Coniacian to Santonian times (5). The final blocking of the north-dipping subduction below the Eurasian margin, and insight of the collision at 73– 71 Ma is constrained by the study of blueschists in the Stepanavan area.

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inalterable patience and Analytical data were acquired with the help of the Geosciences Azur Laboratory, in which we thank M. Fornari and G. Fe´raud for their involvement during data acquisition. We also thank the support of the French Embassy at Yerevan for the MAE PhD grant granted to G. Galoyan. This paper was improved by detailed reviews undertaken by R. Hebert and P. Agard.

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New and revised radiolarian biochronology for the sedimentary cover of ophiolites in the Lesser Caucasus (Armenia) T. DANELIAN1,2*, G. ASATRYAN1,3, L. SAHAKYAN3, GH. GALOYAN3,4, M. SOSSON4 & A. AVAGYAN3 1

Universite´ Pierre-et-Marie-Curie (Paris VI), CNRS-UMR 5143 Pale´obiodiversite´ et Pale´oenvironnements, C. 104, 4 place Jussieu, 75005 Paris, France

2

Present address: Universite´ Lille 1, CNRS-UMR 8157 Ge´osyste`mes, UFR des Sciences de la Terre – baˆtiment SN5, 59655 Villeneuve d’Ascq cedex, France 3

Institute of Geological Sciences, National Academy of Sciences of Armenia, 24a Baghramian avenue, Yerevan, 0019, Armenia

4

Universite´ de Nice-Sophia Antipolis, CNRS-UMR Ge´osciences AZUR., Parc Valrose, 28 Avenue de Valrose, 06108 Nice Cedex 2, France *Corresponding author (e-mail: [email protected]) Abstract: In order to improve our understanding of the palaeogeographic and geodynamic evolution of the Tethyan realms preserved in the Lesser Caucasus we here review the existing data for the sedimentary cover of ophiolites preserved in Armenia. Particular attention is given to those dated sedimentary rocks that are in direct genetic contact with ophiolitic lavas, as they provide constraints for submarine oceanic activity. The oldest available ages come from the Sevan–Akera suture zone that point to a Late Triassic oceanization. Data from both the Sevan and Vedi ophiolites provide evidence for Middle Jurassic (Bajocian) submarine activity, that continued until at least the Late Jurassic (Mid/Late Oxfordian to Late Kimmeridgian/Early Tithonian), as dated recently in Stepanavan and in this study for the Vedi ophiolite.

Radiolarian biochronology has become invaluable in the last 30 years for understanding the geodynamic and palaeoenvironmental evolution of Tethyan oceanic basins and continental margins (De Wever & Dercourt 1985; Karakitsios et al. 1988; Bragin & Tekin 1996; Al-Riyami et al. 2000, 2002; Tekin et al. 2002; Vrielynck et al. 2003; Bragina & Bragin 2006; Danelian et al. 2008b). More particularly, palaeontological dating of the sedimentary cover of ophiolites allows us to constrain the timing of opening of the various parts of Tethys (De Wever et al. 1987; Chiari et al. 1997, 2000; Danelian et al. 2000; Bill et al. 2001; Go¨ncu¨oglu et al. 2006). Dating radiolarites overlying ancient oceanic crust preserved in the Lesser Caucasus is key to understanding the geodynamic evolution of the greater area between Eurasia and the South-Armenian Block (Fig. 1), a micro-continent detached from Gondwana during the Late Palaeozoic–Early Mesozoic time (Knipper 1975; Knipper & Khain 1980). Two main ophiolite zones are widely accepted in the Lesser Caucasus (Armenia) and they are linked to the geodynamic evolution of the Neotethys

Ocean (Dercourt et al. 1986). They are believed to correspond to the eastward extension of the IzmirAnkara-Erzinc¸an ophiolitic suture zone (Sengo¨r & Yılmaz 1981): † the Sevan –Akera zone (Knipper 1975; Aslanyan & Satian 1977; Sokolov 1977; Adamia et al. 1981; Knipper et al. 1997), situated in the east and SE of Sevan Lake, and including the Amassia-Stepanavan ophiolites in the northwestern part of the country; † the Vedi ophiolitic unit (Aslanyan & Satian 1977; Sokolov 1977; Knipper & Khain 1980), cropping out in the SE of Yerevan. Here, we shall present new biostratigraphic data for Upper Jurassic radiolarites from Vedi (Fig. 2), which are genetically associated with ophiolitic lavas. We will further review the existing biochronological data for the sedimentary cover of ophiolites in the Lesser Caucasus, in order to synthesize our knowledge on the temporal and spatial development of igneous and sedimentary rocks in the region.

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 383–391. DOI: 10.1144/SP340.16 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Structural map of the collisional zone between Arabia and Eurasia (modified after Avagyan et al. 2005).

Geological setting of the studied section The ophiolitic sequence of Vedi is composed of serpentinites, gabbros and a thick pile of massive and pillowed lava flows and their sedimentary cover. The latter consists mainly of reddish-brownish radiolarites; however, pelagic limestones with rare Callovian–Oxfordian ammonites are also present in some areas (Belov et al. 1991; Kvantaliani & Adamia 1992; Kvantaliani 2000). Lavas consist of alkaline and tholeı¨tic basalts and subordinate trachyandesites (Zakariadze et al. 1983). The ophiolitic rocks are weakly deformed, including some shear zones formed in greenschist facies metamorphic conditions. They occur within a folded

klippe sequence that was thrusted over a Coniacian sedimentary melange reworking the ophiolites (Sokolov 1977) and deposited above Cenomanian – Turonian shallow water carbonates of the South Armenian Block. Reefal limestones cover in general the pillow-lavas of the ophiolitic sequence. However, in some areas it is a conglomerate containing clasts of ophiolitic rocks (gabbros, serpentinites, basalts) that covers the ophiolitic lavas. This conglomerate evolves laterally to the above mentioned reefal limestones, built essentially of late Coniacian Hippurites. This sedimentary sequence lies unconformably over the ophiolitic rocks, the melange and the Cenomanian-Turonian shallow-water

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Fig. 2. Geological map of Armenia, including distribution of ophiolitic complexes. Locality 1, Erakh mountains; locality 2, Zodi pass; locality 3, Mt Karawul; locality 4, Stepanavan ophiolite.

carbonates of the South Armenian Block. Based on this evidence, obduction is considered to have taken place during the Coniacian –Santonian (Knipper & Sokolov 1976). From a small side valley of the river Vedi (N:39. 94378, E:44.98938; Fig. 3) we collected three samples of radiolarian cherts (Vedi06-12, Vedi06-21 and Arm006) that come from an c. 2 m thick sequence of bedded radiolarites overlying stratigraphically c. 2– 3 m thick lavas. The relationship of this 4– 5 m thick radiolarite-lava sequence with the surrounding lithologies is not clear, the general impression on the field is of a tectonically disturbed sequence. Most of the lithologies surrounding the studied sequence are composed of alkaline and subalkaline lavas, but there are also lamprophyres of limited extent, situated near the studied radiolarite-lava sequence. It is unclear whether they are part of a dyke system posterior of the ophiolite obduction or part of a tectonic melange.

New results Moderately to poorly preserved radiolaria were extracted from the three radiolarian chert samples

following routine laboratory processing with diluted hydrofluoric acid (HF c. 4%). All species identified in the three studied samples are provided in Table 1. Age-diagnostic species are illustrated in Figure 4. Taxonomic concepts applied during this study follow those stated by Baumgartner et al. (1995a) and O’Dogherty et al. (2006). According to the biozonation by Baumgartner et al. (1995b) all three samples can be correlated with the Unitary Association Zones (UAZ) 9 to 11 (Mid/Late Oxfordian to Late Kimmeridgian/Early Tithonian), based essentially on the presence of Zhamoidellum ovum Dumitrica and co-occurrence of Cinguloturris carpathica Dumitrica and Podocapsa amphitreptera Foreman. According to the recent biozonation by Beccaro (2006) all samples can be correlated with an Oxfordian (or younger) age. More particularly, sample Vedi06-21 could be correlated with the late (?) Oxfordian –Early Kimmeridgian (zone E) based on the co-occurrence of Loopus doliolum Dumitrica and Podocapsa amphitreptera Foreman. However, it should be mentioned that L. doliolum is also reported by Dumitrica et al. (1997) to occur in Berriasian strata.

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Fig. 3. Geological map of the Vedi area (after Sosson et al. 2010). Locality A, studied outcrop; locality B, Middle Jurassic outcrop studied by Danelian et al. (2008a).

Discussion Palaeontological data provided for the sedimentary cover of ophiolites are of key importance to understand the palaeogeographic evolution of Neotethys in the Lesser Caucasus. It is worth recalling briefly the available biochronologic evidence from the sedimentary cover of ophiolites preserved in the Lesser Caucasus area (Fig. 5). Biostratigraphic data known previously for the sedimentary cover of the Vedi ophiolite are published by Belov et al. (1991). These authors report on the presence of Tithonian–Berriasian and rare Callovian–Oxfordian ammonites (identifications in Kvantaliani & Adamia 1992; Kvantaliani 2000) in thinly-bedded pink-grey pelagic limestones

overlying ultrabasic breccia in the Erakh mountain (Figs 2 & 5). Radiolarite lenses situated at the top and beneath picrodoleritic pillow lavas along the Mankouk River are dated as Kimmeridgian – Middle Tithonian (Belov et al. 1991, sample 89-32, identifications by N. Bragin). According to the biozonation by Baumgartner et al. (1995b) this sample can be more specifically assigned to the Late Oxfordian/Early Kimmeridgian to Late Kimmeridgian/Early Tithonian time interval (UAZ 10-11), based on the co-occurrence of Dicerosaturnalis dicranacanthos (Squinabol) and Tetraditryma pseudoplena Baumgartner. Belov et al. (1991) also report on radiolaria yielded from bedded cherts studied along the Khosrov river, which are indicative of mainly Late

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Table 1. Occurrence table of identified radiolarian species in the three studied samples. The age of these three samples is given by means of UAZ (Unitary Association Zones) according to the zonation by Baumgartner et al. (1995b) and the zones identified recently by Beccaro (2006) Sample Vedi06-12 Acaeniotyle diaphorogona Foreman gr. Angulobracchia biordinalis Ozvoldova Archaeodictyomitra patricki Kocher Archaeodictyomitra apiarium (Ru¨st) Archaeospongoprunum imlayi Pessagno Cinguloturris carpatica Dumitrica Emiluvia peteri Beccaro Emiluvia premyogii Baumgartner (Riedel & Sanfilippo) Eucyrtidiellum ptyctum (Riedel & Sanfilippo) Loopus doliolum Dumitrica Pessagno Paronaella mulleri (Pessagno) Tethysetta mashitaensis (Mizutani) Podobursa vannae Beccaro Podocapsa amphitreptera Foreman Protunuma japonicus Matsuoka & Yao Saitoum trichylum De Wever Pseudodictyomitrella ulivii (Chiari, Cortese & Marcucci) Syringocapsa sp. A in Beccaro 2006 Triactoma foremanae Muzavor Zhamoidellum ovum Dumitrica UAZ Baumgartner et al. (1995b) Zones Beccaro (2006)

Sample Vedi06-21

Sample Arm006

P* Cf.†

UAZ Baumgartner et al. (1995b)

Zones Beccaro (2006)

4-22

Cf. Cf. P P

Cf. P

P

7-11

D-F D-F

Cf. Cf. P

D-E Cf.

P

P P P P Cf.

P 9-11 D-F

P Cf. P 9-11 E

8-15 P P Cf. P

9-18 7-12

C-F E-F B-F

B-F P 9-11 E-F

9-11

D-F

*P, confident identification. † Cf, doubtful identification.

Jurassic and Early Cretaceous ages (identifications by V. Vishnevskaya). The reported Jurassic assemblage may be correlated with the Late Bathonian/ Early Callovian to Mid Callovian/Early Oxfordian time interval (UAZ 7-8) based on the co-occurrence of Mirifusus fragilis Baumgartner and M. mediodilatatus Ru¨st gr. Regarding the Lower Cretaceous radiolarian assemblage, it can be now correlated with the latest Valanginian/earliest Hauterivian (UAZ 18), based on the co-occurrence of species Thanarla elegantissima (Cita) and Hemicryptocapsa capita Tan. However, it is noteworthy that for all of these dated cherts reported by Belov et al. (1991) there is no mention of any genetic (stratigraphic) relation with ophiolitic lavas. The discovery of Middle Jurassic (Bajocian) radiolaria in radiolarites intercalated with lavas (Fig. 5; Danelian et al. 2008a) in an outcrop situated at the southern part of the river Vedi (Fig. 3) establishes that this submarine volcanic activity started at least since the Bajocian for that part of the Tethyan oceanic crust preserved at Vedi.

Regarding the main ophiolite suture in the Lesser Caucasus, the Sevan–Akera zone, Zakariadze et al. (1983) reported on radiolaria extracted from bedded cherts of the Pechaniz river (east of Lake Sevan) which appear to be associated with ophiolitic lavas. Although the authors claimed for Early Cretaceous (and Late Jurassic?) ages, this cannot be confirmed based on the 13 species names they provided. Indeed, in addition to the fact that Mesozoic radiolarian taxonomy and biostratigraphy were still in their infancy in the early 1980s, no information is given for the identified assemblages nor any of the identified fauna is illustrated. More recently, Late Triassic (Carnian) and Early Jurassic (Toarcian) radiolaria were extracted from cherts intercalated with ophiolitic breccias in the Zod (Sotk) pass area (Vishnevskaya 1995; Knipper et al. 1997), SE of Lake Sevan (Figs 2 & 5). Moreover, further to the east, at Mt. Karawul (Fig. 2), Late Middle Jurassic (Early/Middle Bajocian to Late Bajocian/Early Bathonian, UAZ 3-5) radiolaria were yielded from radiolarites overlying basaltic lavas (Fig. 5; Vishnevskaya 1995).

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Fig. 4. Scanning electron micrographs of Upper Jurassic radiolaria from the three chert samples of the studied section at Vedi. Figs a–c, sample Vedi06-12; (a) Cinguloturris carpathica; (b) Emiluvia peteri; (c) Zhamoidellum ovum. Figs d– j, sample Vedi06-21; (d) Cinguloturris carpathica; (e) Loopus doliolum; (f) Parvicingula mashitaensis; (g) Podocapsa amphitreptera; (h) Protunuma japonicus; (i) Syringocapsa sp. A; (j) Zhamoidellum ovum. Figs k– l, sample Arm 06; (k) Stichocapsa ulivii, (l) Zhamoidellum ovum.

Finally, with respect to the Stepanavan ophiolite, that can be considered as the northward extension of the Sevan–Akera ophiolites, Danelian et al. (2007) established recently a Late Jurassic age (Mid/Late Oxdordian to Late Kimmeridgian/Early Tithonian, UAZ 9-11) for radiolarites overlying ophiolitic lavas (Figs 2 & 5).

Conclusion In the Vedi area, the age provided by Danelian et al. (2008a) represents the oldest known evidence for the sedimentary cover of the Vedi ophiolite. Oceanic crust in this part of Neotethys was formed during or most likely before the Bajocian. In this

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Fig. 5. Radiolarian ages determined in this study, together with published ages from various other ophiolite zones in the Vedi area, Stepanavan and the Sevan–Akera suture zone (Vishnevskaya 1995; Knipper et al. 1997; Danelian et al. 2007, 2008a). Time scale after Gradstein et al. (2004).

study we establish that dating thus a late Jurassic submarine volcanic activity. We here establish that volcanic activity continued until at least the Oxfordian since Upper Jurassic (Middle/Upper Oxfordian to Upper Kimmeridgian/Lower Tithonian) radiolarian cherts overly stratigraphically volcanic extrusives. Pelagic carbonate sedimentation was installed in places during the Callovian –Oxfordian and Tithonian –Berriasian, probably on relative bathymetric highs, while radiolarian cherts accumulated at deeper environments until the Hauterivian. Our results are consistent with the development of a Mesozoic oceanic realm (Neotethys), between the active Eurasian margin (to the North) and the South-Armenian micro-continent (to the South), for which spreading started since the Late Triassic

and carried on through most of the Jurassic into the Early Cretaceous. Part of this oceanic crust formed during the early Middle Jurassic (Bajocian or earlier) is now preserved at Vedi, following Late Cretaceous (Coniacian– Santonian) obduction of ophiolites. Funding from the Middle East Basins Evolution (MEBE) programme, the French Ministry of Foreign Affairs (ECO-NET grant to T. Danelian) and the French Embassy in Yerevan (student grants to G. Asatryan and G. Galoyan) are gratefully acknowledged. Fieldwork was facilitated by the support of the Institute of Geological Sciences (Armenian Academy of Science). Constructive remarks from M. Chiari and N. Bragin improved the initial manuscript. C. Abrial and A. Lethiers helped with drafting.

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References Al-Riyami, K., Danelian, T. & Robertson, A. H. F. 2002. Radiolarian biochronology of Mesozoic deep-water successions in NW Syria and Cyprus: implications for south-Tethyan evolution. Terra Nova, 14, 271–280. Al-Riyami, K., Robertson, A. H. F., Xenophontos, C., Danelian, T. & Dixon, J. E. 2000. Tectonic evolution of the Mesozoic Arabian passive continental margin and related ophiolite in Baer-Bassit region (NW Syria). In: Panayides, I., Xenophontos, C. & Malpas, J. (eds) Proeedings of the 3rd International Conference on the Geology of the Eastern Mediterranean, 61–81. Adamia, S., Chkhotua, T., Kekelia, M., Lordkipanidze, M., Shavishili, I. & Zakariadze, G. 1981. Tectonics of the Caucasus and the adjoining regions: implications for the evolution of the Tethys Ocean. Journal of Structural Geology, 3, 437–447. Aslanyan, A. T. & Satian, M. A. 1977. On the geological features of Transcaucasian ophiolitic zones. Izvestia Acad. Sci. Armenian SSR, Nauki o Zemle, 4 –5, 13–26 (in Russian). Avagyan, A., Sosson, M. et al. 2005. Neogene to Quaternary stress field evolution in Lesser Caucasus and adjacent regions using fault kinematics analysis and volcanic cluster data. Geodinamica Acta, 18, 401– 416. Baumgartner, P. O., O’Dogherty, L. et al. 1995a. Radiolarian catalogue and systematics of Middle Jurassic to Early Cretaceous Tethyan genera and species. In: Baumgartner, P. O., O’Dogherty, L., Gorican, S., Urquhart, E., Pillevuit, A. & De Wever, P. (eds) Middle Jurassic to Lower Cretaceous Radiolaria of Tethys: Occurrences, Systematics, Biochronology. Me´moires de Ge´ologie (Lausanne), 23, 37– 685. Baumgartner, P. O., Bartolini, A. et al. 1995b. Middle Jurassic to Early Cretaceous radiolarian biochronology of Tethys based on Unitary Associations. In: Baumgartner, P. O., O’Dogherty, L., Gorican, S., Urquhart, E., Pillevuit, A. & De Wever, P. (eds) Middle Jurassic to Lower Cretaceous Radiolaria of Tethys: Occurrences, Systematics, Biochronology. Me´moires de Ge´ologie (Lausanne), 23, 1013–1048. Beccaro, P. 2006. Radiolarian Biostratigraphy of Middle-Upper Jurassic pelagic siliceous successions of Western Sicily and the Southern Alps (Italy). Me´moires de Ge´ologie (Lausanne), 45, 120. Belov, A., Bragin, N., Vishnevskaya, V., Satian, M. & Sokolov, S. 1991. New data on the age of Vedi ophiolite (Armenia). Comptes Rendus Acade´mie URSS, 321, 784–787 (in Russian). Bill, M., O’Dogherty, L., Guex, J., Baumgartner, P. O. & Masson, H. 2001. Radiolarite ages in AlpineMediterranean ophiolites: constraints on the oceanic spreading and the Tethys-Atlantic connection. Eclogae geologicae Helvetiae, 90, 43– 54. Bragin, N. Y. & Tekin, U. K. 1996. Age of radiolarianchert blocks from the Senonian Ophiolitic Me´lange (Ankara, Turkey). The Island Arc, 5, 114– 122.

Bragina, L. G. & Bragin, N. Yu. 2006. Stratigraphy and Radiolarians of Upper Cretaceous sedimentary cover of the Arakapas Ophiolite Massif (Cyprus). Stratigraphy and Geological Correlation, 14, 507– 523. Chiari, M., Cortese, G., Marcucci, M. & Nozzoli, N. 1997. Radiolarian biostratigraphy in the sedimentary cover of the ophiolites of south-western Tuscany, Central Italy. Eclogae Geologicae Helvetiae, 90, 55–77. Chiari, M., Marcucci, M. & Principi, G. 2000. The age of Radiolarian Cherts associated with the ophiolites in the Apennines (Italy) and Corsica (France): a revision. Ofioliti, 25, 141–146. Danelian, T., Lekkas, S. & Alexopoulos, A. 2000. De´couverte de radiolarites triasiques dans un complexe ophiolitique a` l’extreˆme sud du Pe´loponne`se (Agelona, Lakonie, Gre`ce). Comptes Rendus de l’Acade´mie des Sciences, Paris, 330, 639– 644. Danelian, T., Galoyan, G., Rolland, Y. & Sosson, M. 2007. Palaeontological (Radiolarian) Late Jurassic age constraint for the Stepanavan ophiolite (Lesser Caucasus, Armenia). Bulletin of the Geological Society of Greece, 40, 31–38. Danelian, T., Asatryan, G., Sosson, M., Person, A., Sahakyan, L. & Galoyan, G. 2008a. Middle Jurassic (Bajocian) Radiolaria from the sedimentary cover of the Vedi ophiolite (Lesser Caucasus, Armenia). Comptes Rendus PalEvol, 8, 327– 334. Danelian, T., De Wever, P. & Durand Delga, M. 2008b. Revised Radiolarian ages for the sedimentary cover of the Balagne ophiolite (Corsica, France). Implications for the palaeoenvironmental evolution of the Balano-Ligurian margin. Bulletin de la Socie´te´ Ge´ologique de France, 179, 169–177. Dercourt, J., Zonenshain, L. P. et al. 1986. Geological evolution of the Tethys belt from the Atlantic to the Pamirs since the Lias. Tectonophysics, 123, 241– 315. De Wever, P. & Dercourt, J. 1985. Les radiolaires triasico-jurassiques marqueurs stratigraphiques et pale´oge´ographiques dans les chaıˆnes alpines pe´rime´diterrane´ennes: une revue. Bulletin de la Socie´te´ Ge´ologique de France, 8, 653–662. De Wever, P., Danelian, T., Durand-Delga, M., Cordey, F. & Kito, N. 1987. Datations des radiolarites post-ophiolitiques de Corse alpine a` l’aide des Radiolaires. Comptes Rendus de l’Acade´mie des Sciences, Paris, 305, 893– 900. Dumitrica, P., Immenhauser, A. & Dumitrica-Jud, R. 1997. Mesozoic Radiolarian biostratigraphy from Masirah Ophiolite, Sultanate of Oman. Part I: Middle Triassic, Uppermost Jurassic and Lower Cretaceous Spumellarians and multisegmented Nassellarians. Bulletin of the National Museum of Natural Science, 9, 1– 105. Go¨ncu¨og˘lu, M. C., Yalıniz, M. K. & Tekin, U. K. 2006. Geochemistry, tectono-magmatic discrimination and Radiolarian ages of basic extrusives within the Izmir-Ankara suture belt (NW Turkey): time constraints for the Neotethyan evolution. Ofioliti, 31, 25–38. Gradstein, F. M., Ogg, J. G., Smith, A. G., Bleeker, W. & Lourens, L. J. 2004. A new Geologic time scale, with special reference to Precambrian and Neogene. Episodes, 27, 83–100.

BIOCHRONOLOGY OF RADIOLARITES FROM THE LESSER CAUCASUS (ARMENIA) Karakitsios, V., Danelian, T. & De Wever, P. 1988. Datations par les Radiolaires des Calcaires a` Filaments, Schistes a` Posidonies supe´rieurs et Calcaires de Vigla (zone ionienne, Epire, Gre`ce) du Callovien au Tithonique terminal. Comptes Rendus de l’Acade´mie des Sciences, Paris, 306, 367–372. Knipper, A. L. 1975. The Oceanic Crust in the Alpine Belt, Tr. GIN NAS USSR, Edition 267 (in Russian). Knipper, A. L. & Khain, E. V. 1980. The structural position of ophiolites of the Caucasus. Ofioliti, Special Issue, 2, 297–314. Knipper, A. L. & Sokolov, S. D. 1976. Vedi ophiolites (Armenia): autochtone or allochton? Geotectonics, 4, 54– 66 (in Russian). Knipper, A. L., Satian, M. A. & Bragin, N. Yu. 1997. Upper Triassic-Lower Jurassic Volcanogenic and Sedimentary Deposits of the Old Zod Pass (Transcaucasia). Stratigraphy, Geological Correlation, 3, 58–65 (in Russian). Kvantaliani, I. V. 2000. Middle Jurassic-Early Cretaceous Ammonites of Vedi Ophiolite Zone of Lesser Caucasus. Tbilisi, pub AS Georgia. Kvantaliani, I. V. & Adamia, Sh. A. 1992. Paleotectonic and Late Jurassic-Early Cretaceous Ammonites in Vedi Ophiolite Zone of Lesser Caucasus (Armenia). Georgian Academy of Sciences, Institute of Geology of A.I. Djanelidze. O’Dogherty, L., Bill, M., Gorican, S., Dumitrica, P. & Masson, H. 2006. Bathonian radiolarians from an ophiolitic me´lange of the Alpine Tethys (Gets Nappe, Swiss-French Alps). Micropaleontology, 51, 425–485. S¸engo¨r, A. M. C. & Yılmaz, Y. 1981. Tethyan evolution of Turkey: a plate tectonic approach. Tectonophysics, 75, 181–241.

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Sokolov, S. D. 1977. The Olistostroms and Ophiolitic Nappes of the Lesser Caucasus. Izdatelstvo Nauka, Moscow (in Russian). Sosson, M., Rolland, Y. et al. 2010. Subductions, obduction and collision in the Lesser Caucasus (Armenia, Azerbaijan, Georgia), new insights. In: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 329–352. Tekin, U. K., Go¨ncu¨og˘lu, M. C. & Turhan, N. 2002. First evidence of Late Carnian radiolarians from the Izmir-Ankara suture complex, central Sakarya, Turkey: implications for the opening age of the Izmir-Ankara branch of Neo-Tethys. Geobios, 35, 127– 135. Vishnevskaya, V. 1995. Jurassic and Cretaceous Radiolarians from the Lesser Caucasus (Zod Pass, Mount Karawul and site 22 in the Koshuni River Basin). In: Baumgartner, P. O., O’Dogherty, L., Gorican, S., Urquhart, E., Pillevuit, A. & De Wever, P. (eds) Middle Jurassic to Lower Cretaceous Radiolaria of Tethys: Occurences, Systematics, Biochronology. Me´moires de Ge´ologie (Lausanne), 23, 701– 708. Vrielynck, B., Bonneau, M., Danelian, T., Cadet, J. P. & Poisson, A. 2003. New insights on the Antalya Nappes in the apex of the Isparta angle: The Isparta C¸ay unit revisited, Geological Journal, 38, 283– 293. Zakariadze, G. S., Knipper, A. L., Sobolev, A. V., Tsamerian, O. P., Dmitriev, L. V., Vishnevskaya, V. S. & Kolesov, G. M. 1983. The ophiolite volcanic series of the Lesser Caucasus. Ofioliti, 8, 439– 466.

Recent tectonic stress evolution in the Lesser Caucasus and adjacent regions ARA AVAGYAN1, MARC SOSSON2*, ARKADI KARAKHANIAN3, HERVE´ PHILIP4, SAMIRA REBAI5, YANN ROLLAND2, RAFAEL MELKONYAN1 & VAHAN DAVTYAN3 1

Institute of Geological Sciences, National Academy of Sciences of Armenia, 24a Baghramian avenue, Yerevan, 0019, Armenia

2

University of Nice-Sophia Antipolis, CNRS, Ge´oazur, UMR 6526, 250 rue A. Einstein, 06560 Valbonne, France 3

GEORISK Scientific Research Company, 24a Baghramian avenue, Yerevan, 375019, Armenia

4

Laboratory of Geophysics, Tectonics and Sedimentology, UMR 5573, University of Montpellier II, France

5

IMS RN, Agence Mediterranee 2214, Bd de la Lironde Bat 21, Parc Agropolis, 34980 Montferrier sur Lez, France *Corresponding author (e-mail: [email protected]) Abstract: The stress indicators describing the recent (provided by active tectonics framework) and palaeo-stress (provided by micro-fault kinematics and volcanic cluster) patterns show the scale and temporal changes in stress states since the beginning of Arabian– Eurasian collision. The recent stress derived from the active fault kinematics in the Lesser Caucasus and adjacent area corresponds to a strike–slip regime with both transtension and transpression characteristics. The kinematics of active structures of various scale are conditioned by tectonic stress field with general north–south compression and east– west extension. The distribution of Neogene to Quaternary volcanic cluster geometries and micro-fault kinematic data evidence the time and orientation variability of the stress field since the beginning of the Arabian–Eurasian collision. In addition to the general north– south compression orientation, two other – NW–SE and NE–SW – secondary orientations are observed. The first one was dominant between the Palaeogene and the late Early Miocene and the second one has prevailed between the Late Miocene and the Quaternary. Since the continental collision of Arabia with Eurasia the tectonic stress regime in the Lesser Caucasus and adjacent area changed from compression (thrusting and reverse faulting) to transtension-transpression (strike–slip faulting with various vertical components).

Since Jurassic times the alpine (sensu lato) tectonic evolution of the Caucasus mountain belts and to the south of the Armenian Highland is the result of subductions of the Neotethys Ocean and associated obductions and collisions (Dercourt et al. 1986; Nikishin et al. 1998; Robertson 2002; Golonka 2004). Oligocene compressional forces induced by the continental subduction of the Arabian lithosphere beneath Eurasia affect a wide domain from the peri-Arabic frontal thrusts to the northern Greater Caucasus (Adamia et al. 1981; Sengo¨r & Yilmaz 1981; Dercourt et al. 1986; Ricou 1994; Yilmaz et al. 2000; Robertson 2002; Golonka 2004; Faccenna et al. 2006). Moreover, due to the occurrence of rigid oceanic lithosphere remnants in the Black and Caspian seas, the continental

lithosphere of the Caucasus and Armenian Highland areas between them is mainly deformed by a dense active fault network of various strikes and kinematics (Fig. 1). This wide range of active faults (Figs 1 & 2) (e.g. Jackson & McKenzie 1984; Rebaı¨ et al. 1993; Philip et al. 2001; Trifonov et al. 2002) is related to the north–south to NNE –SSW trending Arabian and Eurasian plates convergence (McKenzie 1972; Sengo¨r & Kidd 1979; Jackson & McKenzie 1984; Dewey et al. 1986; Taymaz et al. 1991; Philip et al. 2001; Karakhanian et al. 2004). As previously evidenced (Philip et al. 1989; Karakhanian & Balasanian 1992; Karakhanian et al. 2004a; Avagyan et al. 2005) four coeval existing types of neotectonic major faults are present comprising: (1) NE trending

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 393–408. DOI: 10.1144/SP340.17 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Active tectonics north of the Arabian plate (from Avagyan et al. 2005). 1: Major strike–slip faults. 2: Major thrust faults. 3: Relative motion of blocks with respect to Eurasia. 4: Focal mechanisms of Mw . 4.8 earthquake (CMT Harvard). 5: Instrumental seismicity of 3 , Mb , 4.9 (USGS-NEIC). Ar, Armenia; D, Dagestan; T, Talish; E.A.F., East Anatolian fault; N.A.F., North Anatolian fault; P.S.S.F., Pambak – Sevan-Sunik fault, Z.F., Zagros fault; N.T.F., North Tebriz fault; G.S.F, Geltarechka-Sarikhamish fault.

sinistral strike –slip faults; (2) NW trending dextral strike–slip faults; (3) east-trending thrusts; and (4) north-trending normal faults (Figs 1 & 2). The major active faults form large top to the North structural arcs. Some of these structures correspond to the Sevan –Akera suture zone. Some other major crustal faults are inherited from normal faults in the basement (Sosson et al. 2010). During the last 100 Ma this region imprinted several tectonic processes which deformed it and led to its structural pattern. The main one is the Paleocene to Middle Eocene collision of a Gondwanianderived terrane (the South Armenian Block, SAB, Knipper & Kain 1980) and the eastward extension of the Taurides– Anatolides one (e.g. Barrier & Vrielynck 2008). The northward continental subduction of the SAB beneath Eurasia conducted to a huge reverse faulting and folding of the Eurasian and the SAB crusts (Sosson et al. 2010). Moreover, the Arabian plate subducted beneath the southern margin of the Taurides –Anatolides, the SAB

and the Iranian terranes. In the Lesser Caucasus, due to this second phase of collision, the Sevan– Akera ophiolitic suture zone was reactivated by some major thrusts and strike –slip faults (Avagyan et al. 2005). Normal faults of the northern palaeo-passive margin of the SAB were reactivated by reverse and strike–slip faults (Sosson et al. 2010). Consequently the studied region is enclosed between mainly rigid plates (Arabia and Eurasia), and escaping secondary blocks or microplates (Iran and Taurides –Anatolides). One of the important questions is: Do the crustal structures within this area reflect the regional recent tectonic stress field? In order to answer this, we present the evolution of the recent stress field in the Lesser Caucasus and adjacent regions. Accordingly our study of the tectonic stress field was focused on the active structures, the micro-fault kinematics, and the volcanic cluster data in order to reveal the stress space –time characteristics.

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Fig. 2. Active fault map in the Lesser Caucasus and adjacent area. A1–A2 is the GPS profile line (Fig. 10). The sites of Figures 4, 6, 7 and 13 and areas of Figures 5 and 9 are indicated. 1, strike–slip faults; 2, thrust faults; 3, normal faults.

Kinematic setting The recent Arabian –Eurasian convergence rate (estimated at 20 –30 mm a21 by DeMets et al. 1990) is the result of collision between these two plates and of tectonic escape of the Taurides– Anatolides and Iranian terranes. The central region in front of the Arabian plate was submitted to an intense shortening (McKenzie 1972; Sengo¨r & Kidd 1979; Jackson & McKenzie 1984; Dewey et al. 1986; Philip et al. 2001). As previously summarized by Avagyan et al. (2005) three main changes in plate motion are known. (i) The Arabian plate moved to the NW with the onset of the Arabian –Eurasian collision during the Early Oligocene (35 Ma) (Savostin et al. 1986). (ii) A change from a north–south to NW –SE trend of convergence occurred during the Early Miocene (Zonenshain et al. 1987). (iii) In the Middle Miocene the movement of the Arabian plate again became top to the North and in eastern part, top to the NNE (Zonenshain et al. 1987). Since the Middle Miocene

the stress field probably remained relatively constant (e.g. Rebaı¨ et al. 1993).

Active faulting Major active structural pattern In the Caucasus and Armenian Highland the structural pattern analysis allows constraining the regional tectonic stress of this active orogenic belt, and furthermore comparing it with local scale data obtained from micro-fault kinematics analyses. Numerous active strike–slip and reverse faults outline wide structural wedges with tips oriented to the north and NE and bounded by dextral strike– slip faults on the eastern branch and sinistral strike– slip faults on the western one (Figs 1 & 2). At least three large wedges have been distinguished: the Daguestan, the North Armenian and the Talish (Philip et al. 1989, 2001; Avagyan 2001) (Fig. 1). The first and the second ones are directly related to the Greater Caucasus belt. The widest Daguestan wedge is formed by a large NE Anatolian sinistral strike –slip fault (NEAF) and dextral one localized between Greater Caucasus belt and Caspian basin

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(Fig. 1). In the modern structural pattern, the NEAF zone (so-called the Agrakhan-Tbilisi-Levantine fault, Khain & Koronousky 1997) is the longest (about 650 km length) strike –slip fault crossing the Lesser Caucasus belt (Fig. 1). The NEAF is morphologically and structurally much less distinct than the North Anatolian or East Anatolian Fault Zones (Fig. 1). It extends through the Alpine belt from Karliova junction (crossing point of North and East Anatolian faults) up to the north of the Greater Caucasus and is formed by two large fault zone segments with left step (Fig. 1). The northern one is called the Borjomi-Kazbek fault (Philip et al. 1989) or Kazbek-Tskhinvali fault (Trifonov 1999). The southern segment is formed by several faults of secondary order, most significant are the Oltu (so-called Olur fault), Dumlu, Erzurum faults and to the South limiting eastward the Chat basin (Trifonov 1999; Dhont & Chorowicz 2006) the Dumlu fault zone (Kocyigit et al. 2001). Contrary to the debatable activity of the northern BorjomiKazbek segment, the southern segment manifests high seismic activity (02 January 1978, Mw ¼ 5.6; 13 May 1986, Mw ¼ 5.8; 16 December 1990, Mw ¼ 5.5, Taymaz et al. 2004) with earthquake focal mechanisms evidencing sinistral strike –slip kinematics of the fault (Fig. 1). The earthquakes of magnitude less than five follow all strike of the fault including the Borjomi-Kazbek segment (Fig. 1) (Khain & Koronousky 1997). According to Philip et al. (1989) the crest of the Greater Caucasus was displaced of about 50–60 km left laterally due to activity of this faults zone. The Neogene to Quaternary volcanism of the area is mainly exposed to the East of the NEAF, and it limits the seismically more active eastern part from the western one less active (Khain & Koronousky 1997). Moreover, Reilinger et al. (2006) consider the NEAF zone as the northeastern limit of the Caucasus block. To the east of the NEAF, in the study area, two structural arcs form the North Armenian wedge (Fig. 1). The outer one formed by the GeltorechkaSarighamish Fault (GSF) to the west and the Pambak-Sevan-Sunik fault (PSSF) to the east. The inner part of the arc is defined by the Akhourian fault to the east (AF), and the Garni fault (GF) to the West (Fig. 2) (Karakhanian et al. 2004a). The most significant structure of the wedge is the 490 km length PSSF. The PSSF is located in the Lesser Caucasus mountain belt and consists of five large segments, each separated by overstep zones (Avagyan 2001; Philip et al. 2001). The tracks of the northern ones are aligned along the Sevan– Akera suture zone (Figs 1, 2 & 3). The southern segments to the North cross a wide Quaternary volcanic area which overlies the southern branch of the suture zone (Knipper et al. 1997). The PSS strike– slip

fault is inherited from major reverse faults (Fig. 2) which deformed the belt since the continental collision of the South Armenian Block with the Eurasian margin during Paleocene to Early Eocene (see Sosson et al. 2010). Along the PSSF zone, kinematics change examples are in the Vanadzor trough (Figs 2 & 4), where the pre-existing reverse faulting and younger flower structures of strike –slip fault are distinguished in Palaeogene formations (Avagyan 2001). The 16 km long and up to 3 km wide Vanadzor trough is located in the overstep zones of the 90 km long Arpi– Vanadzor and the 115 km long Vanadzor –Artanish fault segments between the mountain range of Bazum to the north and that of Pambak to the south (Figs 2 & 5). The western half of this trough is occupied by the Vanadzor city (the third city of Armenia). In the depression the fault is subdivided into several branches controlling the almond-shaped basin. The depression is composed by Upper Cretaceous to Eocene sedimentary rocks (limestones, tuffs, tuffobreccia, sandstones) and post Oligocene plutons. The youngest 145 m thick sediments are made of clay and sands and occupy the bottom of the depression (Milanovski 1968). Morphological features such as pressure ridges, deflected derange, offset streams, triangular facets, and sag pond evidence a recent tectonic activity. The main fault segment dextrally offsets the stream 400 m (see 1 on Fig. 5) and deflect more than 1.3 km the NE-trending valley of the Pambak river (see 2 on Fig. 5) to the north of Vanadzor city. In the east part of the Vanadzor trough, the fault scarps and the deformed morphological markers are better conserved than in the west owing to lesser volume of the recent covering sediments and to less human occupation. The valleys which go down from the Bazum Mountain are well developed in the relief with dip V shape erosions, but they stop very roughly at the limit of the basin which presents one segment of the fault. There, triangular facets are well developed indicating a relatively normal component along the fault. This transition slope/ basin underlined by springs corresponds to the fault plane limiting the depression from the north. Towards downstream the incisions are rather shallow or absent. Somewhere recent small alluvial fans related to important valleys are observed. An elongated ridge with an abandoned valley on its central part dams the streams from the north forming sag pond (see 3 on Fig. 5). It is characterized by a flat surface which corresponds to fine sediments accumulation in a temporary lake. The dead valley incision restarted owing the accumulated water level uplift. Along the PSSF in the Sunik massif and Vanadzor depression area the overstep zones create favourable conditions for recent volcanic activity.

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Fig. 3. Geologic cross section of the Lesser Caucasus in Armenia and Azerbaijan, modified from Sosson et al. (2010). 397

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Fig. 4. The outcrop to the east of Vanadzor trough in Middle Eocene formations where the pre-existing reverse faulting and younger flower structures of strike– slip fault are distinguished. The site is indicated in Figure 2.

In the Vanadzor depression the strike– slip faults with normal component can be explained by segment about 308 clockwise change with respect to fault general strike. This fault geometry with

right bending along the right lateral strike–slip fault with general north–south compression creates conditions of local east– west extension responsible for a pull apart structure. The volcanic

Fig. 5. (a) Active fault map superposed on the 3D topographic model (modified from Avagyan 2001). 1, active and inferred strike– slip faults; 2, reverse faults; 3, normal faults; 4, basalt andesitic rocks; 5, landslides; 6, triangular facets; 7, regional compression axes; 8, local extension of right bend; 9, peopled area. (b) Topographical profile (the I– II line of profile is shown on numerical model) with vertical exaggeration, faults and basalts andesitic rocks situation are shown. The situation of the figure area is shown on Figure 2.

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Fig. 6. The Sev Litsh lake site (situation is shown on Fig. 2) on the southern segments of PSSF manifested mainly oblique slip with normal component. The cross section in (c) through the lake (b) is shown in (a). Small black arrows indicate the faults.

activity considered Quaternary in age (Bagdasaryan & Jrbashyan 1970) is evidenced by andesitic basalts related to this local extension tectonics (Fig. 5). The strike–slip faults in the Lesser Caucasus area show variable vertical kinematics not only in local but also in regional scale. Along its southern segments, PSSF generally manifests mainly oblique slip with normal component (Figs 6 & 7b), whereas along the northern ones is exposed reverse component (Fig. 7a). This is explained by progressive change of the fault orientation from the N1058 (in the north) to N1558 (in the south) in correlation with stress field.

The same kinematics variation was observed along the Garni fault (Fig. 2). The Spitak earthquake of 1988 ruptured the northernmost segment of the fault. Along its northern part the rupture has reverse kinematics with dextral slip component. To the south, the Spitak rupture shows principally right-lateral strike– slip kinematics (Karakhanian & Balasanian 1992; Philip et al. 1992). Moreover, the GPS data confirm the mainly strike –slip character of the Garni fault central segment (south from Spitak rupture) (Davtyan et al. 2007). The 1988 Spitak earthquake rupturing process reconstruction from the fault plane striation of

Fig. 7. Examples oblique slip with reverse component on the northern 1088 trending (a) and with normal one on the southern 1558 trending (b) segments of PSSF. This is explained by progressive change of the fault orientation in correlation with stress field. The situations of the sites are shown on Figure 2.

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various orientation and shapes allows us to confirm the multi-phase main event and shows the nearsurface dual rupturing mechanism. The first motion of the superficial secondary rupture (earliest striation with sinistral component occurred before the rupture front reached the surface) and is not consistent with the resulting dextral component of the displacement (on sinistral striation the dextral ones are superimposed) (Fig. 8). It shows possible nearsurface kinematic variations during the single strong earthquake rupturing process, which were not accompanied by any general dynamic change at depth (Avagyan et al. 2003). The phenomena can be explained by the law of momentum conservation. During the late stage of collision the most important actives faults kinematics of the Lesser Caucasus area corresponds to strike– slip movements followed the bending of the originally rectilinear shape of the Lesser Caucasus –Pontides (Bazhenov & Burtman 2002) due to the northward motion of the SAB.

Within wedges structures The active structures within the wedges are characterized by general east –west striking reverse faults, NE or NW-trending strike–slip faults, and north– south striking normal faults of local signification (Philip et al. 2001). The example of normal faulting

is well expressed in the Gegam volcanic massif crest area and on its eastern slope (Fig. 9). In the crest area the faults are segmented and show morphological evidence of oblique slip with dextral and normal association (Fig. 9). The Gavaraget and Noratus normal faults are also segmented and are generally north– south trending. The first one dips to the West and the other dips to the east (Fig. 9). Some segments are restricted to 1–3 km long tilted blocks. By extending the slope line of the massif up to the escarpment the vertical cumulated displacement of 250 m has been estimated, which suggests about 50 m thick mainly alluvial deposits. The fault activity perturbed the hydrographic network. At least three abandoned valleys are located on the horst delimited by the faults of Gavaraget and Noratus (Fig. 9). The escarpment of the Gavaraget fault has dammed the rivers that went down from the Gegam massif. At first a lake was formed and lacustrine and volcano-lacustrine sediments were accumulated, indicating a coeval volcanic activity. Then the spillway intense erosion leads to water emptying and as the result the lake sediments were profoundly indented. The C14 dating of charcoal samples (realized in the laboratory of Hydrology and Isotopic Geochemistry of the University of Paris – South by J.-L. Michelot) found 2 m below the surface level, lacustrine sediments

Fig. 8. Two steps (a) and (b) of rupturing process in the case of the Spitak earthquake deduced from striation analysis. Block diagrams correspond to the site of the striated samples.

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Fig. 9. The fault system of the Gegam massif crest and East slop area (shown in Fig. 3). (a) The area on the SPOT satellite image and (b) the interpretation with 1, faults; 2, supposed fault; 3, I, normal faults, II, block tilting orientation; 4, cliffs; 5, limits of the reliefs of different ages; 6, lakes; 7, lacustrine sediments; 8, I, volcanoes, II, abandoned valleys on the Gavaraget fault foot wall; 9, settlement. (c) Cross-section [section line shown on (b)] with 10, Lower and pre-Quaternary basement; 11, Quaternary lava flow; 12, lava flow of Late Quaternary; 13, volcano-lacustrine sediments; 14, fluvio-lacustrine formations; 15, Sevan lake sediment. The situation of the figure area is shown on Figure 2.

giving about 30 000 year PB (Fig. 9b). Near the Gavar city, the Gavaraget fault sharply changes its orientation (N110) and kinematics, which become dextral strike –slip (Fig. 9). The amplitude of the normal component of the Noratus fault has been estimated from the borehole performed at 500 m east from the scarp. There it was observed 100 m of normal displacement within the same formation along the Noratus fault (Milanovski 1952). This fault was active during the Pleistocene and even Holocene (Milanovski 1952). Consequently in the Lesser Caucasus area observations made at various scales on the major active structures and the faulting from the within wedges evidence a tectonic stress field with general north–south compression and east –west extension. Moreover,

the P axis data (Tovmassian et al. 1997) with some variation (orientation of north –south, north– NW and NNE) and USGS-NEIC focal mechanisms (Fig. 1) confirm it. The observed active fault kinematics is characteristic of strike– slip stress regime. The analysis of GPS data obtained from network of 22 points deployed since 1998 in Armenia (in the framework of collaboration between the University of Montpellier-II France, GEORISK Scientific research company and Institute of Geological Sciences, Armenia) indicates a general NNE shortening and east–west dilatation (Davtyan 2007) consistent to strike–slip stress regimes. In the Gegam massif and Southern Sevan basin area (Figs 2, 9 & 10) the extension rate is of 2.4 + 0.9 mm a21 with

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Fig. 10. The extension is shown on GPS profile A1-A2 (the line is shown on Fig. 2) with Eurasia-fixed reference frame (Davtyan 2007). 1, linear correlation for all points; 2, separate linear correlation; 3, polynomial correlation; 4, error limits and station names. GF, Garni fault; PSSF, Pambak – Sevan–Sunik fault.

the azimuth of about N608. According to GPS data the extension is concentrated on the Gegam massif crest area, where to the east the velocity abruptly increase (Fig. 10). It consistent with active fault (Fig. 9) and volcanic cluster orientation data (see below).

Volcanic cluster analysis In the study area, many works concerned magmatism (since the Jurassic times) and its tectonic origins (e.g. Melkonyan 1976; Innocenti et al. 1980; Karapetian 1985; Adiyaman et al. 1998; Karapetian et al. 2001; Karakhanian et al. 2002, 2003). Since the onset of the continental collision, volcanism has evolved with bimodal character comprising both basaltic and rhyolitic lavas (Karapetian et al. 2001). In Armenia two stages of intracontinental volcanism are identified: one during Oligocene to Early Miocene and the other from Late Miocene to Quaternary. During the later stage widespread volcanic massifs and cones were formed in all volcanic provinces of the Armenian Highland (Karapetian et al. 2001) exceeding 1 km in thickness in places and ranging in ages from 11 Ma to present (Pearce et al. 1990; Keskin 2003). The volume of volcanic

products was estimated of 35 000 km3 (Milanovski 1977). The most recent volcanic activity of the Holocene times comprises the volcanism of Nemrout, Sipan, underwater volcanoes of Van lake, Tondourek and Ararat volcanoes in Eastern Turkey (Innocenti et al. 1980; Dewey et al. 1986; Yilmaz et al. 1998; Haroutiunian 2001; Karakhanian et al. 2002), and in Sunik, Vardenis and Gegam massifs in Armenia (e.g. Karakhanian et al. 2002). As Haroutiunian (2001) mentioned, these volcanoes form the 400 km active volcanoes belt of the Armenian Highland. The volcanic alignments of same age are used as indicators of the tectonic regime and stress axis orientation. Such an approach has been used by several researchers for the Armenian Highland (e.g. Karapetian 1985; Karakhanian et al. 2004b; Avagyan et al. 2005) and other regions (Nakamura 1969; Jackson & Shaw 1975; Opheim & Gudmundsson 1989; Chorowicz et al. 1997; Koyaguchi & Takada 1994; Bellier & Sebrier 1994; Adiyaman et al. 1998). Most volcanoes in Armenia and adjacent areas form clusters of different lengths from several kilometres up to 50 km and are clearly rooted on tension fracture. As reported by Avagyan et al. (2005) structural data from satellite images reveals an extensional

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and strike –slip tectonics related to the tensional fractures, pull-apart, tail-crack or horsetail structures (e.g. Adiyaman et al. 1998; Karakhanian et al. 2002; Aghamalyan 2004). In the Gegam massif Karakhanian et al. (2004b) generally distinguished three clusters zones Eastern, Central and Western. The orientations of volcanic clusters range from NW–SE to NE–SW in the studied area. Taking into account the cluster dimensions we suggest that the geological structures controlling the big Gegam and Sunik major clusters influence the entire thick of the crust. This is supported by occurrences of basalts, of probable mantle origin, within the fissured areas (Karakhanian et al. 2002). On the other hand, within smaller clusters several km in length magmas with crustal origin are present (Innocenti et al. 1982; Adiyaman et al. 1998). The big clusters are oriented principally NW– SE in central and southern parts of Armenia and are either emplaced on deeply rooted faults of the

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same orientation or are parallel to them (Fig. 11). The same observed in the Ararat volcano area with some relatively more recent volcanoes forming NNW orientated cluster (Fig. 12) which is different with the NE orientation of Ararat massif (Great and Small Ararats). If we consider data for the area including Armenia, Southern Georgia and Eastern Turkey (Adiyaman et al. 1998) we observe another permanent cluster orientation of NNE –SSW sited principally to the west (Fig. 12). Thus the volcanic cluster in the study area forms a cluster wedge with top to the north (Fig. 12). It is important to note that the active faults in the same area also form similar structural wedges with left-lateral strike –slip faulting to the west and rightlateral faulting to the east (Philip et al. 2001). Consequently there is a good spatial correlation between active faults and volcanic activity. Moreover some clusters are controlled by pull-apart and tail-crack structures (Adiyaman et al. 1998; Karakhanian et al. 2002).

Fig. 11. Neogene to Quaternary volcanic clusters on the SE of the Armenian territory. 1, Elongated volcanic cones groups; 2, volcanic clusters well recognized; 3, supposed volcanic clusters; 4, small volcanic clusters; 5, volcanic cones; 6, active faults.

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Fig. 12. (a) Horizontal component of the principal compression (s1) stress orientations from 1, microtectonic data; 2, volcanic clusters (microtectonic data from Philip et al. (1989) for southern Georgia and clusters data of Adyaman et al. (1998) for Eastern Turkey are also used). (b) Rose diagram of s1 stress orientations observed in geological formations of different ages (D-P, Devonian to Permian; J-K, Jurassic and Cretaceous; P, Palaeogene; N-Q, Neogene and Quaternary).

Micro-fault kinematics

Discussion and conclusion

The micro-fault kinematics analyses (Tayan 1980; Philip et al. 1989; Pistre et al. 1995; Arthaud et al. 1998; Avagyan et al. 2005) obtained from the Greater and the Lesser Caucasus area indicate three principal orientations of the compression axis: north –south, NE–SW and NW–SE (Fig. 12). The palaeo-stress orientation variability is observed in the same micro-blocks as in the case of the earthquake P axes data (Avagyan et al. 2005). Neogene to Quaternary formations have only been affected by the last stress phases related to the more developed state of the Arabia –Eurasia collision. In the Neogene to Quaternary formations the dominant directions of compression are north– south and NE–SW (Fig. 13). It is NE –SW and NW–SE for the Palaeogene formations. It is north–south and NW–SE for the Jurassic to Cretaceous formations and it is NE– SW with general north–south orientations for Devonian to Permian formations.

According to the data presented above, the tectonic stress indicators describing the recent (provided by active tectonics and micro-fault kinematics), and the palaeo (also provided by micro-fault kinematics and volcanic clusters) stress fields show scale and temporal changes since the Gondawanian terranes –Eurasia collision and then since the Arabian–Eurasian collision. According to micro-fault kinematics data the NW– SE compressive stress orientation is quite significant in the Jurassic to Palaeogene formations but absent in the Neogene to Quaternary ones (Fig. 12b). Previous fault kinematics results from the Eurasian margin (Greater Caucasus) show evidence for a NW–SE trend of compression in the Mesozoic formations mainly due to an Eocene transpressional stress field (Sebrier et al. 1997). The same compression orientation is described in the northwestern Greater Caucasus at the Paleocene –Eocene boundary (Saintot & Angelier

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Fig. 13. Analysis of microtectonic data (Schmidt’s projection, lower hemisphere) from the Azat river area Quaternary formations (location is shown on Fig. 2). Palaeostress axis (circle, s1; square, s2; triangle, s3) have been calculated by inversion of slip vectors measured on fault planes using the method proposed by Etchecopar et al. (1981).

2002). On the other hand our fault kinematics data are derived from the Lesser Caucasus area and concern the Gondwanian terrane (SAB) and the Mesozoic Eurasian margin series in the North Lesser Caucasus, which collided together during Paleocene to Early Eocene times. Consequently, we assume that this NW –SE compression direction characterizes the stress field due to the first step of the collision (Sosson et al. 2005, 2010). Micro-fault kinematics data from the Neogene to Quaternary series are consistent with a change in the convergence direction during the Middle Miocene, which became sub-meridian to NNE –SSW in the East (Zonenshain et al. 1987). The NE–SW compression orientation in Palaeogene formations can be due to farther Neogene –Quaternary stress field (Fig. 12). The volcanic cluster data clearly predict three compressional orientations in the whole area. In contrast to the micro-fault kinematic data the volcanic cluster orientations show a space distribution with general regularity. They form a wedge with tip to the north like an active fault. This is evidence for the genetic relationship between active faults and volcanism. One part of Neogene to Quaternary

volcanism is rooted on normal fault systems or on strike –slip fault pull-apart structures (Innocenti et al. 1980; Adiyaman et al. 1998; Yilmaz et al. 1998; Karakhanian et al. 2002) as in the case of Vanadzor depression (see above). GPS data evidence a NNE–SSW shortening and an east –west dilatation in the region consistent with active tectonic data showing strike– slip stress regime corresponding to general north– south compression and east –west extension. The general cluster orientation in Gegam and Vardenis massifs is consistent with Davtyan et al.’s (2007) GPS new data showing dilatation with the azimuth of about N608 (Fig. 10). If we take into account the scarce microtectonic stations in the Neogene series (14 stations from 16 measured in Quaternary rocks) showing north –south and NE– SW compression orientation (Fig. 12b) and volcanic clusters suggesting a NNW– SSE shortening one (Fig. 11), we can suppose that Quaternary structures controlling volcanism in the Lesser Caucasus area are inherited from former NW–SE Palaeogene to Neogene faults. These inherited structures also influence the farther stress field.

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While an overall north–south compression orientation can be derived from fault kinematics and volcanic cluster data, more secondary NW– SE and NE–SW compression is also noted. We interpret these two orientations as being the result of two distinct tectonic phases when they were dominant and part of related structures continuing to influence to the farther stress field in various scale. The first one was dominant between the Late Palaeogene and the end of the Early Miocene and the second one has prevailed between the Late Miocene and the Quaternary. Newly obtained data (Faccenna et al. 2006) about tectonic reorganization which occurred in the Late Miocene– Early Pliocene in the Arabia –Eurasian collision area, when the North Anatolian Fault Zone formed, is consistent with our suggestion about tectonic stress axes change. Probably at this time the tectonic regime changed from predominantly compressional (thrusting and reverse faulting) to wrench (transtensiontranspression producing strike–slip faulting with various vertical component) and followed the bending of the originally rectilinear shape of the Lesser Caucasus– Pontides due to the northward motion of the SAB. This review of data and our new results show that stress patterns have evolved with time. These changes in the stress states are related to the changing regime imposed by Gondwanian terrane – Eurasia collision and by the Arabian – Eurasian plate convergence. At a smaller tectonic scale the stress orientation variation is explained by the influence of a dynamics of secondary blocks, structural particularities and clearly inherited structures. This work was supported by the NATO post-doctoral award and the MEBE programme. The authors would like to thank R. Jrbashian (Institute of Geological Sciences, Armenia) E. Barrier and M. Gaetani (leaders of the MEBE programme), for their useful support, J. Angelier and F. Bergerat are also thanked for their constructive and valuable remarks.

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Late Cretaceous to Recent kinematics of SE Anatolia (Turkey) ¨ Z2, P. ERTEPINAR3 & A. KOC¸1 N. KAYMAKCI1*, M. INCEO 1

Middle East Technical University, Department of Geological Engineering, 06531-Ankara, Turkey

¨ niversitesi Mu¨hendislik Faku¨ltesi Jeoloji Mu¨hendislig˘i Bo¨lu¨mu¨, 23119 Elazıg˘, Turkey Fırat U

2

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Forthoofddijk Paleomagnetic Lab. Department of Earth Sciences, Utrecht University, Budapestlaan 17, 3584 CD Utrecht, the Netherlands *Corresponding author (e-mail: [email protected]) Abstract: Five different deformation phases have been recognized in the SE Anatolian orogen and the Arabian Platform based on palaeostress inversion studies using fault-slip data sets. The timing and duration of these phases are determined using various criteria including the age of the affected strata, syndepositional structures, cross-cutting structures and overprinting slickensides. The oldest deformation phase is characterized generally by NE–SW-directed extension. The extension is thought to have resulted from slab-roll back processes during the Maastrichtian to Middle Eocene interval (c. 60 Ma to 40–35 Ma). The second deformation phase is characterized by east– west to NW– SE-directed compression and thought to result from cessation of roll-back processes possibly due to subduction of younger oceanic crust or increase in the convergence rate between Africa and Eurasia during the post-Middle Eocene to Late Oligocene interval (c. 40– 35 Ma to 25 Ma). The third deformation phase is characterized by east– west to NW– SE-directed extension possibly due to slab detachment that initiated in Iran and migrated westwards during the latest Oligocene to Middle Miocene period (25–11 Ma). The fourth deformation phase is characterized by approximately north–south-directed compression due to collision and further northwards indentation of Arabian Plate by the end of Middle Miocene (11– 3.5 Ma). The fifth and present deformation phase is characterized by NE–SW compression which might result from tectonic re-organization in the region since the Middle Pliocene (c. 3.5 Ma to recent).

The Late Cretaceous to recent tectonic development of SE Turkey is related to the closure of the Neotethys Ocean along two trenches namely the southern and northern branches (S¸engo¨r & Yılmaz 1981) while the evolution of these subduction systems partly overlapped in space and time. The Izmir-Ankara-Erzincan Suture Zone (Fig. 1) marks the former position of the northern branch of the Neotethys which separated the Pontides (Eurasian affinity) in the north and the Taurides (Gondwana affinity) in the south. In eastern Turkey, the Taurides are represented by the Keban-Malatya-Bitlis Block (KMBB). The Bitlis–Zagros Suture Zone marks the former position of the Southern Neotethys. During the Mesozoic to pre-Late Miocene, it separated the Taurides from the African– Arabian Plate. There is still an ongoing debate on the exact location and timing of closure of these two oceanic domains. Nevertheless, it is generally accepted that the closure of the northern branch took place, along the ˙Izmir-Ankara-Erzincan Suture Zone, at the end of the Late Cretaceous (Tu¨ysu¨z 1999; Kaymakcı et al. 2003, 2009) to Early Cenozoic (S¸engo¨r & Yılmaz 1981; Go¨ru¨r et al. 1984; Robertson et al. 2005) and gave way to a widespread southwards

thrusting, ophiolitic nappe emplacement and wholesale metamorphism of the KMBB and northern margin of the Taurides. These metamorphic massifs which belong to the northern margin of the Taurides are also known as Anatolides (Ketin 1966). The closure of the southern branch took place along the Bitlis– Zagros Suture Zone during the Middle Miocene (S¸engo¨r et al. 1985; Yig˘itbas¸ & Yılmaz 1996a, b; Huesing et al. 2009; Kus¸cu et al. 2010). However, some argued that its closure took place in the Late Cretaceous and the Neogene deformation is related to post-collisional convergence (e.g. Yazgan & Sussex 1991; Bayarslan & Bingo¨l 2000). Two contrasting styles of thrusting and nappe emplacement define the current structural architecture of eastern and south-eastern Anatolia in addition to a number of generally NE –SW to NNE –SSW striking strike– slip faults (Fig. 2). These are generally south-verging thrust faults along which much of the ophiolitic masses and slivers belonging to the northern Neotethys Ocean were emplaced onto the KMBB during the ¨ zkaya Campanian –Maastrichtian (Perinc¸ek & O 1981). Other important structures are the northverging thrust faults related to the successive

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 409–435. DOI: 10.1144/SP340.18 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Regional tectonic units around Turkey (modified from Okay et al. 1994; Stampfli & Borel 2002). Note that Menderes, Kırs¸ehir, Alanya, Keban-Malatya and Bitlis units are also known as Anatolides (Ketin 1966) which are the metamorphosed equivalents of Taurides (S¸engo¨r & Yılmaz 1981). Note also that Sakarya-Tokat-Kargı, Agvanis-Pulur massifs are part of Rhodope-Pontide Fragments. The Baskil, Maden and Urimeh-Doktar belts belongs to Late Cretaceous to Eocene arc and back arc settings of Bitlis–Zagros Suture zone.

collision of an island arc chain (Afs¸in-Baskil arc) during the Late Cretaceous, located between the KMBB and the Arabian Block (Figs 1 & 2c), within the southern branch of the Neotethys. Terminal collision of the Arabian Plate produced approximately north–NE-striking transcurrent faults which translated Late Miocene to Middle Pliocene thrust faults more than 100 km along their strike. Due to bivergent thrusting, which is due south in the south and due north in the north, gave way to pop-up-like uplift of the KMBB during the Late Miocene to Middle Pliocene (Fig. 2c). The recent studies indicated that the evolution of the region is rather complex. Various ophiolitic and magmatic units in the region are originated from different tectono-magmatic settings including processes possibly related to slab roll-back, slab detachment, supra-subduction zone ophiolite generation during the northwards subduction of southern

branch of the Neotethys (e.g. Yig˘itbas¸ & Yılmaz 1996a, b; Robertson 2002; Parlak 2004; Parlak et al. 2004, 2006; Robertson et al. 2005; Faccena et al. 2005; Kus¸cu et al. 2010). The active tectonic scheme of the region is related to (1) the detachment of the subducting oceanic lithosphere at the northern end of the Arabian plate; (2) development of a Subduction Transform Extension Propagator (STEP) fault (cf. Govers & Wortel 2005) along the northwestern margin of the Arabian Plate (Facenna et al. 2005); and (3) collision and further northwards convergence of the Arabian Plate by the end of the Middle Miocene (c. 11 Ma) (S¸engo¨r et al. 1985; Dewey et al. 1986; Huesing et al. 2009). These processes gave way to inversion of most of the pre-existing structures and development of a number of NNE – SSW to ENE –WSW striking strike –slip fault systems (Perinc¸ek et al. 1987; Westaway & Arger

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Fig. 2. (a) Outline active tectonics of Turkey. (b) Major palaeotectonic (red) and neotectonic (black) structures and Cenozoic basins in central and southeastern Turkey. (c) Simplified regional cross-section along the line XY.

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2001; Kaymakcı et al. 2006) in response to STEP faulting, north–south shortening, crustal thickening (Dewey et al. 1986) and westwards escape of the Anatolian Block (S¸engo¨r et al. 1985). The aim of this study is to present Late Cretaceous to recent kinematic evolution of SE Anatolian Orogen and the northwestern part of the Arabian Platform in order to constrain the timing, mechanism and extend of processes related to subduction of the southern Neotethys Ocean and collision and further northwards convergence of Arabian Plate. The SE Anatolian Orogen, here, refers to the area north of the Bitlis–Zagros Suture and south of Izmir-Ankara-Erzincan Suture Zone in Turkey (Yılmaz 1993).

Tectono-stratigraphy The tectonic units that played a role in the development of the region include the Arabian Platform, Bitlis Massif, Keban-Malatya Platform, non-metamorphosed Tauride Platform units, and various ophiolitic and magmatic complexes comprising Afs¸in-Baskil arc and Maden complex (Figs 1 & 2). These units are tectonically imbricated along north dipping thrust faults and are unconformably overlain by marine to continental carbonates

and clastics deposited during the Cenozoic (Figs 2–4).

Stratigraphy of the Arabian Platform The Arabian Platform comprises sedimentary successions ranging from Precambrian to Recent (Fig. 3). It is generally accepted that it was located at the northern margin of the Gondwana from the Precambrian to the Late Palaeozoic. The Arabian Platform constituted the southern passive margin of the Southern Neotethys Ocean after rifting and northwards drifting of the Taurides from its northern margin during the Mesozoic to Palaeogene (S¸engo¨r & Yılmaz 1981). The oldest compressional structures on the platform are related to the emplacement of ophiolitic units during the Campanian to Maastrichtian ¨ zkaya 1981). These structures and (Perinc¸ek & O thrust contacts are sealed by the Late Maastrichtian to Paleocene and younger units. The important youngest structures that shape the present structural grain of the platform are related to the late Miocene collision and preceding foreland basin development. In this study, we concentrated mainly on the post-Maastrichtian kinematic events within the platform.

Fig. 3. Generalized Cretaceous to Recent tectonostratigraphical column of Arabian Platform in Turkey (modified ¨ zkaya 1981; Yılmaz 1993). Absolute ages correspond to the beginning of deformation phases from Perinc¸ek & O mentioned in the text.

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Fig. 4. Generalized Maastrichtian to Recent tectono-stratigraphical column of SE Anatolian orogen (modified from Yig˘itbas¸ & Yılmaz 1996a).

The sampled interval on the Arabian Platform ranges in age from Maastrichtian to Pliocene and includes a facies association, from bottom to top, few tens of metres thick continental red clastics (Terbu¨zek Formation), chalky marls (Besni Formation), a turbiditic sequence (Germav Formation) composed of alternation of sandstone, siltstone and shale. These facies are overlain by clastics (Gercu¨s¸ Formation) and about ten metres thick neritic limestone (Becirman Formation) (Yılmaz 1993). These units are unconformably overlying the older units and seals the thrust contact between the late Cretaceous emplaced ophiolitic melanges. The thickness of the turbidites (Germav Formation) increases northwards as the grain sizes become coarser and comprises olistostrommal conglomerate lenses. In the southwestern part of the study area, the Paleocene rocks unconformably overlies the ophiolitic units (Karadut & Koc¸ali complexes) which indicates that these units are deposited after the Campanian–Maastrichtian emplacement of the ophiolitic units onto the

Arabian Platform. The Paleocene association is unconformably overlain by Eocene to Lower Miocene rock units. The Eocene to Lower Miocene units have very rapid lateral and vertical facies changes in different parts of the Arabian Platform. Therefore, in the literature, various local names are used for these units. For simplicity we followed the nomenclature of Sungurlu (1974) and Perinc¸ek (1979) who collectively named the Eocene clastic sequence as the Gercu¨s¸ Formation and overlying carbonates as the Midyat Limestone. The Eocene clastic sequence (Gercu¨s¸ Formation) is composed of various clastics including conglomerates, sandstones and mudstones gradually fining upwards as the depositional environments are changing from continental to marine settings. The overlying carbonates (Midyat Limestone) comprise three distinct levels. The lower level is composed of dolomitic limestones that gradually grade into limestones with cherty intervals and limestones with cherty nodules. These levels are also known as Gaziantep Formation (Wilson & Krummenacher

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1957). The cherty levels upwards grade into the medium- to thick-bedded marl, chalky limestone alternations with occasional cherty nodules. This uppermost chalky limestone part of the Midyat Formation is also named as Fırat Formation along the northern edge of the Arabian Plate and is early to Middle Miocene in age (Perinc¸ek 1979). The Fırat Formation also comprises tuff layers in its lower parts and grades northwards laterally into turbidites (Lice Formation) (Tuna 1973; Sungurlu ¨ zkaya 1981). 1974; Perinc¸ek 1979; Perinc¸ek & O The Lower to Middle Miocene Turbidites (Lice Formation, Tuna 1973) comprises northwards and upwards coarsening turbidite facies. It is developed mainly in front of the south-directed thrust faults at the northern edge of the Arabian Platform along the leading edge of the Bitlis–Zagros Suture. It is generally accepted that these turbidites (Lice Formation) mark the terminal stage of subduction of the Southern Branch of the Neotethys and beginning of collision of the Arabian Plate (S¸engo¨r & Yılmaz 1981). They are conformably overlain by continental red clastics which are known also as S¸elmo (Perinc¸ek 1979) and Adıyaman (Tuna 1973) formations (Fig. 3) in the northeastern and northwestern parts of the Arabian Platform respectively. A very widespread isolated alkaline basaltic lava flows are emplaced onto the Arabian Platform. These basaltic lava flows can be subdivided into three distinct groups. The oldest lava flows are observed around the western margin of the Arabian Platform and have lateral gradations with the Fırat Formation. The K– Ar ages of these volcanic rocks range between 21.24 + 2.04 to 9.22 + 0.2 Ma (Yoldemir 1987; Tatar et al. 2004). The second group is exposed in the central part of the Arabian Platform and are unconformably overlying the Fırat Formation and their K –Ar ages range between 12.1 + 0.4 to 7.02 + 0.07 Ma (Yoldermir 1987; Ulu et al. 1991). The third and youngest group belongs to the Karacadag˘ Volcanic Complex and their K –Ar ages range between 0.94 + 0.33 to 0.83 + 0.88 (Pearce et al. 1990).

Stratigraphy of the SE Anatolian orogen The sampled intervals within the SE Anatolian orogen range in age from Maastrichtian to Early Quaternary. The Eocene volcanic and volcaniclastic rocks of the Maden Complex were not studied for palaeostress analysis. The reason for this is because, in the volcanic fields local stress variations are unpredictable and they reflect local stress fields rather than regional stress configurations due to rapid deposition and extensive local vertical movements during volcanic eruptions. However,

contemporary sedimentary sequences were studied in detail (Fig. 4). The Maastrichtian to Paleocene sequences have lateral and vertical facies changes ranging from turbidites (Gu¨rer & Aldanmaz 2002) to pelagic limestones to shallow continental red clastics (Aziz et al. 1982; Perincek & Kozlu 1984). Around the Darende Basin (Fig. 2b), at the bottom, these sequences rest on ophiolitic units with an unconformity and are composed of turbiditic facies comprising conglomerate, sandstone, siltstone and shale alternations which grades upwards into shallow marine limestones. In the southern margin of the Malatya Basin, these units are Maastrichtian in age and are composed of alternation of pelagic limestones and marls intercalated with thin sandstones. The Upper Maastrichtian sequences are unconformably overlain by the Upper Paleocene continental red clastics comprising red to purple conglomerates and sandstone sequences. The lower part of the Upper Paleocene sequence is always associated with normal growth faults (Fig. 5). In the southern margin of the Malatya Basin, this sequence is about 100 m thick and rapidly grades into boulder conglomerates and sandstones which subsequently grades upwards into turbiditic facies alternating with nummulites bearing limestone horizons. This sequence is overlain by Oligocene shallow marine limestones only to the western margin of the Malatya Basin. However, within and in the eastern margin of the Malatya Basin, Oligocene sequences are missing. In these areas, Lower Miocene sequences directly rest either on the metamorphic basement or onto the Eocene sequences. The Lower Miocene sequences start at the bottom with few metres thick shallow marine shales and marls which continues upwards with few tens of metres thick condensed sequence of shallow marine algal limestones. The condensed sequences are conformably overlain by Middle Miocene lacustrine facies in the Malatya and Gu¨ru¨n basins (Figs 2 & 4). The lacustrine sequences grade upwards into fluvio-lacustrine clastics intercalated with coal seams and into a volcano-clastic sequence in the northern part of the Malatya Basin. Based on rodent fauna (Kaymakcı et al. 2006) and 15.2 – 15.9 Ma (Arger et al. 2000) K –Ar ages from the inter-layered lava flows these sequences are of Middle to Late Miocene in age. In the Kangal Basin (Fig. 2), deposition of these sequences continued until the Early Pliocene. All these sequences are unconformably overlain by the Upper Pliocene to Pleistocene fluvial clastics.

Palaeostress inversion Palaeostress analysis aims at the reconstruction of palaeostress configurations using fault-slip data

KINEMATICS OF SE ANATOLIA (TURKEY)

415

Fig. 5. Growth faults developed within the Paleocene clastics (location: Site 28 in Fig. 6, view to NE) and their palaeostress configuration. Note that north– south extension of the growth faults.

sets (Angelier 1994). The back bone of the method is based on the assumption that the slip occurs parallel to the maximum resolved shear stress along an optimally orientated plane of weakness (Wallace 1953; Bott 1959). Basically, there are two types of palaeostress inversion techniques. The graphical approaches generally based on Anderson’s (1951) fault theory which states that the maximum principal stress is the acute-bisetrix of conjugate fault sets while minor stress is the obtuse-bisetrix and intermediate stress is parallel to the intersection of conjugate fault planes. This approach is based on the fact that the earth’s surface is one of the principal planes so that one of the stresses is always vertical. This implies that oblique-slip faults are always re-activated planes of weaknesses. In this approach, the most reliable results can only be obtained in the case of conjugate faults which are, most of the time, very difficult to determine during field studies. On the other hand, the numerical methods are based mainly on the reconstruction of the stress ellipsoid by using the Wallace (1953)–Bott (1959) assumption. In this approach slip occurs along the maximum resolved

shear stress along a pre-existing plane of weakness. Therefore, maximum resolved shear stress direction corresponds to the slip direction which is manifested on the plane as slickenlines (Angelier 1994). During the construction of the palaeostress configurations, the maximum shear stress direction is determined by calculating the axes of the stress ellipsoid by using the shape factor (f ¼ s2 2 s3/s1 2 s3). In which s1, s2 and s3 correspond to maximum, intermediate and minor principal stress magnitudes, respectively. In this study, we have applied only the numerical method developed by Angelier (1988). The method is based on the calculation of a reduced stress tensor which means that the relative magnitudes and orientations of the principal stresses are determined by using the shape factor and orientation of at least four fault slip data belonging to the same deformation phase (see Angelier 1994 for a full account of the method). In addition to direction and magnitudes, one of the other key issues in palaeostress analysis is the dating of the constructed stress configurations. This is accomplished by using various criteria that

416 N. KAYMAKCI ET AL.

Fig. 6. Major faults and sample locations of the study area. Fault names are as in Figure 2.

KINEMATICS OF SE ANATOLIA (TURKEY)

417

Fig. 7. Palaeostress configurations belonging to Paleocene to Middle Eocene extension (phase 1). Note that sites 91–107 belong to the Arabian Platform, all others belong to SE Anatolian orogen. Equal area, lower hemisphere projection.

include, in decreasing order of reliability, synsedimentary structures and faults, the age of the youngest faulted strata and age of the oldest sealing strata, age of the youngest sequences displaced by a specific fault or fault sets (if no seal is

discernable), any incompatible compressional structure overprinted or displaced by an extensional structure (or vice versa), overprinting kinematic indicators within the same fault plane, and overprinting slickensides (Kaymakcı et al. 2000).

418

N. KAYMAKCI ET AL.

Deformation phases Arabian Platform The oldest syn-sedimentary structures are observed with in the Germav Formation in the southwestern part of the study area (site 3 in Fig. 6). In this area, a number of synsedimentary normal faults with graben-horst morphology are observed. Most of the horsts and grabens were tilted possibly due to successive tectonic events. During the palaeostress analyses these structures were tilt corrected. During correction, mean bedding plane attitudes were used as the palaeo-horizontal surface. In the case of listric normal faults, along which differential rotation about horizontal axes is expected, bedding planes of the footwall blocks were used. From the Upper Maastrichtian to Eocene sites, 17 palaeostress configurations are constructed (Fig. 7, Table 1). Based on this data it is concluded that the region has experienced NE –SW extension during the Paleocene to Eocene. This phase of extension is accepted to be the first phase of deformation after the Campanian –Maastrichtian emplacement of the ophiolitic units on to the ¨ zkaya 1981). Arabian Platform (Perinc¸ek & O

Since Midyat Formation was deposited during the Eocene to Early Miocene interval, the upper boundary of this phase could not be constrained precisely and it is accepted roughly as Eocene based on the correlation of stress configurations with the other regions (see below). In most of the sites overprinting slickensides indicating compressional deformation were also observed. These slickensides were constrained mainly to the Upper Eocene–Oligocene (Gaziantep Formation) and older units. In addition to this, a new set of extensional structures characterized generally by growth faults are also observed within the chalky limestone and volcanogenic parts of the Fırat Formation (Fig. 8). The orientation of palaeostress configurations constructed from the growth faults within the Fırat Formation yielded different extensional stress configurations compared to the ones observed in the Upper Cretaceous to Paleocene units. These stress configurations, as being postdating compressional structures observed in preOligocene units are developed within the Lower part of the Fırat Formation (Lower Miocene). This indicates a new phase of extension. In addition, in the Malayat Basin it is well documented by Kaymakcı et al. (2006) that this extensional

Table 1. Geographical coordinates and palaeostress orientations for the phase 1 Site code 11A 14 16 25B 28A 30B 31 42 52D 53 56 63 64 65 79 80B 81A 83A 83D 84B 87B 88A 95A 100A

X*

Y*

s1 (D/P)†

s2 (D/P)

s3 (D/P)

F‡

N§

39.161 38.804 38.509 37.919 37.231 38.357 38.357 38.097 37.373 37.442 37.456 36.334 36.330 36.442 34.761 34.813 34.792 34.916 34.916 34.990 35.141 35.235 38.008 38.155

38.674 38.741 38.978 38.383 37.254 38.348 38.348 39.425 39.015 38.968 38.634 39.046 39.055 38.904 37.260 37.209 37.322 37.497 37.497 37.649 38.172 38.200 38.829 38.317

309/69 231/77 107/71 153/83 138/67 320/61 146/66 029/65 268/45 296/15 290/64 237/66 223/69 217/69 228/62 203/68 341/73 70/48 13/80 153/69 329/65 166/52 221/68 176/47

130/21 330/02 02/05 313/07 263/14 141/29 312/24 162/18 139/32 68/68 118/26 140/3 353/14 124/2 348/15 338/16 134/15 185/21 103/0 309/20 130/24 342/38 73/19 003/42

40/00 61/13 270/19 044/02 357/19 051/00 44/04 258/17 30/28 202/15 26/3 49/24 87/16 33/21 85/23 72/15 226/8 290/35 193/10 42/8 223/7 73/2 339/11 34.13

0.51 0.84 0.16 0.65 0.34 0.56 0.54 0.68 0.33 0.97 0.98 0.15 0.20 0.16 0.55 0.86 0.41 0.54 0.83 0.68 0.47 0.86 0.48 0.11

14 11 14 7 10 8 8 9 6 16 14 12 17 21 8 9 31 35 7 7 10 15 8 9

*X,Y, coordinates in decimal degrees. † D/P, direction/plunge. ‡ F, shape factor. § N, number of faults.

KINEMATICS OF SE ANATOLIA (TURKEY)

419

Fig. 8. Growth faults developed within the Upper Oligocene– Lower Miocene Fırat Formation. Note increase in throw from a –a0 to c– c0 . (Site 94 in Fig. 6, view to west).

deformation phase started at the end of Oligocene and lasted until the Middle Miocene. Based on this information, it is concluded that the extensional structures belong to deformation phase 3 which took place possibly latest Oligocene to Middle Miocene. Therefore, the compressional structures predating phase 3 and postdating phase 1 are classified into deformation phase 2 and thought to have taken place during the post-Middle Eocene to preLatest Oligocene. The palaeostress configurations and stress orientations for phases 2 and 3 are depicted in Figures 9 and 10 and Tables 2 and 3.

Within the Middle to Upper Miocene Adıyaman Formation a very well developed mesoscopic thrust fault occasionally with overprinting slickensides was observed. The older slickensides were related to generally east –west-striking conjugate thrust faults dipping 458 to 208 due north and due south. The hanging wall and footwall blocks of these thrust faults were displaced locally by NE–SWtrending strike–slip faults. The palaeostress configurations for these thrust faults, older overprinted slickensides, and the younger overprinting slickensides were expectedly found to be different

420

N. KAYMAKCI ET AL.

Fig. 9. Palaeostress configurations belonging to Late Eocene– Oligocene compression (phase 2). Note that sites 91– 104 belong to the Arabian Platform, all others belong to SE Anatolian Orogen. Equal area, lower hemisphere projection.

(Figs 11 & 12 and Tables 4 & 5). Based on this information, it is interpreted that the reverse faults predate the strike– slip faults and they belong to a new phase of deformation (phase 4) that took place during the deposition of the S¸elmo and Adıyaman formations in the Late Miocene to Pliocene. Both the hanging wall and footwall blocks of the thrust faults were displaced by these strike –slip faults, therefore, they can not be tear faults of the thrust faults. Based on this observation it is concluded that the strike –slip faults and younger overprinting slickensides should belong to the younger phase of deformation (phase 5) which has been active since the Middle Pliocene (c. 3.5 Ma) as proposed by Kaymakcı et al. (2006).

SE Anatolian orogen In the western and southern margin of the Malatya Basin and southeastern margin of the Ulukıs¸la Basin (Fig. 2) well-developed growth faults were observed where Paleocene to Middle Eocene sequences directly rest on to the older units. For example in sites (27 –30 and 39) at the contact between the basement metamorphics and the

Maastrichtian to Paleocene red clastics a number of growth faults were observed (Fig. 5). Also in the southern margin of the Malatya Basin around sites 30– 31 growth faults are observed at the contact between the Paleocene and Eocene units (Fig. 13). Similar relationships are also observed in various localities throughout the region. Based on these information it is concluded that Maastrichtian to Middle Eocene rocks in the region are deposited during a phase of extensional deformation (phase 1) which might have started during the end of Maastrichtian and continued until the Middle Eocene. The stress configurations for each sites for the deformation phase 1 is depicted in Figure 7 and Table 1. Within the Oligocene units a number of mesoscopic faults were observed. Some of these faults have overprinting slickensides. However, no diagnostic features that could be used for dating the activity of these faults were not observed during the field studies. On the other hand, a number of normal growth faults were observed within the Lower to Middle Miocene units. For example, in the western margin of the Malatya Basin, at an unconformable boundary between the

KINEMATICS OF SE ANATOLIA (TURKEY)

421

Fig. 10. Palaeostress configurations belonging to latest Oligocene to Middle Miocene extension (phase 3). Note that sites 93–110 belong to the Arabian Platform, all others belong to SE Anatolian orogen. Equal area, lower hemisphere projection.

Eocene limestones and Lower Miocene limestones, numerous normal faults were observed. Some of the normal faults displaced the Eocene units while they do not propagate into the Miocene units. Palaeostress configurations constructed from the faults displacing only the Eocene units and the ones displacing both Eocene and Miocene units expectedly yielded completely different stress configurations. In addition, the stress configurations for the ones sealed by Miocene and those displaced the Miocene rocks are similar to the orientations of the deformation phase 1 and to deformation phase 3 configurations, respectively. At this locality and elsewhere overprinting slickensides indicating reverse faulting are observe within the Eocene and older units. The palaeostress orientations

constructed from the younger slickensides and those observed within the Oligocene units found to be similar. In addition to this, around the southern margin of the Gu¨ru¨n Basin (Fig. 2) a number mesoscopic thrust faults were observed between the Mesozoic and Eocene units. These thrust faults are sealed by Middle Miocene lacustrine deposits of the Gu¨ru¨n Basin. Similar relationships are also observed around the sites 50– 90 around Gu¨ru¨n, Sarız and Ecemis¸ faults (Fig. 2). Considering extensional nature of deformation during the Paleocene to Middle Eocene and age of the sealing lithologies as being Miocene, it is concluded that compressional deformation observed within the Eocene and Oligocene units indicates that a phase of compressional deformation took place

422

N. KAYMAKCI ET AL.

Table 2. Geographical coordinates and palaeostress orientations for the phase 2 Site code 7B 8B 23 26A 32B 36B 48 52 54 55 56B 58 59B 62 67B 71C 72 78B 82A 96C 104E

X*

Y*

s1 (D/P)†

s2 (D/P)

s3 (D/P)

F‡

N§

37.813 37.856 37.865 37.787 38.807 36.957 37.976 37.373 37.334 37.427 37.456 37.437 37.359 37.289 36.455 36.306 36.465 34.749 34.877 38.328 38.497

38.991 38.972 38.189 36.996 38.427 37.183 38.745 39.015 38.938 38.688 38.634 38.622 38.705 38.770 38.665 38.322 37.902 37.287 37.470 38.644 37.368

120/17 291/16 164/13 121/05 103/28 096/22 205/10 242/6 298/38 269/19 227/6 297/7 271/25 27/14 231/22 78/26 256/8 224/33 149/43 261/22 263/3

301/73 152/70 269/50 029/21 305/60 243/65 352/781 335/24 150/47 43/63 94/82 199/47 97/65 291/21 321/1 254/64 165/9 62/56 358/43 164/18 355/33

210/00 25/12 064/37 223/68 198/10 001/13 114/06 138/65 41/16 173/18 318/6 33/42 2/3 148/64 54/68 348/2 25/78 319/8 254/15 38/61 168/57

0.36 0.92 0.32 0.12 0.88 0.92 0.79 0.25 0.65 0.58 0.63 0.82 0.44 0.43 0.33 0.60 0.67 0.56 0.86 0.34 0.30

9 8 6 11 10 10 5 12 5 17 8 5 6 8 5 6 16 6 13 5 18

*

X,Y, coordinates in decimal degrees. D/P, direction/plunge. ‡ F, shape factor. § N, number of faults. †

after Eocene (phase 1) and before the Miocene (phase 3) extension phases. Therefore, phase 2 was concluded to have taken place during Late Eocene –Oligocene interval. Early Miocene to recent palaeostress evolution of Malatya Basin is discussed in Kaymakcı et al. (2006). Therefore, in this study we are concentrated on the regions to the west of Malatya Basin. Similar to Malatya Basin within the Late Miocene to recent units a number of large scale and of mesoscopic scale faults with reverse and strike–slip character are observed. One of the most prominent of these structures is observed at the southern margin of Kangal Basin which is delimited by a thrust fault along which the Upper Cretaceous opholitic melange thrust over the Upper Miocene (Messinian) to Lower Pliocene infill of the Kangal Basin (Fig. 14). In addition to this, along the Sarız Fault a number of strike–slip faults with reverse components are observed within the Upper Miocene to Pliocene units. Similar relationships are also observed within the Ecemis¸ Fault where strike –slip faults displaced the Upper Miocene units are sealed by Plio-Quaternary deposits. Likewise, the Plio-Quaternary deposits are locally displaced by mesoscopic faults. The ages of these units are not well constrained. Therefore,

the fault-slip data collected from these units are separated and analysed independently. They produced completely different stress configurations compared to the ones observed within the Upper Miocene to Pliocene units. Therefore, we followed the deformation phase schema outlined in the previous section and concluded that the deformation phase 4 corresponds to Late Miocene to Early Pliocene interval and the deformation phase 5 commenced at the Middle Pliocene (c. 3.5 Ma) and is still active (see also Kaymakcı et al. 2006).

Discussion As discussed above, in the eastern and southeastern Anatolia five different deformation phases have been recognized (Figs 15 –19). The duration of these phases are constrained by using various stratigraphical, structural and kinematic criteria as outlined in Angelier (1994) and Kaymakcı et al. (2000). The oldest deformation phase is characterized by generally NE –SW directed extension both in the Arabian Platform and north of it (phase 1). Together with stratigraphical information, this information suggests that the extensional deformation is not a local phenomenon but extends from the Arabian Platform to the eastern Turkey

KINEMATICS OF SE ANATOLIA (TURKEY)

423

Table 3. Geographical coordinates and palaeostress orientations for the phase 3 Site code 1B 11C 18 19 21 22B 24B 25A 26B 29 30A 41 45B 61 67 68 73D 78 81B 83 83C 85A 85C 86B 88B 93C 94B 96B 100B 102B 103F 105A 107A 110A

X*

Y*

s1 (D/P)†

s2 (D/P)

s3 (D/P)

F‡

N§

37.321 39.161 38.415 38.333 37.992 37.849 37.957 37.919 37.787 37.231 38.267 38.097 38.489 37.458 36.455 36.432 36.447 34.749 34.792 34.916 34.916 34.995 34.995 35.054 35.235 36.965 38.210 38.702 38.155 37.919 38.628 38.325 38.217 38.005

39.184 38.674 39.028 39.008 38.213 38.139 38.357 38.383 36.996 37.254 38.278 39.425 39.283 38.467 38.665 38.585 37.891 37.287 37.322 37.497 37.497 37.663 37.663 37.849 38.200 36.846 38.329 38.602 38.317 38.383 37.306 37.474 37.536 37.707

220/66 206/66 265/67 028/69 321/66 094/47 261/69 096/79 090/73 029/71 052/67 127/86 255/86 182/29 92/66 22/56 40/63 301/67 210/65 327/64 135/80 340/55 243/48 186/47 200/65 13/46 317/17 124/75 124/71 56/70 086/42 212/3 264/56 359/23

47/23 68/18 37/16 223/20 060/04 220/29 052/18 212/05 213/09 216/19 236/23 014/02 045/04 347/60 356/2 190/34 188/23 80/18 62/21 190/19 344/09 205/27 34/38 27/41 45/23 268/13 154/72 232/4 226/4 292/11 243/36 104/80 6/8 243/46

316/02 333/15 132/16 131/05 152/23 328/29 145/09 303/10 305/14 125/02 145/02 284/04 135/02 88/6 265/24 284/6 284/12 175/14 328/12 94/16 253/05 104/21 136/15 288/11 311/10 167/41 49/5 323/14 318/18 199/16 349/08 302/10 101/33 106/35

0.47 0.76 0.78 0.62 0.01 0.56 0.55 0.87 0.43 0.66 0.60 0.58 0.69 0.38 0.94 0.67 0.76 0.44 0.31 0.63 0.33 0.59 0.22 0.41 0.46 0.78 0.98 0.52 0.11 0.31 0.42 0.47 0.35 0.39

8 11 7 8 8 7 6 13 8 7 10 14 6 12 17 13 5 12 9 21 24 12 5 6 8 6 5 10 9 6 9 25 8 9

*

X,Y, coordinates in decimal degrees. D/P, direction/plunge. F, shape factor. § N, number of faults. † ‡

(Fig. 15). The extension in the region postdates the Campanian to Maastrichtian emplacement of ophiolitic nappes, in the eastern Anatolia, over the ¨ zkaya 1981; PeriTauride platform (Perinc¸ek & O nc¸ek & Kozlu 1984; Yig˘itbas¸ & Yılmaz 1996a, b). Geographically, regional extension during the Paleocene to Eocene reaches as far north as to the Eastern Black Sea Basin. Similarly, Spadini et al. (1996) have proposed that Eastern Black Sea Basin was opened as a back-arc basin during the Paleocene to Eocene interval. Therefore, we propose that Paleocene to Eocene extension in the eastern and southeastern Turkey is related to increase in the subduction rates of the southern Branch of the Neotethys (cf. Facenna et al. 2001) and decrease in the rate of convergence between Africa and Eurasia c. 60 Ma (Dewey et al. 1989;

Allen et al. 2004). This, possibly, gave way to slab roll-back processes and widespread extension, mainly during the Eocene (Robertson et al. 2005; Kuscu et al. 2010). This extension caused the opening of the Eastern Black Sea Basin and NE– SW stretching of the Tauride Block which gave way to exhumation of metamorphic massifs including Bitlis-Po¨tu¨rge, Keban and Malatya metamorphics onto which Paleocene to Eocene units are deposited non-conformably (Fig. 20a). On the contrary, Kaymakcı et al. (2000, 2003) have reported that during the Late Paleocene to Oligocene interval, western and the north Central Turkey was under compression during which a number of piggy-back basins in the Central Pontides were developed (Gu¨rer & Aldanmaz 2002; Hippolyte et al. 2010). This relationship implies

424

N. KAYMAKCI ET AL.

Fig. 11. Palaeostress configurations belonging to the Upper Miocene to Pliocene compression (phase 4). Note that sites 91– 108 belong to the Arabian Platform, all others belong to SE Anatolian orogen. Equal area, lower hemisphere projection.

that the tectonic processes in the Eastern Turkey and central north to western Turkey were different during the Paleocene to Eocene interval. The Paleocene to Eocene extension is replaced by compressional deformation during the Late Eocene to Oligocene (phase 2). During this phase the direction of compression was relatively

uniform and orientated NW–SE in the Arabian Platform while it changed from east –west to NNW– SSE north of it (Fig. 16). We believe that the change in the compression directions in the north of the Arabian Plate within the Taurides was possibly due to inversion of some of the Paleocene –Eocene extensional structures and

KINEMATICS OF SE ANATOLIA (TURKEY)

425

Fig. 12. Palaeostress configurations belonging to late Pliocene to recent compression (phase 5). Note that sites 91–107B belong to the Arabian Platform, all others belong to SE Anatolian orogen. Equal area, lower hemisphere projection.

pervasive internal deformation of the region. Uniform compression directions in the Arabian Plate might indicate lack of major internal deformation which may indicate relative rigidity of the plate.

We speculate that the resumption of compression (Fig. 16) during the Late Eocene– Oligocene interval (Fig. 20b) might be due to subduction of mid-oceanic ridge and/or younger oceanic crust (cf. Hafkenshade et al. 2006). It is

426

N. KAYMAKCI ET AL.

Table 4. Geographical coordinates and palaeostress orientations for the phase 4 Site code 1A 7A 8A 9A 11B 12 17 22A 23 24A 32A 38 45A 46 47 52B 55B 57 59 60 66 66B 69 71 72B 73 74 75 75B 76 77B 80 84 85B 86 87C 90 91B 93B 94A 99A 100C 102A 103A 103D 104A 104B 104C 108A 108B *

X*

Y*

s1 (D/P)†

s2 (D/P)

s3 (D/P)

F‡

N§

37.321 37.813 37.856 37.981 39.161 38.959 38.487 37.849 37.865 37.957 38.807 38.097 38.489 38.100 38.081 37.373 37.427 37.413 37.359 37.458 36.450 36.450 36.503 36.306 36.465 36.447 36.407 35.835 35.835 35.903 35.873 34.813 34.990 34.995 35.054 35.141 35.354 37.282 36.965 38.451 38.266 37.666 37.919 38.628 38.628 38.497 38.497 38.497 38.160 38.160

39.184 38.991 38.972 38.742 38.674 38.714 39.012 38.139 38.189 38.357 38.427 39.425 39.283 38.641 38.646 39.015 38.688 38.559 38.705 38.467 38.683 38.683 38.426 38.322 37.902 37.891 37.822 37.539 37.539 37.823 37.856 37.209 37.649 37.663 37.849 38.172 38.152 37.022 36.846 38.770 38.271 36.991 38.383 37.306 37.306 37.368 37.368 37.368 37.720 37.720

169/16 164/9 191/19 210/4 338/34 352/14 008/43 330/14 164/13 349/30 343/05 188/12 157/04 161/17 356/01 347/20 144/5 181/24 322/12 182/29 359/13 346/42 346/20 169/9 347/10 135/4 336/32 154/10 139/37 124/25 345/7 154/3 166/8 150/35 162/10 150/20 347/26 156/5 131/19 159/10 179/38 333/14 82/72 72/19 151/32 262/12 333/21 318/12 36/2 146/11

359/74 294/76 312/57 309/66 147/55 207/73 197/46 062/07 269/50 167/60 102/80 359/78 038/82 341/73 091/81 106/53 46/57 299/46 220/45 347/60 101/43 157/47 141/68 265/35 81/20 39/60 85/27 246/14 329/52 273/62 247/50 244/1 57/65 246/8 65/39 304/68 157/64 254/58 327/70 322/79 357/52 237/26 214/12 335/18 247/ 9 358/26 91/50 48/1 306/4 40/55

259/3 73/11 91/26 119/23 245/05 84/09 102/05 178/74 064/37 259/01 252/09 097/02 248/07 251/00 266/09 245/29 237/33 73/34 63/42 88/6 256/44 252/4 253/8 67/54 233/68 227/29 207/45 31/73 233/5 28/13 80/39 353/87 259/23 346/53 264/50 57/9 255/4 63/32 223/5 68/3 88/1 90/60 307/13 204/63 350/57 149/61 229/32 145/78 148/85 243/33

0.56 0.63 0.64 0.64 0.61 0.47 0.62 0.12 0.32 0.53 0.63 0.52 0.63 0.35 0.40 0.34 0.42 0.43 0.09 0.38 0.50 0.67 0.59 0.38 0.64 0.39 0.41 0.21 0.02 0.81 0.26 0.01 0.61 0.45 0.45 0.43 0.52 0.19 0.57 0.41 0.70 0.29 0.59 0.58 0.72 0.56 0.35 0.77 0.78 0.44

13 25 10 11 10 11 15 8 6 13 32 11 23 10 16 14 5 8 28 16 18 7 22 7 7 22 13 27 7 14 8 24 33 6 16 12 14 5 8 9 16 8 7 7 17 15 8 7 22 8

X,Y, coordinates in decimal degrees. D/P, direction/plunge. F, shape factor. § N, number of faults. † ‡

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427

Table 5. Geographical coordinates and palaeostress orientations for the phase 5 Site code 2 4 5 10 13 15 22A 23 27A 35 36A 38 44 48B 51 52C 53B 59B 70 76B 77 82A 83D 87 87D 88C 88D 91B 93B 93E 94A 94B 95B 99A 102A 104B 108C

X*

Y*

s1 (D/P)†

s2 (D/P)

s3 (D/P)

F‡

N§

37.473 37.668 37.703 38.047 38.817 38.692 37.849 37.865 37.512 37.151 38.328 38.097 38.489 37.976 37.356 37.373 37.442 37.359 36.406 35.903 35.873 34.877 34.916 35.141 35.141 35.235 35.235 37.282 36.965 36.970 38.451 38.210 38.618 38.266 37.919 38.497 38.160

39.166 39.082 39.057 38.722 38.736 38.827 38.139 38.189 37.059 37.303 38.644 39.425 39.283 38.745 39.028 39.015 38.968 38.705 38.281 37.823 37.856 37.470 37.497 38.172 38.172 38.200 38.200 37.022 36.846 36.807 38.770 38.329 38.839 38.271 38.383 37.368 37.720

11/8 25/6 41/25 17/26 34/5 199/14 330/14 164/13 004/59 008/02 018/02 188/12 226/00 205/10 357/14 29/0 213/0 271/25 194/8 190/18 358/27 149/43 13/80 1/30 220/38 230/1 5/46 156/5 131/19 10/17 159/10 317/17 158/58 179/38 82/72 333/21 146/11

117/62 121/49 184/62 168/60 301/26 58/72 062/07 269/50 178/31 098/15 110/43 359/78 136/01 352/781 100/40 119/21 121/82 97/65 73/74 311/59 169/63 358/43 103/0 174/60 45/52 321/44 106/10 254/58 327/70 171/60 322/79 154/72 199/46 357/52 214/12 91/50 40/55

277/27 290/41 305/15 280/12 134/64 292/11 178/74 064/37 269/03 272/75 286/47 097/02 338/89 114/06 252/46 298/69 303/8 2/3 286/14 91/25 266/4 254/15 193/10 269/3 312/3 139/46 205/42 63/32 223/5 275/8 68/3 49/5 102/7 88/1 307/13 229/32 243/33

0.31 0.04 0.16 0.34 0.16 0.38 0.12 0.32 0.50 0.40 0.30 0.52 0.16 0.80 0.37 0.42 0.51 0.44 0.50 0.60 0.42 0.86 0.83 0.59 0.77 0.32 0.80 0.20 0.57 0.49 0.41 0.98 0.46 0.70 0.58 0.35 0.44

11 10 6 7 8 7 8 6 30 7 18 11 9 5 13 9 6 10 13 9 22 12 7 16 6 24 7 5 8 8 9 5 7 16 7 8 8

*

X,Y, coordinates in decimal degrees. D/P, direction/plunge. ‡ F, shape factor. § N, number of faults. †

important to note that the Late Eocene –Oligocene compression is very widespread and extends from eastern Turkey to Aegean region (S¸engo¨r & Yılmaz 1981; Okay et al. 2001) and further west into Greece (Van Hinsbergen et al. 2005). As an indirect evidence for the Late Eocene –Oligocene compression, during this time interval, most areas in Turkey is characterized generally by erosion and period of local continental deposition (Kaymakcı 2000; Gu¨rer & Aldanmaz 2002) which might be coupled with aridification and sea level drop (cf. Dupont-Nivet et al. 2007). Starting from the latest Oligocene, the Early Miocene is characterized by widespread shallow

marine deposition in eastern and south-eastern Turkey which was previously covered by Upper Eocene to Oligocene continental deposits. During this time interval, most of the north– NE- to south – SW-trending normal faults were developed under east –west to NW–SE orientated extension (phase 3, Figs 17 & 20c). The tectonic cause of Lower Miocene extension is thought to relate with the westwards migration of detached slab which started in Iran during the Oligocene and reached to southeastern Turkey during the Middle Miocene. This gave way to a short period of localized extension that the topography of eastern Turkey was about sea level. Topographical lowering is followed by uplift

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Fig. 13. Growth faults developed within the Lower Eocene units and resultant stress configuration (sites 11A and 30B in Fig. 6, view to north).

Fig. 14. Close-up view of the southern boundary thrust fault of the Kangal Basin (location 51 in Fig. 6).

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429

Fig. 15. Major structures and extension directions (i.e. horizontal component of s3) in south-eastern Turkey for the Late Paleocene to Middle Eocene period which corresponds to deformation phase 1. Deviations from NE–SW extension are possibly due to block rotations in the successive deformation phases.

Fig. 16. Late Eocene to Oligocene compression (large arrows) and extension (small arrows) directions (i.e. horizontal components of s1 and s3) in southeastern Turkey.

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Fig. 17. Latest Oligocene–Middle Miocene extension directions (i.e. horizontal component of s3) in southeastern Turkey.

Fig. 18. Upper Miocene to Pliocene (phase 4) compression (large arrows) and extension (small arrows) directions (i.e. horizontal components of s1 and s3) in southeastern Turkey.

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431

Fig. 19. Upper Pliocene to Recent (phase 5) compression (large arrows) and extension (small arrows) directions (i.e. horizontal components of s1 and s3) in southeastern Turkey. As compared to Figure 18 there is slight but significant variation in the s3 directions. This variation is possibly due to tectonic reorganization in eastern Turkey related to eastwards jump of the junction between the north and east Anatolian fault Zones as proposed by Westaway & Arger (2001).

as the detachment migrated westwards (cf. Buiter et al. 2002). This gave way to continental deposition in Malatya (Kaymakcı et al. 2006; Tu¨rkmen et al. 2007), and adjacent basins in the north of the Bitlis Suture while foreland flysch deposition ¨ zkaya 1981; continued south of it (Perinc¸ek & O Karig & Kozlu 1990; Huesing et al. 2009). The deformation phase 4 (Fig. 18) is related to the collision and further northwards convergence of the Arabian Plate into the Eurasian Plate. The beginning of the collision is generally accepted as the end of Serravalian (11 Ma) based on the youngest marine deposition in south– east Anatolia (S¸engo¨r & Yılmaz 1981) which coincides with Middle Miocene climate transition (Huesing et al. 2009). The collision gave way to approximately north–south compression; crustal thickening (Dewey et al. 1986) and westwards escape of Anatolian Block (S¸engo¨r et al. 1985). During the Middle Pliocene (c. 3.5 Ma) a tectonic reorganization occurred in the region which might be related to the complete emergence of the North and the East Anatolian fault zones as independent lithospheric structures. This resulted in slight but

relatively significant change in the stress configurations (phase 5, Fig. 19) coupled with the westwards escape of the Anatolia and eastwards jump of the junction between the north and east Anatolian fault Zones as proposed by Westaway & Arger (2001).

Conclusions This study has reached following conclusions. † The Arabian Platform and the SE Anatolian orogen have evolved into five different deformation phases. † The oldest deformation phase took place during the Paleocene to Middle Eocene interval and is characterized by NE –SW extensional deformation resulted from roll-back of northwards subducting southern branch of the Neotethys lithosphere. † The second deformation phase took place during the Late Eocene to Oligocene interval and is characterized by east –west to NW–SE compression resulted from termination of roll-back

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Fig. 20. Two dimensional approximately north–south orientated conceptual cross-sections from Eastern Black Sea to Arabian Platform explaining the Paleocene to Early Miocene evolution of the region. (a) Slab roll-back process during the Paleocene to Middle Eocene in eastern Turkey resulted in far reaching extensional deformation in eastern Turkey which also resulted in the opening of the Eastern Black Sea Basin. During this time interval extension on the Arabian Platform was due to slab-pull forces of the down-going slab. (b) During the Late Eocene to Oligocene convergence between Pontides and Arabian Plate resumed and due to subduction of relatively younger and lighter oceanic crust gave rise to compressional deformation in the region (both on the Arabian and over-riding plates). (c) Towards end of Oligocene slab started to detach, this gave way to short time subsidence (as evidenced by Early Miocene marine deposits i.e. Adilcevaz Fm.) followed by uplift of the over-riding plate that resulted in uplift of the basins while extensional deformation was prevailed. This gave rise to continental deposition characterized mainly by fluvio-lacustrine environments (e.g. Gu¨ru¨n, Malatya, Kangal and Sivas basins). In the south, marine deposition along the trench continued until the end of Middle Miocene (Kaymakcı et al. 2006; Huesing et al. 2009).

KINEMATICS OF SE ANATOLIA (TURKEY)

possibly due to subduction of younger oceanic lithosphere. † The third deformation phase took place during the Latest Oligocene to Middle Miocene interval and characterized by NW –SE directed extension resulted from west wards migrating detachment of the Neotethyan lithosphere. † The fourth and fifth deformation phases are relate to terminal subduction that gave way to collision and indentation of Arabian Plate into the Eurasian Plate and complete emergence of north and east Anatolian faults zones as independent lithospheric structures. This study is supported by Middle East Basins Evolution (MEBE) project between 2003– 2005. We would like to thank Ilkay Kus¸cu for fruitful discussions about the evolution of Tethys system.

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Magmatism in the southeastern Anatolian orogenic belt: transition from arc to post-collisional setting in an evolving orogen ILKAY KUSCU1,2*, GONCA GENCALIOGLU KUSCU1,2, RICHARD M. TOSDAL2, THOMAS D. ULRICH2 & RICHARD FRIEDMAN2 1

Department of Geological Engineering, Mugla University, TR-48100, Turkey 2

Mineral Deposit Research Unit, Department of Earth and Ocean Sciences, University of British Columbia, V6T 1Z4 Vancouver, BC, Canada *Corresponding author (e-mail: [email protected]; [email protected])

Abstract: Late Cretaceous to Middle Eocene calc-alkaline to alkaline magmatic rocks emplaced within the southeastern Anatolian orogenic belt, the most extensive magmatic belt in Turkey, result from the complex collision between the Afro-Arabian and Eurasian plates and the subduction of the southern and northern Neotethyan oceanic basins beneath the Eurasian continental margin during the Alpine– Himalayan orogeny. In a transect in east-central Turkey extending from Baskil (Elazig) to Divrigi (Sivas) to the north, and from Copler (Erzincan) to Horozkoy (Nigde) to the SW, these magmatic rocks vary in time, spatial distribution, and composition. 40 Ar/39Ar ages supplemented by a few U– Pb ages geochronology from major plutons demonstrate a general younging of magmatism in the transect from c. 83 Ma in the south (Baskil) to c. 69 Ma in the north (Divrigi-Keban), followed by a c. 44 Ma scattered magmatic complex now found along a NE trending arcuate belt between Copler and Horoz. In general, trace element and rare earth element (REE) geochemistry in the magmatic rocks suggest two main sources for the melts: (1) a mantle-wedge and subducted oceanic lithosphere producing arc-type magma; and (2) metasomatized lithospheric mantle modified by subduction producing magmatic rocks with more metasomatized mantle and within plate signatures. The combination of geochemical and geochronological data presented herein provides a basis to reconstruct the temporal and spatial transition from subduction-related to post-collision and to late-orogenic magmatism in the eastern Mediterranean region. Subduction-related magmatism is rooted to closure of the Neo-Tethyan Ocean whereas post-collision and late orogenic-within plate-related magmatism is driven by the collision of a northern promontory of the SE Anatolian orogenic belt with northerly derived ophiolitic rocks. The magmatic transition occurs regionally in northerly to northwesterly trending belts in the southeastern Anatolian orogenic belt. The magmatism exhibit a clear shift from deep seated arc-type to late-orogenic from south (Baskil) to more deeply eroded mid-crustal plutons at the north (Divrigi), then to magmatism related to incipient slab-rupture from northeast (Copler, Kabatas, BizmisenCalti) to SW (Karamadazi and Horoz). The age progression follows a south-to-north geochemical trend of decreasing crustal input into mantle-derived magmas, and is explained as a consequence of slab roll-back after the collision/obduction of northerly ophiolites followed by slab steepening and incipient rupture leading to transtensional block faulting and subsidence, and thus to the preservation of near-surface magmatic products along a NE trending belt.

Complex earth movements have shaped the Anatolia (Asia Minor) region of Turkey for millions of years. Except for a small part of the country along the Syrian border that is a continuation of the Arabian Platform, Turkey is geologically part of the great Alpine –Himalayan belt, known also as Tethyan collage, which extends from the Atlantic Ocean to the Himalayan Mountains. This belt, which is about 7000 km long, was formed during the Late Cretaceous to Quaternary period, as the Arabian, African, and Indian continental plates began to interact with the Eurasian plate. Caught between colliding continents, the Mesozoic and Cenozoic Tethyan collage stretching from Europe

across southern Eurasia is an extremely complex geological terrane. Nonetheless despite the complexity, the geodynamic setting of the Tethyan collage is reasonably well known. Within the Tethyan collage, the southeastern Anatolian orogenic belt (SEAOB) in east-central Turkey resulted from the closure of former Mesozoic ocean basins, including the incorporation of a number of smaller continental microplates, during the final collision of Arabia with Eurasia. The SEAOB is an arc-shaped orogenic belt (Fig. 1a–c) that extends across Turkey towards western and northwestern Iran (Fig. 1b). In detail, northward subduction of the African –Arabian plate and

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 437–460. DOI: 10.1144/SP340.19 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Terrane map of Alpine belt showing the correlation of major tectonic units (Modified from http://www-sst. unil.ch/research/plate_tecto/present_day.htm). (b) Simplified geological map of eastern Turkey showing the main rock units within the Baskil-Divrigi transect.

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related basins of the Neo-Tethys Ocean beneath the Eurasian plate of the eastern Taurides, sensu stricto, formed the SEAOB (Aktas & Robertson 1984; Yılmaz 1993; Yılmaz et al. 1993; Beyarslan & Bingol 2000; Sengo¨r et al. 2003). Subduction initiated during the Late Cretaceous (c. 90 Ma) and culminated at the Middle Miocene (Parlak 2005; Robertson et al. 2005). Whereas the overall complicated tectonic sequence is reasonably well understood, the magmatic evolution and how these events fit into the tectonic evolution are open to conflicting interpretation. Placing constraints on the magmatic complexes and how they evolve with time and space is the focus of this chapter. Magmatic processes in rapidly shifting convergent tectonic environments are complicated. In general, the formation of magmas having a subduction-related geochemical signature result from: (1) subduction of oceanic lithosphere (Hawkesworth et al. 1993; Pearce & Peate 1995); (2) melting of metasomatically enriched, subcontinental lithosphere with an inherited subduction signature (Turner et al. 1996); or (3) extensive crustal contamination of MORB-like (mid-ocean-ridge basalt) magmas (Turner et al. 1999). Generation of intraplate-type magmas in continental settings is generally related to upwelling sub-lithospheric mantle material. In Turkey, the subduction-related and late orogenic to within plate-type magmatic rocks are commonly related to kilometre-scale continental extensional terranes. Such a linkage on a global basis is thought to reflect lithospheric response to rapid changes in subduction geometry or reorganization of upper mantle (Bird 1979; Turner et al. 1996, 1999; Duggen et al. 2003, 2005). Large-scale reorganization of the upper mantle may be related to: (1) roll-back and detachment of subducted oceanic lithosphere (Innocenti et al. 1982; Keller 1982; Davies & Blanckenburg 1995; Wilson & Bianchini 1999; Wortel & Spakman 2000); (2) detachment or convective thinning of subcontinental lithosphere (Pearce et al. 1990; Platt & England 1993; Turner et al. 1999; Lopez-Ruiz et al. 2002); or (3) delamination (peeling-off ) of subcontinental lithosphere (Bird 1979; Serri et al. 1993; Duggen et al. 2003). The processes responsible for uplift and the generation of volcanic areas with geochemically diverse subduction-related and post-collisional to late orogenic-type igneous rocks are still a matter of debate, and poorly understood. As the SEAOB preserves a temporal and spatial transition from convergent margin arc to late-to-post orogenic magmatic events followed by successive extensional periods, it presents a challenge to better constrain the temporal and spatial associations of diverse magmatic events. We address this challenge through a combination of systematic geochemical

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and geochronological data across a transect between Baskil and Divrigi, which lies normal to the northerly to northwesterly trending magmatic belts within the complex arcute belt lying parallel to the main trend of Bitlis suture zone (Fig. 1c). The data permits the construction of the temporal and spatial framework to place the transition from post-collisional subduction related to late-orogenic magmatism in this eastern Mediterranean region. We summarize herein the magmatic associations based on field criteria, their geochemical compositions supported by the age of magmatism. Our goal is to reconstruct the geological evolution of the Late Cretaceous to Middle Eocene in the Baskil-Divrigi transect, and to identify the temporal and spatial associations of the magmatic events. The transect includes ophiolites originating in the Vardar and Neo-Tethyan Oceans, as well as several calc-alkaline and alkaline magmatic belts including Baskil (Baskil magmatics), Keban (Keban syenite, Elazig), Hasancelebi (Hasancelebi volcanics and syenitoids, Malatya), Divrigi (Murmano pluton, Sivas) Copler-Yakuplu (Copler, Kabatas, Yakuplu pluton, Erzincan) and Bizmisen, Calti (BizmisenCalti plutons, Sivas) (Fig. 1) regions. Although, they are not included in the Baskil-Divrigi transect post-collisional magmatic rocks in Karamadazi (Karamadazi granitoid, Kayseri) and Horozkoy (Horoz granite, Nigde) were also evaluated for complementary to spatial and temporal associations related to Late Cretaceous to Eocene events in Bitlis –Zagros subduction zone. The new geochemical and geochronological data presented herein for igneous rocks from this transect in southeastern Turkey are also integrated into a tectonic model.

Regional tectonic evolution and geodynamic setting The SEAOB, a microcontinent rifted from Gondwanaland during Early Mesozoic time (Sengo¨r & Yılmaz 1981; Robertson & Woodcock 1982), contains the remains of a Neo-Tethyan Ocean basin that was closed and consumed by a northward subduction beneath the eastern Tauride sedimentary platform. Following subduction, the amalgamated SEAOB and Tauride units were accreted to the Arabian margin by mid-Cenozoic time, possibly accompanying transtensional deformation (Robertson 1998; Stampfli 2001). Just how SEAOB was assembled is the subject of ongoing discussion especially for the Late Cretaceous and Early Cenozoic. The main debate is the palaeogeographical position of the Neo-Tethyan Ocean with respect to Eurasian and/or Arabian plates, and the number and evolution of the convergent plate margins that evolved between Arabian and Anatolide –Tauride

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platform, which forms the southeastern margin of Eurasian plate. Some authors (Hall 1976; Aktas & Robertson 1984; Dewey et al. 1986; Yazgan & Chessex 1991; Yılmaz 1993; Yılmaz et al. 1993) argue that the geological events in the SEAOB are rooted in a northerly single ocean that was closed in Late Cretaceous time along a northdipping subduction zone, followed by thick-skinned re-activation of the suture zone during the Eocene to Miocene. This subduction zone is considered to have generated supra-subduction ophiolites (Goksun, Ispendere, Komurhan, Guleman, etc.) to the south and the Baskil magmatic arc rocks on the active continental margin, the Keban platform, to the north. Conversely, Robertson et al. (2005) noted that the Goksun and other ophiolites (Komurhan, Ispendere and Guleman) represent well-developed oceanic crust generated during an inferred mature stage of a suprasubduction zone life cycle, and that the Baskil arc has a large volume of intrusive rocks cutting the Malatya –Keban platform. Based on these constraints, they argue that it is unlikely a single subduction would be sufficient to generate two well-developed tectonomagmatic events in the region. Based on these arguments, others (Robertson 1998, 2000, 2002; Beyarslan & Bingol 2000; Parlak et al. 2004; Robertson et al. 2005) propose a multiphase convergence along two subduction zones within the Neo-Tethys Ocean. In this model, the ophiolite and related units in SE Turkey formed during these subduction and collisional events related to the consumption of the NeoTethyan Ocean along two subduction zones, one a supra-subduction or an intra-oceanic subduction zone and the second a north dipping subduction zone beneath the continental margin of the Malatya –Keban platform to the north. Intra-oceanic subduction on the south accreted the Goksun and other ophiolites (Komurhan, Ispendere and Guleman) to the overriding oceanic plate. Subduction beneath the Malatya –Keban platform generated the Baskil magmatic arc.

Geological framework of southeastern Anatolian orogenic belt in Baskil-Divrigi transect The north- to northwesterly-trending Baskil-Divrigi transect, the focus of our investigation, lies between Baskil (Elazig) and Divrigi (Sivas) (Fig. 1). The transect obliquely crosses aligned general NE-trending belts of magmatic rocks in the eastern Tauride terrane. The transect lies oblique to the young sinistral Yakapinar-Goksun fault zone, and is normal to the Bitlis–Zagros suture zone (Figs 1 & 2). Five Late Cretaceous to midCenozoic tectonomagmatic terranes make up the

southeastern Anatolian orogenic belt (Fig. 2). These are separated from one another either by major north-dipping thrust faults or by NE-trending strike –slip faults. These terranes are: (a) Late Cretaceous high-grade metamorphic rocks along the major shortening structures; (b) Late Cretaceous ophiolitic rocks; (c) Late Cretaceous active margin units (subduction related magmatic and sedimentary rocks); (d) Late Cretaceous to Eocene post-collisional to late orogenic magmatic rocks; and (e) Paleocene –Eocene to Miocene volcanicsedimentary units. High-grade metamorphic rocks referred to as the Bitlis-Poturge massif and the Malatya – Keban, Engizek, and Binboga metamorphic rocks (Yılmaz et al. 1993; Yılmaz 1993) of Late Palaeozoic – Mesozoic age are the oldest rocks in the basement to the Late Cretaceous and younger rocks. The Bitlis-Poturge massif, exposed mainly at the southern part of the transect (Figs 1 & 2), consists mainly of polydeformed gneiss, amphibolite, and mica schist metamorphosed at greenschist to amphibolite facies conditions (Helvaci & Griffin, 1984; Michard et al. 1984, Yazgan & Chessex 1991; Erdem 1994). Amphibole and mica from amphibolites and mica schist within the Bitlis-Poturge massif yielded minimum K –Ar ages ranging from 85 –76 Ma and 74–56 Ma, respectively, which represent cooling of the terrane (Yazgan & Chessex 1991). The Bitlis –Pu¨tu¨rge massif is generally regarded either as the northernmost extension of the metamorphic basement of the Arabian margin to the south (Yazgan 1984; Yazgan & Chessex 1991) or as one or several continental fragments rifted from Gondwana in the Triassic during the opening of the Neo-Tethyan oceanic basin (Hall 1976; Sengo¨r & Yılmaz 1981; Robertson & Dixon 1984; Yılmaz 1993; Yılmaz et al. 1987; Robertson 1998). Early Cenozoic shallow-water carbonate sequences overlie the massif (Yazgan & Chessex 1991). Ophiolitic rocks in Baskil-Divrigi transect compose two main groups based on the interpreted palaeogeography of the eastern Mediterranean region, that is the Vardar (Izmir-Ankara-Erzincan) or Neo-Tethyan Oceans. These two basins are separated by continental fragments and unconformably overlain by platform carbonate rocks. Emplacement of ophiolitic rocks and ophiolitic me´langes along north dipping thrust sheets within Divrigi and Keban region probably indicate that these are derived from the northerly ocean, Vardar or Izmir-Ankara-Erzincan Ocean (Fig. 1). This ophiolite complex was emplaced on top of the Keban metamorphic rocks by the latest Cretaceous (c. 75 Ma) based on the age of intruding granitoids

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Fig. 2. Geological map of Baskil-Divrigi transect (modified from 1/500 000 geological map of Turkey).

(Boztug et al. 2005; this study) (Fig. 2). Late Cretaceous to Cenozoic sedimentary and volcaniclastic rocks bury the ophiolite. Ophiolites mainly in the southern margin of the transect as well as other parts of the SEAOB from Maras to Bitlis are interpreted to have originate as suprasubduction zone (SSZ) ophiolites (Parlak et al. 2004, Robertson et al. 2005) during the closure of the Neo-Tethyan Ocean. Some of these ophiolites have been metamorphosed to amphibolite facies metamorphism and deformed (Yazgan & Chessex 1991), but nonetheless preserve a coherent ophiolite stratigraphy (Beyarslan & Bingol 2000). Other unmetamorphosed ophiolites, the Guleman and Ispendere complexes, preserve a complete ophiolite pseudostratigraphy, from ultramafic rocks underlain tectonically by a metamorphic sole to volcanic rocks. The high-grade metamorphic sole rocks, preserving an inverted metamorphic gradient, are observed in the Dogansehir (Malatya) region in tectonic contact with overlying ophiolitic units.

According to Parlak et al. (2004), the metamorphic sole is likely to be equivalent of the Berit metaophiolite (Perincek & Kozlu 1984; Genc et al. 1993), found farther to the SW in the Goksun– Afsin (Kahramanmaras) region. K –Ar isotopic age determination on the amphibole from the amphibolite facies yielded a cooling age of 90 + 7 Ma (Parlak et al. 2004). The Malatya – Keban metamorphic terrane, outcropping only at the central and northern part of the transect (Fig. 2), is underlain by lowmetamorphic grade rocks consisting of slate, phyllite, quartzite, metapelite, calcareous schist, and marble. They represent the Late Palaeozoic to Mesozoic (Ozgul et al. 1981) carbonate rock dominated platform sedimentary rock succession that formed part of the Tauride Platform located along the northern margin of the Neo-Tethys. These rocks were metamorphosed to greenschist facies beneath ophiolitic nappes and ophiolitic melange derived from the Vardar (Izmir-Ankara-Erzincan)

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Ocean and then exhumed by latest Cretaceous time (Robertson et al. 2005). The fault separating these two terranes is interpreted to be a remnant of the subduction zone. Late Cretaceous –Cenozoic post-collisional granitoids intrude the metamorphic terrane. The active margin, the Baskil arc (Aktas & Robertson 1984; Yazgan & Chessex 1991) or Elazig magmatic suite (Beyarslan & Bingol 2000), consists predominantly of calc-alkaline volcanic and plutonic rocks of Coniacian to Early Campanian age (Yazgan & Chessex 1991; Parlak et al. 2001), that locally intrude the suprasubduction zone ophiolites, the structurally overlying Malatya –Keban metamorphic rocks, Bitlis-Poturge metamorphics and the fault contact between the ophiolitic and metamorphic rocks (Perincek & Kozlu 1984; Robertson et al. 2005). These granitoids thus postdate ophiolite accretion and must result from north-dipping subduction beneath the Malatya –Keban metamorphic rocks and the Tauride platform (Yazgan & Chessex 1991; Beyarslan & Bingol 2000; Parlak et al. 2004; Robertson et al. 2005). Published K – Ar geochronology of the granitoid rocks range from 85.76 + 3.17 to 77.49 +1.91 Ma (Parlak et al. 2004; Robertson et al. 2005), an age range confirmed by 40Ar/39Ar geochronology summarized below. Tonalite, granodiorite, diorite, gabbro and quartz monzonite intruded by younger microdiorite, porphyritic quartz microdiorite, and by aplite and lamprophyre dykes compose the remnants of the Baskil arc. The late- to post-orogenic and within-plate magmatic rocks in the transect consist, in general, of late Cretaceous calc-alkaline to alkaline volcanic and intrusive rocks, along with derivative volcaniclastic rocks, and Early to Middle Eocene late-orogenic calc-alkaline magmatic rocks. These rocks postdate ophiolite obduction and unconformably overlie or intrude the metamorphic massifs mainly central and northern parts of the transect. Syenite, syenodiorite, monzonite, phonolite, granodiorite, diorite, monzodiorite and gabbro compose the varied intrusive complexes (Kuscu et al. 2007). Spatially, the calc-alkaline suites are generally older and outcrop along the northern margin of the transect. The alkaline suites, however, are younger and generally accompanied extension-related basin formation and deposition of volcaniclastic sequences where they occur as sills or dykes and plugs. These are most common in the Hasancelebi-Hekimhan basin and within Keban region. In general, the nature of magmatism shifts from calc-alkaline to alkaline from north to south (Kuscu et al. 2007), a progression of magmatic composition which might be explained as a consequence of gradual change in the geometry of subduction and slab roll-back followed by a possible slab rupture during the

advanced stages of subduction and ophiolite obduction (see below). Volcanic-sedimentary sequences in the BaskilDivrigi transect are included in two distinct units, the Late Cretaceous to Eocene rocks of the Hekimhan Basin and the Middle Eocene Maden group. The Hekimhan basin (Fig. 2) referred to as a piggy-back basin (Gurer & Aldanmaz 2002) formed above the obducted ophiolitic nappe that roots in the Vardar (Izmir-Ankara-Erzincan) Ocean. Extensional basin formation is considered to have been driven by slab roll-back following the collision and obduction of ophiolites from Vardar (Izmir-Ankara-Erzincan) Ocean. Late Cretaceous to Paleocene continental to shallow marine sedimentary rocks overlain by Paleocene to Eocene marine sequences fill the gradually subsiding basin. Late Cretaceous alkaline sill and dyke complexes and pyroclastic rocks are present in the lower parts of the basin. In contrast, the younger Maden group consists of deep-marine sedimentary rocks, volcaniclastic rocks, and olistostromes, possibly representing a short-lived extensional basin floored by the metamorphic basement (Yılmaz et al. 1993; Yigitbas & Yılmaz 1996; Beyarslan & Bingol 2000). Basin formation is generally considered to have formed in either a back-arc basin (Yılmaz et al. 1993) or pull-apart basin (Aktas & Robertson 1984). Regardless of the tectonic scenario, it appears as if the Neo-Tethyan Ocean basin remained open with subduction continuing in southeastern Turkey into Cenozoic times (Kaymakci et al. 2006).

Temporal and spatial evolution of magmatism in the Southeastern Anatolian Orogenic Belt The tectonic setting, age, type and composition of the magmatic rocks vary temporally (Table 1) and spatial along the Baskil-Divrigi transect. Calc-alkaline, active margin-type rocks (Baskil arc) predominate mainly at the south. Igneous rock compositions progressively resemble postcollisional, late-orogenic to within-plate magma (see below) towards the northern part of the transect. The magmatic activity in the BaskilDivrigi transect initiated at the southern border of the SEAOB with the intrusion of large monzodiorite, granodiorite-granite body into Malatya –Keban metamorphic complex and the suprasubductiontype ophiolitic rocks. 40Ar/39Ar geochronology on hornblende and biotite from these intrusive rocks yielded an age from 82.9 + 0.4 to 79.4 + 0.6 Ma (Table 1), similar to published K– Ar ages between about 85 to 77 Ma (Parlak et al. 2004; Robertson et al. 2005). A 3 Ma period of apparent

Table 1. Ar/Ar, K –Ar and U–Pb dating on selected magmatic rocks in the Baskil-Divrigi transect (shown are weighted mean or plateau and concordia ages; uncertainities +2s) Magmatic rock/pluton

Rock type

Ar/Ar ages Hornblende

Biotite 54.3 + 1.7 50.44 + 0.28

Granodio. Q-monzodiorite Gradio-Q-monzonite

47.17 + 0.69

Karamadazi

Granodio-monzodiorite Diorite-monzodiorite enclave

48.74 + 0.67 46.58 + 0.82

Monzodiorite Granite porphyry Granite Q-porphyry Granite Q-porphyry with enclave Q-monzodiorite

81.1 + 1.0 80.15 + 0.51 79.43 + 0.58

Baskil

Keban

Hasancelebi

Divrigi

Q-syenite Granite Monzonite-monzodiorite Monz.-monzodior Granite with biotite phenocrysts

Copler

Microdiorite-diorite porphyry

Kabatas

Sericitized granodiorite weakly seritized microdio. Porph.

Zircon

Biot-hornb.

72.26 + 0.62 82.90 + 0.43 69.9 + 0.5

Syenite felds. porphyry with metamorphic xenoliths Syenite felds. porphyry with F veining Trachyte-trachy andesite Syenite felds. porphyry Syenite porphyry Diabase dyke Syenite porph. intruding ophiolites

Sericite

K-Ar

74.08 + 0.

71.85 + 0.5

76.84 + 0.6 71.85 + 0.5 70.48 + 0.42 74.40 + 0.51

71.3 + 0.5

74.32 + 0.4 69.0 + 0.4 76.6 + 0.6† 77.2 + 1.8† 73.48 + 0.40 73.50 + 0.40 73.40 + 0.39

44.43 + 0.61 48.63 + 0.62 48.51 + 0.34 48‡

Dog˘ans¸ehir C¸altı

Q-diorite

44.7 + 0.9*

Bizmis¸en

Granodiorite, tonalite

48.8 + 0.9* 43.7 + 0.9* 46.3 + 0.4*

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*Data from Onal et al. 2005; † Boztug et al. 2007; ‡ Parlak et al. 2006.

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Horozkoy

K-feldspar

U-Pb

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Table 2. Representative geochemical analyses of the magmatic rocks in the Baskil transect Sample 534664 534667 534673 534674 534677 534688 534693 534698 535304 535312 535316 535317 Horozkoy (wt%) SiO2 TiO2 Al2O3 Fe2OT3 MnO MgO CaO Na2O K2O P2O5

70.58 0.27 15.24 3.59 0.03 0.68 2.25 4.20 3.08 0.10

Total (ppm) Ba Rb Sr Cs Ga Tl Ta Nb Hf Zr Y Th U Cr Ni Co V Cu Pb Zn W Mo Ag La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

66.71 0.35 16.93 5.25 0.04 1.03 3.33 4.24 2.58 0.14

Karamadazi

Baskil

Keban

62.29 0.71 15.14 8.62 0.10 3.66 6.24 4.63 0.75 0.21

63.49 0.62 17.52 4.81 0.05 2.4 5.62 6.1 0.2 0.23

75.83 0.13 12.79 0.90 0.01 0.02 0.48 3.71 4.55 0.01

67.01 0.27 15.8 6.55 0.08 0.95 2.47 3.78 4.80 0.07

73.00 0.28 13.9 4.57 0.06 0.75 2.17 3.86 1.65 0.04

50.13 0.92 19.12 19.2 0.17 3.44 9.94 2.19 0.25 0.06

72.06 0.31 14.15 4.49 0.06 0.76 3.03 3.64 1.84 0.04

63.14 0.28 18.47 2.53 0.04 0.15 2.20 5.08 5.32 0.06

65.44 0.34 17.25 3.57 0.03 0.48 1.89 4.49 4.35 0.09

65.44 0.37 17.00 3.25 0.03 0.28 2.25 4.57 4.32 0.12

99.54 100.05

100

100

98.46

99.96

100.05

99.35

99.41

99.39

100

99.92

558 592 91.5 62 450 605 1.28 0.78 16.1 17.6 0.5 0.5 1.2 1.1 13.4 14.7 4 5.1 124.5 175.5 14.9 18 26.3 14.8 5.58 2.49 20 20 9 12 4.1 5.6 35 52 5 11 5 5 29 33 6 11 2 4 1 1 52.7 33.8 86.5 59.6 8.02 6.19 24.9 21.7 3.14 3.46 0.75 0.98 3.37 3.52 0.39 0.48 1.7 2.47 0.4 0.56 1.48 1.82 0.22 0.28 1.74 2.12 0.3 0.35

238 16.2 825 0.34 15.4 0.5 1.1 17.8 4.5 161 23.5 9.5 2.97 80 45 15.7 113 14 5 46 3 2 1 27.2 53.1 6.13 23 4.39 1.18 4.5 0.64 3.78 0.78 2.43 0.33 2.27 0.37

164 3.6 1325 0.1 16.6 0.5 0.9 16.6 4.1 148.5 19.8 9.47 3.09 40 17 8.7 88 7 5 57 5 2 1 32.6 60.1 6.55 24.1 4.05 1.13 4.26 0.56 3.04 0.65 2.05 0.28 1.88 0.3

28.6 232 20.2 4.52 15.2 0.5 3.4 27.1 4.5 91.9 10.2 33.9 6.7 20 10 1 9 6 11 18 4 2 1 19.4 30.3 2.44 6.5 0.73 0.16 1.1 0.14 0.76 0.23 0.9 0.17 1.54 0.31

438 204 187 4.45 15.8 0.5 1.5 16.6 5.7 203 24.7 26.7 4.34 30 16 7.7 42 14 52 54 5 2 1 26.4 46.2 4.81 16 2.9 0.56 3.24 0.53 3.07 0.78 2.44 0.4 2.89 0.47

227 39.6 102.5 0.32 12.2 0.5 0.1 1.7 3.2 96.9 17.1 1.88 0.73 20 11 5.6 62 8 19 35 1 2 1 6 11.8 1.44 6.1 1.62 0.42 1.84 0.37 2.18 0.51 1.68 0.26 2.08 0.4

52.4 4.2 188 0.07 19 0.5 0.1 0.8 1.9 54.1 24.3 0.35 0.2 30 21 28.7 380 108 5 94 1 2 1 2.7 5.8 0.91 5.1 1.87 0.79 2.56 0.53 3.73 0.82 2.45 0.34 2.36 0.4

231 41.9 116 0.79 13.2 0.5 0.1 1.5 2.8 74.9 17.1 1.9 0.89 30 9 5 48 7 5 43 3 2 1 6.3 12.4 1.48 6.1 1.46 0.49 1.9 0.37 2.27 0.53 1.72 0.24 2.08 0.34

magmatic quiescence (Table 1) appears to have been followed by the resumption of post- to lateorogenic magmatism (Table 1) producing smallto medium-sized isolated intrusions into the

2990 2480 2360 180.5 152 164.5 1680 1375 1495 3.66 2.83 4.37 23.8 23.2 21.9 0.7 ,0.5 0.5 1.9 2.7 2.6 52.9 51.7 49 11 11.3 11.2 337 398 335 16.1 22.3 20.8 63.7 82.4 77.8 17.85 17.25 14.7 20 20 20 10 12 14 2.4 4.9 4 52 72 59 6 5 17 27 11 21 49 35 44 5 8 10 21 3 3 1 ,1 1 61.5 83.4 82 124 160 151 13.8 16.95 16 47.1 57.9 54.5 6.92 8.67 8.38 1.3 1.72 1.66 5.8 7.31 7.19 0.56 0.77 0.8 2.34 3.71 3.31 0.46 0.66 0.62 1.36 2.06 2 0.14 0.29 0.24 1.04 2.1 1.8 0.17 0.3 0.29

Malatya –Keban metamorphic complex and structurally overlying ophiolitic rocks and melange in Divrigi and Hasancelebi regions. These magmatic rocks post-date the ophiolite obduction, and are

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT

Sample 535318 h.celebi 535321 535338 535340 535452 535454 535457 Hasancelebi

Divrigi

C1

445

C6

Copler

K3 Kabatas

(wt%) SiO2 TiO2 Al2O3 Fe2OT3 MnO MgO CaO Na2O K2O P2O5

58.10 1.32 16.24 5.61 0.03 2.33 4.98 3.40 6.94 0.36

70.78 0.91 16.50 3.33 0.05 1.24 4.62 4.92 5.29

60.55 1.08 15.96 2.42 0.04 1.39 4.75 2.69 8.59 0.26

58.96 1.08 17.98 7.9 0.04 2.15 4.29 4.64 4.91 0.32

58.47 1.19 18.08 8.84 0.04 2.20 4.43 4.16 4.76 0.35

64.09 0.63 17.33 3.93 0.02 1.30 2.84 3.88 5.61 0.21

65.47 0.52 16.76 5.36 0.02 1.30 2.67 3.76 5.02 0.19

58.91 1.05 17.65 7.19 0.04 2.65 4.63 4.13 4.59 0.36

60.05 0.62 17.39 4.97 0.09 1.91 8.00 2.16 3.29 0.26

61.91 0.65 16.76 6.38 0.04 1.76 4.15 2.54 2.26

61.93 0.59 17.41 7.82 0.02 2.57 2.93 4.01 1.35 0.15

Total

99.95

99.73

99.52

100.05

99.86

99.33

99.81

99.62

99.94

99.52

98.66

2940 1892.33 2500 99.8 69 102.5 326 252 257 2.34 0.88 0.19 21.1 23 23.1 0.5 0.5 0.5 2.5 2.7 2.6 40.8 42 40.2 8.9 8.33 9.6 341 348 361 44.3 40 38.4 20.4 27.33 23.1 13.05 11.3 9.41 30 20 20 38 21 18 9.6 5 4.6 113 62 54 21 13 12 14 8 6 45 30 29 8 5 1 20 9.33 6 1 1 1 70.2 53.7 73.5 145.5 107.4 139.5 15.55 11.46 13.85 52.5 39.53 45.8 9.16 7.28 7.92 2.64 2.08 2.38 9.03 7.34 8.04 1.3 1.07 1.08 7.2 6.27 6.13 1.5 1.34 1.29 4.54 4.11 3.9 0.62 0.58 0.59 4.38 3.95 3.97 0.64 0.61 0.63

1215 153 352 1.58 20.3 0.5 2.4 37 5.5 224 29.1 22.8 7.77 30 12 9 98 7 13 44 2 2 1 40 77.2 8.75 33.8 6.45 1.36 6.24 0.88 5.23 1.03 3.1 0.41 2.91 0.46

1330 162 389 1.98 20.6 0.5 2.1 31.9 4.6 188.5 29.6 23.5 6.04 30 13 9.8 108 10 30 72 4 2 1 44.9 93.9 10.95 39.7 6.9 1.45 6.58 0.92 4.98 0.98 2.93 0.42 2.94 0.4

975 160.5 311 1.35 19.6 0.5 2.7 29.1 4.8 186.5 26 30.5 7.5 30 15 5.7 67 5 10 32 3 2 1 43.6 87.6 9.87 34.9 6.24 1.22 6.06 0.83 4.48 0.88 2.69 0.38 2.65 0.4

914 161.5 302 1.38 19.4 0.5 2.5 26.8 6.4 257 23 28.4 8.29 30 15 6.9 57 5 16 41 5 2 1 36.7 69.8 7.6 26.5 4.76 0.92 4.38 0.69 3.54 0.73 2.27 0.32 2.27 0.37

(ppm) Ba Rb Sr Cs Ga Tl Ta Nb Hf Zr Y Th U Cr Ni Co V Cu Pb Zn W Mo Ag La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

thus post-collisional intrusions. Although there is a temporal overlap in products of post-collisional magmatism, 40Ar/39Ar geochronology (Table 1) suggests that a Campanian (74.40 –73.40 Ma)

1080 865 136 74.4 380 715 1.46 3.42 20.3 19 0.5 0.5 2 0.6 30.9 8.6 7.2 3.4 303 127 30.5 20.8 21.2 7.49 5.86 1.76 50 20 41 10 15.8 7.9 106 168 26 7 11 5 39 42 1 3 2 2 1 47 45.8 23.3 95.6 44.9 11.1 5.2 39.7 21.2 7.36 4.24 1.47 1.26 7 3.99 0.94 0.57 5.08 3.46 1.06 0.68 3.02 2.06 0.44 0.3 3.04 2.1 0.44 0.31

679.33 402 49 32.5 470 300 1.6 0.41 18 17.3 0.5 0.5 0.63 0.7 9.8 10.5 3.33 2.8 124 100.5 16 19.5 8.22 8.98 1.74 1.72 20 20 11 12 9 18.4 161 134 971 2420 8 7 94 207 2.67 3 10.67 26 16.3 1 24 21.8 43.37 36.2 5.09 4.43 19.2 16.6 3.51 3.52 0.98 1.01 3.36 3.74 0.46 0.58 2.7 3.6 0.52 0.72 1.57 2.02 0.23 0.3 1.6 2.05 0.24 0.31

calc-alkaline magmatic phase was followed by a shift to alkaline magmatism during the Maastrichtian (71–69 Ma, U –Pb zircon ages) (Table 1). Slightly younger calc-alkaline magmatic rocks

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with 40Ar/39Ar ages of 74.40 + 0.51 to 73.40 + 0.39 Ma (Table 1) also intruded the Malatya –Keban metamorphic complex, Tauride–Anatolian Platform, and ophiolitic rocks. Alkaline intrusions into the Malatya –Keban metamorphic complex and northerly derived ophiolitic rocks in the Hasancelebi and Keban regions are dated as 71.0 + 1.0 to 69.9 + 0.5 Ma by U –Pb geochronology on zircon. The same rocks have been dated as 74.1 to 71.8 + 0.5 Ma based on 40Ar/39Ar dating on biotite (Table 1). Still younger felsic to intermediate calc-alkaline shallow intrusions present mainly along left-lateral strike–slip faults (CAFZ and YGFZ and MOFZ) associated with sub-basins bounded by these faults or their splays at Copler, Kabatas, Bizmisen-Calti, Karamadazi, Horoz regions yielded 40Ar/39Ar ages at 54.3 + 1.7 to 44.43 + 0.61 Ma (Table 1). These rocks crop out as isolated plutons that postdate collision of the Pontides with Tauride– Anatolian Platform and obduction of northerly ophiolites from the Vardar Ocean. The available geochronology from this study and other studies by K –Ar geochronology (Onal et al. 2005; Parlak et al. 2006; Table 1) also show Eocene magmatism is essentially continuous from Late Paleocene to Eocene (early Ypresian to late Lutetian), and contemporaneous with basin evolution and sedimentation. The Eocene magmatism in the SEAOB is not confined to Copler, Kabatas, Horoz and Karamadazi granitoids, but is generally widespread, although volumetrically less than the older igneous complexes such as Dedeyazi-Polat (Dogansehir) and Bizmisen-Calti granitoids. The K –Ar geochronology on magmatic hornblende/biotite yielded an age of c. 48 Ma (Parlak et al. 2006), and 43.7–46.3 Ma (Onal et al. 2005), respectively. Onal et al. (2005) dated the Eocene postcollisional granitoids in Bizmisen and Calti (Divrigi, Sivas) plutons as 42–56 Ma to 40– 49 Ma. Similar Eocene magmatic rocks are also reported in Iran along the eastern continuation of the Bitlis –Zagros subduction zone where UrumiehDokhtar magmatic belt formed as a consequence of slab break off during the Middle Eocene (Bird 1978; Ghasemi & Talbot 2005). Yigitbas & Yılmaz (1996) and Robertson et al (2005) also described Eocene magmatism in the Maden Complex and Helete volcanics, and postulated an arc and back-arc setting for their formation, respectively. Eocene magmatism discussed herein is located north of the arc-related rocks of Late Cretaceous age, and is unrelated to the Maden Complex and Helete volcanics. In summary, the 40Ar/39Ar geochronology (Table 1) along with geochemical characteristics of the granitoids (see below) indicate a geographic shift in the magmatic loci was coupled with a

transition from arc-type calc-alkaline to postcollisional calc-alkaline to alkaline then to withinplate alkaline magmatism with time from 82.9 + 0.4 to 79.9 + 0.6 Ma, 74.40 + 0.51 to 73.40 + 0.39 Ma to 71.0 + 1.0 to 69.9 +0.5 Ma from south to north, respectively. These igneous complexes are subsequently overprinted by postcollisional calc-alkaline magmatism at 54.4 – 44.3 Ma (Table 1) mainly along a NE-trending belt located in the central part of the transect. Available geochronologic data suggest each individual magmatic period has a relatively short lifetime of c. 2– 5 Ma.

Geochemical characteristics of the magmatic rocks in Baskil-Divrigi transect Magmatic rocks in Baskil-Divrigi transect are subalkaline and alkaline but have significant scatter due to wide range of compositional diversity (Table 2; Fig. 3). Baskil magmatic rocks have both a felsic and a mafic end-member. The mafic and felsic end-members appear to follow a tholeitic and calc-alkaline trend, respectively (Fig. 3b). In general, the Horoz, Karamadazi, Copler and Kabatas (Yakuplu) granitoids have subalkaline characteristics with more calcic affinities (Fig. 3c) whereas Keban and Hasancelebi volcanic-plutonic rocks have alkaline characteristics with a calcic to calc-alkaline affinity. Variable A/CNK [molar Al2O3/(CaO þ Na2O þ K2O)] and A/NK [molar Al2O3/(Na2O þ K2O)] characterize the complexes. The magmatic rocks in Baskil and Copler-Kabatas are peraluminous whereas in Hasancelebi, Divrigi, Karamadazi and Horoz regions, these rocks are metaluminous (Fig. 3d). In general, the granitoids in Baskil, Copler and Kabatas are medium-K suites whereas the Hasancelebi and DivrigiMurmano plutons are high-K to shoshonitic (Fig. 3e). The Divrigi-Murmano pluton has compositions that suggestion both alkaline and calc-alkaline; field relations and petrographical observations indicate the extreme compositional heterogeneity in these is due to alkali metasomatism (scapolite-albite alteration overprinted by Kfeldspar and phlogopite) of a 20  5 km2 region that is associated with a large (.108 tonnes) IOCG deposit (Kuscu et al. 2002, 2007). The magmatic rocks range from basalt-gabbro to syenite and trachyte with a common overlap between the compositional fields of Pearce (1996) (Fig. 4a) and Winchester & Floyd’s (1977) diagrams (Fig. 4b). In the Baskil region, basalt, diorite, gabbro-diorite and granite to quartz diorite are present. Diabasic dykes at Baskil have distinct compositions characterized by the lowest Zr/TiO2 and Nb/Y ratios recorded in the transect, and plot in

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT FeOT

(b)

(a)

447

18

Na2O + K2O

14 10

Tholeiitic

Alkaline

6 Subalkaline

Calc-Alkaline

2

40

50

60 SiO2

70

80

(c) 15

(d) Alkalic

A-C C-A

2.8

Calcic

Metaluminous

Peraluminous

2.4 2.0

9 A / NK

Na2O + K2O

12

6

1.6 1.2

3 0 40

0.8 Peralkaline 50

60 SiO2

70

0.4 0.5

80

1.0

A/CNK

1.5

2.0

(f) 1.2

(e) 10 9 8 7 6 5 4 3 2 1 0 50

Shoshonitic

Log(Fe2O3/FeO)

K2O

MgO

Na2O + K2O

High-K Medium-K Low-K

60

70

80

SiO2 Kabatas Copler Horozkoy Karamadazi

Post-collisional, calc-alkaline (Early–Middle Eocene, 50.4–44.3 Ma)

1.1 1.0 Very strongly oxidized 0.8 0.7 0.6 0.5 0.4 Ox=0.8 0.2 Strongly oxidized 0.1 0.0 0 1 2 3 4 5 6 7 8 9 10 FeOT Baskil

Active margin (Late Cretaceous, 82.9–75 Ma)

Divrigi

Post-collisional, calc-alkaline, (Late Cretaceous, 75–74 Ma)

Keban Hasancelebi

Post-collisional-late orogenic, alkaline (Late Cretaceous, 71–69 Ma)

Fig. 3. Major element plots of the magmatic rocks. (a) Total alkalis v. silica (TAS); (b) AFM plots of Irvine & Baragar (1971); (c) Peacock (1931) index; (d) the Shand’s Index of Maniar & Piccoli (1989); (e) Gill (1981); and (f) oxidation state diagrams of (Blevin & Chappell 1992).

basalt field (Fig. 4a). The Baskil rocks are furthermore characterized by higher SiO2 composition than magmatic rocks in Keban, Hasancelebi and Divrigi regions (Fig. 4b), perhaps reflecting higher degrees of fractionation. Although, the Baskil rocks have compositions similar to other magmatic rocks within the transect, they form a distinct group

with low- to medium-K and strongly oxidized compositions (Fig. 3f). In contrast to the more normal calc-alkaline rocks at Baskil, the magmatic rocks near Keban are mainly syenite and syenodiorite in composition. They have the highest Zr/TiO2 and Nb/Y ratios in the transect (Fig. 4a). At Hasancelebi, trachyte,

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(a)

I. KUSCU ET AL.

(b) 80

2 1

Alkali Rhyolite

Rhyolite

Rhyolite+Dacite

0.1

Com/Pan

70 Rhyodacite-Dacite

Trachyte Tephri-Phonolite TrachyAnd

SiO2

Zr/TiO2*0.0001

Phonolite

Trachyte

60

TrAn

Andesite

Phonolite

And/Bas-And

0.01

Foidite Basalt

0.002 0.01

0.1

Baskil Divrigi Keban Hasancelebi

Sub-Ab

Alk-Bas

1 10 Nb/Y Active margin (Late Cretaceous, 82.9–75 Ma) Post-collisional, calc-alkaline, (Late Cretaceous, 75–74 Ma)

Ab

50

40 0.001

100

Post-collisional-late orogenic, alkaline (Late Cretaceous, 71–69 Ma)

Kabatas Copler Horozkoy Karamadazi

Bas-Trach-Neph

0.01

0.1 Zr/TiO2*0.0001

1

10

Post-collisional, calc-alkaline (Early–Middle Eocene, 50.4–44.3 Ma)

Fig. 4. (a) Modified Zr/TiO2-Nb/Y plot (Pearce 1996) of Winchester & Floyd (1977), and (b) Zr/TiO2-SiO2 plot (Winchester & Floyd 1977).

trachyandesite and syenite to syenodiorite form volcano-plutonic complexes intruding the Late Cretaceous to Paleocene Hekimhan sedimentary basin. The syenites occur as plugs or stocks intruding the ophiolitic and trachytic rocks within the basin. Magmatic rocks in Copler-Kabatas plutons and Divrigi-Murmano share similar compositional ranges, from gabbro, diorite, monzodiorite to granodiorite in composition. Similar compositional ranges

Magmatic rocks in the Baskil transect contain moderate to low HFSE (high field strength elements such

Sample/Ocean Ridge Granite

Sample/Primitive Mantle

Trace element-REE geochemistry

(b) 200 100

(a) 3000 1000

100 10 1 0.1 0.01 0.001

are present in the Karamadazi and Horoz regions (Fig. 4a). However, those in the Horoz region tend to contain higher Zr/TiO2 ratios and SiO2 compared to Karamadazi, Copler and Kabatas regions.

10

1 0.2

Cs Rb U K Nb Sr Zr Tb Cr Ni Ta Ba Th Pb Li Ce Ti Y Zn Baskil Divrigi Keban Hasancelebi

Active margin (Late Cretaceous, 82.9–75 Ma) Post-collisional, calc-alkaline, (Late Cretaceous, 75–74 Ma)

Post-collisional-late orogenic, alkaline (Late Cretaceous, 71–69 Ma)

Cs Rb U K Nb Sr Zr Tb Cr Ni Ta Ba Th Pb Li Ce Ti Y Zn

Copler Horozkoy Karamadazi

Post-collisional, calc-alkaline (Early–Middle Eocene, 50.4–44.3 Ma)

Fig. 5. (a) Rock/primitive mantle-normalized, and (b) rock/ocean ridge granite-normalized (Pearce et al. 1984) spidergrams.

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT

Fig. 6. Rock/chondrite normalized average REE plot of the magmatic rocks from the central-eastern sector.

as Nb, Ta, Ga, Zr, Hf, Y) indicating an arc to postcollisional/late-orogenic setting for their generation (Figs 5a, b). Primitive mantle-normalized extended trace element spiderplots (Fig. 5a) are characterized by large ion lithophile element (LILE) enrichment with respect to HFSE, and strong negative K and Cr-Ni anomalies. The LILE enrichment in the magmatic suites other than Baskil arc is more pronounced. Ocean ridge granite-normalized plots, used to compare arc-related rocks with rocks in the eastern-central transect (Fig. 5b), illustrates a moderate negative Y, Zr and Nb (HFSE) anomalies, and positive Rb and Th anomalies in all rocks. In general, the Copler and Kabatas plutons, Karamadazi and Horoz granitoids are characterized by depleted Nb and Zr contents (Fig. 5b). In contrast, the Divrigi-Murmano pluton and Hasancelebi volcano-plutonic rocks are characterized by moderate Nb and Y contents. Syenitoids in the Keban region have the highest Nb and Zr contents measured in the transect (Fig. 5a, b). The negative Nb anomalies in the Baskil transect rocks are common features of rocks associated with arc lavas and incipient back-arc basin basalts due to the influx of low field strength element (LFSE) into the mantle source regions from the subducted slab (Wilson 1989). As is commonly accepted, arc magmatism derives mainly from the melting of subducted oceanic crust, overlying mantle wedge and lower continental crust. Such magma will essentially reflect the inherited geochemical signatures of these source regions,

449

as these source materials can also be melted in a post-collisional environment. Other trace element characteristics of subduction-related magmas include enrichment in Ba, Th and Pb, and depletion in Nb in addition to Ti and Ta (Fitton et al. 1988; Saunders et al. 1980, 1988). Of them, Nb and Ta are also highly sensitive to crustal contamination. Mantle-derived magmas, which might be contaminated by continental crustal rocks during their ascent to the surface, have marked Nb and Ta negative anomalies (Wilson 1989). Therefore, the negative Nb anomaly on these plots need not reflect an arc environment. Nonetheless, the trace element geochemistry does indicate a subducted slab component is present in these rocks. The rocks are generally enriched in light rare earth elements (LREE) relative to chondrites (Nakamura 1974), but to differing extents (Fig. 6). All samples are characterized by significant flattening of heavy rare earth element (REE) patterns relative to the LREE, and a marked depletion in Eu. La/Sm in Baskil magmatic rocks is less compared to other igneous suite, suggesting a distinct environment for the generation of these rocks. Likewise, the Eocene Horoz, Copler, Kabatas and Karamadazi display almost the same REE patterns implying a common mechanism or source for their generation. Rocks in Hasancelebi, Divrigi and Keban have similar LREE patterns but distinct HREE (heavy rare earth element) patterns with a flatter HREE pattern for Hasancelebi. In addition, REE also show three distinctive groups, principally based on LREE compositions (Fig. 6); these groups are almost the same as those defined by trace elements (see above), and differ from each other by their relative LREE enrichment with respect to HREE and to chondrite (Fig. 6). The Late Cretaceous rocks from the Baskil region define a group with the lowest rock/ chondrite ration, and are characterized by a slight enrichment in terms of LREE. The Eocene magmatic rocks in Copler, Kabatas, Karamadazi and Horoz regions constitute the second group, and have a moderate LREE enrichment. The third group has higher LREE enrichment, and includes the latest Late Cretaceous magmatic rocks from Keban, Hasancelebi and Divrigi regions. Combining with the distinct ages for the magmatic rocks, these chemical groups provide further evidence that different magma generation events were active at different times, and these also likely reflect different tectonic environments (see below). Ta/Yb, Th/Yb, Th/Y, Nb/Y, and Nb/Zr ratios and their comparison on suitable binary plots may be useful in assessing the role of source heterogeneities, metasomatism and crustal contamination (Pearce 1983) (Fig. 7a, b) in the generation of magmas. All magmatic rocks within the transect have trends in Th/Yb v. Ta/Yb that parallel the

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(a) 100

Th/Yb

(b) 10

ica

m

s ati ome) s ta t SZ me pla le intra t an ( M

1

Subduction zone enrichment

SZ

1

Within plate trend

Th/Y

rea inc crust

10

il gs sin

OIB

0.1 SZ

0.1 average N-MORB

0.01

0.1

Andean rocks

W

SZ: subduction zone W: within-plate 1 10 Ta/Yb

0.01 0.1

Nb/Y

1

10

(c) 500 Within plate (Zr/Nb < 10) Post-collisional (Zr/Nb > 10)

100

10 CA 2 20

Low K 100

Divrigi Keban Hasancelebi Kabatas Copler Horozkoy Karamadazi

Zr = 500 ppm Peralkaline

High K

Nb

Baskil

10 /N 20 Zr b= N / Zr b=

Zr

1000

Active margin (Late Cretaceous, 82.9–75 Ma) Post-collisional, calc-alkaline, (Late Cretaceous, 75–74 Ma) Post-collisional-late orogenic, alkaline (Late Cretaceous, 71–69 Ma) Post-collisional, calc-alkaline (Early–Middle Eocene, 50.4–44.3 Ma)

4000

Fig. 7. (a) Th/Yb v. Nb/Y diagram showing the rocks from Andean subduction zone, and from the eastern-central transect; (b) Th/Yb v. Ta/Yb diagram (Pearce 1983) showing the effect of mantle metasomatism (crust average for central Anatolia were taken from Ilbeyli 2005); and (c) Nb-Zr plot of Leat et al. (1986) showing within-plate component.

mantle metasomatism array, but are displaced towards higher Th/Yb ratios. In addition, both ratios increase from the calc-alkaline through alkaline intrusive rocks. The alkaline syenitic rocks in Keban region contain higher Ta/Yb and Th/Yb ratios compared to Hasancelebi region, and to the calc-alkaline plutonic rocks in Copler, Kabatas, Karamadazi, Horoz, and Divrigi regions. The magmatic rocks in Baskil plot at low Th/Yb and Ta/Yb ratios with lowest Ta/Yb ratios that may be indicative of subduction-related setting (Pearce 1983). The remaining magmatic rocks, display moderate Th/Yb and Ta/Yb ratios suggesting a derivation from a metasomatized mantle inherited from an earlier subduction, likely reflecting a postcollisional or late-orogenic setting. The Th/Y v. Nb/Y diagram (Fig. 7b) distinguishes within-plate source from subductionrelated source regions. Mantle that has been enriched by subduction components and/or crustal components shift to higher Th/Y ratios with respect to Nb/Y, whereas mantle previously enriched by a small volume of partial melt is displaced along the melt enrichment and within-plate trend in the direction of high Th/Y and Nb/Y

ratios (Pearce 1983) (Fig. 7b). Most felsic magmatic rocks at the Baskil region plot in mantle metasomatism array with increasing Th/Y and Nb/Y ratios suggesting melt enrichment and within plate signature most probably derived from a lithospheric mantle source carrying an older subduction component. Of particular importance, the Horoz, Karamadazi and Copler intrusions have a weak subduction zone enrichment array (Fig. 7a, b), that we infer to be inherited from an older subduction event, or melting of a source that could be of an older subducted slab, or mantle wedge contaminated by a subduction component. A similar behaviour is observed in an Nb v. Zr diagram (Fig. 7c) which has been proposed to characterize the geochemical and tectonic setting characterization of felsic rocks (Leat et al. 1986). Depletion in Zr and Nb concentrations coupled with increasing Zr/Nb ratios is typical of arc-related rocks. Likewise, a gradual increase in Nb and corresponding decrease in Zr/ Nb ratio cause shift toward post-collisional setting to within plate settings (Fig. 7c). The Baskil arc-type rocks plot in high Zr/Nb with ratios. Copler, Kabatas and Horoz granitoids plot in moderate Nb and Zr concentrations, which indicate

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT

this magmatism is related to a source modified by subduction. This magmatic association has been described in provinces such as Snowdonia, Avoca and Parys Mountain, within the Southern British Caledonides (Leat et al. 1986). The magmatic rocks in Divrigi, Hasancelebi, Keban and to some extend Karamadazi regions have Zr/Nb ratios ,10 with highest Nb and Zr concentrations which collectively suggest an anorogenic setting. However, as discussed previously, the alkalinity and consequent geochemical signature in DivrigiMurmano pluton is due to a pervasive hydrothermal alteration (alkali metasomatism) that affected the entire pluton. Alkaline syenitoids in Keban always plot in the highest Nb and Zr corner of the diagrams suggesting these were derived from within plate setting modified by a subduction zone component. Magmatic suites in Baskil-Divrigi transect are clustered in three overlapping groups based on compositions and ratios of Nb, Zr, Th, Y, Ta and Yb in Figure 7a–c. These may represent grouping of granitoids formed during different tectonic settings. The high Th/Yb-Ta/Yb and Th/Y-Nb/Y and high Zr/Nb ratios (group 3) in Keban region are typical for within plate settings whereas, low Th/Yb-Ta/Yb, Th/Y-Nb/Y and high Zr/Nb (group 1) ratios are typical for low-K arc-related granitoids exposed mainly in the Baskil region. The other magmatic rocks from the Karamadazi, Horoz, Divrigi, Hasancelebi, Copler and Kabatas regions are clustered at moderate Th/Yb-Ta/Yb, Th/Y-Nb/Y and Zr/Nb ratios (group 2) in between the within-plate and the arc-related settings. Some of the samples from Baskil have Th/Y and Nb/Y that is similar to rock defining an Andean rock field (Pearce 1983).

Petrogenesis and tectonomagmatic classification Possible tectonic environments for the Baskil magmatic rocks can be generally inferred based on trace and REE characteristics using classical tectonomagmatic discrimination diagrams (Fig. 8e, f ), coupled with other geological evidence. Although not completely conclusive in discriminating the complex tectono-magmatic settings, trace element and REE characteristics coupled with multi-element variation diagrams suggest three main petrotectonic settings for the generation of the magmatic rocks in Baskil-Divrigi transect. These are: (1) subduction and arc setting and (2) post-collisional and extensional setting (exhumation, probably due to roll-back) followed by (3) late-orogenic extension, probably due to crustal scale strike –slip faulting or subduction transfer edge propagator (STEP)faulting (as defined in Wortel & Spakman 2000). The magmatic rocks from the Baskil, Copler,

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Kabatas, Karamadazi and Horoz all plot mainly within the VAG and pre-plate collision fields with a marked input from the mantle by their narrow Y and Y þ Nb, and low Rb and Ta compositions (Fig. 8a–f; Pearce et al. 1984; Pearce 1996; Batchelor & Bowden 1985). Magmatic rocks in the Baskil region are of arc-type (Hall 1976; Aktas & Robertson 1984; Dewey et al. 1986; Yazgan & Chessex 1991; Yılmaz 1993; Yılmaz et al. 1993; this study) that post-date the collision between Malatya – Keban metamorphic rocks and ensimatic arc/ophiolite (Beyarslan & Bingol 2000; Parlak et al. 2004; Parlak 2005; Robertson et al. 2005). Northward subduction of SSZ-type oceanic crust during the evolution of the southern branch of the Neo-Tethyan Ocean generated the arc (Hall 1976; Aktas & Robertson 1984; Dewey et al. 1986; Yazgan & Chessex 1991; Yılmaz 1993; Yılmaz et al. 1993; Beyarslan & Bingol 2000, Parlak et al. 2004; Robertson et al. 2005). Subduction of the northern margin of Arabian platform beneath the southern passive margin of Tauride–Anatolide platform during latest Cretaceous time marked the cessation of arc magmatism. Late Cretaceous magmatic rocks in Divrigi, Hasancelebi, and Keban have compositions suggestive of a post-collisional setting with a common overlap to within plate setting. Commonly, lateorogenic extensional regimes generated during collision and collision-induced roll-back of the subducting slab will produce post-collisional granitoids of diverse type and origin; these are the postCOLG and late-orogenic granitoids on the Pearce (1996) and Batchelor & Bowden (1985) diagrams, respectively (Fig. 8b– d). In eastern Turkey, these Late Cretaceous granitoids intrude the Malatya – Keban metamorphic rocks and structurally overlying ophiolitic nappes, and are post-collisional. Hinge retreat or roll-back, which likely followed collision, appears to be active when continental lithosphere of the Arabian plate subducted beneath the southern margin of the Anatolide– Tauride platform (continental collision). The syenitic rocks in Keban and Hasancelebi region show geochemical features characteristic to within-plate setting (Fig. 8a–f ). The ternary Nb-Y-Ga diagram (Fig. 8e) of Eby (1992) suggests these magmas were derived from partial melting of mantle material. As melting conditions in mantle require generally excess water and rapid decompression, an influx of fluids from subducted slab and rapid removal of pressure by slab roll-back mechanisms appear to be required. Implicit in this model is the hypothesis that generation of alkaline rocks accompanied extensional tectonism related to advanced stages of slab roll-back. An arc origin can be inferred for the Eocene magmatic rocks in Copler, Kabatas, Karamadazi

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Fig. 8. (a) Nb-Y; (b) Rb 2 Nb þ Y; (c) Ta-Yb diagrams of Pearce et al. (1984), Pearce (1996); (d) R1-R2 diagram of Batchelor & Bowden (1985); (e) Nb-Y-Ga diagram of Eby (1992); and (f ) Ta v. Nb diagram of Harris et al. (1986).

and Horoz regions, based on the tectonomagmatic discrimination diagrams (Fig. 8a– d). Several lines of geological evidence, however, suggest that the arc setting may not be correct for any rocks except for those the Late Cretaceous Baskil granitoids. All of the Eocene magmatic rocks intruded obducted ophiolitic nappes. The relationship

suggests that, at a minimum, their emplacement post-dates collision and obduction of the ophiolitic rocks. Furthermore, none of the granitoids have contemporaneous fore-arc sequences or accretionary prism, and no significant high-pressure metamorphism rooted to a subduction event can related in time and space, unlike older arc terranes. Therefore,

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT

concluding an arc origin for the Eocene magmatic rocks based on the geochemical discrimination diagrams appears inconsistent with the available geological evidence. Instead, an arc signature might have been inherited from older subduction events that producing Baskil arc magmatic rocks, or could be attributed to melting of a mantle wedge contaminated by older subduction. If the Eocene rocks are related to subduction, then the available geochronologic constraints requires two subduction events, one in the Late Cretaceous, and the younger one in early to middle Eocene times. Late Cretaceous subduction is well established in the SEAOB. In contrast, the case for subduction in Eocene is subject to conflicting interpretations (Sengo¨r & Yılmaz 1981; Yigitbas & Yılmaz 1996). The Eocene subduction zone proposed by Yigitbas & Yılmaz (1996) would have been located south of Maden complex, which has been interpreted to be a remnant of a back-arc basin filled with volcanic rocks and lacking significant intrusions. Robertson et al. (2005) have also proposed that the Eocene Helete volcanics, deposited within an extensional basin above the metamorphic rocks, record subduction but suggest a back-arc setting to be more likely. Richards (2003) suggested that the UrumiehDokhtar magmatic belt in Iran, which may be the southeastern continuation of the magmatic rocks in SEAOB, formed by Eocene subduction of southern Neo-Tethyan Ocean beneath the Eurasian plate. In contrast, Ghasemi & Talbot (2005) proposed that the Urumieh-Dokhtar magmatic belt formed as a consequence of slab break-off during middle Eocene. Onal et al (2005) also suggested that the Calti and Bizmisen plutons have been formed due to slab break-off. However, an origin via slab break-off for the Eocene magmatism seems to be contradictory in SEAOB. There are several lines of evidence against slab break-off in easternsoutheastern Turkey; Keskin (2003) and Sengo¨r et al. (2003) proposed that a slab break-off event related to Afro-Arabian and Eurasian collision took place during middle Miocene (after 11 Ma), significantly after intrusion of the Eocene granitoids. They also put constraints on the magma chemistry and postulated that the magmatism related to slab break-off is more-like plume or OIB-type, which is significantly different from the Eocene magmatism in SEAOB. Furthermore, surface waveform velocity data by Maggie & Priestly (2005) showed that the slab-break-off and resultant delamination processes within the eastern Mediterranean or western Tethys took place at about 12 Ma. Besides, seismic tomography and tectonic reconstructions by Hafkenscheid et al. (2006) are all agreement with subduction of Arabian continental lithosphere underneath Eurasia and slab break off around 12 Ma in the

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Arabian– Tethyan region. Therefore, these are sound evidence ruling out a possible slab break-off event during Eocene in the SEAOB, and a possible slab break-off would not be a driving mechanism for the Early– Middle Eocene magmatism in easternsoutheastern Turkey. Therefore, we suggest that the Early–Middle Eocene magmatism in the SEAOB should have been formed as result of extensional tectonics and crustal scale fractures due most likely to initial slab rupture or STEP faulting during the final re-organization of plate configurations and construction of the NeoTethyan subduction front during Early–Middle Eocene (Stampfli et al. 2002; Moix et al. 2008). Moix et al. (2008) favour a bending of the subduction front from Cyprus to Erzincan mainly at 48 Ma. It is highly likely that the further bending, and continuous north– south push of the Arabian plate caused a rupture at the overriding plate, or tearing along NE– SW trending crustal scale STEP faults, and left-lateral southwards displacement of the subduction front. This displacement is likely to be accommodated by several NE-trending strike–slip faults such as central Anatolian, Sariz-Goksu, and MalatyaOvacik fault zones since the Early Eocene. This event also resulted in the extensional regime and asthenospheric windows to create magmas with typical arc signatures. The present spatial distribution of the Eocene magmatism in the central, eastern and southeastern Anatolia suggest that these faults acted as loci for the generation and emplacement of Early–Middle Eocene magmatic rocks in the SEAOB.

Magmatic sources In general, the geochemical signatures (major, trace and REE elements) and regional geological constraints of all Late Cretaceous – Cenozoic igneous rocks in the Baskil transect mimic the characteristics of magmatic-arc and post-collisional to late-orogenic magmas of subduction to postsubduction-related origin. Overall, the geochemical characteristics of the magmatic rocks suggest a role of metasomatized mantle in their generation and point to major changes in subduction and mantle geometry beneath the subduction-collision zone between Afro-Arabia and Eurasia from 82.9 to 69.9 Ma then to 44 Ma. Wortel & Spakman (2000), Duggen et al. (2005), McQuarrie et al. (2003), Govers & Worlet (2005), Von Quadt et al (2005) and Neubauer et al. (2005) argue that Alpine magmatism is a direct consequence of collision between Afro-Arabia and Eurasian plates, and that changes in the subduction geometry resulted in diverse magma types at different time periods. The rapidly changing geochemical

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characteristics of the magmas in the Baskil transect are consistent with this model. The proposed large-scale mantle and subduction-collision processes are associated with large-scale extension and contemporary magmatism possessing characteristics of metasomatized mantle signatures. Continental to marine clastic rocks filled the temporally related sedimentary basins. Metasomatized mantle sources with fundamentally distinct geochemical compositions should have been partially melted in the course of the large-scale reorganization of the upper mantle. Potential sources in the transect could be (1) a shallow mantle wedge consisting of depleted asthenosphere and attenuated continental lithospheric mantle enriched by fluids or melts from subducted oceanic lithosphere beneath the Anatolide – Tauride platform at 82.9– 77.5 Ma and (2) metasomatized subcontinental mantle lithosphere beneath the Malatya –Keban platform that had been modified by fluids or melts from the subduction zone or an older subduction event. Possible magma generation models explaining spatial-temporal evolution of magmatism in SEAOB are: (1) subduction of oceanic lithosphere; (2) roll-back of subducted oceanic lithosphere; and (3) initial rupture or step-faulting of subcontinental mantle lithosphere. The field geological, geochemical and geochronological evidence suggest that these models acted successively from Late Cretaceous to Middle Eocene. Simple subduction can only explain the Late Cretaceous arc-type magmatism in a frontal part of the Baskil arc, on the south. The calc-alkaline to alkaline magmatic rocks mainly in the northern and central part of the transect have geochemical features more akin to late-orogenic settings and metasomatized mantle source, leading us to postulating a roll-back model at 74.40 + 0.51 to 69.9 + 0.5 Ma. The timing of magma emplacement indicated that the roll-back resulted in two successive pulses of magmatism generating calc-alkaline as the first phase and alkaline as the second phase (Fig. 9). The change is proposed to reflect an increase in the angle of bending (hinge retreat) of the subducted slab beneath the Malatya –Keban platform, which would likely enhance regional extension leading to formation of alkaline rocks with more within plate affinity during 71.0 + 1.0 to 69.9 + 0.5 Ma. Resumption of calc-alkaline magmatism between 54.4 and 44.3 Ma is preserved as isolated shallow-level intrusions lying parallel to left lateral Yakapinar-Goksun, Malatya-Ovacik, and central Anatolian fault zones that might be originated as crustal scale faults during a continental lithospheric subduction. The present spatial distribution of the Eocene magmatism in the central, eastern and southeastern Anatolia suggest that these

faults acted as loci for the generation and emplacement of Early–Middle Eocene magmatic rocks in the SEAOB. These also resulted in the extensional regime and generation of asthenospheric windows to create magmas with typical arc signatures inherited from older subduction that created Baskil arc.

Geodynamic-tectonic evolution The tectonic environments defined above are related to the northward subduction of African– Arabian plate beneath the Eurasian plate (Anatolide – Tauride platform sensu stricto) initiated during the Late Cretaceous (c. 90 Ma) and culminating in the Middle Miocene (Robertson et al. 2005). Regional geological data suggest that subduction proceeded in two stages; an older episode of northward subduction of oceanic crust of the Neo-Tethyan Ocean beneath eastern Taurides, now represented by the Malatya –Keban metamorphic complex. Ophiolitic rocks and an ensimatic marine arc formed in the supra-subduction zone environment (c. 90 Ma). The younger subduction phase beneath these previously accreted oceanic crust and metamorphic complexes produced arc-related magmatism (82.90–79.9 Ma) intruding the ophiolites and metamorphosed platform rocks of Anatolide – Tauride platform. It appears likely that at an advanced stage of subduction either the geometry of the subduction changed or the rate of convergence of Afro-Arabian plates decreased with respect to Eurasian plate. These changes in convergence initiated a roll-back of the subducting slab during the latest Cretaceous leading to regional extensional due of thickened continental crust, and formation of syntectonic basin formation in the overriding continental plate (Fig. 9). Basinal sedimentary rocks are intruded by Paleocene to Eocene calc-alkaline and alkaline magmatic rocks. The Sivas, Hekimhan, Malatya, Ulukisla and Central Kizilirmak basins are good examples for these extensional basins in easterncentral Anatolia (Gurer & Aldanmaz 2002; Yılmaz & Yılmaz 2006; Kaymakci et al. 2006). These basins formed on top of either ophiolite or ophiolite me´lange nappes or a composite basement of ophiolite and rocks of the Tauride –Anatolide Platform, following the terminal closure of the Vardar and Neo-Tethyan Oceans. The Late Cretaceous to Early Cenozoic basins are filled by basal, terrestrial and shallow marine deposits succeeded by deeper turbiditic sequences of Paleocene to Eocene age. Local basin margin thrust faults emplace ophiolite fragments into the sedimentary basin during sedimentation (e.g. Haymana, C¸ankırı, Sivas and Hınıs). Shallow marine,

MAGMATISM IN THE SOUTHEASTERN ANATOLIAN OROGENIC BELT

Fig. 9. Schematic diagram illustrating the spatial and temporal evolution of magmatism in Baskil-Divrigi transect (error bars represent uncertainties at +2s). 455

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marine–deltaic, fluvial –lacustrine and deltaic – lacustrine deposits filled the basin during the Oligocene to Miocene. Alkaline and calc-alkaline, syn- to post-collisional volcanic products intercalated with the basin in-fills during various stages of basin development. For example, syenitic rocks in Hasancelebi intrude latest Cretaceous to Paleocene sedimentary and volcaniclastic infill of the Hekimhan basin. These may indicate that the extension (probably due to roll-back of the subducted slab of the Neo-Tethyan Ocean) reached its maximum creating relatively deep basins and transgression by marine onlap, then, the incipient rupture of the slab may possibly have resulted in invasion of hot asthenospheric mantle through the slab-window towards shallow and low pressure levels in the lithosphere, and partial melting in the mantle. These melts, then, are either emplaced into the marine basins as sill or dykes and volcanic volcano-sedimentary sequences, or as intrusive stocks. The Arabian/Indian and African continents collided with Eurasian continent along the Bitlis– Zagros in the eastern-southeastern Turkey leading to a thickening crust which underwent incipient slab rupture of the downgoing oceanic plate. Collision appears to have been accompanied by large-scale postorogenic crustal-scale strike –slip faulting and bimodal volcanism during the Early to Middle Eocene. Cessation of subduction likely led to a density increase in the downgoing slab and consequent sinking into the mantle, possibly leading to an initial slab-tear or incipient rupture, leading to the invasion of hot asthenospheric mantle into the overlying mantle wedge (Fig. 9). Although some studies suggest a slab break in eastern and southeastern Anatolia during Late Cretaceous (Boztug et al. 2005; Parlak et al. 2006) and Eocene (Onal et al. 2005), mantle tomography apparently preclude slab break-off at this time in SEAOB and eastern Anatolia. Instead, slab-break or delamination appears to have taken place during mid-Miocene (Sengo¨r et al. 2003; Faccenna et al. 2006; Hafkenscheid et al. 2006; Angus et al. 2006) producing wide spread continental volcanism in eastern Turkey (Keskin 2003). It is therefore proposed that post-collisional Eocene magmatic rocks in Copler, Kabatas, Karamadazi and Horoz, which are localized along crustal scale strike– slip faulting, were triggered by an initial rupture in the subducting slab, or by orogenic collapse following the suturing of Afro-Arabian and Eurasian plates (Arabian platform and Anatolides, sensu stricto). In each case, hot asthenopheric material moved up into the lower lithosphere, leading to magma generation. Crucial to the sequence of events proposed herein is the timing of slab break-off in response to the Cenozoic continental collisions between Arabia and Eurasia. Constraints on this event can

be obtained by relative convergence velocity. According to Van de Zedde & Wortel (2001), convergence velocities of about 3 cm a21 as in the Arabian region are likely to cause subducting slabs to break off about 10 Ma after the onset of continental collision and subduction of continental lithosphere. The timing of continental collision in Turkey suggests that the subducting slabs should have broken off at around 12 Ma (Hafkenscheid et al. 2006). Although this time of slab break-off agrees with other studies (Kohn & Parkinson 2002; Keskin 2003; Sengo¨r et al. 2003; Faccenna et al. 2006; Hafkenscheid et al. 2006; Angus et al. 2006), a lower convergence velocity in Arabian plate would have led to a younger break off locally (Van de Zedde & Worte 2001). If continental collision began at about 22 Ma in the Arabian region, Van de Zedde & Wortel (2001) proposed that the slab would have broken off around 12 Ma. They propose an initiation of slab break-off around 30 Ma in the central Arabian region, below the northern Zagros suture zone after which the tear in the slab may have propagated both eastward and westward along the suture zone. This would have led to slab break off around the expected 12 Ma underneath eastern Turkey, as well as below southern Iran. Considering that the granitoids in the Horoz, Karamadazi, Bizmisen-Calti, Copler and Kabatas regions formed in a post-collisional extension environment between 54.4 and 44.4 Ma, it seems possible that their formation and emplacement does not relate to a slab break off or a subduction, but might relate to the initiation of rupture in the subducting slab (Fig. 9). Upwelling of hot asthenospheric mantle into the incipient rupture would have assisted partial melting in a metasomatized mantle previously contaminated by subduction. Magmas having subduction and/or arc signatures would have resulted.

Conclusion Late Cretaceous to Middle Eocene calc-alkaline to alkaline magmatic rocks emplaced within the southeastern Anatolian orogenic belt result from the complex collision between the Afro-Arabian and Eurasian plates and the subduction of the southern and northern Neo-Tethyan oceanic basins beneath the Eurasian continental margin during the Alpine –Himalayan orogeny. During the subduction, the southern branch of Neo-Tethyan Ocean was consumed, and oceanic basin was subducted beneath the eastern Taurides (Eastern Anatolide – Tauride platform) along the Bitlis–Zagros subduction system. Successive collisional and post-collisional events due to slab roll-back and a possible incipient slab rupture played an important role in the magma generation, leading to distinct chemical

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compositions ranging from syenite, syenodiorite, trachyte, monzonite, phonolite, granodiorite, diorite, monzodiorite and gabbro. The progression of calcalkaline to alkaline magmatism within the transect is explained as a consequence of gradual change in the geometry of subduction and slab roll-back followed by a possible incipient slab rupture following the subduction and collision. This led to block faulting and subsidence, and thus to the preservation of near-surface magmatic systems. These vary in time, spatial distribution, and composition. 40 Ar/39Ar ages supplemented by a few U –Pb ages geochronology from major plutons demonstrate a general younging of magmatism in the transect from c. 83 Ma in the south (Baskil) to c. 69 Ma in the north (Divrigi-Keban), followed by a c. 44 Ma scattered magmatic complexes now found along a NE-trending arcuate belt between Copler and Horoz. The geochemistry in the magmatic rocks suggest two main sources for the melts: (1) a mantlewedge and subducted oceanic lithosphere producing arc-type magma, and (2) metasomatized lithospheric mantle modified by subduction producing magmatic rocks with more metasomatized mantle and within plate signatures. The combination of geochemical and geochronological data suggest a temporal and spatial transition from subduction-related to post-collision and to late-orogenic magmatism in the eastern– southeastern Turkey; subduction-related magmatism by the closure of the Neo-Tethyan Ocean whereas post-collision and late orogenic-within plate-related magmatism by the collision of a northern promontory of the SE Anatolian orogenic belt with northerly derived ophiolitic rocks. The work presented here is part of an on-going research project carried out by the authors at the Mineral Deposits Research Unit at UBC, Canada. The work is generously sponsored by TeckCominco Ltd, Barrick Gold Corp., and Tuprag Metal Madencilik Ltd. The geochemical and geochronological data from Hasancelebi region is part of a research project sponsored by Scientific and Research Council of Turkey (TUBITAK)-CAYDAG with grant no: 103Y023. The authors express their sincere thanks to support from TUBITAK. This paper benefited from reviews by E. Aldanmaz and B. Davies.

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Alpine high pressure evolution of the eastern Bitlis complex, SE Turkey ¨ NSLI1*, O. CANDAN2, R. BOUSQUET1, G. RIMMELE3, A. OKAY4 & J. GOFF5 R. OBERHA 1

Institute of Geosciences, Potsdam University, Karl Liebknecht Strasse 24, 14476 Potsdam, Germany 2 ¨ Institute of Geosciences, Dokuz Eylu¨l Universitesi, Mu¨hendislik Faku¨ltesi, Tınaztepe Yerles¸kesi 35160 Buca/I˙zmir, Turkey 3

Schlumberger, Well Service, Clamart, France

4

Istanbul Technical University, Eurasia Institute of Earth Sciences, Maslak 34469 Istanbul, Turkey 5

BP Exploration Operating Co. Ltd, Middlesex, UK

*Corresponding author (e-mail: [email protected]) Abstract: The Bitlis complex, SE Anatolia, constitutes a crystalline complex derived from the north of the Arabian Plate, accreted to the South Armenian block. Metamorphic studies in the cover sequences of the Bitlis complex allow constraining the thermal evolution of the massif by metamorphic index minerals. A regionally distributed low temperature-high pressure (LT-HP) metamorphic evolution is documented by glaucophane, relics of carpholite in chloritoid-bearing schists and pseudomorphs after aragonite in marbles. The metamorphic age of these HP assemblages is constrained by Ar isotope dating as 74 + 2 Ma. This indicates that (i) the Bitlis complex represents a terrane detached from the Arabian indenter that was subducted and stacked to form a nappe complex during the closure of the Neo-Tethys and (ii) that during Late Cretaceous to Cenozoic evolution the Bitlis complex never underwent temperatures over 450 8C. The consequences of the metamorphic evolution of the Bitlis complex (a cold continental block within a hot environment) for the Eastern Anatolian plateau are complied in a crustal section.

The Bitlis complex is situated at the front of the Eurasian plate collage overriding the Arabian platform. This mountain belt, pinched between the Taurids and Zagros (east –west) as well as Caucasus and Arabian plate (north–south), is part of the southern edge of a high plateau that extends northward to the Caucasus. Recent investigations in southern Armenia revealed blueschists along the Sevan–Akera suture zone, the metamorphic age of this high-pressure (HP) metamorphism event is 95 to 90 Ma (Rolland et al. 2008). Recent seismologic work revealed normal crustal thickness (c. 40 km) and a reduction in seismic velocities at depth (Zor et al. 2003; Go¨k et al. 2007). This is interpreted as asthenospheric upwelling and a missing of the lithospheric mantle lid (Sengo¨r et al. 2003; Lei & Zhao 2007). Geophysics showed that this is a critical area for the geodynamic evolution as well as the dynamics of the North Anatolian fault system (Facenna et al. 2006). All geodynamic models, mostly supported by geochemical investigations of volcanic rocks (Keskin 2003) presented so far

(Sengo¨r et al. 2003; Keskin 2003; Lei & Zhao 2007) assume that the Bitlis-Pu¨turge complex was situated in an arc position. As far as geological information is considered, some of the early reports of the geological survey (MTA) are difficult to access for international readers. Go¨ncu¨oglu and co-workers mapped part of the Bitlis metamorphic complex, between Bitlis and Mus¸, during the 1980s. Parts of this work were published as a short compilation (Go¨ncu¨oglu & Turhan 1984). We present new investigations on the metamorphic petrology of the Bitlis complex, gathered in the context of the Middle East Basins Evolution programme (MEBE). The aim of this study is to understand the Alpine geodynamics of the eastern Bitlis complex by adding knowledge to the structural and thermal evolution. The consequences of these new findings for the geodynamic evolution of Eastern Anatolia and the high plateau that built up between the Arabian and the Eurasian plates are discussed. In this paper we report occurrence and age of HP metamorphism indicative minerals

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 461–483. DOI: 10.1144/SP340.20 0305-8719/10/$15.00 # The Geological Society of London 2010.

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in meta-sediments and mafic metamorphic rocks from the Palaeozoic to Mesozoic cover of the Bitlis complex.

Geological setting of SE Turkey The Bitlis complex forms an arcuate metamorphic belt, about 30 km wide and 500 km long, rimming the Arabian Platform in SE Anatolia (Fig. 1). It is separated from the Arabian Platform by a narrow belt of Upper Cretaceous to Eocene flysch and ophiolitic me´lange and Miocene sediments. Along the Northern front of the Arabian indenter a set of collisional autochthonous and allochthonous massifs is known. From south to north, these are: the Great Zap anticlinorium; the Eocene olistostrome of the Hakkari complex overlain by Cretaceous me´langes of the Yu¨ksekova complex; the metamorphics of the Bitlis complex; and finally the Quaternary volcanics north of Lake Van. The Bitlis metamorphic complex comprises Precambrian to Cretaceous rocks and builds up the uppermost tectonic unit of the area. It is covered by Cenozoic sediments and Quaternary volcanics in the north and overlies Cretaceous (Yu¨ksekova complex) and Eocene to Miocene series, the Hakkari and Maden complexes (Baykan and Ziyaret Formations, south of Bitlis), as well as the sediments of the northern margin of the Arabian autochthon. In an early description, Tolun (1953) interpreted the metamorphic rocks of the Bitlis complex as forming the basement of the region. According to Kellogg (1960) and Go¨ncu¨oglu & Turhan (1984) the Bitlis metamorphics are equivalents of the Arabian autochthonous succession and are assigned a Devonian –Upper Cretaceous depositional age to the meta-sediments. First detailed descriptions of the Bitlis complex were given by Boray (1975), Cailavan et al. (1984), Yılmaz, (1978) and Hall (1976). Horstink (1971), Sungurlu (1974) and Sengo¨r & Yılmaz (1981) proposed various geodynamic interpretations. In their paper Go¨ncu¨oglu & Turhan (1984) suggested that the basic geological data used in earlier interpretations were remarkably incomplete. A remark that still holds true for these eastern areas! The legend of Figure 1 adapts the formation names after Perincek (pers. comm. MEBE). For our compilation we used older literature and different formation names.

Lithostratigraphy and metamorphism of the Palaeozoic to Mesozoic Bitlis complex In the following paragraphs we present a synthesis of the geology of the Bitlis complex mainly based on Turkish literature. A generalized

lithostratigraphic section is comprised as follows from bottom to top. † The pre- to infra-Cambrian Hizan group, composed of gneisses, meta-basic rocks and schists separated into three formations: the Andok augengneiss with biotite, muscovite, ¨ naldi Formation with amphiamphibole, the U bolites and garnet-amphibolites with relics of eclogite (Okay et al. 1985) and the Ohin schists containing biotite, muscovite, garnet and amphibole. † Palaeozoic rocks of the Mutki group unconformably overlie the Hizan group. The base of this rock group is made of the Devonian Meydan Formation comprising meta-conglomerates, quartzites and greenschists with limestone interlayers, reef limestones and albite-chloriteactinolite-chloritoid schist’s of probably volcanogenic origin. The Meydan Formation grades into the volcanoclastic C¸esme Formation consisting of felsic meta-volcanic and meta-tuffs. Both formations are intruded by the Mus¸ metagranite and the C¸esme Formation is considered to be the product of this magmatism. The age of the felsic meta-volcanic rocks is reported as c. 454 Ma based on Rb–Sr whole rock analyses (Yılmaz et al. 1981) The leucocratic granitoids have wide exposures north of Hizan, north of Mutki and SW of Mus¸. Their age is badly constrained (Helvaci & Griffin 1984). They are not affected by the Precambrian regional metamorphism but feldspathized meta-volcanic rocks reveal an age of c. 91 Ma while from the Avnik granite an amphibole-whole-rockfeldspar age of 71 Ma, a biotite-whole-rock age of 41 Ma and from micaschist a chloritemuscovite age of 38 Ma are reported (Helvaci & Griffin 1984). These late Alpine mineral ages are interpreted to reflect recrystallization during emplacement deformation. Rb–Sr and K –Ar white mica ages from the Mus¸ granite are Late Cretaceous (73 –107 Ma) (Go¨ncu¨oglu 1984). † All three units, Meydan Formation, C¸esme Fm, and Mus¸ meta-granite are unconformably overlain by the C¸irrik Limestone, which is a sequence of recrystallized limestone inter-bedded with chloritoid schists and graphite schists. This sequence grades up into calcschists and thinbedded recrystallized limestones. A Lower Permian age is assigned to these rocks. On top of these thinly bedded meta-carbonates a sequence of coarsely bedded recrystallized limestones with interlayers of calcschists, metasandstones and chlorite schists of Upper Permian age, the so-called Malato Formation, was deposited.

HP EVOLUTION OF THE BITLIS COMPLEX 463

Fig. 1. Geological map of the Eastern Bitlis complex (modified after MTA 1:5 000 000 maps Cicre and Van). Inset: Sketch of ophiolite zones and crystalline basement salients of the Eastern Mediterranean Province. 1, Salients of the Hercynian granite-metamorphic basement; 2, salients of the Precambrian granite-metamorphic basement; 3, ophiolite complexes and zones of serpentinite me´lange; 4, inferred suture of Palaeozoic Tethys. Abbreviations: WP, EP, Western and Eastern Pontides; WT, CT, ET, Western, Central and Eastern Taurides; MD, Menderes massif; KR, Kirshehir massif; UD, Ulu Dag massif; BT, Bitlis complex; FR, Fore Range zone of the Great Caucasus; B, Bechasine zone; GC, Main Range of the Great Caucasus; TC, Transcaucasian massif (a, Dzirula; b, Khrami; c, Lokhi and Murguz salients); LK, Lokhi-Karabakh zone; SA,Sevano-Akera zone; MS, Miskhan massif; WI, ECI, Western Iran and Eastern Central Iran; ZAG, Zagros; A, Alborz. Boxed area is the region of study.

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† The Triassic rocks of the Tu¨tu¨ Formation form the upper part of the Mutki group, the base of which consists of recrystallized limestones and calcschists grading upward into meta-shales, met tuffs, meta-diabases and meta-basalts and finally meta-conglomerates, meta-mudstones and shales, indicating a drastic change in depositional conditions. The upper part of the Mutki group contains meta-quartzporphyres. They are interpreted as being the result of the opening of the Tethys Ocean. The Bitlis complex has undergone a low to mediumgrade Alpine metamorphism (Sengu¨n 1993). K– Ar ages from the western part of the Bitlis complex near Pu¨tu¨rge gave 71.2 + 3.6 Ma (Hempton 1985). Helvac¸i & Griffin (1984) reported similar mineral ages from the Bingo¨l area in the western part of the Bitlis complex.

The Mesozoic ophiolitic sequences Tectonically underlying the Tu¨tu¨ Formation ophiolitic me´langes are found. They have been termed the Gu¨leman ophiolites (Go¨ncu¨oglu & Turhan 1984) after the Upper Jurassic – Lower Cretaceous sequence found far to the SW of the Bitlis region. Eastward in the Van and Hakkari regions, the term Yu¨ksekova complex is used because it is more of an ophiolitic me´lange than a regular ophiolite. In the ophiolitic me´langes near Mutki glaucophanebearing blocks have been described (Hall & Mason 1972). In the Hakkari-Narlı region it forms large flat-lying klippen over the Eocene-aged Hakkari complex, and is tectonically overlain by the Bitlis metamorphic rocks. In the Bitlis-Baykan, region it forms tectonic slivers between the Bitlis metamorphic rocks and the underlying Maden complex. The Yu¨ksekova complex has a me´langelike internal structure and represents a strongly deformed accretionary complex. It consists of a chaotic jumble of basalt, gabbro, serpentinite, pelagic limestone, radiolarian chert, neritic limestone, granodiorite, sandstone, siltstone, and shale with an estimated vertical thickness of about 2000 metres. The youngest limestone blocks found in the Yu¨ksekova complex give Coniacian –Campanian ages (Perinc¸ek 1990). SE of the me´lange complex in the Cilo mountains the Oramar and Karadas ophio¨ zkaya 1982). Ophiolitic rocks lites are reported (O also crop out north of the Bitlis complex on the shores of Lake Van. This Gevas ‘ophiolite’ is of special importance as it lays directly under the Bitlis metamorphic rocks (Yılmaz et al. 1981), implying large-scale allochthony for the Bitlis complex. The Gevas ‘ophiolite’ is a disordered ophiolite consisting of serpentinite, gabbro, basalt and limestone blocks. Some of the limestone blocks have yielded

¨ zer 1992) showing an Maastrichtian rudists (O Arabian rather than a Taurid affinity.

Lithostratigraphy of Cenozoic complexes From the western parts of the Bitlis complex, me´langes discontinuously overlie non-metamorphic wildflysch and olistostromal units of the Upper Maastrichtian (Kinzu Formation and Kizilgic Formation after Go¨ncu¨oglu & Turhan 1984). The upper contacts of these Eocene formations have been described as tectonic. This indicates that the Bitlis complex acquired its structure after the Late Eocene times. In the area of investigation, a belt of two complex assemblages occurs below the Bitlis complex and the ophiolitic me´langes: the Hakkari complex and the Maden complex. The Hakkari complex covers large areas SE of the Bitlis complex, between Narlı and Yu¨ksekova, where it tectonically overlies the Eocene and Miocene formations of the autochthon. It is equivalent in time to the Maden complex, which crops out widely farther west. The Hakkari complex differs from the Maden complex by the scarcity of the volcanic rocks. The Hakkari complex is divided into two formations, a lower unit called Urse Formation and an upper me´lange-type unit named as the Durankaya complex (Perinc¸ek 1979, 1990; Yılmaz & Duran 1997). The Urse Formation consists predominantly of slightly metamorphosed siltstone, shale and fine-grained sandstone with limestone intercalations. It is well exposed on the Baskale-Hakkari road, where it starts with fine-grained metabasites and passes into a very thick slate series. Near Hakkari, the slates are overlain by medium-bedded dark carbonates. The Durankaya complex tectonically overlies the Urse Formation and consists of blocks of pelagic and neritic limestone, serpentinite, gabbro, basalt and amphibolite in a strongly deformed shale matrix. Some of the pelagic limestone ‘blocks’ probably represent original limestone intercalations. Such limestones have yielded pelagic and neritic foraminifera of Early to Mid-Eocene age (Perinc¸ek 1990; Yılmaz & Duran 1997). The Maden complex is of the same age as the Hakkari complex in regions west of Narlı. It differs from the Hakkari complex by the presence of abundant volcanic rocks. It widely crops out along the southern margin of the Bitlis complex, as tectonic slivers either directly under the Bitlis metamorphic rocks or through an intervening thrust sheet of the ophiolitic me´lange (Yu¨ksekova complex). The Maden complex consists of sandstone, conglomerate, red pelagic limestone, basaltic lava, and tuff. Limestones in the Maden complex yield Lower to Middle Eocene foraminifera. In a

HP EVOLUTION OF THE BITLIS COMPLEX

few localities the Maden complex is reported as lying unconformably over the Bitlis metamorphic rocks. However, in most places it is positioned between the Cenozoic formations of the Arabian Platform and the Bitlis complex. Yigitbas & Yılmaz (1996) regard the Maden complex as products of a short-lived Mid-Eocene back-arc basin, above the northward-dipping subduction zone between the Arabian Platform and the AnatolideTauride Block as represented by the Bitlis complex. Around Baykan, south of Bitlis, the Maden complex (locally named as the Baykan complex by Go¨ncu¨oglu & Turhan 1992) is lithologically highly variable. It ranges from a regular flysch sequence to an ophiolitic me´lange. It is difficult to put a boundary between the flyschoid Maden complex and the overlying ophiolitic me´lange (Yu¨ksekova complex or the Gu¨leman ophiolite). The Kırkgec¸it Formation occupies large areas SE of Van, where it lies unconformably over the Bitlis metamorphic rocks, the Yu¨ksekova complex and the Hakkari complex. It consists of siliciclastic turbidites with extensive olistostrome horizons (Perinc¸ek 1990). Scarce fossils indicate a Late Eocene to Early Miocene age for the Kırkgec¸it Formation. Post-Miocene tectonics has resulted into the imbrication of the Kırkgec¸it Formation with the underlying units.

Autochthonous sequence of SE Anatolia (Arabian Platform) An account of the stratigraphy of the Arabian Platform as exposed in the anticlines south of Hakkari is given in the following. The authochthonous sequence (Ketin 1980) is well exposed in two faulted anticlines along the Zap River between Hakkari and C ¸ ukurca (Rigo de Rhigi & Cortesini 1964) the Great Zap anticline in the north and the C ¸ ukurca anticline in the south (Fig. 1). The anticlines are major regional east –west-trending structures with half wavelengths of 10–15 km, and extend along strike for over 100 km. Their southern margins are cut by thrust faults. The Great Zap and C ¸ ukurca anticlines expose a thick sedimentary sequence from Early Cambrian to Eocene, albeit with major gaps. Clastic rocks dominate the Cambrian to Carboniferous (Janvier et al. 1984) sequence, whereas the Permian (Ko¨ylu¨oglu & Altiner 2001) to Eocene sequence is largely formed by shallow marine carbonates. The lowermost authochthonous sequences in the core of the Great Zap anticline are medium to thickly bedded sandstones and siltstones belonging to the Derik Group. In the Great Zap anticline this group has a minimum thickness of 600 m and, based on scarce fossils, is of Early Cambrian age (Perinc¸ek 1990;

465

Yılmaz & Duran 1997). The Middle Cambrian dolomites and limestones of the Koruk Formation conformably overlie the arenites of the Derik Group. The Koruk Formation is equivalent to the C ¸ altepe Limestone in the Taurides. It is in turn overlain by yellowish brown siltstone, sandstone, and shale intercalation of the Seydisehir Formation of Late Cambrian –Ordovician age, which forms the core of the C¸ukurca anticline farther south. The Seydisehir Formation is unconformably overlain by the strikingly variegated, thickly bedded quartzites of Upper Devonian age belonging to the Yıgınlı Formation. The Late Devonian age is based on fish fossils (Janvier et al. 1984). The quartzites show strong current bedding and have thin shale and siltstone interlayers, and have a measured thickness of 295 m. The quartzites of the Yıgınlı Formation are conformably overlain by the shale, sandstone, sandy dolomite and limestone of the Ko¨pru¨lu¨ Formation. The Ko¨pru¨lu¨ Formation straddles the Devonian-Carboniferous boundary (Perinc¸ek 1990; Yılmaz & Duran 1997) and has a thickness of about 200 m. The Upper Devonian – Lower Carboniferous Ko¨pru¨lu¨ Formation is unconformably overlain by a thick carbonate sequence of Late Permian age. This Tanin Group has a thickness of nearly 1000 m, and consists of dark, bituminous, limestone and dolomitic limestone locally with chert nodules. A rich foraminifera fauna indicates the presence of all the Upper Permian stages from Murgabian to Dorashamian (Ko¨ylu¨oglu & Altıner 1989). The dark Upper Permian carbonates are overlain by purple, green, yellow shale, siltstone, shaley and lithographic limestone of Early Triassic age belonging to the C ¸ ıglı Group. The C¸ıglı Group has a thickness of about 500 m, and is conformably overlain by the thickly bedded, neritic limestone and dolomite of the Cudi Group of Middle Triassic to Early Cretaceous age. The thickness of the Cudi Group increases from west to east, and in the Hakkari-C¸ukurca region is more than 1000 m. The youngest ages from the Cudi Group are Aptian – Albian, however, in many regions the Lower Cretaceous sequence is eroded, and the Upper Cretaceous rests unconformably over the older formations (Perinc¸ek 1990). In the northern margin of the Great Zap anticline the neritic carbonates of the Cudi Group are unconformably overlain by the shaly pelagic limestones of the Sayındere Formation, about 200 m thick. A rich pelagic foraminifera fauna in the limestones indicates a Campanian age for the Sayındere Formation. The Sayındere Formation is unconformably overlain by the Campanian –Lower Paleocene sandstone, shale and marl of the Germav Formation. In several localities south of Hakkari, the Sayındere Formation is eroded and the Germav Formation rests directly on the neritic carbonates of the Cudi Group. The Lower

466

¨ NSLI ET AL. R. OBERHA

to Upper Eocene Midyat Group unconformably overlies the Germav Formation. The Midyat Group starts with red conglomerates and sandstones, and passes up into thinly to thickly bedded limestones locally with chert nodules and interlayers. The Midyat Group is tectonically overlain by the allochthonous Hakkari complex along the frontal thrust. Slivers of Miocene continental sandstone, siltstone and mudstone (Selmo Formation) indicate a Miocene and younger age for the thrusting (Perinc¸ek 1990).

New geological observations, metamorphic data and age constraints Cross-sections in the Eastern Bitlis complex Hakkari section. The easternmost cross-section runs along the main road from Van to Hakkari and C ¸ ukurca. From Van towards the SE, Oligo – Miocene sediments of the Kırkgec¸it Formation are crossed, that exhibit phenomena of late tectonic movements typifying in the whole region. The Cenozoic and recent deformation led to faulting and block tilting. This sequence overlies Cretaceous ophiolitic coloured me´langes, with a serpentinitic and shaly matrix that contain large limestone blocks (Yu¨ksekova fm). The contacts of the Oligo– Miocene sediments on the Cretaceous me´lange are reported as transgressional although near Van thrusting of Cretaceous me´lange and Eocene sequences is evidenced. Between Baskale and Hakkari, SW of the Yu¨ksekova junction, the rocks of the Urse Formation consist of silvery slates that show excellent kink bands (Fig. 2a) deformation patterns, which relate to metamorphism. Nearby, volcanoclastic and volcanic rocks show a bluish tint. While arkosic metasediments did not show any evident macroscopic trace of a low-grade HP-LT metamorphism the meta-volcanics contain blue amphibole. The blueschists of the Urse Formation are overlain by non-metamorphic Eocene flysch-type sediments (Fig. 2b). Gevas – C ¸ atak – Narli section (Fig. 3). Starting from Van, the first observation of the Northern contacts of the Bitlis complex is exposed around Gevas. There, the so-called Gevas ophiolite (Yılmaz 1978) is actually an ophiolitic me´lange with a serpentinitic matrix that contains blocks of gabbros, basaltic rocks, cherts, limestones, and radiolarites. This me´lange clearly dips southwards below the metamorphic sediments of the Bitlis complex with an angle of c. 20–308 (Fig. 4). The rather flat-lying contact is easily recognized by an alignment of

springs. Strongly deformed and brecciated rocks of both complexes dominate the contact: the ophiolitic me´lange and the overlying Bitlis metamorphics. Inspection of the contact at several locations however revealed that between the ophiolitic me´lange and the Palaeozoic marbles of C¸adir dag a typical rock sequence consisting of metasandstones and reddish marly calcitic marbles as well as marbles with chert layers occur. The reddish marly marbles resemble ‘couche rouge’-type sediments. This rock sequence resembles in many aspects the Cretaceous assemblages found in the Western Taurides. A cross-section east of Gevas (Fig. 4b) exhibits radiolarites of the me´lange complex in direct steep contact with mylonitic marbles (Fig. 5a). These marble-mylonites are part of a metamorphic marble-schist sequence that typically occurs at the base of the Triassic sequence. In the investigated area the metamorphic sequence comprises calcareous and dolomitic marble seams in a greyish chlorite albite schist sequence. In some metapelitic layers phengite and chloritoid occur. Upward in the sequence the amount of

Fig. 2. Photographs of metamorphic rocks of the Eocene Urse Formation along the Bashkale Hakkari road. (a) slaty micaschist with kink bands. (b) Tectonic contact between Eocene Urse blueschists and Eocene flysch type clastic sediments.

HP EVOLUTION OF THE BITLIS COMPLEX

467

Fig. 3. Cross-section along the C¸atak valley in the Eastern part of the meta-sedimentary cover of the Bitlis complex. Mafic and metapelitic rocks exhibit relicts of a HP-LT metamorphic event throughout the section.

dolomitic marbles diminishes and thin tuffitic mafic layers are intercalated within the schists. The mafic layers may also become more substantial and are composed of chlorite, epidote, amphibole and albite. In the uppermost part of the schist-marble sequence the mafic layers show intercalation of greenschist and blueschists (Fig. 5b). The

blueschists contain albite, chlorite, glaucophane and epidote (Fig. 5c). The schist-marble sequence is conformably overlain by Megalodon (Fig. 5d) bearing massive grey marbles of possible Triassic age. Another section across the contact shows strongly deformed serpentinites that are capped by lysvenitic layers. Above these, light grey marbles

Fig. 4. (a) Photograph of the inverted contact putting the sediments of the Bitlis metamorphic complex over the me´lange sequence of Gevas. View westward from NE. The village of Gevas is located at the right margin of the image. The box indicates the location of the section (b) taken over the tectonic contact. (b) Section trough the inverted tectonic contact Gevas me´lange–Bitlis complex. At the base mylonitic dolomite marbles tectonically overlies radiolarian cherts. These grade into metabasites, schists, quartzites and, finally meta-conglomerate at the base of the Bitlis marbles. In the schistose part of the section blueschists occur.

468

¨ NSLI ET AL. R. OBERHA

Fig. 5. Photographs (a) Mylonitic marbles at the base contact of the Bitlis with the Gevas complex. (b) Blueschist and greenschists, intercalated within the marbles. (c) Microphotograph of glaucophane rimmed by blue-green amphibole. (d) Triassic (?) Megalodon occurring in marbles of the Bitlis complex (Fig. 4b top left). (e) Calcitic rosetta ¨ zer 2005) (pseudomorphs after aragonite?) from the Bitlis marbles. (f) Cretaceous Rudist of Arabian faunal affinity (O found in limestones of the Gevas complex.

with whitish elongated calcite aggregates occur. Some of these aggregates show fibrous calcite, possibly pseudomorphs after aragonite (Fig. 5e). To some extent this resembles the Upper Cretaceous ‘rosetta’ type marbles with their conspicuous cherty layers. Overlaying these rocks dark grey Permian

marbles occur. A third section exhibits cataclastic contact relations. However, between the Permian Bitlis marbles and the ophiolite complex again a conspicuous sequence resembling closely the Cretaceous Tu¨tu¨ formation occurs. This metamorphic sequence contains relics of carpholite fibres, this

HP EVOLUTION OF THE BITLIS COMPLEX

again points to a low-grade high-pressure metamorphism. The limestone blocks within the Gevas ophiolitic me´lange revealed rudists (Fig. 5f) that ¨ zer (2005) show an Arabic according to O facies affinity. Entering the C¸atak valley the first outcrops of the Palaeozoic marbles show strong cataclastic disruption and earlier ductile folding. These marbles are calcitic but show in some places relict whitish crystals that might have formed as sedimentary aragonite (Fig. 5e). In ductile shear bands around these rosetta-forming aggregates, fibrous calcite replaces metamorphic aragonite. These fibrous features are a clear hint to a low-grade HP-LT overprint close to the base of the Bitlis metamorphic complex. Intercalated with these Palaeozoic marbles, a sequence of black to silvery schist with mafic intercalations occur some 5 km southward, near Kayabogaz. In those schists, we identified the very first occurrence of carpholite relics in Eastern Anatolia. In these rocks carpholite has reacted to form chloritoid and quartz. Sometimes kyanite can be found in these chloritoid bearing rocks. The associated mafic rocks exhibit a bluish tint and glaucophane was found in thin section from these meta-basic intercalations. Further downstream of the C ¸ atak valley the series are strongly disrupted by complex folding and thrusting. The general structure is an open folding superimposed on overturned south vergent folded structures (Fig. 3). Along this cross-section, the Bitlis metamorphic complex exposes only meta-sediments of upper Palaeozoic to Mesozoic sequences. North of C ¸ atak spectacular chloritoid-bearing rocks with crystal sizes up to 2–3 cm are exposed. South of C ¸ atak near Narli chloritoid-garnet-bearing parageneses occur in the metapelitic rocks. These highergrade metapelites and their marble envelope are locally thrust on top of non-metamorphosed Eocene pillow lavas. Below this tectonic sliver, steep and strongly folded Palaeozoic to PermoTriassic marbles forms the southern frontal part of the Bitlis complex. Along the C¸atak River, the marbles contain fresh carpholite without chloritoid.

Metamorphic evolution In the Bitlis cover (Fig. 6) silvery Al-rich metapelitic schists, intercalated with calcareous marbles, generally contain the assemblage chlorite-phengitequartz. Along the frontal (south) and basal parts the assemblage carpholite-chlorite-phengite-quartz occurs. In rare cases pyrophyllite-chlorite-carpholite assemblages testify prograde relicts. In internal parts of the nappe complex most of the carpholite has reacted to form chloritoid and only carpholite

469

relics included in quartz veins and nodules remain. The mineral stable assemblage is chloritoidphengite-quartz-chlorite sometimes associated with paragonite. A few samples contain kyanite and chloritoid, others chloritoid and epidote and one sample containing garnet together with chloritoid, chlorite and phengite was found. Mafic rocks associated with these metapelites contain glaucophane and testify blueschist metamorphic conditions. At regional scale the distribution of Fe, Mg-carpholite and glaucophane documents the extent of high-pressure low-temperature metamorphism all over the meta-sedimentary part of the Eastern Bitlis complex. Up to now we were not able to identify a similar HP metamorphism in the crystalline basement rocks. A first report on eclogites was given from the central part of the Bitlis complex at Mt Gablor (Okay et al. 1985) south of Mus. There, eclogites occur within garnet mica schists and contain kyanite. P–T estimates are reported with temperatures between 600 8C and 650 8C at 1.0 to 2.0 GPa. Based on findings of eclogite pebbles in Lower Palaeozoic microconglomerates a Pan African age was assumed for these eclogites (Go¨ncu¨oglu et al. 1997). Both high-pressure index minerals, glaucophane in metabasites and carpholite in metapelites can only be used for a rough estimate of the P –T conditions (e.g. Oberha¨nsli et al. 1995, 2001). Therefore we apply the multi-equilibrium approach developed and tested for chlorite-phengite-quartz bearing meta-sediments (Vidal et al. 1999; Vidal & Parra 2000; Parra et al. 2002; Rimmele et al. 2005). Fe, Mg-carpholite occurs overall the schist in the Bitlis complex (Fig. 6). Microprobe analysis document a homogenous chemical composition of Fe, Mg-carpholite (Table 1) with a relatively high Mg-content (XMg ¼ 0.65 – 0.70) in marbles (Fig. 7), and a lower Mg-content (XMg ¼ 0.33 – 0.50) in metapelitic schists. Chloritoid (Table 2) always shows significantly higher Fe contents (XMg ¼ 0.05 – 0.35) (Fig. 7). Thus the Fe-Mg partitioning coefficient [KD ¼ ( Fe/Mg )A/( Fe/Mg )B] for a carpholite/chloritoid pair is equal to 8. This KD value is similar to those reported in the literature for the same rock-type and the equivalent P–T conditions (Crete: Theye et al. 1992; Oman: Vidal & Theye 1996, Alps: Bousquet et al. 2002). Using microprobe analyses and recalculated end members of chlorite such as clinochlore, daphnite, sudoite and amesite as well as white mica such as celadonite, pyrophyllite and muscovite (Tables 3 & 4) it is possible to calculate P –T conditions for each chlorite-phengite mineral pair. Calculated P, T conditions indicate P ¼ 8– 10 kb and T ¼ 320 8C for the prograde relicts, P ¼ 10 –11 kb and T ¼ 350 – 400 8C for peak conditions and a wide distribution of temperatures (T: 370 – 480 8C) at

470 ¨ NSLI ET AL. R. OBERHA

Fig. 6. Geological map of the eastern Bitlis complex (modified after the MTA 1:500 000 maps Cicre & Van). The distribution of HP-LT index minerals throughout the Eastern Bitlis complex and in the Eocene Urse Formation is indicated. Right side: Schematic lithostratigraphic column (modified after Go¨ncu¨oglu & Turhan 1997).

HP EVOLUTION OF THE BITLIS COMPLEX

471

Table 1. Representative electron microprobe analyses of HP-LT index minerals Fe-Mg carpholite and glaucophane. All electron microprobe analyses using natural and synthetic mineral standards at standard conditions (15 kV, 20 nA) were performed on Cameca SX 100 at GFZ Potsdam and at CAMPARIS Paris VI Sample

VAN 7 car

VAN 7 car

VAN 7 car

VAN 11 car

VAN 41 car

VAN 41 car

Sample

VAN 55 glauc

VAN 55 glauc

VAN 55 bl-gn a

SiO2 Al2O3 FeO MnO MgO F Total

37.62 31.81 13.58 0.12 5.01 0.11 88.25

38.19 32.08 12.16 0.03 5.62 0.61 88.69

38.44 32.33 11.50 0.16 5.87 0.86 89.15

42.46 29.60 10.45 0.20 5.89 0.08 88.68

39.35 32.29 7.29 0.14 8.78 2.45 90.30

39.32 32.36 7.70 0.10 8.80 1.80 90.07

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Total

58.18 0.09 4.54 16.00 0.15 11.36 0.49 6.95 0.00 97.77

57.84 0.21 2.35 17.56 0.11 11.88 1.17 6.82 0.02 97.97

48.81 0.63 8.98 6.54 0.17 18.46 12.31 1.75 0.21 97.86

Structural formula based on 13 cations

Structural formula based on 5 cations Si Al Fe3þ Fe2þ Mn Mg F

2.00 1.99 0.01 0.60 0.01 0.40 0.02

2.02 2.01 0.00 0.54 0.00 0.45 0.10

2.02 2.02 0.00 0.51 0.01 0.46 0.14

2.23 1.99 0.01 0.49 0.01 0.50 0.01

2.03 1.99 0.01 0.31 0.01 0.68 0.41

2.02 1.98 0.02 0.31 0.00 0.68 0.29

XMg

0.397

0.451

0.473

0.501

0.684

0.681

lower pressures (P: 3–6 kb) for the retrograde evolution (Fig. 8). These observations from Bitlis meta-sediments fit well with those made in Tethyan meta-sediments in Western Turkey: the Lycian nappes (Rimmele et al. 2002) and the Afyon Zone (Candan et al. 2005).

Si Ti Al Fe Mn Mg Ca Na K

8.04 0.01 0.74 1.85 0.02 2.34 0.07 1.86 0.00

8.06 0.02 0.39 2.05 0.01 2.47 0.17 1.84 0.00

6.82 0.07 1.48 0.76 0.02 3.85 1.84 0.47 0.04

Age of metamorphism Phengites from carpholite bearing metasediments from the C ¸ atak valley were dated by laser 40 Ar/39Ar method. Small amounts of pure phengite separated from carpholite fibres in quartz exudates

Fig. 7. Composition of carpholite and chloritoid from the Bitlis meta-sediments. As generally observed chloritoid is more iron rich than associated carpholite. The data presented fits well with data reported from Western Turkey (Lycian nappes and Afyon zone).

472

Table 2. Representative electron microprobe analyses of chlorites with temperature estimates based on Jowett (1991). Mineral associations are indicated; sample with corresponding numbers and sequential number of analyses indicate mineral pairs Sample

Van 10

Assemblage Anal n8

Van 27

chl-chd

Van 29

chd-chl-phg 5

40

44

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O

24.57 0.09 41.06 23.28 0.99 3.42 0.10 0.00 0.01

24.50 0.02 39.39 25.86 0.39 3.38 0.00 0.00 0.08

24.63 0.19 40.91 24.69 0.21 2.86 0.06 0.00 0.02

25.12 0.00 41.48 24.54 0.00 3.10 0.01 0.01 0.03

Sum

93.51

93.61

93.55

94.29

Chd-chl

Sample

gt-chl-chd-phg

Assemblage

gt-chl-chd-phg

54

56

anal n:

47

49

24.37 0.23 39.87 25.09 0.27 2.51 0.07 0.00 0.00

24.65 0.02 41.54 23.99 0.00 2.84 0.00 0.02 0.00

24.76 0.04 41.61 25.27 0.01 3.07 0.00 0.01 0.02

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O

36.72 0.15 21.16 33.78 3.08 0.73 6.08 0.01 0.02

37.49 0.00 21.37 35.73 0.00 2.07 4.91 0.01 0.00

92.41

93.06

94.77

Sum

101.74

101.58

233

Structural formula based on 12 oxygens

Structural formula based on 6 oxygens Si Ti Al Fe3þ Fe2þ Mn Mg Ca Na K TChd/Chl 8C (Vidal_al99)

1.99 0.01 3.93 0.07 1.58 0.07 0.41 0.01 0.00 0.00 416.8

2.00 0.00 3.79 0.21 1.77 0.03 0.41 0.00 0.00 0.01 392.5

2.01 0.01 3.94 0.06 1.69 0.01 0.35 0.00 0.00 0.00 466.2

2.03 0.00 3.96 0.04 1.66 0.00 0.37 0.00 0.00 0.00 488.8

Van 36

2.03 0.01 3.91 0.09 1.75 0.02 0.31 0.01 0.00 0.00 no equil. (750)

2.02 0.00 4.01 0.00 1.64 0.00 0.35 0.00 0.00 0.00 538.1

1.99 0.00 3.93 0.07 1.70 0.00 0.37 0.00 0.00 0.00 533.2

Si Ti Al FeT Mn Mg Ca Na K

2.95 0.01 2.00 2.27 0.21 0.09 0.52 0.00 0.00

2.98 0.00 2.00 2.37 0.00 0.25 0.42 0.00 0.00

¨ NSLI ET AL. R. OBERHA

4

Van 36

Table 3. Representative electron microprobe analyses of phengite associated to chlorite of Table 2. Mineral associations are indicated; sample with corresponding numbers and sequential number of analyses indicate mineral pairs Sample Assemblage

Van 9 chl-phg

Van 10 chl-chd

Van 11 chl-phg-car

Van 12 Chl-phg

Van 27 chd-chl-phg

Van 29 Van 36 Chd-chl gt-chl-chd-phg

Van 113 chl-phg

Van 120 chl-phg

63

65

84

85

42

50

28

37

38

45

231

61

63

61

142

182

192

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O F Sum

26.15 0.02 20.90 25.45 0.24 14.25 0.04 0.01 0.02 0.14 87.22

26.71 0.00 21.24 25.84 0.39 15.02 0.00 0.03 0.00 0.02 89.26

26.66 0.00 24.53 17.80 0.17 18.63 0.03 0.00 0.00 0.05 87.87

26.96 0.00 23.89 17.47 0.15 19.16 0.11 0.01 0.00 0.25 88.00

25.19 0.00 24.88 17.74 0.09 17.81 0.05 0.04 0.01 0.07 85.88

25.65 0.00 25.44 17.73 0.18 18.32 0.05 0.01 0.04 0.37 87.78

25.18 0.06 20.72 29.36 0.25 12.51 0.04 0.00 0.03 0.02 88.17

25.58 0.00 21.02 27.66 0.07 13.76 0.05 0.05 0.10 0.00 88.28

25.72 0.00 23.50 23.85 0.06 15.67 0.07 0.00 0.04 0.59 89.50

28.21 0.09 26.48 19.47 0.05 12.77 0.01 0.00 0.01 0.17 87.25

23.64 0.22 22.03 32.54 0.10 8.92 0.07 0.00 0.05 0.17 87.75

24.24 0.04 23.47 26.46 0.17 13.58 0.00 0.01 0.01 0.02 87.99

25.55 0.07 21.70 26.52 0.08 14.73 0.00 0.04 0.00 0.17 88.85

24.64 0.07 21.95 30.33 0.14 9.88 0.00 0.02 0.13 0.00 87.15

24.83 0.02 22.07 29.79 0.09 10.68 0.03 0.02 0.02 0.00 87.55

27.44 0.02 21.03 21.43 0.19 17.02 0.01 0.02 0.00 0.33 87.49

27.08 0.02 21.22 21.22 0.16 17.45 0.01 0.00 0.00 0.00 87.17

Si 2.77 2.76 2.68 2.70 2.59 2.59 2.70 2.71 2.64 2.85 Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 Al 2.61 2.59 2.90 2.82 3.02 3.03 2.62 2.62 2.84 3.15 Fe 2.26 2.24 1.49 1.47 1.53 1.50 2.63 2.45 2.05 1.64 Mn 0.02 0.03 0.01 0.01 0.01 0.02 0.02 0.01 0.01 0.00 Mg 2.25 2.32 2.79 2.86 2.73 2.76 2.00 2.17 2.40 1.92 Ca 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.01 0.01 0.00 Na 0.00 0.01 0.00 0.00 0.01 0.00 0.00 0.01 0.00 0.00 K 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.00 F 0.09 0.02 0.03 0.16 0.05 0.23 0.02 0.00 0.38 0.11 TChl 8C (Jowett91) 338.3 341.3 357.6 364.8 390.8 392.1 362.4 359.7 379.9 387.9

2.60 0.02 2.86 2.99 0.01 1.46 0.01 0.00 0.01 0.12 392.9

Structural formula based on 14 oxygens 2.56 2.67 2.68 2.68 2.83 2.80 0.00 0.01 0.01 0.00 0.00 0.00 2.92 2.68 2.82 2.81 2.56 2.58 2.34 2.32 2.76 2.69 1.85 1.83 0.01 0.01 0.01 0.01 0.02 0.01 2.14 2.30 1.60 1.72 2.62 2.69 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 0.02 0.11 0.00 0.00 0.21 0.00 406.0 368.6 369.0 370.8 315.8 327.7

HP EVOLUTION OF THE BITLIS COMPLEX

Anal n:

473

474

Table 4. Representative electron microprobe analyses of chloritoid and garnet. For chloritoid associated with chlorite (see Table 2) temperature were estimated following the method of Vidal et al. (1999). Mineral associations are indicated; sample with corresponding numbers and sequential number of analyses indicate mineral pairs Sample Assemblage

Van 9 chl-phg

Van 11 chl-phg-car

Van 12 Chl-phg

Van 27 chd-chl-phg

Van 36 gt-chl-chd-phg

Van 113 chl-phg

Van 120 chl-phg

62

64

43

51

27

36

39

47

64

64

60

143

183

191

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O F Sum

50.16 0.11 29.99 3.90 0.03 2.20 0.15 0.31 9.31 0.24 96.15

49.72 0.11 29.79 4.13 0.00 2.06 0.00 0.34 9.61 0.15 95.75

47.06 0.02 36.07 2.14 0.08 0.48 0.04 0.99 7.48 0.12 94.35

47.26 0.08 37.07 1.83 0.11 0.47 0.16 1.15 7.72 0.18 95.85

49.05 0.26 28.33 4.66 0.00 2.24 0.10 0.17 10.70 0.00 95.50

48.56 0.26 28.56 4.70 0.00 1.93 0.00 0.27 10.88 0.00 95.16

45.92 0.05 29.23 19.02 0.12 2.17 0.03 0.01 0.00 0.03 96.55

47.79 0.13 35.92 1.82 0.07 0.89 0.06 1.04 8.88 0.39 96.60

46.98 0.11 35.84 1.73 0.00 0.65 0.00 1.36 9.26 0.00 95.93

46.98 0.11 35.84 1.73 0.00 0.65 0.00 1.36 9.26 0.00 95.93

46.50 0.14 31.60 3.82 0.04 1.54 0.02 0.52 7.88 0.00 92.06

47.93 0.17 35.78 1.05 0.01 0.72 0.01 0.47 8.75 0.00 94.88

49.27 0.03 34.29 2.07 0.00 0.96 0.00 3.26 3.96 0.00 93.84

47.85 0.06 39.00 1.05 0.00 0.24 0.01 5.52 0.84 0.13 94.57

3.11 0.00 2.81 0.12 0.00 0.05 0.00 0.13 0.63 0.05

3.08 0.00 2.85 0.10 0.01 0.05 0.01 0.15 0.64 0.07

3.31 0.01 2.26 0.26 0.00 0.22 0.01 0.02 0.92 0.00

3.30 0.01 2.29 0.27 0.00 0.20 0.00 0.04 0.94 0.00

3.10 0.00 2.33 1.07 0.01 0.22 0.00 0.00 0.00 0.01

3.11 0.01 2.76 0.10 0.00 0.09 0.00 0.13 0.74 0.16

3.09 0.01 2.78 0.10 0.00 0.06 0.00 0.17 0.78 0.00

3.09 0.01 2.78 0.10 0.00 0.06 0.00 0.17 0.78 0.00

3.19 0.01 2.56 0.22 0.00 0.16 0.00 0.07 0.69 0.00

3.15 0.01 2.77 0.06 0.00 0.07 0.00 0.06 0.73 0.00

3.22 0.00 2.65 0.11 0.00 0.09 0.00 0.41 0.33 0.00

3.06 0.00 2.94 0.06 0.00 0.02 0.00 0.68 0.07 0.05

Structural formula based on 11 oxygens Si Ti Al FeT Mn Mg Ca Na K F

3.32 0.01 2.34 0.22 0.00 0.22 0.01 0.04 0.39 0.10

3.31 0.01 2.34 0.23 0.00 0.20 0.00 0.04 0.82 0.06

¨ NSLI ET AL. R. OBERHA

Anal n:

HP EVOLUTION OF THE BITLIS COMPLEX

475

Fig. 8. Pressure temperature diagram for Al-rich carpholite bearing metapelites (after Oberha¨nsli et al. 1995; Bousquet et al. 2008) compiling the data for the Bitlis metapelites. 1, Prograde assemblages with pyrophyllite relicts; 2, peak assemblages with carpholite and carpholite-chloritoid; 3, retrograde assemblages with chloritoid chlorite garnet and kyanite. The inferred retrograde paths range from isothermal decompression to moderate heating during decompression (see text).

with grain sizes in the range of 30– 80 microns were hand picked and carefully washed by ultrasonic treatment in acetone, ethanol and distilled water. The samples were irradiated in the FRG-1 facility of the research reactor in Geesthacht (Germany). The neutron flux variation over the length of the sample capsule was monitored by Fish Canyon Tuff Sanidine (FC-3, 27.5 Ma; Ishizuka 1998; Uto et al. 1997) and calculated using a linear fit. Interference correction factors were obtained by analysing CaF2 and K2SO4 irradiated together with the samples. After irradiation the samples were loaded on a copper disc in the sample chamber. The system was baked for two days for reduction of atmospheric

argon contamination. Mean blank values during the experiments for 40Ar, 39Ar, 37Ar and 36Ar were 1.46e-4, 7.32e-08, 8.95e-09 and 4.35e-06 respectively. During the experiment Ar was extracted from the samples using a 50 W CO2 laser and isotopes were measured in a Micromass5400 spectrometer. Age spectra were produced from 3 respectively 7 grains that were heated by a moving beam with a diameter of 1600 microns for 90 seconds with 50 microseconds scan sped. Data have been corrected for blank, mass discrimination, 37Ar and 39Ar decay. They have been fitted on 36Ar/40Ar v. 39Ar/40Ar isochron plots (York 1969). Results are presented in Table 5 and Figure 9.

476

Table 5. 40Ar/39Ar data for samples analysed at the University of Potsdam. The uncertainties on the total-gas ages include the uncertainty in the irradiation J parameter and are reported on 2s level. The individual gas fractions are reported with analytical uncertainty on 1s level, including the uncertainty on the J-value Laser output (W)

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar

40

K/Ca

40

39

448,76 + 28,29 27,43 + 0,72 5,15 + 0,11 4,03 + 0,10 2,67 + 0,04 2,35 + 0,05 3,90 + 0,09 6,98 + 0,48 55,82 + 1,24

0,21 10,16 34,16 45,33 99,98 90,49 66,45 10,76 8,02

4,89 66,43 93,00 95,13 96,79 97,22 95,53 92,36 60,40

0,06 2,77 9,33 12,39 27,35 24,76 18,19 2,95 2,20

6,82 + 356,18 16,02 + 7,55 20,20 + 2,25 23,25 + 1,70 23,77 + 0,77 24,31 + 0,86 24,59 + 1,16 24,85 + 7,19 25,14 + 9,66

21,59 + 1120,66 50,32 + 23,39 63,21 + 6,93 72,55 + 5,22 74,15 + 2,38 75,79 + 2,63 76,65 + 3,57 77,44 + 21,93 78,34 + 29,44

185,37 + 1,59 56,58 + 0,56 17,27 + 0,18 6,83 + 0,13 5,26 + 0,10 5,20 + 0,20 4,92 + 0,14 10,71 + 0,53 111,03 + 1,97

7,83 24,95 36,40 51,07 38,88 18,86 22,60 7,42 4,92

23,50 54,61 81,94 92,23 93,92 94,16 94,35 88,63 44,79

3,68 11,73 17,12 23,80 18,31 8,89 10,66 3,50 2,32

16,82 + 9,81 20,11 + 3,09 23,15 + 2,12 23,95 + 1,52 24,02 + 1,99 24,71 + 4,10 24,23 + 3,42 24,58 + 10,42 26,61 + 15,78

52,79 + 30,33 62,94 + 9,50 72,24 + 6,50 74,68 + 4,65 74,91 + 6,08 77,01 + 12,52 75,54 + 10,45 76,63 + 31,81 82,81 + 48,00

Ar*

ArK

Ar*/39ArK

Age (+1s) Ma

Van 75, phengitic mica J ¼ 0.00177

Van 75A, phengitic mica J ¼ 0.00177 0,014 71,59 + 0,37 0,08 + 75,13 0,016 36,83 + 0,10 0,02 + 23,58 0,018 28,25 + 0,08 0,02 + 16,16 0,020 25,96 + 0,16 0,01 + 11,52 0,022 25,57 + 0,08 0,02 + 15,13 0,024 26,24 + 0,09 0,03 + 31,19 0,026 25,68 + 0,05 0,03 + 26,03 0,028 27,74 + 0,22 0,08 + 79,27 t.f.* 59,40 + 0,33 0,12 + 119,66 Plateau age: 74.4 + 2.8 Ma; total gas age: 72.7 + 3.4 Ma *t.f., total fusion.

¨ NSLI ET AL. R. OBERHA

0,012 139,06 + 6,27 2,74 + 2740,20 0,014 24,12 + 0,22 0,06 + 57,91 0,016 21,72 + 0,03 0,02 + 17,22 0,018 24,44 + 0,05 0,01 + 12,98 0,020 24,56 + 0,05 0,01 + 5,88 0,022 25,00 + 0,04 0,01 + 6,50 0,024 25,74 + 0,03 0,01 + 8,85 0,026 26,90 + 0,16 0,05 + 54,66 t.f. 41,63 + 0,24 0,07 + 73,35 Plateau age: 74.5 + 1.5 Ma; total gas age: 73.3 + 2 Ma

HP EVOLUTION OF THE BITLIS COMPLEX

40

Ar/39Ar phengite dating results from two carpholite bearing rocks of the metasedimentary cover of the Bitlis massif near Gevas.

477

Fig. 9.

478

¨ NSLI ET AL. R. OBERHA

Both samples are from the northernmost carpholite locality near Gevas. Both age spectra show a first step that is severely contaminated by atmospheric argon. Five respectively six gas fractions yield concordant apparent ages, which define plateau dates of 74.5 + 1.5 Ma, and 74.4 + 2.8 Ma. Isochron ages are similar to the plateau ages with intercept ages of 73.8 + 7.7 Ma and 73.6 + 4.4 Ma respectively (Fig. 9). Excess argon may hamper the interpretation of 40 Ar/39Ar phengite ages subjected to very highpressure conditions (e.g. Li et al. 1994; Arnaud & Kelly 1995; Ruffet et al. 1995). Strongly deformed, K-poor bulk compositions at low high-pressure conditions close to closure temperatures (350 + 450 8C; Andriessen 1991) are barely suitable to incorporate excess argon in phengites (Oberha¨nsli et al. 1998; Sherlok & Kelly 2002). The Late Cretaceous age for the blueschist metamorphism in the Bitlis complex is compatible with the geological constraints as well as observations from the Lesser Caucasus, where H-P metamorphism is dated at 95-90 Ma (Rolland 2008). It is also younger than the H-P metamorphism of the Tavsanli zone in western Anatolia (e.g. Okay & Kelly 1994; Sherlok et al. 1999) but fits the age (K –Ar; 71.2 + 3.6 Ma; Hempton 1985) of metamorphism from the Pu¨tu¨rge massif.

Discussion As shown in Figure 6, HP-LT metamorphism is distributed overall in the cover sequence of the eastern Bitlis complex. Along the C ¸ atak River, the marbles contain fresh carpholite without chloritoid. This clearly proves that the frontal part at the base of the western Bitlis complex experienced HP-LT metamorphism and that the temperatures never exceeded 450 8C since carpholite remained stable. On the contrary, the northern portion of the basal thrust experienced a slight heating after the HP-LT overprint, as attested by the reaction of carpholite retrogressed into chloritoid and quartz. Figure 8, a petrogenetic grid for Al-rich carpholite bearing metapelites, evidences low temperatures at elevated pressures for samples from the area where pyrophyllite was found. For samples with carpholite and carpholite relicts higher temperatures at high pressures are documented. Chloritoid bearing samples with carpholite relicts in quartz indicate similar conditions. Chloritoid samples without carpholite relicts indicate a wider range of temperatures at lower pressures. The stability of kyanite together with chloritoid indicates temperatures below 560 8C at 11 kb or 480 8C at 5 kb: the characteristic reaction, for Al-rich metapelites, chloritoid þ kyanite ¼ chlorite þ staurolite (Spear & Cheney 1989) was

never overstepped (Fig. 8). Garnet and epidote indicate decompression (Bousquet 2008). Form these observations we can conclude, that the retrogression from high-pressure low-temperature took place under conditions of isothermal decompression or at only a slightly elevated temperature conditions. The recorded temperature in metamorphic rocks of the Bitlis complex never exceeded 450 8C during the Mesozoic and Cenozoic evolution. Along its northern contact of the Bitlis complex, Yılmaz (1978) described the Gevas complex as an ophiolite that was thrust over the Bitlis complex. Our investigations show that the Gevas complex is a me´lange with a serpentinic matrix containing radiolarites and limestone blocks. The limestone blocks exhibit a rudist fauna with Arabic facies affinity, which is different from the rudist faunas ¨ zer 2005). The metamorphic of the Taurids (O sequence does not represent a metamorphic sole as inferred by Yılmaz (1978) but contains either relics of carpholite fibres or glaucophane. This points to a low-grade high-pressure metamorphism typical for cold geotherms only present in subduction related settings. Our investigation clearly shows that the Bitlis complex experienced a late Alpine subduction related history while the Gevas ophiolite material does not show any metamorphic overprint. Therefore the contact must be interpreted as late back thrust of the Bitlis complex towards the North. The faunistic finding of Arabian facies affinity in the limestone blocks of the me´lange contradicts the hypothesis of an obducted ophiolite block of northern provenance. In addition, HP-LT metamorphic conditions (450 8C) evidenced in the Bitlis complex but not in the Gevas ophiolitic me´lange exclude obduction. It is obvious from petrography that the Bitlis complex and some Eocene formations experienced a subduction event and remained cold during its later geodynamic evolution. These facts were not considered in geodynamic evolution schemes published earlier (Yılmaz 1993; Sengo¨r et al. 2003; Keskin 2003). Most of these scenarios do not consider the metamorphic evolution of the Bitlis complex at all. Based on Cenozoic sediment evolution south of the Bitlis complex Yılmaz (1993) assumes during Late Maastrichtian to Early Eocene an intraoceanic subduction between a northern block and the Arabian Plate. This model accounts for Eocene to Oligocene subduction south of the Bitlis complex but does not detail the metamorphic evolution neither in the Bitlis complex not in the underlying Cenozoic nappes. Timing of the geodynamic evolution south of the Bitlis complex is well constrained in this model. However, geodynamically the nappe stacking of the ‘metamorphic massifs’ during Early Eocene is not well constrained (Yılmaz

HP EVOLUTION OF THE BITLIS COMPLEX

1993, fig. 13). Later models (e.g. Sengo¨r et al. 2003; Keskin 2003) focus on the geodynamic and volcanic evolution north of the Bitlis complex. While Sengo¨r (2003) reflects on the tectonic building of the East Anatolian high plateau, Keskin (2003) focuses on its volcanic and magmatic evolution. In Sengo¨rs model (Sengo¨r et al. 2003, fig. 3) the Bitlis complex is thrust over the Arabian platform some time between Late Eocene and Middle Miocene without any clear geodynamic reason. In Keskin’s model (Keskin 2003, fig. 3) the Bitlis complex is exhumed before Early Eocene and is part of a volcanic arc. The collision of the Arabian plate with the Bitlis Arc terminated during Late Oligocene – Early Miocene, while closure of the East Anatolian Accretion complex continued. At 11 –13 Ma slab brake off followed since 6 Ma by asthenospheric upwelling is postulated (Sengo¨r et al. 2003; Keskin 2003). New geophysical observations in Eastern Anatolia infer seismic velocities smaller than expected for lithospheric mantle (Zor et al. 2003; Go¨k et al.

479

2007). Thus similarly to earlier models asthenospheric up welling following a slab break off event is assumed (e.g. Facenna et al. 2006; Barazangi et al. 2006). From seismic data (Go¨k et al. 2007) infers north directed major thrusts that fit well with our observations along the Gevas me´lange. To respect the metamorphic evolution and especially the preservation of HP-LT assemblages we propose a scenario (Fig. 10) that accounts for the Late Cretaceous (c. 74 Ma) metamorphic evolution. The Bitlis complex is routed northwards below the South-Armenian Block while its frontal parts are thrust southward over the Arabian platform and the Cenozoic complexes. Investigations of the Sevan –Akera suture zone in the Lesser Caucasus (Sosson et al. 2010) and new finding of HP assemblages along the ophiolitic suture near Stepanavan (Rolland et al. 2008) and its correlation to the Izmir Ankara Erzincan suture led us to assume that the Bitlis block collided with the South Armenian block during the Late Cretaceous.

Fig. 10. Schematic geodynamic cross-section including data from ‘MEBE Caucasus Group’ by Sosson et al. (2010). In order to maintain cool conditions we consider a strong underthrust of the Arabian platform and separation of areas with asthenospheric upwelling for the Bitlis complex (see text).

480

¨ NSLI ET AL. R. OBERHA

This implies that the Bitlis complex was part of the Arabian platform and cannot be strictly correlated with the Tauride block. It separated from Arabia probably during Jurassic to Cretaceous time. After northward subduction the Bitlis complex had to be exhumed rapidly (supported by pyrophyllite relicts) probably already during Late Cretacous, since the Bitlis metamorphic units are imbricated with non-metamorphic Eocene pillow lava in the frontal part of the nappe complex (Fig. 1). To preserve its subtle HP-LT phases later significant heating must be excluded. This is possible if exhumation is rapid and the Bitlis complex further on remained close to the surface or if subduction processes are ongoing in the south as it is documented by blueschist assemblages in the Eocene Urse formation. The non-metamorphic ophiolitic me´langes of the Yu¨ksekova complex derive from the oceanic realm between the South Armenian and Bitlis blocks and were thrust over the exhuming Bitlis complex (Fig. 10). After collision of the Arabian plate with the Bitlis complex back thrusting led to the exhumation of basement rocks and the northern part of the Bitlis complex along Lake Van and the Yu¨ksekova complex in Gevas. Our Interpretation with an Oligocene subduction of a mid oceanic ridge allows for later asthenospheric up welling without the need of subduction rollback and slab break off.

Conclusions Petrological investigations in the eastern Bitlis complex clearly demonstrate subduction related HP-LT metamorphic conditions. These findings must be taken into consideration when reconstructing the geodynamic evolution of eastern Anatolia in front of the Arabian indenter, and especially for the delamination processes of the South Armenian block and its relation with recent volcanism. A first and obvious result is that the Gevas complex should not be considered as a complete ophiolite sequence but rather composes a serpentinitic me´lange similar the Yu¨ksekova complex. Now overturned, it overlays the Bitlis complex which exhibits relics of HP-LT metamorphism. Faunistic and petrographic investigations support an Arabian/Gondwanian origin for the Bitlis complex rather than a Tauride provenance. The fact that HP-LT parageneses are distributed over the whole of the Bitlis complex demonstrates that this complex experienced a subduction event and remained cold during its later geodynamic evolution. Geophysical data point to material with lower seismic velocities, interpreted as hot asthenospheric mantle situated just north of the Bitlis complex. Despite this, subtle HP-LT metamorphic

assemblages are preserved indicating that after the subduction event no significant rise of temperatures has occurred. The Bitlis complex is part of the complex Alpine belt with a collage of terranes and clearly more than one subduction zone. The blueschist metamorphism (c. 75 Ma) found distributed within the Bitlis complex points to a second subduction zone south of the major Izmir–Ankara –Erzincan– Sevan– Akera suture with a metamorphic H-P age of c. 90 Ma. The finding of blue amphibole and mica with in the Eocene me´langes south of the nonmetamorphic Yu¨ksekova ophiolitic me´lange points to an even younger subduction type feature south of the Bitlis complex. Thus the region forming a high plateau in it status nascendii in eastern Anatolia composes a set of Gondwana derived blocks separated by oceanic domains that successively collided to with Eurasia. We thank M. F. Brunet, E. Barrier and J. P. Cadet, all part of the MEBE team in Paris for their technical and financial support. M. Sudo and M. Timmerman supported work in the Ar lab. M. Sosson and J. Mosar inspired us in many discussions. The reviews by R. Klemd and Y. Rolland are highly acknowledged.

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HP EVOLUTION OF THE BITLIS COMPLEX ¨ nal, R. N., Sengu¨n, M. & YurtsC ¸ ailayan, M. A., O ever, A. 1984. Structural setting of the Bitlis Massif. In: Tekeli, O. & Go¨ncu¨oglu, M. C. (eds) Geology of the Taurus Belt. Proceedings of the International Symposium on the Geology of the Taurus Belt, Ankara, 245–254. Candan, O., C¸etinkaplan, M., Oberha¨nsli, R., Rimmele´, G. & Akal, C. 2005. Alpine high-P/ low-T metamorphism of the Afyon Zone and implications for the metamorphic evolution of Western Anatolia. Lithos, 84, 102– 124. Faccenna, C., Bellier, O., Martinod, J., Piromallo, C. & Regard, V. 2006. Slab detachment beneath eastern Anatolia: a possible cause for the formation of the North Anatolian fault. Earth and Planetary Science Letters, 242, 85–97. Go¨k, R., Pasyanos, M. E. & Zor, E. 2007. Lithospheric structure of the continent –continent collision zone: eastern Turkey. Geophysical Journal International, 169, 1079–1088. Go¨k, R., Sandvol, E., Turkelli, N., Seber, D. & Barazangi, M. 2003. Sn attenuation in the Anatolian and Iranian plateau and surrounding regions. Geophysical Research Letters, 30, 8042. Go¨ncu¨oglu, M. C. 1984. Mus-Kızılagac¸ metagranitinin baikalasımı ve yası. Maden Tektik ve Arama dergisi, 99– 100, 72– 83. Go¨ncu¨oglu, M. C. & Turhan, N. 1984. Geology of the Bitlis metamorphic belt. In: Tekeli, O. & Go¨ncu¨oglu, M. C. (eds) Geology of the Taurus Belt. Proccedings of the International Symposium on the Geology of the Taurus Belt, Ankara, 237– 244. Go¨ncu¨oglu, M. C. & Turhan, N. 1992. Mus – I 33 paftası. 100 000 scale geological map series of Turkey. Maden Tetkik ve Arama Genel Mu¨du¨rlu¨gu¨, Ankara. Go¨ncu¨oglu, M. C. & Turhan, N. 1997. Rock units and metamorphism of the basement and Lower Paleozoic cover of the Bitlis Metamorphic complex, SE Turkey. In: Go¨ncu¨oglu, M. C. & Derman, A. S. (eds) Lower Paleozoic Evolution in Northwest Gondwana. Turkish Association Petroleum Geology, Special Publication, 3, 75–81, Ankara. Go¨ncu¨og˘lu, M. C., Dirik, K. & Kozlu, H. 1997. General Characteristics of pre-Alpine and Alpine Terranes in Turkey: explanatory notes to the terrane map of Turkey. Annales Geologique de Pays Hellenique, 37, 515–536, Geological Society, Greece, Athens. Hall, R. 1976. Ophiolite emplacement and the evolution of the Taurus suture zone, southeastern Turkey. Geological Society, America, Bulletin, 87, 1078– 1088. Hall, R. & Mason, R. 1972. A tectonic me´lange from the Eastern Taurus Mountains, Turkey. Journal of the Geological Society, London, 128, 395– 397. Helvacı, C. & Griffin, W. L. 1984. Rb-Sr geochronology of the Bitlis Massif, Avnik (Bingo¨l) area, S. E. Turkey. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 403–413. Hempton, M. R. 1985. Structure and deformation history of the Bitlis suture near Lake Hazar, southeastern Turkey. Geological Society, America, Bulletin, 96, 233–243.

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HP EVOLUTION OF THE BITLIS COMPLEX me´ridionale du massif de Bitlis (Turquie). Sciences Ge´ologiques Bulletin, Strasbourg, 34, 59– 73. York, D. 1969. Least square fitting of a straight line with correlated errors. Earth and Planetary Science Letters, 5, 320 –324.

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Zor, E., Sandvol, E., Gurbuz, C., Turkelli, N., Seber, D. & Barazangi, M. 2003. The crustal structure of the East Anatolian plateau (Turkey) from receiver functions. Geophysical Research Letters, 30, 8044.

Seismic images of crustal variations beneath the East Anatolian Plateau (Turkey) from teleseismic receiver functions ¨ ZACAR1*, GEORGE ZANDT2, HERSH GILBERT3 & SUSAN L. BECK2 A. ARDA O 1

Department of Geological Engineering, Middle East Technical University, 06531-Ankara, Turkey

2

Department of Geosciences, University of Arizona, Gould-Simpson Building, 1040E 4th Street, Tucson, AZ 85721, USA

3

Department of Earth and Atmospheric Sciences, Purdue University, 550 Stadium Mall Drive, West Lafayette, IN 47907, USA *Corresponding author (e-mail: [email protected]) Abstract: We used teleseismic P-wave receiver functions recorded by the Eastern Turkey Seismic Experiment to determine the crustal structure across an active continent– continent collision zone. Moho depth and Vp/Vs variations in the region are mapped by incorporating crustal multiples and later two-dimsional (2-D) seismic profiles are produced using a common conversion point technique with our crustal Vp/Vs estimates. Moho depths do not correlate with surface topography and reveal a relatively thin crust consistent with the high plateau being supported by hot asthenosphere near the base of the crust. Under the Arabian plate, the crust is thinnest (c. 35 km) and exhibits high Vp/Vs (1.8) associated with mafic compositions. In the east, the crust gradually becomes thicker towards the north and exceeds 45 km in the northeastern side whereas in the west, the crust thickens sharply near the Bitlis suture and displays pronounced Moho topography within the Anatolian plate that suggests the presence of multiple fragments. Vp/Vs variations show an anomalously high Vp/Vs corridor (1.85) along the North Anatolian Fault and near the youngest volcanic units (c. 3 Ma) and support the presence of partial melt. This corridor is spatially limited from both north and south by low Vp/Vs regions implying a change in crustal composition. Near the Bitlis suture, a layered Vp/Vs model points to the source of low Vp/Vs in the lower crust that may be rich in quartz. Furthermore, the seismic profiles indicate a prominent low velocity zone in the lower crust across a large area beneath the plateau that may act as a decoupling zone between the crust and upper mantle.

The tectonic framework of Eastern Turkey is characterized by active continent– continent collision produced by the northward motion of the Arabian plate that is separated from the African plate by the left-lateral Dead Sea Fault (DSF) (Fig. 1). This collision along the Bitlis suture (BS) has resulted in a topographic high (the East Anatolian Plateau) and is partially responsible for the extrusion of the Anatolian plate westward forming the right-lateral North Anatolian Fault (NAF) and the left-lateral East Anatolian Fault (EAF) in the early Pliocene (McKenzie 1978; Sengo¨r & Yılmaz 1981; Bozkurt 2001). The East Anatolian Plateau which exhibits widespread Neogene and Quaternary volcanism over nearly half of the region and has an average elevation of c. 2 km above sea level is considered a part of the Alpine –Himalayan System (Keskin 2003). The Eastern Turkey Seismic Experiment (ETSE), which was conducted across the region in 1999–2001, produced a large high-quality

broadband seismic data set (Sandvol et al. 2003a). Overall results from this experiment suggested that the subducted Arabian plate must have detached allowing asthenosphere near the base of the crust with little or no mantle lithosphere (Sandvol et al. 2003a). Hence the 2 km high plateau is not supported by thick crust but rather by hot mantle (Sengo¨r et al. 2003). In the region, receiver functions provide the only constraint for crustal thickness. Previously, Zor et al. (2003) modeled P-wave receiver functions using constant Vp/Vs (1.73) and found that the Moho depths range between 40 and 50 km beneath the plateau. Considering the sensitivity of receiver functions to Vp/Vs that can vary significantly, these depths have potentially large errors (Zhu & Kanamori 2000). In this paper, we used teleseismic P-wave receiver functions recorded by ETSE and incorporated crustal multiples following the method of Zhu & Kanamori (2000) to more accurately map out Moho depths and Vp/Vs variations

From: Sosson, M., Kaymakci, N., Stephenson, R. A., Bergerat, F. & Starostenko, V. (eds) Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340, 485–496. DOI: 10.1144/SP340.21 0305-8719/10/$15.00 # The Geological Society of London 2010.

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GC

Black Sea

N

LC

42º 41º

NAF PS Anatolian Plate

Aegean Sea

43º

EA

F

East Anatolian Plateau BS

40º 39º 38º

DSF

37º 36º

Arabian Plate

35º

Mediterranean Sea 24º

26º

28º

30º

32º

34º

36º

38º

40º

42º

44º

34º 46º

Fig. 1. A simplified tectonic map of Turkey with topographic relief. Arrows indicate the direction of plate and fault motions. NAF, North Anatolian Fault; EAF, East Anatolian Fault; DSF, Dead Sea Fault; BS, Bitlis Suture; PS, Pontide Suture; LC, Lesser Caucasus; GC, Greater Caucasus.

beneath the plateau. Together, these observations are used to make better seismic images of the crust and provide additional constraints on composition and rheology that are essential to characterize the crustal structure in relation to the geology and tectonic evolution of the region.

East Anatolian Plateau The East Anatolian Plateau is made up of continental fragments accreted during the northward subduction of multiple branches of the Neotethyan Oceanic lithosphere since the Mesozoic (Sengo¨r & Yılmaz 1981; Sengo¨r et al. 2003; Barazangi et al. 2006). During the Eocene, the Arabian platform collided with the Bitlis-Poturge Massif, which is the easternmost part of the Menderes-Taurus block located near the BS (Yılmaz 1993). However, northward subduction beneath Eastern Turkey continued throughout the middle Miocene and resulted in the development of the Eastern Rhodope-Pontide magmatic arc in the north and the East Anatolian Accretionary Complex (EAAC) that covers a large area of the plateau (Sengo¨r & Yılmaz 1981; Yılmaz et al. 1997; Bozkurt & Mittwede 2001; Sengo¨r et al. 2003). Subsequent to the final continental collision and suturing, the plateau started its rapid block uplift (Gelati 1975) accompanied by widespread volcanism starting at about 11 Ma (Keskin 2003) (Fig. 2). In the Early Pliocene, ongoing convergence resulted in the westward

escape of the Anatolian plate along the NAF and the EAF (Bozkurt 2001; Kocyigit et al. 2001). In the region, there are no subcrustal earthquakes, and crustal seismicity is predominantly strike –slip in nature suggesting that the collision is accommodated by tectonic escape (Turkelli et al. 2003; Orgulu et al. 2003). The crustal thickness obtained by P-wave receiver functions reveal a gradual increase from 40 km in the south beneath the Arabian foreland to 50 km in the north beneath the Pontides mountain range (Zor et al. 2003) and on average is thinner than the 55 km crust proposed by Sengo¨r (1980) on the basis of normal isostatic compensation. Tomographic images of the upper mantle (Piromallo & Morelli 2003; Hafkenscheid et al. 2006; Lei & Zhao 2007; Zor 2008) indicate slab-like fast-velocity anomalies that are interrupted from above by a widespread slow-velocity anomaly beneath the EAAC, and suggest slab detachment and concomitant emplacement of the hot asthenosphere at subcrustal depths. Sandvol et al. (2003b) found large shear-wave splitting delay times with fast polarization directions orientated uniformly in the NE–SW direction beneath the region that is not consistent with the surface deformation and suggested that the mantle flow is mostly asthenospheric. The plateau has little or no Sn propagation across it (Gok et al. 2003) and low shear wave (Maggi & Priestley 2005) and Pn velocities (AlLazki et al. 2003; Gans et al. 2009) in the uppermost mantle, which are consistent with the anomalously

CRUSTAL VARIATIONS IN EAST ANATOLIA

487

Fig. 2. Regional tectonic map of Eastern Turkey with topographic relief, station locations of ETSE array (inverted white triangles). Neogene volcanics (pink area) and Holocene volcanoes (red triangles). Arrows indicate the direction of plate and fault motions. NAF, North Anatolian Fault; EAF, East Anatolian Fault; DSF, Dead Sea Fault; MOFZ, Malatya-Ovacık Fault Zone; EAAC, East Anatolian Accretionary Complex; BS, Bitlis Suture; PS, Pontide Suture. Stations mentioned in the text are also labelled.

thin (60–80 km) lithosphere identified by receiver functions (Angus et al. 2006; Ozacar et al. 2008). Furthermore, the presence of asthenospheric material at shallow depths is also supported by observed Bouguer gravity anomalies (Ates et al. 1999) and anomalously low resistivity in the upper mantle identified by a magnetotelluric study (Turkoglu et al. 2008). Based on such findings, Sengo¨r et al. (2003) suggested that the steepening and breakoff of a northward subducted slab that belongs to the northern branch of Neo Tethys allowed hot partially molten asthenospheric material to be emplaced near the base of the crust beneath the EAAC. Slab detachment likely occurred around 11 Ma ago when the rapid uplift and the oldest collision related magmatism started (Keskin 2003; Sengo¨r et al. 2003). Recently, Barazangi et al. (2006) interpreted the southern branch of Neo Tethys as the detached slab. However, Keskin (2007) argues that the slab detachment in the south might be an older event (Eocene – Oligocene) based on the absence

of young volcanism (,11 Ma) along the BitlisPoturge massif.

Data and method The ETSE array consists of 29 broadband PASSCAL stations that were deployed across Eastern Turkey (Sandvol et al. 2003a). These stations were operated for a period of 22 months between November 1999 and July 2001 in an inverted triangle pattern with c. 50 km station spacing (Fig. 2). In this study, we used P-waveforms of over 200 earthquakes (M . 5.3) recorded by the ETSE with epicentral distances between 358 and 908. The abundance of teleseismic earthquakes with good epicentral distance and azimuthal coverage provides a large data set for receiver function analysis in Eastern Turkey. Figure 3 shows the global distribution of these events. Receiver functions are deconvolved waveforms that contain P to S conversions and their multiples from seismic impedance contrasts beneath

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P Ps PpPs PpSs+PsPs Ps

P wave S wave

PpPs

PpSs

PsPs

35

40

1.9

1.8

Vp/Vs

35º

Ps

1.7

PpPs PpSs + PsPs

90º

1.6 15 Fig. 3. Global distribution of teleseismic earthquakes (circles) used in the receiver function study. Star indicates the centre of ETSE array.

the recording station. The deconvolution is accomplished in the time domain through an iterative approach using a cross-correlation technique (Ligorria & Ammon 1999). In the deconvolution algorithm, we computed receiver functions with different cut-off frequencies (0.5, 1.2, 2.4 Hz) by varying the width of the Gaussian-type low-pass filter (1.0, 2.5, 5.0). For our analysis, we only used high quality receiver functions with variance reductions of over 70%, which relates to the quality of the deconvolution. An important limitation of receiver function analysis is the trade-off between thickness and crustal velocities. Sensitivity tests indicate that crustal thicknesses are much more dependent on the Vp/Vs value used than Vp (Zandt et al. 1995). Zhu & Kanamori (2000) calculated that a 0.1 km s21 change in Vp corresponds to only 0.5 km change in thickness, however, a 0.1 change in Vp/Vs can lead to about 4 km change in the crustal thickness. In this respect, it is essential to have good estimates for Vp/Vs to map the spatial variations in crustal thickness. For this purpose, crustal multiples have been used to determine the Moho depth (H ) and the Vp/Vs ratio beneath each ETSE station. We followed the stacking algorithm of Zhu & Kanamori (2000) which sums the amplitudes of receiver functions along theoretical travel–time curves and performs a grid search through H and Vp/Vs space for the largest summed amplitude at the predicted arrival times of direct and multiple converted phases. Figure 4 shows the contribution of these phases to the stacking plot. We applied the

20

25

30

Depth (km) Fig. 4. Converted phases used in the study and their contribution to stacking; taken from Zhu & Kanamori (2000).

technique to receiver functions with a cut-off frequency of 1.2 Hz using a constant crustal Vp of 6.2 km s21 and estimated uncertainties based on the variance of the summed amplitude. For visual inspection, we also plotted receiver functions in bins according to their ray parameters and overlaid the plots with the predicted arrival times for the preferred model (Fig. 5). It is worth to note that Moho depths are measured from station elevation and corresponds to the crustal thickness beneath the station. At some stations, Moho multiples were obscured by other conversions from upper mantle discontinuities causing smearing and/or other local maxima. For a horizontally stratified model, multiple phases have opposite move out with ray parameter and can be distinguished from other phases. In such cases, we used receiver functions with different cut-off frequencies and tested higher variance reduction thresholds. However, we still failed to obtain unique solutions for two stations (KARS, CMCY) that display complicated site responses (Fig. 2). Station KARS was especially complicated and displays a significantly delayed direct P phase indicating a thick sedimentary fill beneath the station.

Observations Crustal thickness and Vp/Vs variations The results of two stations (ERGN, BTLS) that have remarkably different crustal structures are

CRUSTAL VARIATIONS IN EAST ANATOLIA

489

Fig. 5. Receiver functions recorded at stations ERGN and BTLS as a function of ray parameter with predicted arrival times of direct and multiple converted phases and trade-off plots of stacked receiver function amplitudes in parameter space. Red is used for positive and blue for negative amplitudes.

shown in Figure 5. Station ERGN, which has an elevation of 922 m is located at the BS (Fig. 2). Receiver functions recorded at this station show strong midcrustal and Moho discontinuities with exceptionally clear PpPs multiples. In this case, we were able to obtain a layered model that has a low average Vp/Vs (1.68) for the crust (44 km) but a

high average Vp/Vs (1.8) for the upper crust (25 km). Indeed, this indicates a lower crust with anomalously low Vp/Vs (,1.6) and supports a quartz-rich felsic composition in the lower crust at the suture. On the other hand, station BTLS that is located at an elevation of 1722 m near the youngest volcanic rocks (c. 3 Ma) west of Lake Van (Fig. 2),

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Table 1. Moho depth (H) measured from station elevation and Vp/Vs estimates for ETSE stations Station

Latitude

Longitude

Elevation (m)

N

H (km)

Vp/Vs

AGIN AHLT BNGL BTLS BYBT BYKN CMCY DGRL DGSU DYBR ERGN EZRM HAMR HINS HRPT HRSN ILIC IMRL KARS KOTK KRLV KTLN KYPR MRDN MSDY MUSH SILN SIRN UZML

38.939 38.748 38.920 38.431 40.235 38.167 39.921 41.057 39.131 37.823 38.259 40.104 39.614 39.349 38.704 39.945 39.453 39.879 40.622 40.223 39.375 37.953 37.559 37.290 40.462 38.757 38.135 40.202 39.713

38.713 42.477 40.597 42.124 40.266 41.782 43.197 43.327 42.730 40.319 39.729 41.364 42.993 41.697 39.245 42.287 38.569 38.118 43.069 43.009 40.988 41.705 41.169 40.700 37.777 41.483 41.041 39.120 39.716

945 1738 1160 1722 1684 816 1995 2027 1646 619 922 1910 1714 1727 1428 1800 1300 1708 1904 1396 1894 791 1179 674 1152 1554 917 1496 1593

98 70 109 110 41 55 52 103 31 84 99 142 52 88 71 41 66 38 36 55 47 52 85 80 53 40 37 89 83

35.0 + 0.7 *29.5 + 0.8 45.9 + 1.4 38.0 + 1.0 45.2 + 1.1 40.6 + 1.1 – 38.8 + 0.8 39.7 + 0.9 33.7 + 1.0 44.0 + 0.8 46.2 + 1.3 45.7 + 0.8 39.8 + 1.2 39.7 + 1.2 49.2 + 1.1 43.7 + 0.9 41.0 + 1.0 – 42.8 + 1.1 39.2 + 0.9 39.1 + 0.8 33.7 + 1.2 35.3 + 1.1 37.2 + 1.3 37.3 + 1.1 39.8 + 0.9 42.8 + 1.0 39.8 + 1.4

1.80 + 0.03 1.85 + 0.04 1.77 + 0.06 1.90 + 0.04 1.72 + 0.05 1.73 + 0.04 – 1.82 + 0.03 1.86 + 0.04 1.83 + 0.04 1.68 + 0.04 1.76 + 0.04 1.70 + 0.04 1.82 + 0.06 1.76 + 0.05 1.73 + 0.04 1.77 + 0.05 1.87 + 0.04 – 1.76 + 0.04 1.90 + 0.04 1.73 + 0.04 1.92 + 0.06 1.80 + 0.06 1.87 + 0.06 1.90 + 0.06 1.64 + 0.05 1.82 + 0.05 1.92 + 0.06

N, number of receiver functions used for each station. Depth estimate marked by asterisk, is excluded from further analysis.

indicates a thinner crust (38 km) with high Vp/Vs (1.9) that supports the presence of partial melt. Table 1 lists the Moho depth and Vp/Vs values estimated for each station. Overall, the estimated crustal Vp/Vs values vary from 1.64 to 1.92 and the Moho depths with the exception of station AHLT range between 35 and 49 km with an average around 40 km. Station AHLT located near a young volcanic centre (Fig. 2), has a strong converted phase at 4 s. The depth of this phase is identified as c. 30 km with a Vp/Vs of 1.85. Although the estimated Vp/Vs is similar, the Moho depth is up to 10 km shallower than observed at the nearby stations. Therefore, the large phase is interpreted not as Moho but a lower crustal phase originating at the bottom of low velocity zone and the depth estimate is excluded from further analysis (Table 1). In Figure 6, we compare surface elevations with our Moho depth and Vp/Vs values. Comparison between elevation and Moho depth shows a large scatter and lack of correlation suggesting the region is not in Airy isostasy. Based on their geographical distribution, two station subsets are

distinguished from the rest. Stations at the Arabian foreland shows low elevations, relatively thin crust around 35 km and Vp/Vs values that are 1.8. In contrast, stations near the BS also have low elevations but a thicker crust (c. 40 km) and low Vp/Vs ranging between 1.64 and 1.73. The rest of the stations have relatively higher elevations and highly variable Moho depth and Vp/Vs values (Fig. 6). In order to correlate our results with major tectonic features of the region, we produced regional contour maps of Moho depth and Vp/Vs (Fig. 7). In general, crust is thinnest (c. 35 km) beneath the Arabian foreland in the south and increases towards the north beneath the plateau with the thickest crust (.45 km) within the northeastern corner of the region (Fig. 7a). Although the northward crustal thickening away from Arabian margin is consistent with Zor et al. (2003), the new Moho depths are relatively shallower and indicate on average 5 km thinner crust that is similar to those obtained by S-wave receiver functions (Angus et al. 2006). The largest differences from Zor et al. (2003) are observed where the estimated

CRUSTAL VARIATIONS IN EAST ANATOLIA Stations at Arabian margin

2

491

2

Stations near Bitlis suture

Elevation (km)

Elevation (km)

Other stations

1.5

1

0.5 30

35

45 40 Moho depth (km)

50

1.5

Stations at Arabian margin Stations near Bitlis suture Other stations

1

0.5 1.6

1.7

1.8

1.9

Vp/Vs

Fig. 6. Plots of elevation versus Moho depth and Vp/Vs for each station. Two station subsets are distinguished from the rest and shown by different symbols.

crustal Vp/Vs is anomalously high between Lake Van and Karlıova Triple Junction where NAF and EAF merge together (Fig. 7b). In the east, the increase in crustal thickness is more gradual whereas in the west, the crust thickens sharply near the BS and displays pronounced Moho topography across the Anatolian plate (Fig. 7a). Remarkably, the largest short-wavelength (c. 8 km) Moho depth variation occurs across the MalatyaOvacık Fault Zone (MOFZ) located within the plate (Westeway & Arger 2001). This shortwavelength variation in crustal thickness indicate a step-like Moho boundary and provides evidence for the presence of multiple fragments within the Anatolian plate, which was previously considered as a rigid block (McKenzie 1978). On the other hand, across the major strike –slip faults NAF and EAF, Moho depth variations are rather limited and do not provide clear evidence of major vertical steps in Moho boundary. Unfortunately, any network that has large station spacing is not suitable to determine whether these faults extend deep into mantle, or whether deformation is distributed throughout the crust (Wilson et al. 2004). Such analysis can be conducted only by using data across faults from densely spaced transects like the LARSE transect in southern California (Zhu 2002) or from a station located near enough the fault to sample both sides of the fault. Although, crustal Vp/Vs is among the least constrained parameters from both laboratory and field measurements, variations in Vp/Vs can be used to infer compositional variations within the crust. Zandt & Ammon (1995) estimated Vp/Vs for different types of continental crust. Vp/Vs, which has a global average of 1.78, is especially sensitive to variations in quartz, biotite and iron content and

presence of partial melt. Higher values tend to correspond with more mafic compositions and lower values with more felsic compositions that contain more quartz (Fountain & Christensen 1989). In addition, the presence of partial melt has been found to reduce speeds of shear waves more than compressional waves resulting in higher Vp/Vs. Beneath the ETSE array, spatial variations of crustal Vp/Vs values show coherent patterns that can be linked to major tectonic features (Fig. 7b). In general, we observe high Vp/Vs in the Arabian margin and along a corridor along the NAF and the southern section of the EAAC that is dominated by widespread young volcanism (Keskin 2003). Low Vp/Vs values are observed in the northeastern section towards the Rhodope-Pontide fragment and in the south spatially limited to the BS. The distribution of volcanic ages in the region indicates migration of volcanic activity to the south during the past 11 Ma (inset of Fig. 7b). Spatially, the easternmost extent of the high Vp/Vs corridor (1.85) coincides with the youngest volcanic units (c. 3 Ma) in the plateau and implies the presences of partial melt. Towards the NW volcanism diminishes along the NAF but the high Vp/Vs values continue. In orogenic belts, the common spatial and temporal association of granites with crustal-scale shear-zone systems suggests channelized transfer of melt preferentially along high-strain zones, which helps the system to accommodate strain (Brown & Solar 1998). In this section, perhaps the NAF is acting as a weak zone that accommodates some volume of melt emplaced deeper in the crust. Alternatively, schist and serpentinite sequences with high Vp/Vs may be dominating the crustal composition along the NAF that roughly follows the ancient suture belt (Sengo¨r et al. 2005).

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Fig. 7. Contour maps of Moho depth (a) and Vp/Vs (b) variation in Eastern Turkey. Stations that are not used are marked with yellow fill. Locations of volcanoes and age distribution of volcanic units within the plateau (inset fig.) are taken from Keskin (2003). Numbers correspond to ages in Ma.

CRUSTAL VARIATIONS IN EAST ANATOLIA

Most likely, anomalously high Vp/Vs is related to both factors. In the NE, the crust becomes thicker (.45 km) and has relatively low Vp/Vs values (1.76) (Fig. 7). This change occurs close to the boundary between the EAAC and the RhodopePontide fragment and suggests a change in crustal composition between these two different tectonic units. These results suggest the Rhodope-Pontide fragment that is partly overlain by oldest volcanic units (Keskin et al. 1998; Keskin 2003) has a thicker crust and relatively more felsic composition. S-wave receiver functions also indicate that mantle lithosphere beneath this fragment is thicker than beneath the EAAC (Angus et al. 2006). Near the BS, Vp/Vs values (1.64 –1.73) are remarkably low (Fig. 7b). Keskin (2007) points out the lack of post-collisional volcanism in the region and Yılmaz (1993) estimates final suturing between the Bitlis-Po¨tu¨rge Massif in the north and the Arabian margin in south occurred during the Early Eocene. Using a mid-crustal discontinuity with distinct multiples in addition to the Moho discontinuity, a layered Vp/Vs model is obtained (Fig. 5). This aforementioned model implies the lower crust as the source of low Vp/Vs values and indicates the abundance of lower crustal rock assemblages with quartz-rich felsic compositions near the suture. In the south, part of the Arabian margin is covered with ,1 Ma alkaline basaltic flows derived from an ocean island basalt like mantle source (Sen et al. 2004). Minimal crustal contamination of the volcanics is attributed to the relatively thin crust of the Arabian foreland (Notsu et al. 1995), which agrees with our crustal thickness estimates (c. 35 km). Although, the ETSE stations are not exactly situated on the basalt flows, the observed relatively high Vp/Vs (1.8) for the Arabian foreland may be related to the mafic composition of the crust.

Seismic profiles across the plateau We applied the common conversion point (CCP) method (Dueker & Sheehan 1998) with our Vp/Vs estimates to produce two-dimensional (2-D) crosssections of seismic discontinuities across the plateau. Across seismic profile A-A0 along the northern margin of our study area (Fig. 8), the Moho discontinuity dips gently from west to east indicating gradual eastward crustal thickening. At the thickest crustal section (.45 km), an upper crustal discontinuity is also visible around 10 km depth. S-wave receiver functions show similar feature that is most likely associated with contrasting velocities between a thick volcano-sedimentary sequence and underlying crystalline basement (Angus et al. 2006). Unlike the S-wave receiver functions, the higher frequency P-wave receiver

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functions indicate strong spatial variations in the upper crustal structure throughout the region (Fig. 8). Another prominent feature in profile A-A0 is the strong negative phase at mid-crustal depths indicating a pronounced low velocity zone (LVZ) in the lower crust and correlates well with low shear wave velocities found in the lower crust by Zor et al. (2003) and Gok et al. (2007). This LVZ is visible beneath most of the stations north of Lake Van including the ones in the Anatolian plate (Fig. 8). The extent of the LVZ coincides with the region where fast SKS splitting directions (Sandvol et al. 2003b) and Pn fast directions (Al-Lazki et al. 2003, 2004) are parallel to each other but at high angles to the surface GPS motions (McClusky et al. 2000) suggesting that crust and upper mantle deformation are decoupled. Furthermore, large energy on the transverse component receiver functions corresponds to similar depths in the same region (Ozacar et al. 2004). For these reasons, this LVZ is a potential candidate for a decoupling zone that may exhibit ductile flow. For more conclusive evidence, the existing anisotropy observations should be modeled following the approach of Ozacar & Zandt (2004, 2009) to reveal the nature of crustal deformation, particularly in the LVZ. Profile B-B0 runs across the eastern margin of the ETSE array and indicates a relatively thin (c. 35 km) crust beneath the Arabian margin that becomes gradually thicker towards the NE (Fig. 8). Beneath station AHLT, a strong converted phase shows a large offset from the Moho discontinuity. As mentioned before, this large double-sided Moho offset is unrealistic and therefore this phase is interpreted as lower crustal in origin and the Moho beneath it is somehow disrupted or ‘hidden’ by the presence of melt in the upper mantle. Both profiles A-A0 and B-B0 exhibit upper mantle phases that are most prominent in the northeastern part of the study area. The depths of these phases are not reliable since a constant P-wave velocity (6.2 km s21) that is appropriate for crust is used in CCP stacks, but still within the depth range (60– 85 km) of the lithosphere-asthenosphere boundary (LAB) identified by Angus et al. (2006) and Ozacar et al. (2008). Profile C-C0 that images the western side of the ETSE array, exhibits a high-degree of complexity within the crust and upper mantle (Fig. 8). The Moho discontinuity shows remarkable depth variations across the profile that may represent Moho offsets. These short-wavelength fluctuations are not only limited to the plate edges near the NAF, the EAF and the BS but also occur within the Anatolian plate across the MOFZ and suggest that the Anatolian plate is internally fragmented. For conclusive evidence, these features need to be imaged

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Fig. 8. Cross-sections of CCP stacked receiver functions. Profile locations are shown in the left. Red is used for positive and blue for negative amplitudes. Dashed lines indicate Moho discontinuity and in C-C0 also strong mid-crustal phases.

with better resolution. In the south, the Arabian crust is characterized by strong positive converted phases that are abruptly reversed in polarity across the EAF. This indicates a dramatic change in crustal structure between Anatolian plate and Arabia. Across the profile, series of strong upper mantle phases appear just north of EAF, which are not visible elsewhere. These complex phases are difficult to interpret but may be related to the complex lithospheric structure of the BS identified by Angus et al. (2006).

Conclusion In this paper, we map out the Moho depth and Vp/Vs variations across Eastern Turkey using P-wave receiver functions and produced refined seismic images of the crust. Both crustal thickness and

Vp/Vs estimates exhibit patterns that are spatially coherent with main tectonic features. We found northward crustal thickening away from Arabian foreland where the crust is thinnest (c. 35 km). In the east, the crust becomes gradually thicker towards the north and exceeds 45 km in the NE corner, whereas in the west the crust thickens sharply at the BS and exhibits significant Moho topography across the Anatolian plate. Vp/Vs is low at the BS (1.64–1.73) and the northeastern part of the plateau (1.76), high in the Arabian foreland (1.8) and highest (1.85) along the NAF and near the youngest volcanic units west of Lake Van (c. 3 Ma). The lower crust across a large area is also characterized by a LVZ that may act as a decoupling zone between the crust and upper mantle. These results indicate a very complex crustal structure where crust is made up of different

CRUSTAL VARIATIONS IN EAST ANATOLIA

fragments that vary greatly in composition and deformation style. Moreover, anomalously high Vp/Vs estimates in the eastern volcanic region support the presence of partial melt in the crust. In general, there is a lack of correlation between Moho depths and surface elevations and crustal thickness is thinner than expected beneath the plateau, which is inconsistent with a simple Airy compensational mechanism. However, the absence of a significant crustal root beneath the 2 km high east Anatolian plateau is consistent with the slab breakoff model of Sengo¨r et al. (2003) and Keskin (2003) that allows asthenosphere near the base of the crust with little mantle lithosphere. In this model, uplift of the plateau is supported by hot partially molten asthenosphere rather than a thick buoyant crust. Broadband waveforms used in this study obtained from IRIS Data Management Center which is funded through NSF under cooperative agreement EAR-0004370. NSF supported this research under grant EAR-0309838. We thank R. Gok and E. Sandvol for their helpful suggestions and encouragement.

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Index Note: Page numbers in italic denote figures. Page numbers in bold denote tables. Aalenian isotopic elemental data 224 palaeogeographic maps 227 succession sandstone modal data 209 Abino-Gunai Zone 200–201, 206 Abkhazo-Racha Zone 204 Porphyritovaya Series 203 –204 Afro-Arabian Plate 9, 454 Agrakhan-Tbilisi-Levantine fault 395–396 Akhalkalaki formation 254 Akhaltsikhe basin 248– 249 Akhourian fault 396 Akveren Formation 130 Alasani Basin faults v. topographic map 268 tectonic geomorphology 272 Alazani geological profile 250 Alpine belt terrane map 438 Alpine deformation 233 Alpine fold-and-thrust belt 77 Alpine–Himalayan belt 1 –2 schematic structural map 78 Amasra 130 Aptian sediments 131 Early Cretaceous angular unconformity 127 horsts 132 Late Cretaceous angular unconformity 127 Amphiboles 359 Anatolia Arabian Platform autochthonous sequence 465– 466 Anatolian orogen 417, 420, 421, 424, 425 deformation phase 431– 433 Late Cretaceous to recent kinematics 409–433 Maastrichtian to recent tectonostratigraphical column 413 magmatic rocks element plots 447 niobium yttrium 452 rock/chondrite normalized 449 rock/ocean ridge granite-normalized 448 rock/primitive mantle-normalized 448 rubidium niobium yttrium 452 zirconium, titanium, and silicon plot 448 zirconium, titanium, niobium, and yttrium plot 448 Anatolide–Tauride platform 454 Anchkhoi Formation 203 Andean subduction zone thorium, ytterbium, niobium, and yttrium diagram 450 Arabian-Eurasian convergence rate 395 Arabian facies 478 Arabian plate active tectonics 394 collisional zone structural map 384 movement 395

Arabian Platform basaltic lava flow 414 corridor regional tectonic issues 1–3 Cretaceous to recent tectonostratigraphical column 412 cross-section 432 deformation phase 431–433 Late Eocene-Oligocene compression 420 Oligocene to Middle Miocene extension 421 palaeostress configurations 417, 420, 421, 424, 425 Paleocene to Early Miocene evolution 432 Paleocene to Middle Eocene extension 417 Pliocene to recent compression 425 sedimentary basin tectonics 1 –9 turbidites 413–414 Armenia geological map 385 identified radiolarian species 387 Lesser Caucasus 329 –349, 397 ophiolites 383– 389 sketch geological map 355 Armenian Ophiolite 353–380 alkaline compositional field 360 alkaline lavas origin 377 analytical methods 359 Armenian ophiolites geodynamic significance 379 –380 Armenian ophiolites significance MOR or back-arc setting 376– 377 chemical compositions 360 clinopyroxenes 360 geodynamic reconstruction 378– 379 geological setting 355– 358 Harker variation diagrams 373 insights for Jurassic back-arc formation 353–380 Lower Jurassic to Upper Cretaceous periods 378 –379 magmatic rock plots 373 magmatic rock types 370– 371 ophiolite history reconstruction 377– 379 ophiolitic, alkaline and calc-alkaline series 373 plutonic and volcanic ophiolite series 370– 371 results 359– 376 argon isotope dating 375– 376 field relationships 359–360 major trace REE geochemistry 367– 375 major elements 367– 371 REE geochemistry 371– 375 trace elements 371 mineral chemistry 360 –367 petrography 360 –367 alkaline lavas 361–365 calc-alkaline lavas of Stepanavan zone 367 dacitic dyke-like bodies 367 diorites 361 gabbros 360– 361 mesocratic to leucocratic gabbros 361 olivine gabbros 361

498

INDEX

Armenian Ophiolite (Continued) ophiolite plutonic rocks 360 ophiolite volcanic and subvolcanic rocks 361 plagiogranites 361 volcanic rocks 361 websterites 361 trace and REE plots 374 Armenian territory Neogene to Quaternary volcanic clusters 403 Artvin area 314 Jurassic rifted basin 322 Jurassic units 315 Late Palaeozoic– Early Cenozoic tectonic development 281–322 Lower– Middle Jurassic sedimentary succession 309 metamorphic basement 321– 322 reconstruction 318 sequence 315 tectonostratigraphy 288, 322 Tethys closure stages along Eurasia southern margin 281–322 Artvin Basin Early Jurassic ophiolite 321 tectonic development 316, 322 uplift and erosion 321 Upper Jurassic basalts 321 Artvin-Bolnisi unit 313 As¸utka Thrust Sheet 310 Atbas¸ı Formation nannoplankton dating sites 129 Atchichkh Formation 203 Azat river area Quaternary formations microtectonic data 405 Azerbaijan 266 Cenozoic tectonics 274 faults v. topographic map 268 geomorphology 273 Greater Caucasus uplift 261–276 Lesser Caucasus 329–349, 397 Lower Pleistocene coastlines 270– 271 uplift rates 270–271 vertical movement 270– 271 Azov Sea 65, 150–151 active rifting cessation 155 air-loaded tectonic subsidence curves 152 Albian sediment 144 backstripping data 154 Black Sea northern margin 141–145, 152 Black Sea sedimentary basin tectonics 4 –5 description 137 DSS profile 28 travel-time curves 45 Eocene sediments 144 epochs 141–143 horsts, grabens, and half-grabens 154 joint profile 28 velocity model 52 Paleocene sediments 144 regional evolution 152 seismo-geological profile 144 tectonic subsidence rates maps 153 Backstripping 148, 149, 152 Badenian– Sarmatian 163 deposits 170, 179 isopach map 163

sediments 167 sequence 171 unit 159, 176 view 173 Baiburt-Karabakh unit 313 Bajocian–Bathonian 204 age 311 formations 334 palaeogeographic maps 227 Bajocian–Callovian succession sandstone modal data 210 Bajocian isotopic elemental data 224 Balkanides schematic cross-section 84 structural zones 79 Balkanides shelves Black Sea sedimentary basin tectonics 4– 5 Barremian–Aptian Formation stratigraphic contact 123 Base of Pontian 176 Baskil arc 442 Baskil-Divrig˘i transect 439 geological framework 440–442 geological map 438, 441 magmatic rocks 446 –453 magmatic rocks dating 443 magmatic suites 451 magmatism spatial and temporal evolution 455 volcanic-sedimentary sequences 442 Baskil transect magmatic rocks geochemical analyses 444–445 Bathonian compression 234 Bathonian sedimentary succession 233 Bazum Mountains 396 Belaya River 187– 189 Berezov Zone lithic fragments 206 Bitlis complex chloritoid 478 index minerals electron microprobe analyses 471 mylonitic marbles 468 palaeozoic rocks 462 schists 462 sediments 467 Triassic rocks 464 Bitlis meta-sediments carpholite and chloritoid 471 Bitlis-Po¨tu¨rge metamorphics 442 Bitlis suture zone 439, 485 Bitlis– Zagros 456 Bitlis– Zagros Suture Zone 8, 409 Bjala 84 polyphased site 92 stress states 92 Black Sea 57–73, 138 see also Romanian Black Sea shelf asymmetric back-arc basin 19 back-arc basin 131 back-arc deformation 15, 16–17 Cretaceous-Quaternary stratigraphy 142 DSS line 29 travel-time curves 46 DSS Profile 25 travel-time curves 45 fault zones and gas seeps 62– 63 gas seeps 73

INDEX joint Profile 28 velocity model 52 major tectonic domains 140 model ray paths for seismic phases 51 mud volcanoes gas seeps and gas hydrates location 58 palaeotectonic reconstruction 14 progressive asymmetric trench retreat sketches 17 sedimentary basin tectonics 1– 9 Azov Sea 4 –5 Balkanides shelves 4– 5 basin formation 4 –5 Bulgarian shelves 4 –5 central Pontides of Turkey 4 –5 collision stages in Armenia, Azerbaijan, Georgia and Turkey 6– 7 crustal affinity 4– 5 crustal structure 4 East and SE Anatolian suture zone 7 Eastern Pontides 6 –7 Greater Caucasus basin and tectonic evolution 5 –6 inversion 4 –5 Lesser Caucasus 6–7 lithosphere structure 4 neotectonics 4– 5 obduction in Armenia, Azerbaijan, Georgia and Turkey 6 –7 Odessa Shelf 4 –5 regional tectonic issues 1 –3 Romanian shelves 4– 5 back-arc basin extension and magmatism 8 inversion compression and magmatism 8– 9 Pre-Cimmerian and Cimmerian history 7 –8 present configuration and active tectonics 9 tectonic evolution in Russian, Georgia and Azerbaijan 5– 6 tectonic evolution of Peri-Arabian suture zone 7 seismic recordings SP 8 of Profile 25, 47 SP 9 of Profile 25, 47 subsidence history 170 tectonic elements 44 tectonic map 12 tectonic units 24 temperature 70 travel times 51 water level 171 water-level curve 177 wells and pseudo-wells 139 Black Sea back-arc basin 11– 19 asymmetric trench retreat 15–17 basement 12–13 geodynamic extension models 13–15 lithosphere strength 15 symmetric vs. asymmetric opening 13–15 modern analogues geodynamic models 11– 19 origin 11–19 regional tectonic setting 12– 13 Black Sea Basin 23–39, 24, 113–134, 114 Bouguer anomalies onshore 28 Cenozoic sedimentary successions cross-sections 26 dating new nannoplankton ages 113– 134 DSS data 31– 32

499

DSS Profile 25 velocity model 49 earth’s crust and upper mantle structure 23– 39 European scale P-wave seismic tomography model 34 free air anomalies 28 geodynamical implications 130– 133 Barremian to Albian 130–132 Coniacian to Eocene 132–133 geophysical data and tectonic implications 23– 39 gravity field 27– 29, 28 heat flow 30–31 inverted margin in Central Pontides 113–134 lithosphere seismic tomography study 32–36 lithostratigraphic columns 243– 244 magnetic field 29–30, 29 nannofossil assemblages 119–120 nannoplankton age 116– 118 Pontides stratigraphic sequence overview 114 –121 Pontide stratigraphic sequence nannoplankton dating 121 P-wave 36 P-wave velocity model 35 ray paths for seismic phases 48 sedimentary cover main features 25–27 seismicity map 33 syn-rift C¸ag˘layan Group 121– 126 Black Sea coast 121– 122 Ulus Basin 125 –126 Zonguldak Basin 122– 125 tectonic implications 23–39 tectonic scheme 37 tectonic units 29 time residuals 36 travel time 48 Upper Cretaceous-Eocene post-rift 126– 130 sedimentary formations 129–130 volcanic-sedimentary sequence 126–129 upper mantle structure 23–39 velocities 31, 36 Black Sea crustal structure 43– 55 DSS data reinterpretation 46–53 along Profile 25, 46–50 along Profile 28/29, 50–53 Azov Sea 50–53 seismic data 46–49 seismic data and wave field characteristics 50 velocity model 49–50 velocity model along Profile 28/29, 50–53 wave field description 46–49 seismic sounding data 43–55 Black Sea northern margin 137–155 Azov Sea 141– 145, 152 from 1-D subsidence analysis 148– 149 tectonic evolution from seismic data 145 –148 Cretaceous-Neogene tectonic evolution 137–155 inversion timing 154–155 Odessa Shelf stratigraphy 141 –145 Odessa Shelf tectonic evolution 145 –152 main tectonic elements 145 –146 from seismic data 145– 148 tectonic events age 146 regional tectonic setting 138– 141 rifting timing 152– 154

500 Black Sea northern margin (Continued) seismic reflection data 137– 155 tectonic subsidence analysis 137– 155 Borjomi-Bakuriani lava flows 255 Borjomi-Kazbek fault 396 Bottom stimulating reflectors 68 Bouguer anomalies onshore 28 Bouma sequences 296 Budur syncline 272 Bugunja Formation 189–190, 205 accessory minerals 206 Bulgaria see also Eastern Balkanides shelves 4 –5 structural zones 79 Byala Voda 95 Bzerpia Formation 192 C ¸ ag˘layan Formation 118 stratigraphic position 123 C ¸ ag˘layan Group formations 122 nannoplankton dating sites 121 C ¸ angaldag˘ Complex 312 Callovian ammonites 384, 386 deformation 233 Campanian-Maastrichtian siliciclastic sandstone turbidites 310 Campanian marls of Bjala cliffs extension 99 Campanian-Paleocene Akvenren nannoplankton dating sites 129 Cape Emine polyphased site 93 Capidava-Ovidiu faults 53 C ¸ atak River 478 C ¸ atak valley 469 cross-section 467 Caucasus corridor regional tectonic issues 1– 3 Mesozoic evolution 230 sedimentary basin tectonics 1 –9 structural sketch map of belt 330 tectonic map 282 Cemaller Formation 124 Cenomanian– Turonian autochthonous limestones 341 Cenozoic complexes Eastern Bitlis complex lithostratigraphy 464– 465 Cenozoic sedimentary successions cross-sections Black Sea Basin 26 Central Balkan-Forebalkan zone 84 evolution model 107 Central Greater Caucasus volcanic rocks 256 volcanism 226 Central Pontides 132 arc location 114 arc structural sketch 115 Mesozoic– Cenozoic stratigraphy 133 nannofossil investigations 134 stratigraphic formations dating 120 Central Pontides Belt 114

INDEX Chokrakian regiostage 242, 247 Cimmerian ‘Orogeny’ tectonics 2 Cimmerian tectonics 2 Common conversion point 493 Common deep point 43 Cretaceous extensional tectonics 2 Cretaceous stratigraphic gap 128 Cretaceous succession base 24 Cretaceous ‘syn-rift’ sequence 113 Crimea anomaly zone 30 Eastern Pontides regional comparisons 312–313 mountain belt 154 Orogen 138 Dacian, Romanian and Quaternary section details 176 Dacian– Quaternary 163, 165, 179 Dacian section isopach map 165 Dacian unit 174 Danube Canyon 73 Daralagez continental terrane see South Armenian Block (SAB) Darende Basin 414 Dead Sea Fault 485 Deep seismic sounding (DSS) data 4 Black Sea Basin 31–32, 36, 43– 55 Delfin wells 162 Demirkent Intrusive Complex 289– 291 amphibolitic host rocks 290 crystalline basement rocks 290 MORB-normalized spider plots 294 Devrekani Metamorphic Unit 312 Dinaridic-Carpathian-Balkan system 108 Djangur Formation 190 Dnipro Canyon 68, 73 Domuzdag˘ complex 133 DSS see Deep seismic sounding (DSS) data Durankaya complex 464 Dvurechensky volcano 67 gas release within caldera 67 subsurface structure 68 Dzirula high 247 Dzirula massif 313 Early Cretaceous 14 C ¸ ag˘layan Group formations 122 nannoplankton dating sites 121 East Anatolian Accretionary Complex 486 East Anatolian Fault 485 East Anatolian Plateau 485– 495 CCP cross-sections 494 converted phases 488 crustal variations seismic images 485 –495 data and method 487–488 Moho depth 490 contour maps 492 v. elevation 491 observations 488– 494

INDEX crustal thickness and variations 488– 493 seismic profiles 493– 494 receiver functions recorded 489 East Balkan-Srednogorie zone 93, 95 East Balkan Zone polyphased site 93 East Black Sea 23– 39, 54 cross-section 432 crust 55 East Black Sea Basin 43 crust structure according to existing DSS data 31–32 earth’s crust and upper mantle structure 23–39 rifting 113 Eastern Balkanides 77–109 bedding planes 94 brittle deformation 77–109 brittle tectonic analysis methods 84– 93 compression and strike-slip stress states 100 extensional stress states 96 faulting analysis 93–104 compressional stress states 98– 103 Early Alpine tectonics 104–106 extensional stress states 95–98 Late Alpine tectonics 106–109 inversion stages 106 –107 post-orogenic stage 108 syn- and post-rift stages 106 normal faulting 95– 98 reverse faulting 98–103 strike-slip faulting 103–104 strike-slip stress states 103– 104 fault populations and stress states 99, 103, 105 features 78–81 geodynamical evolution 78– 81 palaeostress field reconstruction 77–109 palaeostress states reconstructions 93–104 palaeostress tensors 86–87 populations 105 schematic structural map 81 stress axes 95 stress states 105 structural zones 81– 84 Central Balkan-Forebalkan zone 83– 84 East Balkan zone 83 Sakar-Strandja zone 82 Srednogorie zone 82–83 tectonic events 81–84 tectonic evolution Cenozoic times 77– 109 Mesozoic times 77–109 trending strike-slip stress states 101 Eastern Bitlis complex 461– 480 age constraints 466–469 alpine high pressure evolution 461 –480 argon data 476 argon phengite 477, 478 chlorites electron microprobe analysis 472, 474 Eocene pillow lava 480 garnet electron microprobe analysis 474 geological map 463, 470 geological observations 466– 469 cross-section 466– 469 geological setting 462

501

lithostratigraphy 462–466 Anatolia Arabian Platform autochthonous sequence 465–466 Cenozoic complexes 464 –465 Mesozoic ophiolitic sequences 464 metamorphism 462– 466 age 471 –478 data 466– 469 evolution 469 –471 phengite electron microprobe analyses 473 pressure temperature diagram 475 schematic geodynamic cross-section 479 Eastern Bulgaria tectonic zones and units 80 Eastern Mediterranean region tectonic outline 3 Eastern Pontides 281–322 active continental margin development 315– 319 Carboniferous 315 Jurassic 315 –317 Late Cretaceous 317 Late Jurassic 317 Oxfordian –Berriasian 317 Paleocene–Eocene 318–319 Triassic 315 alternative tectonic models 319–321 ancient comparisons 319 Artvin area 281 –322 depositional and tectonic settings 314– 315 East Pontide Autochthon 301– 302 geological map 283 Hercynian basement and intrusive rocks 288–293 Demirkent intrusive complex 289–291 East Pontide Autochthon 289 East Pontide Autochthon geochemistry 291–293 Izmir-Ankara-Erzincan suture zone 308 Jurassic platformal succession cover 304 Karadag˘ metamorphics 291 Lower Slice Complex Upper Jurassic-Upper Cretaceous 302–304 Upper Slice Complex metamorphic and intrusive rocks 291 Upper Slice Complex Upper Jurassic-Paleocene units 304 –308 Thrust Sheet 1, 304 Thrust Sheet 2, 304– 305 Thrust Sheet 3, 305– 308 Thrust Sheet 4 and 5, 308 Late Palaeozoic-Early Cenozoic tectonic development 281 –322 Lower–Middle Jurassic sequences 293–299 East Pontide Autochthon 293 Jurassic igneous rocks geochemistry 298–299 Lower Slice Complex 293–298 platform succession 298 Middle Jurassic magmatic arc-type unit 299–301 modern comparisons 319 previous work 283 –285 regional comparisons 308–314 Caucasus 313– 314 central and western Pontides 310– 311 Crimea 312– 313 Ku¨re Complex 311– 312

502 Eastern Pontides (Continued) tectonic units 283 tectonostratigraphy and structure 285–288 Tethys closure stages along Eurasia southern margin 281– 322 Upper Jurassic-Upper Cretaceous sequences 301 –308 Upper Palaeozoic basement 281–283 Upper Slice Complex 299– 301 Eastern Turkey 492 tectonic map 487 Eastern Turkey Seismic Experiment 485, 487 East European Platform 29, 49, 138, 141, 181 East Pontide Autochthon 289, 293, 310 element analysis of granitic rocks 291 geochemistry 291– 293 lithologies exposed 289 Lower Jurassic basalts 299 Lower– Middle Jurassic part of Hercynian granitic 295 overlain 285 REE analysis of granitic rocks 291 sedimentary successions 309 trace element analysis of granitic rocks 291 XRF analysis of basaltic rocks 291 Elazig magmatic suite 442 Electron probe microanalysis 359 Eocene–Miocene basins nannoplankton dating sites 131 Eocene Urse Formation metamorphic rocks 466 Erakh mountains 385 Estasadok Formation 203 Eurasia collisional zone structural map 384 Eurasian margin Southern 353 stratigraphic and lithological data 339 subduction 333– 334 synthetic lithostratigraphic log 335 Eurasian plate 329 overthrusting ophiolites 332 structure 333 Euxinian Threshold 161–162, 163 Fault-slip data 92 Feldspars 205 Fish Canyon Tuff Sanidine 475 Fold-and thrust belts 263 Forebalkan-Central Balkan zone 93 Forebalkan Zone polyphased site 92 stress states 92 Fore-Caucasus Triassic succession sandstone model data 202 Western Greater Caucasus stratigraphic succession 187–189 Fore-range Zone 314 Gabbronorite 375 Gagra-Dzhava zone 267, 276 Garni fault 399 Gavaraget normal faults 400

INDEX GC see Greater Caucasus (GC) Gegam massif crest fault system 401 Gegam massif Karakhanian 403 Geltorechka-Sarighamish Fault 396 Georgia 240 Black Sea Basin lithostratigraphic columns 243– 244 horizontal component 404 Late Cenozoic basins 239–256 Lesser Caucasus 329– 349 Neogene–Quaternary volcanic formations 253, 256 palaeogeographic maps 245– 246 tectonic map 241 thrusts 269 uplift rates, vertical movement, and Lower Pleistocene coastlines 270–271 volcanic formations 253–255 Germav Formation 466 Gerpegem formation 191 Gevas-C ¸ atak-Narli section 466 Gevas complex 478 mylonitic marbles 468 Goitkh Zone 199– 200 Golitsin Fault 146– 147 Golubitsky Volcano eruption 65 Greater Caucasus (GC) 261– 276 active tectonics 267 –268 Bajocian-age volcanic activity 233 belt 181 Cenozoic-Recent tectonics 261– 276 convergence and uplift 267– 268, 269 Crimea mountain belt 154 cross-section 182, 275 crustal-scale cross-section 263 earthquakes and active faults 267–269 evolution 229 extension trends 186 geodynamics 262 geological map 182 Jurassic-age volcanic rocks petrography 193– 198 Jurassic stratigraphy 188 linking topography tectonic model 264 –265 Lower and Middle Jurassic rocks 186 Lower Pleistocene coastlines 270–271 map 262 Mesozoic–Cenozoic basin 262 northern part 313–314 Rare Earth Element spidergrams 219–220 regional geodynamic setting 228 regional tectonics and geodynamics 262–267 sampling sites locations 183 sedimentary basin geometry 266–267 strike-slip tectonics 272 subsidence and tectonics v. topography 269– 274 tectonics v. topography 269–274 trace element 219– 220 trace element composition 222 uplift 267– 274 Azerbaijan 261–276 palaeogeography 272

INDEX rates 270– 271 subsidence studies 269 tectonic geomorphology 272–273 topography and thrusts 273–274 uplift-exhumation 269 –272 vertical movement 270–271 Greater Caucasus Basin 19 back-arc rifting 13 Great Zap anticline 465 Gubkin Ridge 152 Gurii Trough 25 Gu¨rso¨ku Formation 130 Gu¨ru¨n Basin 421 Hakkari complex 464 Hauterivian-Aptian formation 106 Hercynian basement 310 Eastern Pontides 288–293 High field strength elements 299 Highstand systems tract (HST) 159, 174, 179, 180 Histria Depression 162, 163, 164, 166 HST see Highstand systems tract (HST) Illaryonov formation 203 Incised valley 176 Indolo-Kuban Depression 146 Indyuk Formation 200 ˙Inpiri Formation 122, 127 Intra-oceanic subduction 335–336 Iranian belts structural sketch map 330 Irmakyani Formation 296, 297, 298 basalt 300 field relations 305 Izmir– Ankara– Erzincan–Sevan–Akera suture zone 8, 308, 309, 480 Jurassic back-arc formation Armenian Ophiolite insights 353– 380 Jurassic basinal succession cover 302 –303 Jurassic Olur magmatic unit field relations 306 Jurassic platformal succession cover 304 Jurassic rifting 104 Kabarjina group 255 Kalamit Ridge 148 Kamennomostsky 190– 191 Kangal Basin 414 thrust fault 428 Kapanbog˘azı Formation red pelagic limestone 126, 126, 128, 129 syndepositional listric normal faults 127 Karadag˘ Metamorphics 291, 298, 315 Karaganian regiostage 242, 247 Karakaya Complex 315 Karamarian anticline faults v. topographic map 268 geomorphology 273 Karamarian Quaternary Anticline 272 Kargı Block 113 Karkinit-Gubkin rift 147, 148 basin 145, 154

inversion 154 post-rift sag basin 154 Karkinit Trough 18, 144, 145, 146 cross-sections illustration evolution 147 Kartli basin 249– 251 folds and thrusts 272 geological profiles 252 Oligocene-Neogene molasses 251 Kastamonu Eocene basin stratigraphic wedging 131 Kazbegi group 255–256 Keban-Malatya-Bitlis Block 409 Keli complex 256 Keli group Gudauri flow 255 Kelkit Formation 293 Kerch Peninsula gas seeps, local uplifts, and submarine canyons 59 Kerch-Taman marine hydro-gas survey 59 offshore 57 shelf 59 Kerch-Taman Trough 34, 38 gas seeps 58 Khumurdo formation 254 –255 Kilian-Zmeinian Zone 145 Kilimli Formation 121, 122, 127 Kilimli-I˙npiri Formation stratigraphic contact 123 Kınalıc¸am Formation 317 basalt 300 field relations 305 Kırs¸ehir Block 113, 114 Kizilcakilise sandstone and gravel 123 Kochkanov formation 200 Kondolovo, Strandja Zone 102 Konkian regiostage 242, 247 Ko¨pru¨lu¨ Formation 465 Kotel polyphased tectonics 98 Kraishte zone 79 Krasnaya Poliana 202 –204 Jurassic stratigraphy 189 Krasnopoliana Zone 203, 233 Krayova Step 145, 148 Kırkgec¸it Formation 465 Krouchevetz, Srednogorie Zone polyphased tectonics 97 Kumurdo formation 254 Kura Basin 247– 248, 269 dextral strike-slip motion 269 folds and thrusts 272 geological profile 250 geomorphology 273 uplift rates, vertical movement, and Lower Pleistocene coastlines 270– 271 Kura plain faults v. topographic map 268 Ku¨re basin 314 Ku¨re Complex 311 Eastern Pontides regional comparisons 311–312 Ku¨re marginal basin 320

503

504

INDEX

Laba-Malka Zone Jurassic stratigraphy 188 Jurassic succession 189 Western Greater Caucasus stratigraphic succession 189–190 Labina zone Callovian–Tithonian succession 189 Laramian compression 102 Large ion lithophile elements (LILE) 298–299, 377, 449 Late Albian Cemaller Formation 127 nannoplankton dating 128 Late Campanian-Maastrichtian period 106 Late Cenozoic basins of Georgia 239–256 basins structure 249 –251 evolution review 239– 256 Late Cenozoic, syn- and post-collisional stages 240 main tectonic zones 239– 240 molassic basins of Transcaucasus 240– 249 Akhaltsikhe basin 248 –249 Kura basin 247 –248 Chokrakian regiostage 247 Karaganian regiostage 247 Konkian regiostage 247 Late Miocene Early and Middle Sarmatian regiostages 247 Late Miocene Late Sarmatian regiostages 247– 248 Late Miocene Meotian and Pontian regiostage 248 Late Pliocene-Pleistocene 248 Middle Miocene Tarkhanian regiostage 247 Oligocene-Early Miocene time 247 oligocene series 248– 249 Rioni basin 240–247 Chokrakian regiostage 242 Dzirula high 247 Karaganian regiostage 242 Konkian regiostage 242 Late Miocene Late Sarmatian regiostages 242 Meotian regiostage 242–244 Middle Miocene Tarkhanian regiostage 242 Pleistocene Gurian regiostage 246– 247 Pliocene 244 –246 Pontian regiostage 244 Neogene-Quaternary volcanic formations 251–256 Greater Caucasus quaternary volcanoes 255 –256 Kabarjina group 255 Kazbegi group 255–256 Keli group Gudauri flow 255 South Georgia volcanic formations 253–255 Akhalkalaki formation 254 Borjomi-Bakuriani lava flows 255 Khumurdo formation 254–255 Kumurdo formation 254 Samsari formation 254, 255 Tsalka formation 254 Upper Miocene-Lower Pliocene, Goderdzi formation 253– 254 Upper Pliocene-Holocene formation 254 Late Cretaceous Kapanbog˘azı nannoplankton dating 126 Late Eocene to Oligocene compression and extension 429

Late Paleocene to Middle Eocene period structures and extension direction 429 Le Maitre volcanics rocks classification 218 Lesser Caucasus 329 –349, 393–406 active faulting 395 –402 major active structural pattern 395– 400 within wedges structures 400–402 active fault map 395, 398 Campanian palaeotectonic map 348 collision 329–349, 341–344 SAB and Eurasia collision 341–343 SAB collision structures and deformation 343 syn-collisional basin 343 timing collision stages 343 –344 cross-section 334 crustal-scale cross-section 263 faults and sample locations 416 general structural setting 333–336 subduction processes evidence 333–335 geodynamic evolution and discussion 344 –347 Coniacian to Paleocene 344– 347 Middle Jurassic to Coniacian 344 Paleocene to Lower Miocene 347 Upper Miocene to present 347 geodynamic model of evolution 346 geographical coordinates 418, 422, 423, 426, 427 geological map 182 geologic cross section 397 GPS profile 402 identified radiolarian species 387 kinematic setting 395 Maestrichrian palaeotectonic map 348 micro-fault kinematics 404 obduction and collision 329– 349 obduction modalities 336–341 ophiolitic unit 339– 341 SAB autochthon 336–339 oblique slip with reverse component 399 ophiolites 383–389 oriental border 349 palaeostress orientations 418, 422, 423, 426, 427 recent tectonic stress evolution 393–406 SAB 331–332 Sevan-Akera 313 structural map 331 subductions 329–349 synthetic lithostratigraphic log 335 Tauride– Anatolides 349 vertical faults 272 volcanic cluster analysis 402–403 Ypresian palaeotectonic map 348 Lherzolite Ophiolite Type (LOT) 346 –347, 379 Light rare earth elements (LREE) 371, 375, 377, 449 LILE see Large ion lithophile elements (LILE) Lomonosov submarine massif 60 Loss on ignition 210 LOT see Lherzolite Ophiolite Type (LOT) Lower Cretaceous hot spot magmatism SAB 353–380 radiolarian assemblage 387

INDEX Lower Eocene growth faults 428 Lower Slice Complex 283, 285 cross-section 301 lithologies exposed 289 Lower Jurassic sequence measured log 297 Lower-Middle Jurassic basalt 300 Lower-Middle Jurassic sequences measured logs 296 Low field strength element 449 Lowstand systems tract 172, 179 Low velocity zone 493 LREE see Light rare earth elements (LREE) Maastrichtian limestone fault populations and stress states 102 Maden complex 464 Madenler 297 Maikop series 65 Main Caucasus Thrust 185, 263, 267, 274 Main Range Zone 313–314 Makarov Formation 199 Malatya Basin 414 Malatya–Keban metamorphic complex 446 metamorphic terrane 441 platform 440, 454 Malatya-Ovacık Fault Zone 491 Malaya Laba area Lower Triassic conglomeratic unit 230 Triassic succession 187, 189 Western Greater Caucasus stratigraphic succession 187 –189 Maliy Tkhach Formation 205 Mariana-Bonin 317 MBSR see Mid-Black Sea Ridge (MBSR) MEBE see Middle East Basin Evolution (MEBE) Programme Mesozoic ophiolitic sequences Eastern Bitlis complex lithostratigraphy 464 Metamorphic core complexes 79 Methane origin 73 Mezmai formation 191 Mid-Black Sea Ridge (MBSR) 23, 31, 32, 38, 43, 54 Middle East Basin Evolution (MEBE) Programme 1, 8, 13, 461 Middle East-Caucasus area tectonic map 354 Middle Eocene formations 398 Middle Jurassic radiolaria 387 Mid-Eocene shortening 3 Mid Ocean Ridge Basalt signatures 353 Mineral Research and Exploration Institute 283 Mnadon lava flow 256 Moesian Platform 12, 29, 77, 78, 84, 94, 104, 108 Moho 49, 50 Mount Elbrus 261 Mount Karawul 385 Mount Tkhach-Belaya River area Triassic stratigraphy 187, 205

505 Triassic succession 187 Western Greater Caucasus stratigraphic succession 187– 189

Nannofossil dating 131 Naujni Formation 199–200 Neogene deposits lithostratigraphic columns 243– 244 Neogene-Quaternary stress field 405 Neotethys Ocean 353, 383, 393, 439, 440 Eocene subduction 453 subducted slab 456 New Britain subduction zone 15 Noratus normal faults 400 North Anatolian Fault 485 Northeast Anatolian sinistral strike–slip fault 395 –396 Northern Black Sea 57–73 carbonate formations as derivatives of gas emissions 60– 61 carbonate formations radiocarbon dating 60 gas hydrate 68–73, 70, 72 gas seepage 58–60, 59, 64 geological environments characterization 57–73 geological-geophysical investigations 61– 65 methane chimney restoration 60 methane geomorphological environments characterization 57–73 mud volcanoes 65– 68 seismic records 63 temperature distribution 70, 72 Northern Dobrogea Orogen 141 Northern Sevan lake area Eocene thrust faults 345 geological cross-section 342 Ocean Island Basalt 354, 379 Odessa Shelf air-loaded tectonic subsidence curves 148, 150– 151 Black Sea sedimentary basin tectonics 4 –5 Cretaceous-Eocene succession 141 cross-sections illustration evolution 147 interpreted seismic profiles 143, 148 location 137 seismic reflection data 143, 154 stratigraphy 141 –145 tectonic evolution 145 –152, 146 tectonic subsidence rates 151 Upper Cretaceous carbonate succession 144 Okinawa Trough 15 Oligocene deposits lithostratigraphic columns 243– 244 Oligocene scheme Mediterranean, Black, and Caspian Sea regions 242 Olur Magmatics 299 Ophiolitic rocks 308 Oxfordian ammonites 384 Oxfordian-Kimmeridgian palaeogeographic maps 227 Palaeo-Tethys Ocean 353 Paleocene clastics growth faults 415

506

INDEX

Paleocene-Eocene Nannofossils dating 333 Pambak-Sevan-Sunik fault 396 Pchich Formation 200 Peridotites 357 Pikhtar formation 201 Pillow lava 357, 384 Pliensbachian palaeogeographic maps 227 Pontian deposits isopach map 164 faulting 166 seismic line crossing 166 Pontian 1, 172 view 173 Pontian 2, 172 Pontian 3, 173 details 175 Pontian 4, 174 details 175 Pontides 130 continental margin 320 cross-section 287 crystalline basement rocks geochemistry 292 geological map 284 Late Jurassic pillow lava geochemistry 307 nannoplankton dating 121 Narlık granite tectonic discrimination 293 plate tectonic reconstruction 320 pull-apart basins 321 stratigraphic sequence overview 114– 121 tectonic map 282 Porphyritovaya Series 211 Pre-Callovian deformation 233 Pre-Cimmerian tectonics 2 Predobrogea Depression 146 Pre-Upper Jurassic Olur magmatic unit field relations and geochemistry 302 Psebai Formation 190 accessory minerals 206 Psekhako-Berezov Zone 191–199 lithic fragments 206 Pshekha formation 201 Quaternary deposits lithostratigraphic columns 243– 244 Romanian 167 Quaternary formations Azat river area microtectonic data 405 Quaternary scheme Mediterranean, Black, and Caspian Sea regions 242 Quaternary section Romanian isopach map 166 Radiolarian biochronology 383–389 geological setting 384–385 new results 385 sedimentary cover of Lesser Caucasus 383–389 Radiolarites 357 Rare earth elements (REE) 285, 359, 449, 453 Ray-tracing modelling 43 REE see Rare earth elements (REE) Reefal limestones 384 Rhodope-Pontide fragment 491

Rift flank uplift 132 Rioni basin molassic basins of Transcaucasus 240–247 River Rioni geological profile 250 Romania BSR 69 Romanian Black Sea shelf chronostratigraphic calibration 172 Mio-Pleistocene seismic units 172 seismo-stratigraphic units 162 subsidence 179 offshore prolongations 161 onshore tectonic structures 161 seismic lines and boreholes location 160 seismic profile crossing 178 seismic sequence stratigraphy 159 –180 classifications of faults 163 –165 faulting activity 163–165 Mio-Pleistocene structures 162–163 Pontian fault 163– 164 spatial faults 164–165 structure 159–180 subsidence history 159– 180 wells lithological description 168, 169 Romanian outer shelf seismic line crossing 167 Romanian-Quaternary deposits 167 Romanian-Quaternary section isopach map 166 Romanian shelves Black Sea sedimentary basin tectonics 4– 5 Russia Early Mesozoic evolution 181– 234 tectonic evolution of Black Sea 5– 6 Ryukyu subduction zone 15 SAB see South Armenian Block (SAB) Sahdag mountain cross-section 266 Sakar-Strandja Zone 93, 95, 104 palaeostress tensors 85 Sakar unit 82 Sakhrai Group 205 Samsari formation 254, 255 Sandstone turbidite succession 297 Santonian formations 334 Sapc¸a Formation 124 Sarız Fault 422 Schmidt’s projection microtectonic data 405 Scythian Platform 12– 13, 43, 49, 50, 53, 54, 138, 183 crust 55 rifting 230 SEAOB see Southeastern Anatolian orogenic belt (SEAOB) Sevan–Akera zone 329, 333, 343, 347, 355, 394 ophiolite 387 radiolarian age 389 structure 333 tectonic units 335–336

INDEX Sevan area 355 dolerites 359 gabbro 359 pillow lavas 359 serpentinites 359 whole-rock analyses 362–366 Sevan Lake area structural sketch map 341 Sevan ophiolite geological logs 369 sketch geological map 357 Sevlievo-Preslav unit 104 Sev Litsh lake site 399 Shatsky Ridge 113, 314 Shipka-Sliven unit 83 Shipka-Teteven unit 84 Shotpoints 45, 46 Sinemurian palaeogeographic maps 227 Sorokin Trough 27, 32, 34, 38, 39, 54, 57 characteristics 66 gas seeps 58 location 66 mud volcanoes 65, 67 Sosnov formation 200 South Armenian Block (SAB) 329, 331 –332, 353 Arabian plate collision 347 autochthon obduction modalities 336–339 continental subduction 344– 345, 348 faulting and folding 394 Lower Cretaceous hot spot magmatism 353– 380 metamorphic basement 339 stratigraphic and lithological data 339 structure 333 subduction zone 379 underthrusting 347 Upper Cretaceous obduction 353 –380 South Caspian Basin 276 Southeast Anatolia 409– 433 Arabian Platform stratigraphy 412–414 Late Cretaceous to recent kinematics 409–433 orogen stratigraphy 414 –417 Palaeostress inversion 414– 417 SE Anatolian orogen 420– 422 tectonostratigraphy 412 Southeastern Anatolian orogenic belt (SEAOB) 437 –457 Baskil-Divrigi transect geological framework 440–442 Baskil-Divrigi transect magmatic rocks 446– 453 geochemical characteristics 446–453 petrogenesis classification 451–453 tectonomagmatic classification 451– 453 trace element-REE geochemistry 448– 451 evolution of magmatism 442 –446 evolving orogen 437– 457 geodynamic setting 439–440 geodynamic-tectonic evolution 454– 456 magmatic sources 453 –454 magmatism and transition 437– 457 post-collisional setting in evolving orogen 437– 457 regional tectonic evolution 439– 440 slab break-off event 453 transition from arc to post-collisional setting 437– 457

507

Southeastern Moesian Platform palaeostress tensors 91 Southern Crimea Orogen 138 Southern Slope 251 volcanism 226 Spitak earthquake 399 rupturing process 400 Srednogorie Zone 106 polyphased tectonics 99 Shipka-Sliven unit 99 Stara Zagora unit 83 STEP see Subduction transform edge propagator (STEP) Stepanavan 355 radiolarian age 389 whole-rock analyses 362– 366 Stepanavan ophiolite 385 geological logs 369 sketch geological map 356 Strandja unit 82 Strandja Zone 85, 93, 95, 102, 104 Stress field analyses 92 Strike–slip 102, 104 Subduction transform edge propagator (STEP) faults 14, 18, 19, 410, 453 Sulina-Tarhankut Fault 146, 147 Sunik massif 396 Suprasubduction zone 441 Syn-rift erosion 154 Syn-rift phase 148 Tasmaca Formation 124, 125 Tauric basin 314 Tauride– Anatolides belt structural sketch map 330 Tauride– Anatolide– South Armenian Block 9 Tauride Platform 441 Tchataltapa Formation 192 –199, 210–211 Tchvejips formation 203 Tectonic subsidence curves 170 Teleseismic earthquakes global distribution 488 Tethys subduction zone 183, 322 Tithonian–Early Valanginian age 311 Tithonian limestone at Straja compression 104 Tithonian-Valanginian formation 106 Toarcian palaeogeographic maps 227 Tortum River cross-section 301 Transgressive systems tract (TST) 159, 174, 179, 180 Triassic Karakaya Complex 315 Triassic limestone 102 Triassic rifting 104 Tsalka formation 254 TST see Transgressive systems tract (TST) Tuapse Trough 25 Tundja-Topolnica unit 83 Turkey 492 see also East Anatolian Plateau active tectonics 411

508 Turkey (Continued) Cretaceous to recent tectonostratigraphical column 412 geological map 438 Late Cretaceous to recent kinematics 409– 433 Late Paleocene to Middle Eocene period 429 Oligocene-Middle Miocene extension 430 regional tectonic units 410 seismic experiment 485, 487 structures and extension direction 429 tectonic map 486, 487 Upper Miocene to Pliocene 430 Upper Pliocene to recent 431 Turovoy Formation 191–192 Tyrrhenian Sea 319 Ukrainian Odessa Shelf pseudo-wells and seismic lines position 138 Ulus Basin 121, 125–126 Unitary Association Zones 385 Upper Cretaceous Nannofossils dating 332 Upper Cretaceous obduction South Armenian Block 353–380 Upper Eocene Midyat Group 466 Upper Jurassic facies 311 Upper Jurassic Olur magmatic unit field relations and geochemistry 302 Upper Oligocene-Lower Miocene Firat Formation growth faults 419 Upper Slice Complex 283, 285 cross-section 301 geological map 286 Upper Jurassic-Lower Cretaceous sequences 306 Vanadzor depression 396 Vanadzor trough 396 outcrop 398 Vavilov Basin 319 Vedi area 355, 355 basal tectonic contact 359 geological cross-section 338 geological map 386 obduction contact 338 ophiolite section 359 pillows 359 radiolarian age 389 structural map 337 Upper Jurassic radiolaria chert sample 388 whole-rock analyses 362 –366 Vedi ophiolites amphibole 368 argon age spectra and isochrons 368 argon dating 367 gabbro amphiboles 367 geological cross-sections 340 geological logs 369 plutonic rocks 372 sedimentary cover 386 sketch geological map 358 Velibey Formation 123, 123 nannoplankton ages 125

INDEX Websterite 375 West Black Sea 23– 39, 53 hydrate-formation zone 71 rifting 134 West Black Sea Basin 23–39, 43 DSS data 31–32 earth’s crust and upper mantle structure 23– 39 gas seeps 66 methane in hydrates 72 mud volcanoes location 66 rifting 113, 132 stratigraphic sequences 113 Western Central Balkan-Forebalkan Zone palaeostress tensors 88– 90 Western Crimean fault heat flow and gas emission 64 Western Greater Caucasus 181–234 Aalenian and Bajocian isotopic elemental data 224 Aalenian succession sandstone modal data 209 Bajocian-Callovian succession sandstone modal data 210 Central Greater Caucasus Abino-Gunai Zone 200–201 Pshekha and Pikhtar formations 201 Zeitun Formation 201 Central Greater Caucasus Goitkh Zone 199– 200 Indyuk Formation 200 Kochkanov formation 200 Makarov Formation 199 Naujni Formation 199–200 Pchich Formation 200 Sosnov formation 200 Central Greater Caucasus Psekhako-Berezov Zone 191–199 Bzerpia Formation 192 Tchataltapa Formation 192–199 Turovoy Formation 191–192 Central Greater Caucasus summary 201– 202 Abkhazo-Racha Zone Porphyritovaya Series 203–204 Krasnopoliana Zone Anchkhoi Formation 203 Atchichkh Formation 203 Estasadok Formation 203 Illaryonov formation 203 Tchvejips formation 203 Southern slope Krasnaya Poliana area 202– 204 description 183 Early Mesozoic evolution 181– 234 Early Mesozoic Greater Caucasus basin development 184– 185 Early Mesozoic history 226– 234 Lower Jurassic 231– 232 Middle Jurassic 232–234 Triassic 229–230 Upper Jurassic 234 elemental isotopic ratios 225 elements variations plot 217 geodynamic evolution 225– 226 isotopic composition 209–225 Jurassic volcanism 225– 226 eruption dynamics 225– 226 Jurassic geodynamics 226

INDEX sources 226 spatio-temporal distribution 225– 226 Jurassic v. Triassic sandstones 207– 208 locations 184 magmatic history 181–234 magmatic signatures 221 major and trace element data 212– 216 palaeogeographic maps 227 petrofacies 207–209 region structure 184 –185 sedimentary and magmatic history 181– 234 siliciclastic formations petrography and composition 204–207 Central Greater Caucasus 206 Fore-Caucasus area Triassic and Jurassic 205– 206 Southern slope Krasnaya Poliana area 206–207 Sinemurian-Toarcian succession sandstone modal data 208 stratigraphic succession 185–191 Belaya River 187– 189 Fore-Caucasus area 187 –189 Laba-Malka Zone Bugunja Formation 189–190 Djangur Formation 190 Psebai Formation 190 Labina Zone Gerpegem formation 191 Kamennomostsky 190– 191 Mezmai formation 191

509

Malaya Laba 187–189 Mount Tkhach 187– 189 trace element ratios 221 Triassic-Jurassic sedimentary and magmatic history 181 –234 volcanics chemical composition 209– 225 major element compositions 209 –211 strontium and neodymium isotopic compositions 223– 225 trace element compositions 211– 223 Wu¨rmian glaciation bottom temperature 71 Yemis¸lic¸ay Formation 130 deposition 129 nannoplankton dating 126 syndepositional listric normal faults 127 Zangezur ophiolites 355 Zangezur zone 355 Zeitun Formation 201 Zeytinlik fault-bound volcanic interval 296 Zmeiniy Island 141 Zmeiniy Uplift 145 Zodi pass 385 Zonguldak 130 horsts 132 Zonguldak Basin 121, 122 –125, 122 local formation 128 nannofossils 122

This wide area of the Alpine –Himalayan belt evolved through a series of tectonic events related to the opening and closure of the Tethys Ocean. In doing so it produced the largest mountain belt of the world, which extends from the Atlantic to the Pacific oceans. The basins associated with this belt contain invaluable information related to mountain building processes and are the locus of rich hydrocarbon accumulations. However, knowledge about the geological evolution of the region is limited compared to what they offer. This has been mainly due to the difficulty and inaccessibility of cross-country studies. This Special Publication is dedicated to the part of the Alpine –Himalayan belt running from Bulgaria to Armenia, and from Ukraine to the Arabian Platform. It includes twenty multidisciplinary studies covering topics in structural geology/tectonics; geophysics; geochemistry; palaeontology; petrography; sedimentology; stratigraphy; and subsidence and lithospheric modelling. This volume reports results obtained during the MEBE (Middle East Basin Evolution) Programme and related projects in the circum Black Sea and peri-Arabian regions.