Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic (Geological Society Special Publication 341)

Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic

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 )

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It is recommended that reference to all or part of this book should be made in one of the following ways: HOMBERG , C. & BACHMANN , M. (eds) 2010. Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341. GARDOSH , M. A., GARFUNKEL , Z., DRUCKMAN , Y. & BUCHBINDER , B. Tethyan rifting in the Levant Region and its role in Early Mesozoic crustal evolution, In: HOMBERG , C. & BACHMANN , M. (eds) 2010. Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 9–36.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 341

Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic

EDITED BY

C. HOMBERG University Pierre et Marie Curie, France

and M. BACHMANN University of Bremen, Germany

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 10000. 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 978-1-86239-306-6 Typeset by Techset Composition Ltd, Salisbury, UK Printed by MPG Books Ltd, Cornwall 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]

Preface This volume belongs to a series of four regional Geological Society of London Special Publications presenting the results of the Middle East Basin Evolution (MEBE) Programme working groups and associated research teams. MEBE was a 4year consortium (2003– 2006) funded by several oil companies (BP, ENI, PETRONAS, SHELL, and TOTAL), the French Research Organization INSU-CNRS, as well as the French University Pierre and Marie Curie (UPMC). This programme was a multi-disciplinary study of the Middle East, spanning the Arabian –Peri-Arabian and Caucasian –Caspian areas and a detailed presentation can be found in the Preface of a companion volume (Brunet et al. 2009). The four MEBE volumes cover the Black Sea –Caucasus (Sosson et al. 2010), the South Caspian –Central Iran (Brunet et al. 2009), the Zagros –East Arabian

margin (Leturmy & Robin 2010), and the Levant (this volume). The Levant volume combines new data coming from several working groups. In 2005, MEBE workers joined colleagues from various institutes and countries leading ongoing research in the Levant. Several workshops occurred and the last one was held in the University Pierre and Marie Curie, Paris, on 14– 15 December, 2006. This volume stems from scientific presentations and active discussions arose during this last Levant Meeting. The editors would like to thank colleagues who choose to publish their data in this volume, as well as others who kindly helped us to review the papers submitted to the volume. The following companies are thanked for funding several research projects under the frame of the MEBE Programme and whose results are shown here.

References BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) 2009. South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312. LETURMY , P. & ROBIN , C. (eds) 2010. Tectonic and Stratigraphic Evolution of Zagros and Makran during the Meso– Cenozoic. Geological Society, London, Special Publications, 330.

SOSSON , M., KAYMAKCI , N., STEPHENSON , R. A., STRAROSTENKO , V. & BERGERAT , F. (eds) 2010. Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform. Geological Society, London, Special Publications, 340.

CATHERINE HOMBERG & MARTINA BACHMANN

Contents Preface

vii

HOMBERG , C. & BACHMANN , M. Evolution of the Levant margin and western Arabia platform since the Mesozoic: introduction

1

GARDOSH , M. A., GARFUNKEL , Z., DRUCKMAN , Y. & BUCHBINDER , B. Tethyan rifting in the Levant Region and its role in Early Mesozoic crustal evolution

9

MOUSTAFA , A. R. Structural setting and tectonic evolution of North Sinai folds, Egypt

37

YOUSEF , M., MOUSTAFA , A. R. & SHANN , M. Structural setting and tectonic evolution of offshore North Sinai, Egypt

65

CAME´ RA , L., RIBODETTI , A. & MASCLE , J. Deep structures and seismic stratigraphy of the Egyptian continental margin from multichannel seismic data

85

BACHMANN , M., KUSS , J. & LEHMANN , J. Controls and evolution of facies patterns in the Upper Barremian–Albian Levant Platform in North Sinai and North Israel

99

FRANK , R., BUCHBINDER , B. & BENJAMINI , C. The mid-Cretaceous carbonate system of northern Israel: facies evolution, tectono-sedimentary configuration and global control on the central Levant margin of the Arabian Plate

133

WENDLER , J. E., LEHMANN , J. & KUSS , J. Orbital time scale, intra-platform basin correlation, carbon isotope stratigraphy and sea-level history of the Cenomanian–Turonian Eastern Levant platform, Jordan

171

MORSI , A.-M. M. & WENDLER , J. E. Biostratigraphy, palaeoecology and palaeogeography of the Middle Cenomanian– Early Turonian Levant Platform in Central Jordan based on ostracods

187

JOSEPH -HAI , N., EYAL , Y. & WEINBERGER , R. Mesoscale folds and faults along a flank of a Syrian Arc monocline, discordant to the monocline trend

211

COLLIN , P.-Y., MANCINELLI , A., CHIOCCHINI , M., MROUEH , M., HAMDAM , W. & HIGAZI , F. Middle and Upper Jurassic stratigraphy and sedimentary evolution of Lebanon (Levantine margin): palaeoenvironmental and geodynamic implications

227

HOMBERG , C., BARRIER , E., MROUEH , M., MULLER , C., HAMDAN , W. & HIGAZI , F. Tectonic evolution of the central Levant domain (Lebanon) since Mesozoic time

245

HENRY , B., HOMBERG , C., MROUEH , M., HAMDAN , W. & HIGAZI , F. Rotations in Lebanon inferred from new palaeomagnetic data and implications for the evolution of the Dead Sea Transform system

269

MU¨ LLER , C., HIGAZI , F., HAMDAN , W. & MROUEH , M. Revised stratigraphy of the Upper Cretaceous and Cenozoic series of Lebanon based on nannofossils

287

AL ABDALLA , A., BARRIER , E., MATAR , A. & MULLER , C. Late Cretaceous to Cenozoic tectonic evolution of the NW Arabian platform in NW Syria

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Index

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Evolution of the Levant margin and western Arabia platform since the Mesozoic: introduction CATHERINE HOMBERG1* & MARTINA BACHMANN2 1

University Pierre et Marie Curie, ISTEP, Case 129, 4 Place Jussieu, 75252 Paris Cedex 05, France

2

University of Bremen, Department of Geosciences, P.O. Box 330440, 28334 Bremen, Germany *Corresponding author (e-mail: [email protected])

Abstract: The Levant area comprises the offshore Levant Basin (LB) (eastern corner of the Eastern Mediterranean) as well as the adjacent continental slopes and platforms of the African and Arabian plates. This area experienced major events of the geodynamical evolution of the Middle East, such as the Late Palaeozoic– Early Mesozoic Pangea break up, the Late Cretaceous – Cenozoic closure of the Neo-Tethys and individualization of the Arabian plate, as well as a set of external factors like global sea-level and climate changes. This volume combines original data from the offshore and onshore Levant in various fields, including sedimentology, palaeontology, sequence stratigraphy, geochemistry, structural geology, stress reconstitution and geophysics (seismic lines, palaeomagnetism). All together, these multidisciplinary approaches allow the review of the development of the LB and gain a better insight on the later geological history and deformation processes of the Levant provinces.

The Levant describes a composite area that includes (1) the offshore Levant Basin (LB) (eastern corner of the Eastern Mediterranean) in the west, (2) the Afro –Arabian continental slope and platform, which is today largely emerged along the eastern Mediterranean coast of Egypt, Israel, Lebanon and Syria, and in Jordan and disrupted by the Dead Sea Fault and (3) the Cyprian Arc and NW Syria collision zones that mark the southern boundary of the Eurasian plate (Fig. 1). During the last decades, acquisition of geophysical and geological data allowed imaging of the crustal structures and the sedimentary infill of the LB. Together with detailed tectonic and sedimentary field investigations in the onshore Levant, these studies discover the basin forming and filling processes. Modern interpretations indicate that the LB resulted from rifting, but much controversy exists on several aspects, such as the age and kinematics of its opening or the nature of its crust. The Late Cretaceous to Present-day period includes several discrete, more or less widespread, tectonic events in relation with the interaction of the Arabian, African and Eurasian plates that resulted in the division of the Levant into several provinces. Depending on the areas, tectonic structures and timing of deformation are more or less constrained. The sedimentation patterns observed in the Meso-Cenozoic sequences of the LB and

surroundings correlate in a high degree with the tectonic processes. However, various further factors acted on top, such as global sea-level or climate changes and the recognition of their relative contribution to the sedimentation processes is a still not finished puzzle. This Special Publication of the Geological Society of London combines new multidisciplinary datasets acquired recently onshore and offshore Levant. A special aim of the volume is to demonstrate new concepts on the tectonic, stratigraphic, sedimentologic and environmental evolution of the LB and African and Arabian platforms. Periods considered cover both the Late Palaeozoic –Early Mesozoic phase of basin development and the Late Cretaceous to Cenozoic deformation period associated with the closure of Neo-Tethys and the incipience of the Dead Sea Transform.

Middle East Basin Evolution Levant Group and Others This special volume contains new data coming from several working groups. From 2003 to 2006, the Middle East Basin Evolution (MEBE) Programme funded multidisciplinary studies in the Middle East, some of them devoted to the Levant. After field data acquisitions, a working group, the

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 1 –8. DOI: 10.1144/SP341.1 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Situation of the Levant and surroundings. Digital topography model SRTM30 from NASA Shuttle Radar Topography Mission, 30 arc-second grids mosaic (http://www.jpl.nasa.gov.srt). Numbers indicate approximate location of areas investigated in this volume. 1, Gardosh et al.; 2, Moustafa; 3, Yousef et al.; 4, Camera et al.; 5, Bachmann et al.; 6, Frank et al.; 7, Wendler et al.; 8, Morsi & Wendler; 9, Hai et al.; 10, Collin et al.; 11, Homberg et al.; 12, Henry et al.; 13, Mu¨ller et al.; 14, Al Abdalla et al. Inset shows geodynamical frame. Ar, Eu and Nu: Arabia, Eurasia, Nubia plates. DSF and EAF: Dead Sea Fault and East Anatolian Fault. RS, CA and CZ: Red Sea rift, Cyprian Arc and Arabia– Eurasia collision zone.

‘Levant Group’ led by C. Homberg, Universite´ Pierre et Marie Curie, was created in which MEBE workers joined colleagues from various institutes and countries leading ongoing research in the Levant. Several workshops occurred and the last one was held in the University Pierre and Marie Curie, Paris, on 14 –15 December, 2006. This volume stems from scientific presentations and active discussions that arose during this last Levant meeting.

General evolution of the Levant It is generally accepted that the LB formed as a result of rifting during Early Mesozoic time, which started perhaps in the Late Palaeozoic period. This is supported by crustal thinning from 30 –35 km on the African and Arabian continents to c. 8 km below the LB overlain by a 10–14 km thick sedimentary pile of Mesozoic (and Permian?) to present age (Makris et al. 1983; Ginzburg &

EVOLUTION OF THE LEVANT MARGIN AND WESTERN ARABIA PLATFORM

Ben-Avraham 1987; Vidal et al. 2000; BenAvraham et al. 2002). The affinity of the crust below the LB is regarded either as a highly stretched continent (Woodside 1977; Hirsch et al. 1995; Vidal et al. 2000) or as oceanic (Makris et al. 1983; Ginzburg & Ben-Avraham 1987; Ben-Avraham et al. 2002). Shallow-marine conditions prevailed in the basin during Triassic and Liasic times, later followed by the formation of the deep-marine basin (Bartov & Steinitz 1977; Garfunkel 1998). These marine conditions were interrupted by a strong period of erosion during the latest Jurassic – Neocomian time, with huge and widespread volcanic activity and deposition of terrestrial, fluvio-deltaic sediments, which became very thick in some places (e.g. Said 1971; Bartov & Steinitz 1977; Hirsch 2005a). The basin and its surroundings went through intense normal faulting during this period. The most common trend recognized in many places is NE– SW (Cohen et al. 1990; Garfunkel 1998; Vidal et al. 2000; Gardosh & Druckman 2006) but others like NW –SE structures exist as well (Garfunkel & Derin 1984; Homberg et al. 2009). Kinematic models (Dercourt et al. 1986; Stampfli & Borel 2002) predict a north–south opening of the basin and some difficulties arise in adaptation of these models with the local extensional fabrics of the basin and surroundings. Another characteristic of the area is the long-lived rift activity, which likely included several rifting pulses, and various ages for the opening of the basin have been proposed, from Triassic or Late Permian to Cretaceous (Freund et al. 1975; Dercourt et al. 1986; Garfunkel 1998; Ben-Avraham et al. 2002; Stampfli & Borel 2002). General coastal onlap and the development of shallow marine facies belts is a main feature of the Levant platform starting during the late Early Cretaceous period, overprinted by second- and third-order relative sea-level changes (Said 1971; Kuss et al. 2003; Rosenfeld & Hirsch 2005). Until the Mid-Albian, second-order sea-level changes generally did not coincide with those ones observed in other parts of the Tethys, but do in parts with those ones documented from other parts of the Arabian plate, suggesting that the southern Levant regional sea-level history was triggered by the interaction of subsidence and sediment supply, with an increasing eustatic influence since the Mid-Albian (Bachmann et al. 2003). A general flooding of the Levant continued during Cenomanian to Senonian time (e.g. Bartov & Steinitz 1977; Kuss et al. 2003). Since the Mid-Cenomanian a palaeorelief developed at several localities on the southern Levant platform and the subsidence control in relation with extensional tectonics is discussed by several authors

3

(e.g. Bauer et al. 2003; Schulze et al. 2005). Since the Senonian period sedimentation was controlled by a variety of tectonic factors. In some places, large grabens developed like in the interiors of the Arabian plate (Euphrates and Azraq grabens), while other areas underwent a strong inversion. The last one is referred to as the Syrian Arc tectonism that produced broad folds and high-angle reverse faults in southern Levant and northern Africa (Israel, Egypt and offshore of the Mediterranean coast), associated with breaks in deposition, angular unconformities and lateral thickness variations within sedimentary sequences (e.g. Bartov et al. 1980; Mimran 1984; Cohen et al. 1990; Lewy 1991; Moustafa & Khalil 1994; Lu¨ning et al. 1998; Gardosh & Druckman 2006). Although the age of faulting is still debated (see references above and Shalar 1994 for a review), there is a general consensus that the Syrian Arc tectonism includes two episodes, a main one during Senonian time and a second during Palaeogene time (possibly continuing up into Early Miocene time). The curved east–west to NE –SW belt is assumed to continue northward into Lebanon and Palmyrides. Further to the north, the Tethyan ophiolites were obducted onto the Arabian platform in Late Cretaceous time and therefore now outcrop occurs in Cyprus and NW Syria. During the Neogene, individualization of the Arabia plate affected the Levant with the development of the Dead Sea (or Levant) Fault (DSF) that runs almost parallel to the eastern Mediterranean coast with a general north– south trend and connects the Red Sea Rift in the south to the collision zone in the north. The c. 1000 km transform plate boundary includes a major NNE–SSW restraining bend in its central part (Lebanon –Palmyrides), which is regarded either as primary (Butler et al. 1998) or as resulting from a later clockwise rotation (Quennel 1984). Since recognition of the 70 – 110 km left-lateral displacement along the DSF (Quennel 1958; Freund et al. 1970), continuous progress has been made in the identification of the onshore and offshore plate tectonic structures and understanding of their relationships. These tectonic features resulted in latest Cretaceous to Neogene sedimentary processes, marked by small-scale structures and a pronounced palaeorelief with basins, sub-basins, swells and local unconformities, observable in many parts of the Levant. A great range of sediments reach from outer shelf to deltaic-fluviatile deposits (e.g. Buchbinder et al. 2005; Hirsch 2005b; Rosenfeld & Hirsch 2005; Kuss & Boukhary 2008).

Main outline of the papers Papers in this volume deal with the LB, adjacent deformed margins, and platforms (Fig. 1). They

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are organized in three sections according to three main study areas, namely the offshore deep basin and its margins, the southern Levant (Egypt, Israel, Jordan), and the northern Levant (Lebanon, Syria). Papers concern one or several of the three key time-periods of the Mesozoic and Cenozoic evolution of the Levant. These periods can be described as follows: (1) rifting and development of the LB since Late Palaeozoic to Early Cretaceous time; (2) later evolution of the basin; and (3) inversion and deformation of the Levant from Late Cretaceous to Present.

Rifting and tectono-stratigraphic sequences of the LB and margins The papers by Gardosh et al., Camera et al. and Yousef et al. present new subsurface data on the southern part of the LB near the Israeli and Egyptian coasts. Gardosh et al. document normal faults onshore and offshore Israel with vertical offsets of several kilometres, bordering NE –SW and NNE– SSW grabens and horsts. These structures occupy a several hundred kilometre wide deformation zone that extends from the inner part of the Levant to the marine basin offshore Israel. Gardosh et al. also show that faulting activity in this area progressed over a period of 120 Ma and took place in three main pulses: Late Palaeozoic (Carboniferous to Permian), Middle to Late Triassic and Early to Middle Jurassic. The last event was the most intense rifting phase. The authors demonstrate an extension discrepancy between the brittle deformation in the upper crust and the amount of total crustal thinning and explain the basin evolution by a depth-dependent stretching, associated with mantle upwelling and removal of lower crustal layers, by decoupling along deep detachment faults. Moustafa describes similar Early Mesozoic rifting structures onshore in northern Sinai. In addition to a gradual northward increase in thickness of the Mesozoic and Cenozoic rocks, abrupt changes in thickness of the Mesozoic rocks (especially documented from the Jurassic) are reported from seismic profiles and boreholes data. Moustafa identifies several sub-basins in North Sinai with Jurassic thicknesses up to c. 3000 m. He underlines the role of the Sinai hinge belt that extends in an ENE– WSW direction from the northern tip of the Gulf of Suez toward the DSF and represented at this period the southern boundary of the tectonically active area to the north. Camera et al. and Yousef et al. recognize a second major rifting phase during Late Jurassic – Early Cretaceous in the light of sub-subsurface data along the Egyptian margin. Yousef et al. show a detailed study of the offshore area of North Sinai in a sector crossed by c. 6 km seismic reflection sections. They document NE-trending normal faults

bounding asymmetrical half-grabens. The syn-rift package consists predominantly of clastic – dominated successions of Early Cretaceous age, unconformably overlying the shallow marine Upper Jurassic carbonates. Seismic profiles shown by Camera et al. are located several hundred kilometres farther north from the Egyptian coast in the deep basin. Camera et al. combine different processing techniques to image the deep structures. The seismic stratigraphy of the area is interpreted in terms of three main tectono-stratigraphic units; a deep unit with low frequency discontinuous reflectors, a middle on average 6 km thick unit with highly faulted reflectors, and an upper 7 km thick and well-layered unit with continuous horizons (but locally disrupted by gravity structures). They are respectively interpreted as the crustal basement, the syn-rift package, and the post-rift sedimentary cover of the basin. A strong angular unconformity marks the end of the last rifting event in this area and an Aptian age is proposed for this end by correlation of the seismic reflectors with ODP data near Eratosthenes seamount. The continental crust is estimated to be 9 to 12 km thick in this sector of the LB and a 23–25 km depth below sea floor of the Moho is proposed. Based on field data, Collin et al., Homberg et al. and Bachmann et al. characterize a similar tectonic influence on the shallow marine sedimentation and the platform architecture in the onshore northern (Lebanon) and southern Levant (Sinai, Egypt) during Late Jurassic and Early Cretaceous times. Collin et al. present sedimentologic, facies, and biostratigraphic analyses of well exposed sedimentary rocks from Lebanon and discuss the relationships between the sedimentological and tectonic evolution during Middle to Late Jurassic time. Their stratigraphical data allow for the timing of the tectonic events. They first evidence the development of a stable epicontinental shelf, with shallow marine environment extending across large parts of the Lebanon, during Bathonian–earliest Kimmeridgian times and argue for a tectonic quiescence during this period. A thick sediment package accumulated owing to intense subsidence. A regional unconformity, regression and block faulting associated with a volcanic event was recognized and dating of sediments below and above suggests a Kimmeridgian age (Middle– Late Jurassic) for the beginning of this rifting phase. Following this, a Kimmeridgian to Tithonian (Late Jurassic) marine transgression induced the development of a new shallow marine carbonate shelf. Lateral thickness variations during this period are interpreted as resulting from the platform morphology formed by the previous block faulting. These marine successions were ended by regression, resulting in erosion and the deposition of continental sandstones of the basal Cretaceous.

EVOLUTION OF THE LEVANT MARGIN AND WESTERN ARABIA PLATFORM

Homberg et al. document syn-sedimentary WNW– ESE to WSW–ENE normal faults active from Hauterivian (maybe earlier) to Albian times in Lebanon, with offsets of up to several hundreds of metres and slipping under a NNE–SSW extension. This Early Cretaceous extensional phase largely controlled the sedimentologic architecture in Lebanon, especially concerning the Neocomian sequence thickness, which reaches its maximum in Central Lebanon whereas it is strongly reduced in northern Lebanon and replaced by lava flows. The authors propose that the WNW –ESE intra-Lebanon basin is a result of an Early Cretaceous rifting phase in the Levant. This event displays the continuation and acme of the precursor Late Jurassic movements described by Collin et al. That the rifting phase was active until the Early Aptian in northern Sinai (Egypt) is discussed by Bachmann et al. on the base of facies, stratigraphical and sequence stratigraphical data from the Barremian to Albian shallow-marine succession in northern Sinai and northern Israel. They indicate the presence of small-scale sub-basins in northern Sinai until the Early Aptian characterized by significant sediment thickness changes and local normal faults. The study of time equivalent sections in northern Israel miss significant influence of extensional tectonic, which is interpreted as a local restriction of tectonic movements during the late Early Cretaceous. Furthermore, the paper investigates the complexity of factors controlling the sedimentation on the Late Barreminan to Albian southern Levant platform. A detailed stratigraphic concept based on benthic foramifers, ammonites and chemostratigraphy is used for detailed dating of the individual processes. Besides the extensional tectonics, regional and global second-order sea-level changes are determined and dated. The driving processes for the observed sedimentation patterns are interpreted. The tectonic frame defined the start-up architecture of the southern Levant Cretaceous platform, which underwent significant changes by the interaction of tectonics, sea-level changes and sedimentation rates. The sedimentation patterns and rates have a major influence on the southern Levant platform geometry. In northern Sinai, a transition from a shallow-shelf structured by sub-basins through a homoclinal ramp into a flat topped platform is recognized, while the sections in northern Israel show a transition from a homoclinal ramp into a fringing platform.

Post-rift evolution of the basin and adjacent areas The Levant platform was influenced by a set of external factors, like climate or sea-level changes, which can be a result of global changes and orbital cycles. The relatively quiet post-rift tectonic frame

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seemed to be an optimal phase for the investigation of those factors and their influences. The post-rift evolution of the southern Levant Platform is investigated in three papers. That tectonic quietness does not match the reality is shown by Frank et al. They focus on the Cenomanian –Turonian development of northern Israel. By detailed sedimentologic and facies analyses of several sections, they described the cyclic sedimentation patterns and the hierarchical organization of those cycles. They document proximalto-distal facies and thickness changes, and the geometry of genetic –stratigraphic sequences. These datasets are interpreted and authors show that the tectono-sedimentary regime of northern Israel represents an east– NE branch-off of the depositional strike from the north –south-striking Levant margin. Frank et al. indicate several productive and demise phases of the Cenomanian – Turonian carbonate ramp and show that the ramp development was strongly influenced by eustatical and palaeoceanographical trends, similar to those observed in the Tethys. Late Cenomanian normal faults and a latest Cenomanian sequence boundary, marked by subaerial exposure, reflect tectonic activity and uplift of the Galilee during this period, which was generally regarded as tectonically quiet. The papers of Morsi & Wendler and Wendler et al. analyse shallow marine platform successions from an intra-platform basin that developed on the southern Levant platform (central Jordan). They highlight another important event characterizing the Mesozoic history: marine sediments, rich in organic matter, occurred during several Jurassic and Cretaceous intervals. They reflect a significant disturbance in the global carbon cycle related to oceanic anoxic events (OAEs). While the effects of OAEs in pelagic to hemipelagic deposits have been studied intensively, the palaeoenvironmental conditions related to these significant events in the near-shore environments still need to be investigated in more detail. Morsi & Wendler use stratigraphic and palaeontologic investigations of ostracods to complete age dating of the midCenomanian through Lower Turonian shallowmarine succession and to analyse palaeoecology and palaeoenvironmental conditions related to OAE2 at the Cenomanian –Turonian (C –T) transition interval and to study the palaeobiogeographical configurations. A detailed taxomical and detailed palaeontologic part forms a broad base for all interpretations. Wendler et al. concentrate on the influence of orbital cycles on the sedimentation during the C –T interval and on stratigraphic aspects concerning the OAE2. A high-resolution stratigraphy is presented by the use of high-resolution calcimetry and stable carbon isotope stratigraphy. They use these methods to correlate the successions to the

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global record (exemplified for the well-dated C –T boundary interval) in order to eliminate ambiguities in the local stratigraphy. They refine the sequencestratigraphic model by constructing an orbital time scale and present a model of orbitally forced sequences. Finally, they indicate that the Jordan platform demise was later than the onset of the OAE2 and discuss the role of biological factors for demise of platform carbonate producers.

Late Cretaceous to Cenozoic deformation in the Levant Moustafa and Yousef et al. discuss the inversion of the southern Levant margin that began in Late Cretaceous time. New data presented by these authors provide details on the development of the Egyptian segment of the Syrian Arc. Moustafa shows results of a detailed mapping and structural analysis of North Sinai. North Sinai comprises a folded zone dominated mainly by NE– SW doubly plunging anticlines, as well as two fault belts, the ENE– WSW Sinai hinge belt and the east– west themed fault. Faults of the area show a preferred NW to WNW orientation, but other trends, like NNW and ENE ones, also exist in the hinge belt. Mechanism of deformation in this area includes reactivation of deep faults bounding Early Mesozoic basins (see previous section) associated with folding, as well as dextral shear with more or less pronounced transpression. This major NW–SE compression started in north Egypt in Late Cretaceous, reached its acme in Campanian time and ended before Early Miocene time. Moustafa also recognizes post-Early Miocene fault reactivation in the Sinai hinge belt and proposes a genetic relation between this Cenozoic deformation and the sinistral slip on the DSF. Yousef et al. observe similar inversion structures in offshore Sinai. They are the Mango, Goliath, North Sinai and Ziv structures. They correspond to NE–SW asymmetrical folds that most commonly show a steep northwestern flank bounded by a steeply dipping fold progation fault. Most of these faults show reverse offsets at their upper ends whereas the deeper layers increase in thickness towards the main bounding fault and are offset in a normal motion (see previous section). These observations indicate that the offshore Sinai structures correspond to inverted Late Jurassic –Early Cretaceous half-grabens. Yousef et al. evidence progressive onlaps on the crest of the elevated structures and data in boreholes cutting the syncompression sequence indicate a Campanian– Maastrichtian age for the inversion. Inversion continued in a middle way until Late Oligocene. A few WSW–ESE normal faults cut the younger sequences and are thought to form in response

to gravity driven extension in Miocene and post-Miocene times. New data on the tectonic an sedimentologic evolution of southern and northern Levant, in Israel, Lebanon and Syria are presented in Al Abdalla et al., Hai et al., Henry et al., Homberg et al. and Mu¨ller et al. Al Abdalla et al. combine a brittle tectonic analysis and stratigraphic study of the NW Arabia platform. They demonstrate that an extensional context prevailed in NW Syria at Senonian time with the development of meso-scale normal faults. The authors propose that this NE– SW extension is associated with the development of the Euphrate graben to the east. Investigation of the Senonian package of Lebanon by Homberg et al. suggests that little or no major compressive deformation occurred there in Late Cretaceous time. Al Abdalla et al. and Homberg et al. also recognize syn-sedimentary normal faults in the Eocene sediments of Syria and Lebanon and fault analysis indicates that they slipped under a NNE – SSW extension. According to these observations, the tectonic evolution of northern Levant implies that the Mesozoic extensional context continued here until Late Eocene, and maybe Oligocene time, with a short interruption during Maastrichtian time when the Tethyan Ophiolites were obducted to the south onto the Arabian platform. Al Abdalla et al. shows that this event is marked in the Maastrichtian sequence of NW Syria by intra-formational unconformities. Al Abdalla et al. detail the Cenozoic tectonic structures and palaeostress evolution in NW Syria. A major phase of shortening occurred here in the Early Miocene. During this NW–SE compresion event, the Baer-Bassit thrusted over the Coastal Range platform along the SE vergent Lattakia Thrust. This major thrusting induced the flexure of the Arabian platform and the formation of the Middle to Late Miocene Lattakia Basin. From the end of the Miocene, and until Present, the region experienced a NNW–SSE directed regional compression. A coeval east –west trending compression associated with the north– south folding of the Coastal Range was also recognized. The authors suggest that it probably corresponds to a stress-field deflection in relationship with the DSF activity. The left-lateral displacement along the northern segment of the DSF since latermost Miocene is estimated to 30 –40 km. Homberg et al. present a re-evaluated tectonic history of Lebanon based on new field observations of tectonic structures. Lebanon experienced two compressional tectonic events during the Cenozoic period. A first east–west compression produced moderate folding and faulting during Early Miocene, probably as a far-field effect of the ongoing collision in the north between Afro-Arabia

EVOLUTION OF THE LEVANT MARGIN AND WESTERN ARABIA PLATFORM

and Eurasia. A second, but much more severe, folding event occurred during Late Miocene time owing to a NNW–SSE compression. The authors interpret this stress reorganization during the Neogene as the signature of the initiation of the transform tectonics in Lebanon, with a pronounced early transpresssional character. A maximum Late Langhian age of the Yammouneh Fault, the main branch of the DSF in Lebanon, is given. Henry et al. present results from a palaeomagnetical analysis in Lebanon. They document two palaeomagnetical directions in Aptian– Albian formations from widespread sites, one of them corresponding to the primary magnetization and the second being acquitted after folding. Comparison of these data with previous palaeomagnetical results for the Jurassic age in Lebanon, and expected directions from African apparent polar wander path, yields evidence of three different counter-clockwise regional rotations: 338 before Aptian deposition, 118 during Late Miocene time, and 178 in postMiocene times. The regional Cenozoic 288 counterclockwise rotation observed in Lebanon rules out the possibility that the Yammouneh fault originally trended north– south similar to the other main fault branches of the DSF. According to the authors, the origin of the deflection of the fault in Lebanon is to be found in crustal rheological constrasts within the Levant, probably formed during the Mesozoic extension. A structural model of the central segment of the DSF is proposed. In this model, about 35% of the transform plate motion was first accommodated during Late Miocene time, mostly by folding in Lebanon and Palmyrides and was then later transferred onto the Yammouneh and associated faults. The first phase is associated with 118 of the counterclockwise rotation and 178 occurred during the second phase. Mu¨ller et al. studied calcareous nannofossils from predominantly marly Senonian– Maastrichtian, Paleogene, and Neogene series from the central Levant platform in Lebanon. This lithological homogenous stratigraphic interval was poorly subdivided on the existing geological maps. Palaeocene, Upper Eocene, Upper Oligocene and Lower Miocene units are identified for the first time in Lebanon. The study presented by Mu¨ller et al. contains the precise determination of hiatus and tectonic events. They document that the relatively complete marine succession was interrupted by an extended Late Cretaceous (Coniacian and Early Santonian) hiatus and several regional hiatus at the top of the Maastrichtian to lowermost Palaeocene, the Lower Oligocene, and the lower part of the Lower Miocene (Aquitanian –lowermost Burdigalian) sequences. Hai et al. combine a fold and fault analysis in Israel along the western limb of the Ramallah

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monocline. Studied folds in Turonian and Senonian rocks have wavelengths ranging between 25 –100 m and amplitudes between 10–30 m and show two superposed axis trends. The minority of the folds are aligned NNE–SSW and are thus compatible with the WNW –ESE shortening trend of the Syrian Arc fold belt. In the contrary, the majority of measured folds are aligned ENE– WSW and are not compatible with this shortening trend. Kinematic analysis of fault attitude indicates NNW– SSE shortening and ENE –WSW extension in accordance with the shortening of the majority of folds. Based on several arguments, authors argue that folds resulted from tectonic shortening and were not formed owing to karstic activity or other processes. They propose that folds compatible with the main trend of the Ramallah monocline are parasitic small folds within the Syrian Arc fold belt whereas the other major fold trend is associated with Miocene to Recent movement along the DSF.

References Bachmann, M., Bassiouni, M. A. A. & Kuss, J. 2003. Timing of mid-Cretaceous carbonate platform depositional cycles, northern Sinai, Egypt. Palaeogeography, Palaeoclimatology, Palaeoecology, 200(1– 4), 131– 162. Bartov, Y. & Steinitz, G. 1977. The Judea and Mount Scopus Groups in the Negev and Sinai with Trend Surface Analysis of the thickness data. Israel Journal of Earth Science, 28, 119– 148. Bartov, Y., Lewy, Z., Steinitz, G. & Zak, I. 1980. Mesozoic and Tertiary stratigraphy, paleogeography and structural history of the Gebel Areif en Naqa area, eastern Sinai. Israel Journal of Earth Science, 29, 114 –139. Bauer, J., Kuss, J. & Steuber, T. 2003. Sequence architecture and carbonate platform configuration (Late Cenomanian-Santonian), Sinai, Egypt. Sedimentology, 50, 387 –414. Ben Avraham, Z., Ginzburh, A., Makris, J. & Eppelbaum, L. 2002. Crustal structure of the Levant Basin, eastern Mediterranean. Tectonophysics, 346, 23–43. Buchbinder, B., Calvo, R. & Siman-Tov, R. 2005. The Oligocene in Israel: a marine realm with intermittent denudation accompanied by mass-flow deposition. Israel Journal of Earth Sciences, 54, 63– 85. Butler, R. W. H., Spencer, S. & Griffiths, H. M. 1998. The structural response to evolving plate kinematics during transpression: evolution of the Lebanese restraning bend of the Dead Sea Transform. In: Holdsworth, R. E., Strachan, R. A. & Dewey, J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 81– 106. Cohen, Z., Kapstan, V. & Flexer, A. 1990. The tectonic mosaic of the southern Levant: implications for hydrocarbon prospects. Journal of Petroleum Geology, 13, 437– 462.

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Dercourt, J., Zonenshai, L. et al. 1986. Geological evolution of the Tethys belt from the Atlantic to the Pamirs since the Lias. Tectonophysics, 123, 241–315. Freund, R., Garfunkel, Z., Zak, I., Goldberg, M., Weissbrod, T. & Derin, B. 1970. The shear along the Dead Sea Rift. Philosophical Transactions of the Royal Society, London, 267, 107–130. Freund, R., Goldberg, M., Weissbrod, T., Druckman, Y. & Derin, B. 1975. The Triassic-Jurassic structure of Israel and its relation to the origin of the Eastern Mediterranean. Bulletin Geological Survey of Israel, 65, 26. Gardosh, M. A. & Druckman, Y. 2006. Seismic stratigraphy, structure and tectonic evolution of the Levantine Basin, offshore Israel. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 201–227. Garfunkel, Z. 1998. Constrains on the origin and history of the Eastern Mediterranean basin; collision-related processes in the Mediterranean region. Tectonophysics, 298, 5– 35. Garfunkel, Z. & Derin, B. 1984. Permian-early Mesozoic tectonism and continental margin formation in Israel and its implication for the history of the Eastern Mediterranean. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 187–201. Ginzburg, A. & Ben-Avraham, Z. 1987. The deep structure of the central and southern Levant continental margin. Annales Tectonicae, 1, 105– 115. Hirsch, F. 2005a. The Jurassic of Israel. In: Hall, J. K., Krasheninnikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant. Volume II: The Levantine Basin and Israel. Jerusalem: Historical Productions-Hall, Jerusalem, 361– 391. Hirsch, F. 2005b. The Oligocene-Pliocene of Israel. In: Hall, J. K., Krasheninikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant (II): The Levantine Basin and Israel. Historical Productions, Hall Publications, Jerusalem, Israel, 459– 488. Hirsch, F., Flexer, A., Rosenfeld, A. & Yellin-Dror, A. 1995. Palinspatic, crustal studies of the eastern Mediterranean. Journal of Petroleum Geology, 18, 149– 170. Homberg, C., Barrier, E., Mroueh, M., Hamdan, W. & Higazi, F. 2009. Basin tectonics during the Early Cretaceous in the Levant margin, Lebanon. Journal of Geodynamics, 47, 218– 223. Kuss, J. & Boukhary, M. A. 2008. A new upper Oligocene marine record from northern Sinai (Egypt) and its paleogeographic context. GeoArabia, 13, 59–84. Kuss, J., Bassiouni, A., Bauer, J., Bachmann, M., Marzouk, A., Scheibner, C. & Schulze, F. 2003. Cretaceous – Paleogene sequence stratigraphy of the

Levant Platform (Egypt, Sinai, Jordan). In: Gili, E., Negra, H. & Skelton, P. (eds) North African Cretaceous Carbonate Platform Systems. Nato Science Series, 171–187. Lewy, Z. 1991. Periodicity of Cretaceous epeirogenic pulses and the disappearance of the carbonate platform facies in the Late Cretaceous times (Isr). Israel Journal of Earth Science, 40, 51–58. Lu¨ning, S., Marzouk, A. M., Morsi, M. & Kuss, J. 1998. Sequence stratigraphy of the Upper Cretaceous of central-east Siani, Egypt. Cretaceous Research, 19, 153–196. Makris, J., Ben-Avraham, Z. et al. 1983. Seimic refraction profiles between Cyprus and Israel and their interpretations. Geophysical Journal of the Royal Astronomical Society, 75, 575– 591. Mimran, Y. 1984. Unconformities on the eastern flank of the Faria anticline and their implications on the structural evolution of Samaria (central Israel). Israel Journal of Earth Science, 33, 1 –11. Moustafa, A. R. & Khalil, M. H. 1994. Structural characteristics and tectonic evolution of north Sinai fold belts. In: Said, R. (ed.) The Geology of Egypt, A.A. Balkema. Rotterdam, 381– 389. Quennel, A. M. 1958. The structure and the evolution of the Dead Sea Rift. The Quarterly Journal of the Geological Society, London, 64, 1 –24. Quennel, A. M. 1984. The western Arabia rift system. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Blackwell, Oxford, 775–788. Rosenfeld, A. & Hirsch, F. 2005. The Cretaceous of Israel. In: Hall, J. K., Krasheninnikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant (II): The Levantine Basin and Israel. Historical Productions, Hall Publications, Jerusalem, Israel, 393–436. Said, R. 1971. Explanatory notes to accompany the Geological Map of Egypt. Geological Survey of Egypt, Papers, 1– 123. Schulze, F., Kuss, J. & Marzouk, A. 2005. Platform configuration, microfacies and cyclicities of the upper Albian to Turonian of west-central Jordan. Facies, 50, 505– 527. Shalar, J. 1994. The Syrian arc system: an overview. Palaeogeography, Palaeoclimatology, Palaeoecology, 112, 125 –142. Stampfli, G. M. & Borel, G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth and Planetary Science Letters, 196, 17– 33. Vidal, N., Alvarez-Marron, J. & Klaeschen, D. 2000. Internal configuration of the Levantine Basin from seismic reflection data (Eastern Mediterranean). Earth and Planetary Science Letters, 180, 77– 89. Woodside, J. M. 1977. Tectonic elements and crust of the eastern Mediterranean Sea. Marine Geophysical Researches, 3, 317– 354.

Tethyan rifting in the Levant Region and its role in Early Mesozoic crustal evolution MICHAEL A. GARDOSH1, ZVI GARFUNKEL2*, YEHEZKEL DRUCKMAN3 & BINYAMIN BUCHBINDER3 1

The Geophysical Institute of Israel, P.O. Box 182, Lod, 71100, Israel

2

Institute of Earth Sciences, Hebrew University, JerusPlealem, Israel, 91904 3

Geological Survey of Israel, 30 Malkhe Israel St. Jerusalem, 95501 *Corresponding author (e-mail: [email protected])

Abstract: At the time of the opening of the Tethys Ocean the northern edge of Gondwana was affected by several rifting events. In this study, we used data from deep exploration wells, seismic profiles, and seismic depth maps to reconstruct the pattern of Tethyan rifting in the Levant region and to investigate its effects on the evolution of the Levant crust. The results show a several hundred kilometre wide deformation zone, comprised of graben and horst structures that extend from the inner part of the Levant to the marine basin offshore Israel. The structures are dominated by sets of NE– SW and NNE– SSW oriented normal faults with vertical offsets in the range of 1–8 km. Rifting was associated with a NW–SE direction of extension, approximately perpendicular to the present-day Mediterranean coast. Faulting activity progressed over a period of 120 Ma and took place in three main pulses: Late Palaeozoic (Carboniferous to Permian); Middle to Late Triassic; and Early to Middle Jurassic. The last, and the most intense, tectonic phase post-dates the activity in other rifted margins of northern Gondwana. Rifting was associated with the modification and stretching of the Levant crust. Our results demonstrate an extension discrepancy between the brittle deformation in the upper crust and the amount of total crustal thinning. Seismic reflection data shows that the Levant Basin lacks the characteristics of typical rifted margins, either volcanic or non-volcanic. The evolution of the basin may be explained by depth-dependant stretching, associated with the upwelling of divergent mantle flow and removal of lower crustal layers by decoupling along deep detachment faults.

The separation of Gondwana and Eurasia and the development of the Tethys Ocean during the Late Palaeozoic to Early Mesozoic period were accompanied by continental breakup, rifting and drifting of various micro-continental blocks (Sengor & Yilmaz 1981; Robertson & Dixon 1984; Robertson 1998; Garfunkel 1998, 2002; Robertson 2007). This fragmentation of northern Gondwana, referred to here as Tethyan rifting, resulted in the formation of several marine basins and extensional margins in the Eastern Mediterranean region. Some of the basins were later consumed during the Cenozoic closure of the Tethys Ocean; the Levant continental margin and basin (Fig. 1) remained however, relatively intact and the Tethyan extensional structure in this area were therefore preserved (Garfunkel 1998; Gardosh & Druckman 2006). Structures that are associated with Tethyan rifting are recognized throughout the Levant onshore, from the Palmyra area in central Syria to the Egyptian Western Desert (Fig. 1) (Freund et al. 1975; Druckman 1984; Garfunkel & Derin 1984; Druckman et al. 1995; Garfunkel 1998; Guiraud & Bosworth 1999; Brew et al. 2001;

Sawaf et al. 2001). Tethyan structures are characterized by thickness variations, faulting and magmatism spanning the Palaeozoic to Early Mesozoic succession. This sedimentary section is typically found at great depth and, therefore, provides limited amount of geological information. Indeed, the study of Tethyan structures in the Levant region was so far based on rather small number of deep boreholes and partial coverage by land seismic data. Several deep exploration wells that were recently drilled in Israel and particularly, a new set of two-dimensional (2D), marine seismic reflection lines (Fig. 2) (Gardosh & Druckman 2006; Gardosh et al. 2006, 2008) add more details on some of the known Tethyan structures and reveal new structures that were previously unknown. Tethyan tectonic activity have been studied in various regions, located along the northern margins of the ancient super-continent of Gondwana. Robertson & Mountrakis (2006) and Robertson (2006, 2007) recently discussed the confusion in geological literature regarding the Tethyan nomenclature. Two terms are commonly used: ‘Palaeo-tethys’ and ‘Neotethys’. The former refers

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 9 –36. DOI: 10.1144/SP341.2 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Main tectonic elements of the Eastern Mediterranean region shown on the background of satellite imagery. The Levant Basin is located on the northeastern edge of the African plate, south of the Cyprian Arc plate boundary (marked by thick white lines). The area of study includes the central part of the basin and its margin onshore and offshore Israel. The outline of a new, seismic reflection survey offshore Israel is marked by dashed blue line. The insert shows a regional imagery of the Eastern Mediterranean and Red Sea (discussed in the text).

to oceanic basins of the Late Palaeozoic to Early Mesozoic age and the later is generally used for an Early Mesozoic, primarily Late Triassic to Early Jurassic ocean. However, these terms are often

misused by researchers that apply the same names to different marine basins (Robertson & Mountrakis 2006). Following Robertson (2007), we apply the more generalized term ‘Tethyan rifting’ for all the

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

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Fig. 2. The Tethyan rift system of the Levant. The map in (a) depicts inferred, Triassic to Early Jurassic faults and structural highs and lows that were formed during Tethyan rifting activity; (b) is a seismic depth map of the ‘top of the crystalline basement’ taken from Gardosh & Druckman (2006); and (c) is a Bouguer gravity map of the Levant from Rybakov & Al-Zoubi (2005). The gravity map (c) shows the transition from light, continental crust inland (negative values), to heavier crust (positive values) in the Mediterranean Sea. The faults and structures in (a) are compiled from various sources: the map in (b), and data published by Roberts & Peace (2007) and by Aal et al. (2000) for the offshore area, seismic and well data (Figs 3– 6) and seismic depth maps of Gelbermann (1995) for central Israel, data published by Moustafa & Khalil (1990) for Northern Sinai and by Walley (1998) for Lebanon. The location of the Palmyra Trough in (a) is restored to its position prior to the c. 100 km of sinistral motion along the Dead Sea Transform. The outlines of magnetic anomalies that are associated with Early Mesozoic magmatic activity in (a) are adopted from Rybakov et al. (1997). A list of well names is in Table 1. Abbreviations for structural elements in (b) and (c) are: ER, Eratosthenes High; JU, Judea Graben; PL, Pleshet Basin; NS, North Sinai Basin; JN, Jonah High; LV, Leviathan High; YM, Yam High.

Late Palaeozoic to Early Mesozoic structures of the Levant, with no reference to a specific ocean system. The main goal of this paper is to present an overview of Tethyan structures and rifting activity in the Levant region based on the integration of old and new well and seismic data. The regional structural pattern is used to reconstruct the style, direction and timing of continental breakup. As one of the notable results of Tethyan rifting activity is the modification of the Levant crust, an additional goal of this study is to examine the relations between faulting and crustal thinning as compared to other continental margins and extensional basins worldwide. The relatively large amount of geological and geophysical data makes the Levant margin a unique location for testing concepts of continental margin evolution.

The main tectonic events that shaped the Levant region The crust of the Levant was stabilized in the Late Precambrian by the Pan-African orogeny c. 1000– 550 Ma (Eyal et al. 1991; Stern 1994; Alsharhan & Nairn 1997). During the Early to Middle Palaeozoic period, the Levant area formed a wide continental platform of the northern part of Gondwana. Thick sections of predominantly siliciclastic sediments of fluvio-deltaic to shallowmarine origin were deposited on this platform south of the Tethys Ocean (Klitzsch 1981; Wolfart 1981; Weissbrod 1981, 2005; Beydoun 1988; Alsharhan & Nairn 1997; Garfunkel 2002). The Palaeozoic succession of northern Gondwana were affected by regional epiorogenical

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movements that produced broad swells and depressions, several hundred kilometres in diameter (Garfunkel & Derin 1984; Gvirtzman & Weissbrod 1984; Alsharhan & Nairn 1997; Garfunkel 2002, 2004). The swells were partly eroded, resulting in significant stratigraphic gaps. Gvirtzman & Weissbrod (1984) and Weissbrod (2005) inferred the formation of a regional swell in the Levant area, which they called the Geanticline of Helez. These authors suggested two phases of uplift: at the beginning of the Carboniferous and in the Early Permian. The Carboniferous erosion presumably produced a concentric outcrop pattern of the older Palaeozoic units, centred on the crest of the Helez Geanticline in the southern Israeli coastal plain (Gvirtzman & Weissbrod 1984; Weissbrod 2005). Garfunkel & Derin (1984) and Garfunkel (1998) found a more complex pattern of vertical motions with a structural high extending from Jordan to the Israeli coast. They further stressed the role of the mid-Palaeozoic erosion phase, which is supported by fission track dating (Kohn et al. 1992). Some authors attributed the Palaeozoic epeirogenic movements to compression of the Arabian plate that was associated with collision and subduction in the northern margin of Gondwana (Hercynian Orogeny) (Boote et al. 1998; Guiraud & Bosworth 1999; Weissbrod 2005). Others (Garfunkel 1998, 2004) think that during the Palaeozoic period most of the northern Gondwanian margin, including the Levant, was affected by extension rather than subduction, and this allowed the detachment of Gondwanian terrains that were incorporated in the Hercynian orogenic edifice in Europe. A profound change in the pattern of vertical motions and sediment distribution that took place during the latest Palaeozoic to Early Mesozoic period reflects a new tectonic regime (Garfunkel 1998). The Permian, Triassic and Lower Jurassic rock section of the Levant is dominated by shallowmarine, mostly carbonates and less common siliciclastic rocks. Considerable localized thickness and facies variations in these strata were identified in outcrop, well and seismic data from southern and central Israel (Goldberg & Friedman 1974; Druckman 1974, 1977; Freund et al. 1975; Druckman 1984; Garfunkel & Derin 1984; Gelbermann & Kemmis 1987; Bruner 1991; Druckman et al. 1995; Gelbermann 1995; Garfunkel 1998; Gardosh & Druckman 2006). Thickness variations are also documented in the Palmyrides and other parts of Syria (Al-Youssef & Ayed 1992; Brew et al. 2001; Sawaf et al. 2001) and in northern Sinai (Alsharhan & Salah 1996; Hirsch et al. 1998). These widely spread phenomena indicate repeated differential motions and extension on the edge of the Gondwanian plate during the latest

Palaeozoic to Early Mesozoic period (Freund et al. 1975; Garfunkel & Derin 1984; Druckman et al. 1995; Garfunkel 1998; Gardosh & Druckman 2006). The extensional pulses that were accompanied by magmatic activity resulted in the separation of the Tauride, Eratosthenes, and probably other small continental blocks from the AfroArabian craton (Biju-Duval & Dercourt 1980; Sengor & Yilmaz 1981; Robertson & Dixon 1984; Le Pichon et al. 1988; Robertson et al. 1996; Garfunkel 1998), and led to the opening of an ocean north and west of the Levant margin. The Early Mesozoic rifting activity was followed by subsidence and the formation of the deep marine Levant Basin in the eastern Mediterranean area (Garfunkel & Derin 1984; ten Brink 1987; Garfunkel 1988). Starting from the Middle Jurassic, a continental shelf and slope developed along the eastern margin of the basin (Fig. 1). This continental margin slope, originally identified in deep wells near the present-day Israeli coastline (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977; Garfunkel 1988), formed a narrow transition zone, termed by Bein & Gvirtzman (1977) as the ‘hinge-line’. It separated the Jurassic and Cretaceous shallow-marine shelf on the east from the deep marine basin on the west. The Mesozoic shelf-edge is identified near the modern Mediterranean coastline from northern Egypt to western Lebanon (Harms & Wray 1990; Jenkins 1990; Walley 1998, 2001; Garfunkel 1998). Late Cretaceous convergence of the Eurasian and Afro-Arabian plates resulted in the formation of a northward-dipping subduction zone at the present-day areas of Cyprus and southern Turkey (Fig. 1) (Le Pichon et al. 1988; Robertson 1998). Far-field stresses related to this subduction caused contraction of the Levant margin and generated the development of anticlines and synclines throughout the Levant region known as the ‘Syrian Arc’ (Krenkel 1924) fold belt. The contractional activity continued in the Levant area through the Cenozoic period (Eyal & Reches 1983; Walley 1998; Gardosh & Druckman 2006). Well and seismic data suggest that the Syrian Arc deformation is associated in many places with inversion of Late Palaeozoic to Early Mesozoic normal faults that were reactivated in a reverse motion (Freund et al. 1975; Druckman et al. 1995), although in some structures these relations are not clearly observed (Druckman 1981). The last, major tectonic event that affected the Levant area was the breakup of the African – Arabian Craton and the formation of the Dead Sea fault zone (Fig. 1) (Freund et al. 1970; Garfunkel 1981, 1997). This new plate boundary is, in fact, a transform fault that connects the spreading centre of the Red Sea with a collision zone in the Taurus

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

Mountains of southeastern Turkey (Fig. 1). The Dead Sea transform offset some of the older Tethyan structures. The sinistral, strike –slip motion along this fault system amounts to about 100 km (Freund et al. 1970; Garfunkel 1981).

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and lows; and (d) the Eratosthenes High. These zones include symmetric and asymmetric grabens, horsts, and down-stepping blocks, all controlled by sets of normal faults generally trending NE –SW.

The inner basins Data sources The distribution of Tethyan rift structures in the subsurface of the Levant (Fig. 2a) is compiled from the following sources: data from deep boreholes that penetrated the Palaeozoic to Lower Jurassic section in Israel (Table 1); seismic interpretation maps of the Levant offshore (Gardosh & Druckman 2006; Gardosh et al. 2006, 2008); seismic interpretation maps of central and southern Israel (Gelbermann 1995), outcrop and well studies by Freund et al. (1975), Garfunkel & Derin (1984), Druckman et al. (1995) and Garfunkel (1998). The seismic mapping of Gardosh et al. (2006, 2008) was based on the interpretation of a highquality set of 2D seismic reflection lines acquired in 2001 (Figs 1 & 2b). These regional marine lines extend from the coastal area of Israel to a distance of 150 –200 km offshore and penetrate the subsurface down to the top of the crystalline basement at depth of 10 –15 km below mean sea level (MSL). The excellent coverage allows the interpretation and mapping of the deep part of the Levant Basin that was not imaged in older seismic datasets. Two important seismic markers were used to map Tethyan rift structures: (a) top of the crystalline basement (Fig. 2b) and (b) Middle Jurassic unconformity of an assumed Bathonian age (Gardosh et al. 2006, 2008). Three regional, geological cross-sections (Fig. 3) were constructed from the seismic depth maps (Levant Basin and central Israel) and results of 18 deep exploration boreholes located onshore and offshore Israel (Table 1). The structure in the northern part of section 3 (Fig. 3c) is considered somewhat speculative because the area of northern Israel is not sufficiently covered by seismic lines. The main chronostratigraphic boundaries in the wells (Table 1) follow Derin et al. (1980), Fleischer & Varsahvsky (2002), Zion Oil and Gas Inc. (2005) and Gardosh et al. (2008).

Rifting structures in the subsurface of the Levant Basin and margin The Tethyan extensional structures of the Levant area were divided here into four zones of deformation, distributed, more or less, parallel to the Mediterranean coast (Fig. 2). These are from east to west: (a) the inner basins; (b) the highs along the Mediterranean coastline; (c) the offshore highs

The central part of Israel is occupied by the Judea Graben (Figs 2a & 3) (Freund et al. 1975). This c. 150 km long and c. 30 –50 km wide basin formed as a Triassic and Early Jurassic depocentre, bounded by elevated horst blocks (Fig. 2a). The axis of the Judea Graben is marked by the thick Triassic sections of the Ramallah-1 (1700 m, base not reached) (Derin et al. 1980) and Devora-2A wells (2600 m, base not reached) (Derin & Gerry 1979). The Palmyra Trough is a similar depocentre filled by more than 2 km of Triassic sediments that is identified by well and seismic data in southern Syria (Garfunkel 1998; Brew et al. 2001; Sawaf et al. 2001). Restoring the sinistral motion along the Dead Sea transform shifts the Palmyra Trough 100 km southward (Fig. 2a). In this position, it correlates to the northeastern part of the Judea Graben, suggesting a link between these two structures in Triassic time (Garfunkel 1998; Flexer et al. 2005). The Judea Graben changes direction and widens towards the north, but its structure north of latitude 318500 (Fig. 2a) is poorly constrained by the available subsurface data. Important evidence to the structural pattern in northern Israel is the c. 2.5 km thick, latest Triassic and earliest Jurassic series of the Asher volcanics that were penetrated by Atlit-1 Deep (Fig. 2a, Table 1) and several other wells in northern Israel (Derin et al. 1982; Gvirtzman & Steinitz 1983; Kohn et al. 1993). Stratigraphic analysis of theses wells shows that the Asher extrusive volcanics are overlain by a continuous and relatively flat, shallow-marine, Middle Liassic section, indicating that the volcanic rocks accumulated in an active, fault controlled depression (Fig. 3c) (Garfunkel & Derin 1984; Garfunkel 1989). This depression is termed here the Asher Basin. The Asher Basin is reconstructed as a several tens of kilometre long graben that extends from the Haifa Bay on the northern coast of Israel towards the SE (Fig. 2a). Owing to the limited amount of seismic and well data, this postulated graben is largely based on the shape of the magnetic field in the area. The magnetic map shows a conspicuous, positive anomaly centred NE of the Atlit-1 Deep well that is interpreted to be caused by the deeply buried, Asher basalt flows (Rybakov et al. 1997, 2000). The elongated, NW–SE oriented magnetic anomaly is assumed to reflect the shape of a fault bounded depression that was filled with the Jurassic extrusive rocks (Garfunkel & Derin 1984;

14

Table 1. Summary of well data used in this study. The depth intervals of chronostratigraphic units (in meters from KB) are taken, for most of the wells, from Fleischer & Varsahvsky (2002). Data for well 9 is taken from Derin et al. (1980); for well 11 from Zion Oil and Gas Inc. (2005); and for the Liassic to Bathonian section in wells 18 and 19 from Gardosh et al. (2008). The Cretaceous depth interval in wells 17, 18 and 19 includes the Tithonian succession Well no. Well name

Depth and thickness (in metres) of stratigraphic units Palaeozoic (Permian)

Precambrian Basement Depth Atlit-1 Deep Ashqelon-2 Betarim-1 Bessor-1 David-1 Deborah-2 Gaash-2 Gevim-1 Halal-1 Helez Deep-1 Maanit-1 Deep Meged-2 Nevatim-1 Pleshet-1 Ramallah-1 Deep Talme Yafe-4 Yam-2 Yam West-1 Yam Yafo-1 Yinnon-1

þ Base of unit was not penetrated. – Unit is missing.

Depth

Thick.

4567 –4465

102

4465 – 4230 235 5998 – 5455 543þ

4620 –4588

32

4588 – 4078

510

6093 –5978

115

5978 – 5726

252

5344 –5315

29

5315 – 4788

527

Jurassic (Liassic – Bathonian)

Jurassic (Callovian – Kimmeridgian)

Cretaceous (Neocomian – Turonian)

Depth

Thick.

Depth

Thick.

Depth

Thick.

Depth

Thick.

6531 – 5390

1141þ 65 1645 2585þ 913þ 208 685þ 926 1419þ 1599þ

4788 – 4200 6361 – 4672

588 1689þ

3525 161þ 1449þ 1762 1540 1582 1614 1564 2814 2830 937 1482 434þ 1595 3162 1493 142þ 350þ 437þ 364þ

1865 – 1446 – 1790 – 1241 2403 – 2060 2270 – 1956 1480 – 646 2981 – 2506 2306 – 1876 815 – 457 1970 – 1858 2341 – 1856 2119 – 1770 1760 – 1340 2605 – 2131 1510 – 901 – 5235 – 4765 4900 – 4400 5350 – 4410 2616 – 2210

418

4230 – 4165 5455 – 3810 5647 – 3062 5508 – 4595 4078 – 3870 4314 – 3629 5726 – 4800 4697 – 3278 5200 – 3601

5390 – 1865 4076 – 3915 3239 – 1790 4165 – 2403 3810 – 2270 3062 – 1480 4595 – 2981 3870 – 2306 3629 – 815 4800 – 1970 3278 – 2341 3601 – 2119 2194 – 1760 4200 – 2605 4672 – 1510 4204 – 2711 5377 – 5235 5250 – 4900 5787 – 5350 2980 – 2616

1447 – 60 3915 – 2249 1241 – 10 2060 – 628 1956 – 279 646 – 23 2506 – 803 1876 – 567 457 – 0 1858 – 455 1856 – 3 1770 – 10 1340 – 24 2131 – 615 901 – 4 2711 – 1199 4765 – 2730 4400 – 2874 4410 – 2380 2210 – 394

1367 666 1231 1432 1377 623 1703 1309 457 1403 1853 1760 1316 1516 897 1512 2035 1526 2030 1816

549 343 314 834 475 430 358 112 485 349 420 474 609 470 500 940 406

M. A. GARDOSH ET AL.

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

Thick.

Triassic

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN Fig. 3. Geological cross-sections showing the main stratigraphic successions, faults and structural blocks on the Levant Basin and margin. Thickness variations in the Palaeozoic, Triassic and Lower Jurassic (Liassic to Bathonian) reflect syn-tectonic deposition, related to the development of an extensive graben and horst system during Tethyan rifting activity. The location of the sections is shown in Figure 2a. 15

16

M. A. GARDOSH ET AL.

Garfunkel 1989). The reduced thickness of the Asher volcanics in Devora-2A (Garfunkel 1989) may indicate the existence of a large, normal fault west of the well that formed the northeastern limit of this depression (Fig. 3c). The Asher Basin seems to form a north– west extension of the Tethyan, Judea Graben (Fig. 2a). It was later modified by the c. 1 km vertical offset of the Carmel Fault that is associated with the Cenozoic activity along the Dead Sea plate boundary (Fig. 3c). In Early to Middle Jurassic time, the Judea Graben may have also extended northward, as indicated by a relatively thick Bathonian section (600 m) found in outcrops at the Hermon area (Hirsch et al. 1998; Hirsch 2005). The postulated Hermon Basin (Fig. 2a) was probably limited in size and did not extend eastwards to the Palmyra Trough where the Jurassic section is only 100 m thick (Mouty 1997; Hirsch 2005). A set of NE–SW oriented faults found in the southern part of the Judea Graben (Figs 2a & 3b) formed intra-graben fault blocks. These were later inverted during the Syrian Arc contractional deformation of the northern Negev (Freund et al. 1975; Gelbermann 1995; Druckman et al. 1995). The Judea Graben is delimited in the SE by a series of small, positive structures, that is, the Massada and Heimar Highs, located on the western margin of the Dead Sea (Fig. 2a) (Goldberg & Friedman 1974; Druckman 1974; Freund et al. 1975). The Ramon High, located near the Ramon fold (Druckman 1977; Buchbinder & Le Roux 1993), delimits the graben further to the south (Fig. 2a). The Judea Graben extends from the central Negev westward into northern Sinai. A Triassic tectonic depression in this area, termed here North Sinai Basin (Fig. 2a), is indicated by a relatively thick Carnian section at the bottom of the Halal-1 well (Fig. 2a, Table 1) (Garfunkel & Derin 1984; Derin & Gerry 1986a). Subsidence of the North Sinai Basin continued during the Early to Middle Jurassic. This activity is reflected by the thick Liassic to Bathonian interval found in outcrops at Gebel Maghara (1700 m) (Alsharan & Salah 1996; Hirsch et al. 1998) and penetrated in the Halal-1 well (2600 m) (Derin et al. 1976). The shape and internal structure of the North Sinai Basin are not well delineated owing to insufficient seismic and well data. It was most likely formed by down-to-the-north sets of fault blocks, oriented to the NE–SW and east –west (Fig. 2a). This trend is inferred from the direction of Syrian Arc folds that are interpreted as inverted Early Mesozoic structures (Moustafa & Khalil 1990; Alsharhan & Salah 1996). Small intra-basin blocks and localized highs and lows, such as those found within the Judea Graben, may have also existed in this area.

The highs along the Mediterranean coastline A series of structural highs along the Mediterranean coastline forms the western boundary of the Judea Graben. The most prominent structure is the NE – SW trending Gevim High (Figs 2a, 3b, 4 & 5). This narrow structure was penetrated by the Gevim-1 and the Bessor-1 wells (Figs 3c & 4). Both wells reached Precambrian rocks in depth of 4460 m and 4293 m below sea level, respectively, thus indicating that the Gevim High is the highest basement block in central Israel. The Gevim High is bounded by two sets of faults: the Helez Fault on the NW and the Pleshet Fault on the SE (Fig. 4). The stratigraphic relations in the Gevim-1 and the Helez Deep-1 wells, located respectively on the footwall and hanging wall of the Helez Fault, indicate 1740 m of vertical offset during the Triassic and Early to Middle Jurassic (Gardosh & Druckman 2006). The stratigraphy of the Pleshet-1 well (Derin et al. 1985), located on the downthrown side of the Pleshet Fault, indicates about 700 m of vertical offset on an Upper Permian level that can be attributed to the same tectonic activity. Further evidence for the activity on the Helez Fault is found in Helez Deep-1. The well penetrated c. 300 m thick section of coarse-grained, polymictic conglomerate, within the Triassic succession [the Erez Conglomerate of Druckman (1984), or the Or Haner Conglomerate of Derin (1979) and Garfunkel & Derin 1984]. This clastic unit was interpreted as an alluvial fan that accumulated at the foot of the Helez Fault escarpment (Fig. 4) (Druckman 1984; Garfunkel & Derin 1984). In the central and northern part of Israel the Gaash-Meged and Maanit structures (Figs 2a, 3a, c & 6) delineate the western boundary of the Judea Graben and probably also of its NW extension, the Asher Basin. The two highs are inferred from the reduced thickness of the Jurassic and Triassic sections in the wells that penetrated these structures (Table 1). The Liassic to Bathonian interval reaches 1614 m in Gaash-2 (Derin et al. 1981; Gvirtzman et al. 1984), 1482 m in Meged-2 (Givot Olam Oil Ltd 1995), and 937 m in Maanit-1 Deep (Zion Oil and Gas Inc. 2005) (Fig. 3a, c, Table 1). The same stratigraphic interval is more then 3000 m thick in the Ramallah-1 well within the Judea Graben (Fig. 3a, Table 1) (Derin et al. 1980; Fleischer & Varsahvsky 2002). A similar trend is observed in the Triassic section. In Gaash-2, the Scythian to Carnian interval is only 913 m thick (Derin et al. 1981), whereas in Ramallah-1, within the Judea Graben it reaches 1700 m (Fig. 3a, Table 1) (Derin et al. 1980). Seismic data indicate that the axis of the coastal highs changes its direction from southern Israel

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

17

Fig. 4. Interpreted, seismic profile DS-512 (time migrated), across the Gevim High and the western edge of the Judea Graben. The structure is bounded by two fault zones (Helez and Pleshet) that were intermittently active from Late Palaeozoic to Middle Upper Jurassic time. The Middle Triassic Erez conglomerate (a) is an alluvial fan that accumulated at the foot of the Helez fault escar/minent. The Helez and Pleshet faults were re-activated in a reverse motion during the Late Cretaceous to Cenozoic Syrian Arc folding phase. Note the location of the Upper Jurassic to Middle Cretaceous, rimmed shelf-edge that developed on the western edge of the Gevim High. The ages of the formation tops are: Basement, Precambrian; Gevim, Permo-Triassic; Arkov, Permian; Erez, Anizian; Saharonim, Anizian to Carnian; Mohila, Carnian; Qeren, Aalenian; Shederot, Bathonian; Zohar, Callovian; Beer Sheva, Oxfordian; Yagur, Albian. Abbreviations for chronostratigraphic ages are: L., Lower; M., Middle; U., Upper. Double arrows refer to reactivation of older, normal faults in a reverse motion. The location of the profile is shown in Figure 2a.

northward. The normal faults that bound the Gevim High trend NE– SW whereas in the area of the Gaash High the main faults trend north –south (Fig. 2) (Gelbermann 1995). A north–south to NNW–SSE direction is inferred also for the axis of the Maanit High further to the north, although the seismic coverage of this structure is limited. The northward shift of the structural grain appears to be a fundamental characteristic of the central Levant area and is reflected also by the direction of two younger geological elements: the Jurassic – Cretaceous shelf –edge and the Syrian Arc fold belt.

The offshore highs and lows Palaeozoic to Early Mesozoic structures in the Levant offshore are revealed by the interpretation and mapping of two regional seismic markers: (a) the top of the crystalline basement and (b) the Middle Jurassic unconformity (Figs 7–10) (Gardosh & Druckman 2006; Gardosh et al. 2006, 2008). The top of the crystalline basement is

identified by significant change in seismic character from reflection-free zones below to parallel, continuous reflection series above (Fig. 8). This transition is assumed to reflect the contact of magmatic and metamorphic rocks with the overlaying Palaeozoic to Mesozoic sedimentary section. Similar characteristics of the ‘top basement’ reflector were interpreted by Vidal et al. (2000) in the northern part of the Levant Basin. In seismic profiles, the Middle Jurassic unconformity (Fig. 8) marks a change in seismic character, from high amplitude and continuous reflections below, to low amplitude and discontinuous reflections above. In the offshore Yam West-1 well, this marker is correlated to the top of the Bathonian, Shederot Formation (Fig. 8) (Gardosh et al. 2006, 2008). The change in seismic character is interpreted to reflect a transition from Lower Jurassic, shallow-marine carbonates to Middle–Upper Jurassic deepwater strata (Gardosh & Druckman 2006; Gardosh et al. 2006, 2008). Thickness changes, dipping reflections, and discontinuity of seismic

18

M. A. GARDOSH ET AL.

Fig. 5. Interpreted, seismic profile DS-3589 (time migrated), showing a Palaeozoic basin in the northeastern edge of the Gevim High. Onlapping of the Permian, Triassic and Lower Jurassic strata (marked by open arrows) is associated with subsidence and faulting. Correlation of seismic reflections to the Gevim-1 well suggests that this activity may have started prior to the deposition of the Permian Arkov Formation. The ages of the formation tops are in Figure 4. Abbreviations for chronostratigraphic ages are: L., Lower; M., Middle; U., Upper. The location of the profile is shown in Figure 2a.

events found throughout the offshore, between the top of the crystalline basement and the Middle Jurassic seismic markers, were interpreted to reflect syn-tectonic rifting activity (Figs 7 –10).

A large structure, revealed by the depth of the crystalline basement and the increased thickness of the Palaeozoic to Middle Jurassic interval, is the Pleshet Basin (Figs 7–9). This structure is a

Fig. 6. Interpreted seismic profile DS-559 (time migrated) across the Gaash-Meged High. Discontinuity of seismic reflections within the Triassic and Permian intervals indicates vertical motions during these times, depicted by the normal faults at the eastern part of the profile. The Gassh-Meged High was an elevated structure between the Pleshet Basin to the west (not shown on this profile) and the Judea Graben to the east. An apparent tilt of the Gaash area in the western part of the profile is a pull-down effect, created by the low-velocity Cenozoic cover (compare to section 1 in Figure 3a). The ages of formation tops are: Zafir, Scythian; Mohila, Carnian; Shefayim, Norian; Lower Haifa, Bajocian; Upper Haifa, Oxfordian. Abbreviations for chronostratigraphic ages are: L.-Lower, M.-Middle, U.-Upper. The location of the profile is shown in Figure 2.

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

c. 150 km long and c. 50 km wide graben that extends along the southeastern part of the Levant Basin. The Pleshet depocentre is found about 50 km west of the coastline, opposite the Gaash– Meged High (Figs 2a, b & 3a). The existence of a deep Mesozoic basin in this area was first suggested by Cohen et al. (1988) based on potential field data that indicates a magnetic basement at great depth (Domzhalski 1986). The interpretation of the new seismic data show a structural low that more or less coincides with the area of the Pleshet Basin as outlined by Cohen et al. (1988) (Fig. 2b). This basin structure is interpreted as a Tethyan Graben, similar in size and shape to the Judea Graben further to the east (Figs 2a & 3a, b). The Pleshet Basin is bounded on the SE by a 5 – 10 km wide horst, oriented in a NE–SW direction, termed here the Yam High (Figs 2, 3 & 8). A set of normal, down-to-the-west step faults displace the top of the crystalline basement marker west of the Yam High. The normal faulting is associated with a significant increase of thickness and onlapping of the Palaeozoic to Middle Jurassic section (Fig. 8). Some of the faults, such as west of the Yam West-1 well (Fig. 8), were later reactivated as high-angle thrust faults. The Yam High probably extended further to the NE converged with the

19

Gaash High near the coastline. The structural configuration of this area near the Yam Yafo-1 well is, however, highly complex and difficult to interpret owing to the younger contractional deformation (Fig. 3a). The Pleshet Basin is delimited to the NW by the Jonah High (Figs 2 & 3a). This feature, which was first identified by Folkman & Ben Gai (2004), is a conspicuous basement high, about 80 km long and 10 –20 km wide, controlled by NE–SW oriented sets of normal faults (Figs 2b, 7 & 9) (Gardosh & Druckman 2006; Gardosh et al. 2006, 2008). The Jonah High is likely underlain by a large magmatic intrusion (Fig. 10). A magnetic anomaly that is located on this narrow horst (Fig. 2a) may be associated with an extrusive volcanic body of Early Mesozoic age (Fig. 7) (Gardosh et al. 2006, 2008). A conspicuous asymmetric graben is observed NW of the Jonah High. It is characterized by southeasterly dipping seismic events and increased thickness of the Palaeozoic to Middle Jurassic interval (Figs 9 & 10). The western limit of this graben is a large basement structure, about 100 km long and 20 –40 km wide, termed here the Leviathan High (Figs 2, 3a & 9). This wide and flat horst is controlled by several sets of normal faults. It may also be associated with a magmatic intrusion in

Fig. 7. Interpreted, composite marine seismic profile (time migrated) showing Tethyan extensional structures in the Levant Basin, offshore. Thickness variations in the Palaeozoic to Middle Jurassic intervals were controlled by normal faulting and reflect syn-tectonic deposition. Most, but not all of the normal faults were reactivated in a reverse motion during the Syrian Arc contractional phase. The Jonah High is a fault controlled basement structure that bounds the Pleshet Basin on the east, and a smaller asymmetric graben on the west. The Jonah High may have been the loci of Early Mesozoic, extrusive volcanic activity (a) (Gardosh et al. 2008). Abbreviations for chronostratigraphical ages are: L., Lower; M., Middle; U., Upper. Double arrows refer to reactivation of older, normal faults in a reverse motion. The location of the profile is shown in Figure 2b.

20

M. A. GARDOSH ET AL.

Fig. 8. Interpreted marine seismic profile (time migrated), showing the southeastern edge of the Pleshet Basin. This fault-controlled, rift basin is characterized by a high-amplitude, divergent reflection package that onlaps the Yam High. The base of the syn-rift sequence is correlated to the contact between chaotic and well-layered seismic packages, interpreted as the top of the crystalline basement. The top of the sequence is a continuous reflection that is correlated to the Bathonian, Shederot Formation in the Yam West-1 well. The continuous, sub-parallel, high-amplitude seismic reflections within the Pleshet Basin are interpreted as continental to shallow-marine strata. The maximal thickness of the basin-fill in this profile is 4 –5 km (calculated with interval velocity ¼ 6000 m s21). Some, but not all of the Tethyan normal faults were reactivated as high-angle thrusts during Syrian Arc folding, observed at the edges of the profile. Abbreviations for chronostratigraphic ages are: L., Lower; M., Middle; U., Upper. Double arrows refer to reactivation of older, normal faults in a reverse motion. The location of this profile (an enlarged part of Fig. 7) is shown in Figure 2.

depth. The conspicuous, positive Bouguer gravity anomaly found in the Leviathan area (Fig. 2c), suggest the existence of such an intrusive body (Gardosh et al. 2006, 2008). Several, small basement highs, about 10 km wide and 20 km long were identified south of the Jonah and Leviathan horsts (Fig. 2a, b). The belt of deep-seated highs and lows that is identified offshore Israel extends further to the NE and SW parts of the Levant Basin. Seismic and gravity data indicate the existence of Early Mesozoic structures below the Cenozoic, Nile Delta cone (Fig. 2c) (Aal et al. 2000; Bentham et al. 2007). The Rosetta fault that extends from the Nile Delta to the eastern side of the Eratosthenes high (Fig. 2a) (Aal et al. 2000) is oriented in NE– SW direction, similar to the direction of the

Leviathan and Jonah horsts. Triassic and Jurassic highs that were recently identified on seismic reflection profiles offshore Lebanon (Fig. 2a) (Roberts & Peace 2007) fit well to the regional structural pattern.

The Eratosthenes High The Eratosthenes seamount is a large structure, located at the NW of the Levant Basin (Figs 1 & 2). It is characterized by distinct geophysical and geological properties. The seamount overlies a positive magnetic anomaly that is interpreted to be associated with a voluminous igneous body in depth (Fig. 2a) (Ben-Avraham et al. 1976; Garfunkel & Derin 1984). Wide-angle seismic refraction tests show that the Eratosthenes area is

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

21

Fig. 9. Interpreted, composite marine seismic profile (time migrated), showing Tethyan extensional structures in the Levant Basin, offshore. Three structural highs: Jonah, Leviathan and Eratosthenes, are underlain by elevated basement blocks that display a chaotic seismic character and may comprise magmatic intrusions. These fault-controlled structures bound three rift basins that are characterized by continuous, divergent seismic reflection packages. The Pleshet Basin on the right-hand side of the profile is bounded by the Gaash-Meged High, located east of the profile. Abbreviations for chronostratigraphic ages are: L., Lower; M., Middle; U., Upper. Double arrows refer to reactivation of older, normal faults in a reverse motion. The location of the profile is shown in Figure 2b.

Fig. 10. Interpreted marine seismic profile (time migrated), showing the Jonah High and the asymmetric graben on its western flank This Tethyan, fault-controlled rift basin is characterized by a high-amplitude, divergent reflection package that thickens towards the Jonah structure. The base of the syn-rift sequence is correlated to the contact between chaotic to well-layered seismic packages, interpreted as the top of the crystalline basement. Its top is a high-amplitude reflection that is correlated to the Middle Jurassic unconformity. The maximal thickness of the syn-rift sequence in this profile is 7 –8 km (calculated with interval velocity ¼ 6000 m s). Abbreviations for chronostratigraphic ages are: L., Lower; M., Middle; U., Upper. The location of this profile (an enlarged part of Fig. 8) is shown in Figure 2b.

22

M. A. GARDOSH ET AL.

underlain by a low velocity crystalline crust (Vp ¼ 6 km/s) that is interpreted as a continental or intermediate type (Makris et al. 1983; BenAvraham et al. 2002). A borehole drilled on the northern slope of the seamount penetrated Lower Cretaceous, shallow-marine limestones similar to Cretaceous rocks found in the Levant margin (Mart & Robertson 1998). Based on these lines of evidence, the Eratosthenes is interpreted by most authors as a continental block that was detached and drifted from the nearby Afro-Arabian plate during the Early Mesozoic break-up of Gondwana (Garfunkel & Derin 1984; Kempler 1998; Mart & Robertson 1998). This interpretation is generally supported by the new, offshore, seismic profiles that show a basement high and a series of fault blocks of an assumed Mesozoic age SE of the seamount (Figs 2b & 9). This structure probably represents part of the larger Eratosthenes high that is delimited on the east by the NE –SW trending Qattara – Eratosthenes (Rosetta) fault zone (Fig. 2a) (Aal et al. 2000). The Cretaceous limestones found on the seamount supports the assumption that during Late Mesozoic times this area was part of an elevated, shallowmarine shelf located at the northwestern side of the deep-marine Levant Basin (Gardosh & Druckman 2006). The Eratosthenes high was further uplifted by reactivation of the Mesozoic faults during the Cenozoic convergence of the AfroArabian and Eurasian plates (Robertson 2000). Zverev & Ilinsky (2000, 2005) recently studied the deep structure of the Eratosthenes by a new set of seismic refraction profiles. They found a complex pattern of high and low velocity layers at the upper crust, below the seamount. The different velocities were attributed to alternating layers of lavas, volcanic sediments and basic magmatic rocks. Based on these findings, Zverev & Ilinsky (2000, 2005) suggested that the Eratosthenes seamount is a long-lived volcanic edifice associated with pre-Cretaceous magmatic activity, possibly similar to Early Jurassic volcanic rocks found in northern Israel. These authors further suggested that the crust underneath the Eratosthenes is similar to the crust of the adjacent Levant Basin. Although it is not clear whether the Eratosthenes seamount overlies a drifted continental block (Garfunkel 1989) or an Early Mesozoic volcano edifice (Zverev & Ilinsky 2000, 2005), in either case its formation appears to be the result of Tethyan rifting activity.

Timing of the rifting activity The structures found in the deep subsurface of the Levant region, onshore and offshore, appear to

have been active during several pulses of extension and faulting, followed by periods of relative tectonic quiescence. Parts of these rift structures were active in varying intensity during different times. Well data indicate three main phases of activity: Late Palaeozoic, Middle to Late Triassic and Early to Middle Jurassic.

Late Palaeozoic Direct evidence for Late Palaeozoic continental breakup is found in the Palmyrides area of central Syria (Fig. 1). Well and seismic data show up to 1 km thick Permian succession in a fault-controlled trough that is interpreted as an intra-continental Permian rift (Garfunkel 1998; Brew et al. 2001; Sawaf et al. 2001). More evidence for this tectonic phase is found near the Mediterranean coastline. In the Gevim High (Fig. 2a), the Gevim-1 and Bessor-1 wells record a reduced Permian section of several tens to a few hundred metres, overlaying the Precambrian basement (Fig. 3c). These stratigraphic relations indicates at least two major tectonic events: a pre-Permian, possibly Carboniferous or earlier uplift that resulted with erosion of the entire Early to Middle Palaeozoic section, and an uplift that resulted with deposition of a thin Permian section. Palaeozoic vertical motions in the area of the Gevim High are evident also on the seismic profile in Figure 5, which shows significant thickness increase and onlapping of Palaeozoic strata on the northeastern flank of the structure. Correlation of seismic reflections to the Gevim-1 well (Fig. 5) suggests that the subsidence may have started during pre-Permian time. Permian tectonic event in this area is supported by facies change of the Saad and Arkov Formations. In the Gevim-1 well (Salhov 1996), these rock units are predominantly siliciclastic and sand rich, whereas in the David-1 well (Derin 1995), located about 40 km north of the Gevim High (Fig. 3b), they are comprised of shallow-marine carbonate. In the offshore, the Pleshet Basin, and the a-symmetric graben located between the Jonah and the Leviathan highs contain 4– 8 km thick, preMiddle Jurassic section (Fig. 3a, c) (Gardosh & Druckman 2006). The maximal thickness of the Triassic to Middle Jurassic section in the onshore Deborah-2 and Ramallah-1 wells is 4 to 5 km respectively. It is, therefore, possible that part of the excessive thickness in the offshore basins is associated with greater Triassic and Palaeozoic subsidence. This scenario may be supported by the similarity of the onlapping pattern observed on seismic profiles at the edge of the Pleshet Basin (Fig. 8) and north of the Gevim High (Fig. 5). Considering the examples described above, it is evident

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

that extension and rifting already took place in the Levant region during the latest Palaeozoic period, although the dimension and configuration of this tectonic event are not well constrained.

23

shelf deposits. It is, therefore, likely that the Triassic Tethyan rifting did not result with the development of a deep-water oceanic realm in the studied area.

Early to Middle Jurassic Middle to Late Triassic The Early part of the Triassic was a period of relative tectonic quiescence (Garfunkel 1998). Evidence for Middle Triassic faulting and vertical motions is found in several locations. An uplift of the Gevim High is interpreted from the Gevim-1 and Bessor-1 wells that penetrated a thin section of the Ladinian (Upper–Middle Triassic) Saharonim Formation overlying the Permian (Fig. 3c). Middle Triassic faulting event is interpreted from the Erez (Or Haner) fault breccias of an assumed Anisian age (Druckman 1984; Garfunkel & Derin 1984). Further evidence for this phase is inferred from thickness changes of the Anisian, Gevanim Formation in the northern and central Negev (Druckman 1974). A second, Middle to Upper Triassic faulting event is observed in the Judea Graben. The Deborah-2 and Ramalah-1 wells (Fig. 2) (Derin & Gerry 1979, 1986b) penetrated a c. 1 km of the Carnian, Mohilla Formation, which is only several tens of meters thick in other parts of the country. The evaporitic facies of the Mohilla Formation is recorded by wells and outcrops in the northern Negev (Druckman 1974). This facies is associated with restricted water circulation and the development of evaporitic conditions in tectonically active depressions (Druckman 1974) at the southeastern part of the Judea Graben. Upper Triassic subsidence also took place within the North Sinai basin as indicated by the relatively thick Carnian section at the bottom of the Halal-1 well (Garfunkel & Derin 1984; Derin & Gerry 1986a). The accumulation of a thick sedimentary section in the Pleshet Basin and in the graben NE of the Jonah High (Fig. 3a, b) may be partly attributed to Triassic tectonic activity although no well control is available to confirm this hypothesis. It is further unknown whether the Levant Basin was a deep marine basin during Triassic time. Pelagic Triassic strata is found in outcrops at Antalya (southern Turkey) and in the Mammonia blocks (Cyprus) (Druckman et al. 1982; Robertson 2000; Robertson & Xonophontos 1993), suggesting the development of deep-marine conditions north of the Levant (Garfunkel 1998; Robertson 1998, 2007). On the other hand, the Triassic section of the Levant onshore is characterized by shallow-marine to continental environments. The continuous, high-amplitude reflections that characterize the Palaeozoic to Middle Jurassic interval on offshore seismic profiles (Figs 7–10) may be interpreted as shallow-marine,

Activation of Tethyan rift structures during the Early to Middle Jurassic took place throughout the Levant region. More than 3 km of Liassic strata were deposited at the central part of the Judea Graben near the Ramallah-1 well (Figs 2a & 3a). Further to the NW the Asher Basin (Figs 2a & 3c) accumulated c. 2.5 km thick section of Liassic basalts and pyroclasts (Asher volcanics). The Deborah-2 well, located on the edge of the Asher Basin (Figs 2a & 3c), penetrated c. 200 m of the Asher volcanics and only c. 1 km thick Liassic to Bathonian section. Whereas, increased thickness of the Bajocian and Bathonian strata is found in outcrops at Mount Hermon (Goldberg et al. 1981; Hirsch et al. 1998). These thickness variations suggest segmented subsidence at the northern part of the Judea Graben (Fig. 2a) during Early to Middle Jurassic time. Reduced thickness of the Liassic to Bathonian sections in the Maanit, Gaash-Meged and Gevim highs (Fig. 3) indicates uplifting of the coastal area. The vertical offset on the Helez Fault, at the northwestern edge of the Gevim High amounts to 1.2 km (Figs 3b & 4) (Gardosh & Druckman 2006). In the southern edge of the Judea Graben, thickness changes of the Ardon and Inmar formations (Goldberg & Friedman 1974; Druckman 1977; Buchbinder & le Roux 1993) are associated with Early–Middle Jurassic tectonic activity. Further to the SW, in the North Sinai Basin the Liassic to Bathonian section is more then 2000 m thick (Halal-1 well and the Gebel Magahra outcrops; Druckman 1977; Derin & Gerry 1986a; Alsharhan & Salah 1996; Hirch et al. 1998), indicating significant subsidence during the Early to Middle Jurassic period. The thick sedimentary section in the Pleshet Basin and in the graben NE of the Jonah High (Figs 2a & 3a, b) may be partly attributed to this faulting phase. Similar to the Triassic conditions it is unknown whether the Levant Basin was a deep-marine basin during Early Jurassic time. The lithology of the Middle Jurassic Shederot Formation in the Yam West-1 well offshore indicates shallow marine conditions (Gardosh et al. 2006, 2008). Likewise, shallow-marine to fluvio-deltaic depositional environments characterize the Lower to Middle Jurassic section throughout the Levant onshore. Continuous, high-amplitude seismic reflections found below the Middle Jurassic unconformity in offshore seismic profiles are interpreted as shelf carbonate (Fig. 8). This raises the possibility that

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M. A. GARDOSH ET AL.

the Jurassic rift structures of the Levant area developed on a continental to shallow-marine platform. Deeper-marine conditions may have locally prevailed in parts of the Pleshet Basin and in other offshore structural lows, accompanied by growth of carbonate buildups on the adjacent Jonah, Leviathan and Eratosthenes highs. The Early Jurassic rifting phase was associated with extensive magmatic activity. Extrusive magmatic rocks were penetrated by wells in the Asher Basin at northern Israel (Asher Volcanics) (Dvorkin & Kohn 1989; Kohn et al. 1993). Magnetic anomalies that are interpreted to be associated with Early Jurassic volcanism are found in the Eratosthenes and Jonah highs offshore (Fig. 2) (Zverev & Ilinsky 2000, 2005; Gardosh et al. 2006, 2008). A large magnetic anomaly found within the Judea Graben in central Israel (Fig. 2a) is also interpreted to be associated with Early Jurassic magmatic activity, probably of a more intrusive nature (Garfunkel 1989; Rybakov et al. 1995). The activity of the Levant rift system ceased during the Middle to Upper Jurassic period. The Oxfordian –Cimmerian and the younger, Cretaceous section in the inner part of the Levant show minor thickness variations indicating tectonic quiescence. At this time conspicuous, carbonate shelf-edge developed on the eastern part of the Levant Basin, along the present-day Mediterranean coast, and large volumes of deepwater siliciclastics and carbonates were deposited on the slope further to the west (Cohen 1976; Bein & Gvirtzman 1977; Gardosh 2002). This accumulation of deepwater sediments within the Levant basin is associated with post-rift subsidence, caused by cooling of newly formed crustal units (Garfunkel & Derin 1984; ten Brink 1987; Gardosh 2002).

Tectonic reconstruction The Tethyan tectonic reconstruction of the Levant region is controversial. Some important aspects, such as the timing and location of rifting, are still under debate (Robertson & Mountrakis 2006). Several models have been proposed for the rifting direction and structural style (Fig. 11). Dewey et al. (1973), Bein & Gvirtzman (1977), Robertson & Dixon (1984), Stampfli & Borel (2002), and others, suggested a NE –SW opening in the Levant Basin that was accommodated by a major north– south oriented transform fault along the eastern Mediterranean coast (eastern transform margins) (Fig. 11a). According to Robertson (1998), the Levant margin represents either an orthogonally rifted or a transform margin. Garfunkel & Derin (1984) proposed rifting and extension in a NW– SE direction that resulted in separation of the Eratosthenes and Tauride blocks from the Gondwanian

Craton and was accommodated by a transform fault along the Sinai coast (southern transform margin) (Fig. 11b). Hirsch et al. (1995) described a north – south oriented extension on a Gondwanian shallow shelf with no reference to a specific fault pattern. According to the present analysis, the dominant direction of the normal faults offshore and onshore Israel and in northern Sinai is NE– SW to ENE – WSW (Fig. 2a). This direction indicates extension that is perpendicular to the strike of the normal faults, in a NW– SE and NNW–SSE direction as proposed by Garfunkel & Derin (1984) and Garfunkel (1998) (Fig. 11d). The NW –SE direction of the extension in the Levant Basin implies an accommodating transform fault somewhere along its southern margin. It is unlikely that such a fault extended along northern Sinai, as this area appears to have been dominated by a NW–SE extension (Fig. 2a). Alternatively, a southern transform fault may have been located along the African coast west of the Nile Delta (Fig. 11d), as shown by Bentham et al. (2007). Minor, transfer fault zones may also have existed within the Levant Basin (Bentham et al. 2007). An apparent, left-stepping offset in the northern edge of the Jonah ridge (Fig. 2b) can be explained by such a fault. The existence of eastern transform margin, which requires north– south direction of extension in the Levant Basin and a strike-slip fault along the Levant coast (Fig. 11a) (Dewey et al. 1973; Stampfli & Borel 2002), is not consistent with the structures and fault directions that are identified within the Levant Basin. Furthermore, this model can not explain the formation of the coastal highs and the interior basins of the Levant that are located east of the presumed Levant transform fault. An additional important aspect of the tectonic reconstruction is the timing of activity of Tethyan structures. Previous views advocated a time span ranging from Carboniferous to Permo-Triassic (Stampfli et al. 2001), Late Permian to Late Triassic (Robertson 1998), Triassic to Liassic (Garfunkel 1998) and Triassic and Late Jurassic to Early Cretaceous (Hirsch et al. 1995). The present analysis indicates three main phases of activity. Vertical motions started in several locations during the Late Palaeozoic, either in the Permian or the Carboniferous (Fig. 11c), continued during the Middle and Late Triassic, and climaxed in the Early Jurassic (Fig. 11d) when most of the structures were active and magmatic activity was widespread. Although the intensity of the faulting significantly decreased towards the Middle Jurassic it appears that some activity continued through the Bathonian and possibly later, till the early Late Jurassic. According to the time frame described above, Tethyan rifting in the Levant was pulsated, and extended over a period of more then 120 Ma. Its initiation may

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

(a)

(b)

(c)

(d)

25

Fig. 11. Several, alternative tectonic reconstructions of Tethyan rifting in the Levant region. Two, previously proposed models are: (a) after Dewey et al. (1973) and Stampfli & Borel (2002), showing north– south extension with eastern transform margin and (b) after Garfunkel & Derin (1984) and Garfunkel (1998), showing NW–SE extension with southern transform margin. The reconstructions in (c) and (d) are based on the present study. Tethyan rifting activity on the northern edge of Gondwana was pulsed and progressed from the Late Palaeozoic (c) to Early Jurassic (d). The southeastern part of the Mediterranean region was not affected by sea-floor spreading during rifting. Oceanic lithosphere was formed further to the north and NW, between the Eratosthenes block and the Taurid micro-continent, as indicated by ophiolitic complexes found in Antalya, Mamonia and Baer-Bassit. ER, Eratosthenes block; AN, Antalya; MA, Mamonia; PA, Palmyra; BB, Baer-Bassit.

have been contemporaneous with the time of opening of the ‘Palaeotethys’ whereas its most active phase took place during the development of the ‘Neotethys ocean’ (Robertson 2007). The available well and seismic data suggest that break-up took place in an intra-continental setting. There is no evidence for the development of a deepmarine basin in the Levant region throughout the Late Palaeozoic to Middle Jurassic. Deep marine conditions were established within the Levant Basin in the Late Jurassic and Early Cretaceous, a period considered by most authors to be associated with post-rift cooling and subsidence (Garfunkel & Derin 1984; ten Brink 1987; Gardosh 2002). On the other hand, north of the study area (Cyprus, NW Syria and southern Turkey), there is a record

of deep-marine conditions already in the Triassic (Robertson 1998; Garfunkel 1998). Available data do not allow assessing the nature of the transition from these areas to the Levant. Perhaps the most problematic aspect of the Tethyan tectonic reconstruction is the effect of rifting processes on the shaping of the Levant crust. This subject is discussed in the following chapter.

Shaping of the Levant crust during rifting The deep structure of the Levant continental margin The deep structure of the Levant was investigated by several wide-angle seismic refraction tests, shot

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M. A. GARDOSH ET AL.

onshore (Ginzburg & Folkman 1980; Weber et al. 2004) and offshore Israel (Makris et al. 1983; Ben-Avraham et al. 2002; Netzeband et al. 2006). Although differing in some details, their results provide a consistent picture of the crustal properties, summarized in Figure 12. The refraction profiles show that the depth to the Moho decreases from about 30 –35 km underneath southern Israel to about 20–22 km in the central part of the Levant Basin offshore (Fig. 12) (Makris et al. 1983; Ben-Avraham et al. 2002; Netzeband et al. 2006). A corresponding change was observed in the seismic velocities and thickness of the crust below the sedimentary cover. In southern Israel the crust is 25– 35 km thick and is separated to an upper (Vp ¼ 6.2–6.4 km/s) and a lower (Vp ¼ 6.7 km/s) layer (Fig. 10) (Makris et al. 1983; Ben-Avraham et al. 2002; Weber et al. 2004), whereas in the Levant Basin it is only 8 –10 km thick. Makris et al. (1983) and Ben-Avraham et al. (2002) interpreted one crustal unit in the Levant Basin with seismic velocity of 6.7 km/s. Netzeband et al. (2006) recently interpreted a thin (2–3 km) upper (Vp ¼ 6.0–6.3 km/s) and a lower (Vp ¼ 6.5–6.8 km/s) layers (Fig. 12). A change in crustal properties is observed further to the west. In the Eratosthenes area, the depth to the Moho increases to 26–28 km and its thickness to NW Eratosthenes Seamount 0

Levant Slope

Levant Basin

Cenozoic 6 km/s

Plz.-Tr. 6.0–6.3 km/s Depth (km)

Jur. Volcanic(?) Intruded upper crust 7.8–8 km/s

Mantle 0

40

SE

Negev

0

Jur.

Precambrian

6.2–6.4 km/s

50 km

10

Upper continental crust

6.5–6.8 km/s

20 6.7 km/s

30

coastline

Pliocene Jur.-Cret.

Messinian 10

23 km (Fig. 12) (Makris et al. 1983; Ben-Avraham et al. 2002; Netzeband et al. 2006). Makris et al. (1983) interpreted an upper (Vp ¼ 6.0 km/s) and lower (Vp ¼ 6.7 km/s) layers underneath the seamount (Fig. 12), whereas Zverev & Ilinsky (2000, 2005) suggest a more complex layering pattern of high and low velocity layers ranging from 5.3 to 7.7 km/s. In the Galilee area of northern Israel, BenAvraham & Ginzburg (1990) interpreted a 20 –25 km thick crust below the sedimentary cover. The relatively thinner crust of this area, in comparison to southern Israel, is associated with a uniform seismic velocity of 6.1–6.5 km/s (Ginzburg & Folkman 1980). The change in the crustal properties is indicated also by gravity data. The regional Bouguer gravity map (Fig. 2c) shows negative anomaly associated with a lighter, continental type crust south and east of the Mediterranean coastline and a positive anomaly associated with a denser crust in the Levant Basin (Rybakov et al. 1997; Rybakov & AlZoubi 2005). A similar, although less pronounced, change is observed from southern Israel northward (Fig. 2c). It is reasonable to assume that the variations in thickness and velocity of the Levant crust are related to the Tethyan break-up and rifting activity.

20

Lower continental crust 6.7 km/s

30

7.9 km/s 40

Fig. 12. Crustal scale section across the Levant Basin and margin, from southern Israel (SE) to the Erathosthenes seamount (NW); adapted from Garfunkel (1998). Structure and stratigraphy in the upper crust are taken from sections a and b in Figure 3 of this study; crustal thickness and seismic velocities are based on wide-angle seismic refraction tests (see details in text). The total crustal thinning along this section is considerably larger than the extension produced by the high-angle faults in the upper crust, indicating an extension discrepancy (see details in text). Depth-dependant-stretching derived by upwelling, divergent mantle flow (schematically shown by curved arrows) is associated with mobilization of the lower crust and part of the upper crust, along low-angle detachment surfaces (thick dashed line). The thinned crust in the centre of the Levant Basin is interpreted as the remaining, attenuated Upper continental crust that is heavily intruded by mantle material. In the Eratosthenes seamount, relatively thick crust with low seismic velocity and conspicuous magnetic anomaly may be explained by the existence of a large volcanic edifice of Early Mesozoic age. See location of the cross-section in Figure 2c.

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

Crustal thinning, subsidence, and accumulation of thick sedimentary wedges that are observed in the Levant area, characterize rifted, continental margin settings. It is, therefore, useful to discuss the evolution of the Levant margin with regards to other margins worldwide.

A global perspective on passive continental margin formation It is generally accepted that passive continental margins are created as a result of rifting causing extension and often also magmatism that leads to breakup of continents and separation of their fragments, followed by the formation of deep marine basins and an oceanic crust along the margins of the separating continental fragments. Rifted continental margins form transition zones between thick, relatively undisturbed continental crust and a newly formed, thin oceanic crust outboard. Based largely on the study of the central and northern Atlantic, two margin types were distinguished: volcanic and non-volcanic (Robertson 2007). Volcanic margins contain large volumes of intrusive and/or extrusive igneous rocks and are characterized by seaward-dipping seismic reflectors (SDR). The Greenland and the Norwegian margins of the north Atlantic are good examples (Planke et al. 2000). Non-volcanic margins generally lack extrusive igneous rocks, contain exhumed Upper mantle material, and are characterized by large rotated fault blocks. A typical example is the Iberian margin of the eastern Atlantic (Montadert et al. 1979). Breakup and separation of continental fragments, which move at large angles to the trend of the initial rift, provide a suitable tectonic setting for thinning of the lithosphere along newly formed continental margins of these two types. Recent studies of many rifted continental margins, both volcanic and non-volcanic, reveal an important characteristic; a so-called ‘extension discrepancy’ (Reston 2007). Studies from the Newfoundland and Iberia margins of the north Atlantic (Reston 2007), the northwestern Australian margin (Driscoll & Karner 1998), the Norwegian margin (Kusznir et al. 2004), and other areas, show that the extension and thinning of the whole crust and lithosphere greatly exceeds that of the upper crust.

Extension discrepancy in the Levant margin The variations in thickness of the Levant crust are well documented by wide-angle refraction profiles, as described above (Fig. 12). The new, offshore seismic reflection data allow imaging the entire Levant basin-fill down to the top of the crystalline

27

basement (Figs 7 & 9). With this new information, it is possible to estimate the amount of extension that was produced by brittle deformation during Tethyan rifting activity. The amount of extension is estimated from the geometry of the faults that displace the Palaeozoic and Mesozoic strata along the regional cross-section shown in Figure 12. There are 26 main fault zones in the area extending from southern Israel to the eastern edge of the Eratosthenes block (Fig. 12). These fault zones are comprised of normal faults that bound Tethyan grabens and horsts and reverse faults that are assumed to be reactivated Tethyan rift structures. By using the amount of apparent dip– slip on these faults we estimate the cumulative lateral displacement to be in the range of 6–7 km. The length of the section is 360 km (Fig. 12); therefore, the amount of extension is about 2%. On the same section the crystalline crust thins from southern Israel to the centre of the Levant Basin by 75% (Fig. 12). Clearly, the overall crustal thinning largely exceeds an extension produced by brittle deformation. It should be noted that our 2D estimation of fault related extension is highly simplified and it is reasonable to assume that more faults actually exist. However, even increasing the amount of faults and their vertical displacement by a factor of 2–3 would not fundamentally change the above relations. It is, therefore, concluded that, similar to many other margins worldwide, the Levant continental margin displays an extension discrepancy. Extension discrepancy at continental margins may be explained by partitioning of extension with depth, also termed depth-dependent-stretching (DDS) (Driscoll & Karner 1998; Reston 2007). Davis & Kusznir (2004) and Kusznir & Karner (2007) proposed that the mechanism that produces DDS is upwelling and divergence of continental lithosphere and asthenosphere. These authors applied a fluid-flow model for deformation of continental lithosphere and based on their modelling results concluded that an upwelling, divergent flow may cause thinning, followed by breakup, rifting and sea-floor spreading. Variations in the form and velocity of the divergent flow may lead to diversity of a rifted continental margin structure (Kusznir & Karner 2007). Can the DDS model be used to explain the extension discrepancy of the Levant continental margin? To further discuss this issue two regions are distinguished. The first region is the present-day coastal area and nearby slope (Fig. 12). There, the crystalline crust resembles the crust further inland in its two-fold seismic-velocity structure, although the upper and lower crustal units become notably thinner in this area compared with the normal continental crust farther inland. The second region is

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the basin outboard the base of the continental slope (Fig. 12). There, the crystalline crust is thin and its average Vp is higher than the Vp of a typical continental crust, although in a relatively thin layer at its top the Vp may resemble that of an upper continental crust (Netzeband et al. 2006).

The coastal area and nearby slope Here the crust is thinned by a factor of 1.2–1.3. If thinning was the result of uniform stretching of the entire crust, then this should have been expressed by fault structures in its brittle upper part. Seismic reflection data reveal faulting (i.e. the coastal highs and the Pleshet Basin) but as noted previously the extension that it produces is quite inadequate to account for the observed thinning. As the uppermost crust was hardly stretched, one should look for processes at greater depth. Igneous activity is a likely process. Early Mesozoic volcanic rocks were found in several wells close to the Levant margin and offshore magnetic anomalies are probably caused by coeval volcanic activity that is widespread in the Eastern Mediterranean region. An important indication for the nature of this activity is provided by the Gevim quartzporphyry in the Helez Deep-1well (Fig. 4). This is c. 200 m thick, fine-grained volcanic unit of PermoTriassic age (Segev 2005). The significant feature of these rocks is their high initial 87Sr/86Sr. Steinitz (1980) found I ¼ 0.7100 + 0.029 (age 244 + 44 Ma); Segev and Eshet (2003) reinterpreted these data (rejecting one of the results) and proposed I ¼ 0.7075 + 0.0034 (age 275+ 47 Ma). These values show that the volcanics were formed by heating and melting of older mid-crustal rocks by a large intrusion of basic magma. Magma ascent was probably arrested below the cold lowdensity upper crust. Although the size of this presumed igneous body cannot be determined, it seems unlikely to have been an isolated feature, because magma formation is expected, by its nature, to occur in a region of substantial size. These considerations raise the possibility that hidden igneous intrusions are present in other places along the Levant margin. Their distribution can not be inferred from the known extent of volcanism. Neither can it be revealed by seismic reflection or refraction data. Deep intrusions are expected to heat and thus weaken the lower crust and underlying mantle. The existence of magma intrusions on the Levant margin supports the upwelling mantle flow model of Kusznir & Karner (2007). The weakened lower crust may have been stretched and thinned beneath the little deformed, cold uppermost crust resulting with DDS (Fig. 12). Reston (2007) noted that explaining the extension discrepancy by DDS

requires that the lower crust has somehow been displaced away. Such displacement may be produced by shear along low-angle detachment (Reston 1993; Driscoll & Karner 1998) or by necking or boudinage of the lower crust (Reston 2007). Decoupling of the crust by shear has been inferred from seismic reflection data in the North Sea (Reston 1993). The existence of deep shear zones on the Levant margin is postulated (Fig. 12), though it can not be proved or disproved owing to insufficient seismic data.

The basinal domain The main features of this domain, at the central part of the Levant Basin (Fig. 12) are the small thickness of the underlying crystalline crust and its high seismic velocity compared with normal continental crust. Shaping of this crust by uniform stretching is not compatible with either the minor brittle deformation observed in its upper part or with its high average Vp. Here also, crustal thinning may be explained by upwelling of a divergent mantle flow, as suggested by the DDS model (Fig. 12). Kusznir & Karner (2007) showed that a thinned upper crust lying directly on the mantle, as observed in the centre of the Levant Basin, is predicted by their upwelling divergent flow model, and can occur with pausing of the flow field prior to its reaching the surface. Similar to the coastal area, the complete displacement of the lower crust within the basin may have been produced by various types of decoupling mechanisms (Fig. 12) (Reston 2007). The higher velocity of the crust in this area is an expected result of the addition of basic intrusions from the postulated upwelling mantle, since their Vp is higher than that of normal lower continental crust.

The Red Sea analogue The processes that shaped the present-day northern Red Sea (Fig. 1) may provide modern analogues for the formation of the Levant basin. The Red Sea is an active rift system that formed by break-up of the continental lithosphere beginning in the Late Oligocene, leading to sea-floor spreading by about 5 Ma (Cochran 2005; Cochran & Karner 2007; and references therein). Two parts of this rift system are distinguished: (a) the southern Red Sea, where an oceanic sea-floor spreading centre that produces new crust with oceanic magnetic anomalies has developed, and (b) the northern Red Sea where organized sea-floor spreading is not observed (Gaulier et al. 1988; Martinez & Cochran 1988). Seismic refraction data show that the northern Red Sea is characterized by a thin crust (7–8 km)

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

with low, continental type seismic velocities (Vp ¼ 6.2–6.3 km/s), that overlies the mantle (Gaulier et al. 1988). Large-amplitude, linear gravity highs and lows observed in the northern Red Sea are interpreted as series of crustal-scale fault blocks that are oriented sub-parallel to the trend of the rift. Small, scattered magnetic anomalies are interpreted as axial volcanoes (Martinez & Cochran 1988; Cochran & Karner 2007). The physical characteristics of the Levant Basin are similar to those of the northern Red Sea, except for the somewhat higher velocity of the crust. It is, therefore, possible that the two basins were formed in a similar manner. Martinez & Cochran (1988) suggest that extension of the northern Red Sea is accommodated by block rotation at the upper crust and ductile flow within the lithosphere but no explanation for the complete removal of the lower crust, as observed by refraction data (Gaulier et al. 1988) is presented. Given the large separation between the two sides of the northern Red Sea rift that amounts to 130 –150 km and the small number of fault blocks (Martinez & Cochran 1988; Cochran & Karner 2007), it is most likely that ‘extension discrepancy’ occurs in this basin. An alternative model for the evolution of the northern Red Sea may be DDS, associated with some brittle deformation in the upper crust and removal of the lower crust by an upwelling mantle flow. Upwelling of magma in the basin is supported by the existence of small magnetic anomalies that are interpreted as extrusive igneous rocks, exposures of young basalts along segments of the basin axis, and by the very high-heat flow (Cochran & Karner 2007). It is worthwhile discussing the relations between the north and south parts of the Red Sea. Martinez & Cochran (1988) and Cochran (2005) proposed that the two parts of the rift represent different stages of evolution and with continued extension and magmatism, organized sea-floor spreading will eventually take place in the northern Red Sea. An alternative model suggested by Cocharan & Karner (2007) postulates that the two regions develop differently owing to differences in lithosphere rheology and it is possible that an oceanic spreading centre will not develop in the northern part. The two models reflect on the regional context of the analogous Levant Basin. Similarly, this basin may have developed differently than other rifted Tethyan domains owing to an inhomogeneous lithosphere.

Discussion and conclusions Rift structures of the Palaeozoic to Early Mesozoic age were previously recognized in well and seismic

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data in the inland part of the Levant region. New, 2D seismic reflection profiles reveal similar structures in the Levant Basin offshore Israel. Integration of the onshore and offshore data allows reconstructing the pattern of this important rifting phase. Several fault belts that are generally oriented in a NE–SW direction, more or less parallel to the presentday coastline, are identified (Fig. 2a): (a) the inner basins, (b) the highs along the Mediterranean coastline, (c) the offshore highs and lows and (d) the Eratosthenes high. The fault belts comprise a broad zone of extension and rifting, more than 400 km wide, from southern Israel to the far offshore (Fig. 2a). Seismic and gravity data indicate that this zone extends further to the SW and north, below the Nile Delta and northern Egypt (Aal et al. 2000; Bentham et al. 2007) as well as onshore and offshore Lebanon (Walley 1998; Roberts & Peace 2007). The extensional structures of the Levant form part of a series of rifted margins that developed on the northern edge of Gondwana and are found today in a wide region extending from North Africa to northern India (Robertson 2007). The distribution of structures within the Tethyan rift belt of the Levant suggests an extension in a NW –SE direction. This extension was probably accommodated by the transform-rifted margin in western Egypt and North Africa (Garfunken & Derin 1984; Bentham et al. 2007). An alternative model of north –south oriented extension and transform-rifted margin along the Levant coast (Dewey et al. 1973; Robertson & Dixon 1984; Stampfli & Borel 2002) is not compatible with the direction and style of deformation observed throughout the Levant. Tethyan rifting in the Levant was pulsated and accentuated during three periods: Late Palaeozoic, Triassic and Early Jurassic (Garfunkel & Derin 1984; Garfunkel 1998). Evidence for Permian differential motions is found in well and seismic data from the Palmyra Trough (Sawaf et al. 2001) and the Gevim High near the Mediterranean coast (Figs 2a, 4 & 5). A small basin found NE of the Gevim structure (Fig. 5) may have formed prior to the deposition of the Upper Permian Arqov Formation, possibly implying Carboniferous to Early Permian time of activity. The great thickness of the sedimentary fill in the Pleshet Basin and other Tethyan lows offshore may be better explained by initiation of subsidence during the Late Palaeozoic period. In other parts of northern Gondwana, evidence for Permian breakup and rifting that was followed by the creation of oceanic crust is found in Oman and North India (Robertson 2007). Robertson (2006) recently identified deepwater, siliciclastic turbidites of Mid–Late Carboniferous to Early

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Permian age in western Sicily and Crete and suggested that a deep-marine rift opened along the northern margin of Gondwana during this time. In summary, Late Palaeozoic rifting in the Levant area may have been more extensive and started earlier then previously considered. This rifting pulse produced vertical motions and some extension, but there is no evidence for significant magmatic activity and the formation of an oceanic crust as suggested by Stampfli et al. (2001). A second rifting pulse took place in the Triassic, activating the inner basins, the coastal highs, and the offshore highs and lows (Figs 2a & 3). Well data indicates two faulting episodes: during the Anizian (i.e. the syn-tectonic Erez Conglomerate in Helez Deep-1 well) and the Carnian (i.e. the Mohilla Evaporites). Although, these may reflect one pulse that gradually intensified towards the end of the Triassic. Minor magmatic activity is indicated by Triassic volcanic rocks found in wells at the Negev and Palmyrides areas (Segev 2005). In other parts of the Eastern Mediterranean, Triassic deep-water sediments and extensive volcanism, in places of mid-ocean range (MOR) type were found in the Baer– Bassit ophiolitic section of northern Syria, in the Antalya Complex of SW Turkey and in the Mamonia Complex of western Cyprus (Fig. 11) (Robertson 1998, 2007). These indicate the formation of well developed rifted continental margins and deep-marine basins. The Baer–Bassit section was located on the northern Gondwana margin whereas the Antalya and Mamonia Complexes formed the southern margin of the Tauride micro-continental block, at the opposite side of the Southern Tethys Ocean (Fig. 11) (Robertson 2007). In the Levant area, there is no indication for MOR type volcanism and the existence of a deep-marine basin during the Triassic. The lack of continental margin characteristics suggests that a deep Tethyan ocean did not extend as far south as the Levant. The most significant rifting pulse took place in the Levant during the Early Jurassic. Evidence for large-scale vertical motions and alkaline volcanism (i.e. Asher volcanics) are found throughout the Levant rift belt (Figs 2, 3 & 11d ), except for the Palmyra Trough that was not active during this time. The Early Jurassic tectonic pulse appears to be particularly significant in the Levant region. Other rifted margins of northern Gondwana that were formed in the Permo-Triassic (i.e. Baer–Bassit, Antalya, Mamonia) continued to subside, although rift-related activity and volcanism generally ceased (Robertson 2007). The time and style of rifting activity in the Levant are relatively well constrained by outcrop, well and seismic data. The effects of this activity on the shaping of the Levant crust are less well

understood and still being argued. The Levant Basin and adjacent coastal area are described by many authors as Mesozoic, passive continental margin (Bein & Gvirtzman 1977; Makris et al. 1983; Garfunkel & Derin 1984; ten Brink 1987; Ben-Avraham et al. 2002; Gardosh 2002). There are two main lines of evidence: (a) the marked westward thinning of the crust (Fig. 12) and (b) the formation of a shelf-edge and the accumulation of deepwater section within the Levant Basin during the Late Jurassic to Middle Cretaceous. Breakup and rifting that was followed by post-rift subsidence provides a tectonic framework to explain the formation of the Levant continental margins. In this framework two alternative models may be considered: (a) rifting–drifting and (b) DDS. The rifting–drifting model postulates breakup and separation between the African craton and the Eratosthenes seamount, which is assumed to be rooted in a continental crust. This separation was presumably followed by formation of a new crust with oceanic affinities in the centre of the Levant Basin. The rifting –drifting model raises several difficulties. There is no convincing geophysical evidence for the existence of an oceanic crust in the basin centre. The seismic velocity of the thin crust within the Levant Basin is not typically oceanic and in its upper part the velocity is similar to that of an upper continental crust (Fig. 12) (Netzeband et al. 2006). Linear magnetic anomalies that are characteristic of sea-floor spreading in other oceans are not observed. On the other hand, seismic reflection data show that Tethyan rift structures are preserved and the seismic characteristics of an oceanic crust are missing (Gardosh & Druckman 2006). Additionally, the coastal and slope area lack features that are found in other rifted continental margin such as: seaward-dipping-reflections (volcanic margin) or rotated continental blocks (non-volcanic margin). Finally, a comparison of the amount of brittle deformation in the upper crust to the overall crustal thinning reveals a considerable ‘extension discrepancy’ that is not explained by the rifting–drifting model. An alternative model for the evolution of the Levant margin may be DDS, derived by an upwelling, divergent mantle flow (Fig. 12) (Kuzsnir & Karner 2007). Mantle flow may explain thinning by removal of the lower crust and part of the upper crust (Hirsch et al. 1995). However, it also requires decoupling and shearing between the displaced crustal units (Reston 2007). Some brittle deformation in the remaining upper crust is compatible with the DDS model, assuming a pause of the flow field prior to its reaching the surface (Kuzsnir & Karner 2007). Intrusions from the mantle flow may explain the higher velocity at the lower part of the crust in the centre of the basin. Early

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN

Mesozoic intrusive and extrusive rocks found in wells along the Levant margin further indicate upwelling of mantle material. A modern analogue for the Mesozoic Levant Basin is the northern part of the Red Sea. In this rift basin, the DDS model may be similarly applied to explain crustal thinning in the absence of sea-floor spreading. Cochran & Karner (2007) speculated that, unlike the southern Red Sea, an oceanic spreading centre may not develop in its northern part. Similarly, in the Levant Basin, rifting activity stopped at an early magmatic stage (Gardosh & Druckman 2006). It is unclear why seafloor spreading that took place further to the north and west (Robertson 1998, 2007) did not occur in the Levant area. Possibly, it is related to variations in lithospheric properties or in the pattern of deep mantle convection. In any case, given the above considerations, the Mesozoic Levant Basin may be considered as a ‘failed breakup basin’ (Kusznir & Karner 2007) somewhat similar to the ‘aborted rift system’ proposed by Hirsch et al. (1995). Application of the DDS model implies that the amount of drifting in the Levant Basin was limited. In this framework, breakup from the Gondwanian craton and northward motion of the Eratosthenes block is not required. The seamount may constitute a Mesozoic volcanic edifice rooted within the same modified continental crust that is found in other parts of the basin (Zverev & Ilinsky 2005). The DDS model may be applied also to the areas of the Galilee and southern Lebanon. The reduced thickness (c. 14 –16 km) and high velocity of the crystalline crust and the higher Bouguer gravity anomaly in this area (Fig. 2c) bear similarity to the Levant Basin. The modification of the northern Levant crust during Tethyan rifting (Asher and Hermon basins) is an alternative model to accretion of oceanic terranes that is related to the closure of the Palaeo-Tethys, as proposed by Ben-Avraham & Ginzburg (1990). Well data from southern and central Israel show that during the Late Palaeozoic, Triassic and Early Jurassic rifting phases the Levant region was dominated by continental to shallow-marine environments. The seismic character of the marine reflection lines and the lithology of the Middle Jurassic strata in the offshore Yam West-1 well suggest that shallow marine conditions prevailed also during rifting in the area of the Levant Basin. On the other hand, the attenuation and thinning of the crust within the basin, either by rifting–drifting or DDS, should have resulted with subsidence and increase in water depth. At the present state of knowledge, this apparent disagreement is not resolved and requires further analysis of the basin’s subsidence history.

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Finally, Robertson (2007) noted that none of the Tethyan-rifted margins that developed in northern Gondwana show the ideal margin characteristics that are observed in the modern oceans. The deep structure of the Levant area and its complex evolution further demonstrates the diversity of margin forming processes that remain to be studied in both recent and ancient examples. The authors are grateful to Y. Mimran and A. Honigstein from the Israeli Petroleum Commissioner office for allowing the use the offshore seismic data and for supporting this study. The comments and suggestions of P. Bentham, D. Frizon de Lamotte, and an ammoniums reviewer are greatly appreciated. The authors thank M. Rybakov for his contribution to this study.

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K. C., Richter, C. & Camerlenghi, A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 160, 701 –708. Martinez, F. & Cochran, J. R. 1988. Structure and tectonics of the Red Sea: catching a continental margin between rifting and drifting. Tectonophysics, 150, 1– 32. Montadert, L., Roberts, D. G., de Charpal, O. & Guennoc, P. 1979. Rifting and subsidence of the northern margin of Bay of Biscay. Initial Report Deep Sea Drilling Project, 48, 1025–1060. Moustafa, A. R. & Khalil, M. H. 1990. Structural characteristics and tectonic evolution of north Sinai fold belt. In: Said, R. (ed.) The Geology of Egypt. A. A. Balkema, Rotterdam, 381–389. Mouty, M. 1997. Le Jurassique de la chaine de Palmyrides (Syrie centrale). Bulletin de la societe Geologique de France, 168, 181– 186. Netzeband, G. L., Gohl, K., Hu¨bscher, C. P., Ben-Avraham, Z., Dehghani, G. A., Gajewski, D. & Liersch, P. 2006. The Levantine Basin – crustal structure and origin. Tectonophysics, 418, 167– 188. Planke, S., Symonds, P. A., Alvestad, E. & Skogseid, J. 2000. Seismic volcanostratigraphy of large-volume basaltic extrusive complexes on rifted margins. Journal of Geophysical Research, 105(B8), 19335–19351. Reston, T. J. 1993. Evidence for extensional shear zones in the mantle, offshore Britain, and their implications for the extension of the continental Lithosphere. Tectonics, 12, 492 –506. Reston, T. J. 2007. The formation of non-volcanic rifted margins by the progressive extension of the lithosphere: the example of the West Iberian margin. In: Karner, G. D., Manatschal, G. & Pinheiro, L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 77–110. Roberts, G. & Peace, D. 2007. Hydrocarbon plays and prospectivity of the Levantine Basin, offshore Lebanon and Syria from modern seismic data. GeoArabia, 12, 99–124. Robertson, A. H. F. 1998. Mesozoic-Tertiary tectonic evolution of the easternmost Mediterranean area: integration of marine and land evidence. In: Robertson, A. H. F., Emeis, K. C., Richter, C. & Camerlenghi, A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 160, 723–782. Robertson, A. H. F. 2000. Mesozoic-Tertiary tectonicsedimentary evolution of a south Tethyan oceanic basin and its margins in southern Turkey. In: Bozkurt, E., Winchester, J. A. & Piper, J. D. A. (eds) Tectonics and Magmatism in Turkey and the Surrounding Area. Geological Society, London, Special Publications, 173, 97–138. Robertson, A. H. F. 2006. Sedimentary evidence from the south Mediterranean region (Sicily, Crete, Peloponnese, Evia) used to test alternative models for the regional tectonic setting of Tethys during Late Palaeozoic–Early Mesozoic time. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 91–154.

TETHYAN RIFTING IN THE LEVANT BASIN AND MARGIN Robertson, A. H. F. 2007. Overview of tectonic settings related to the rifting and opening of Mesozoic ocean basins in the Eastern Tethys: Oman, Himalayas and Eastern Mediterranean regions. In: Karner, G. D., Manatschal, G. & Pinheiro, L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 325– 338. Robertson, A. H. F. & Dixon, J. E. 1984. Introduction: aspects of the geological evolution of the eastern Mediterranean. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 1– 74. Robertson, A. H. F. & Mountrakis, D. 2006. Tectonic development of the Eastern Mediterranean region: an introduction. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 1– 9. Robertson, A. H. F. & Xenophontos, C. 1993. Development of concepts concerning the Troodos ophiolite and adjacent units in Cyprus. In: Prichard, H. M., Alabaster, T., Harris, N. B. W. & Neary, C. R. (eds) Magmatic Processes and Plate Tectonics. Geological Society, London, Special Publications, 76, 85–119. Robertson, A. H. F., Dixon, J. E. et al. 1996. Alternative tectonic model for the Late Paleozoic–Early Tertiary development of Tethys in the Eastern Mediterranean. In: Morris, A. & Tarling, D. H. (eds) Paleomagnetism and tectonics of the Mediterranean region. Geological Society, London, Special Publications, 105, 239– 263. Rybakov, M. & Al-Zoubi, A. 2005. Bouguer gravity map of the Levant – a new compilation. In: Hall, J. K., Krasheninnikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant, Vol. II, the Levantine Basin and Israel, 539–542. Rybakov, M., Fleischer, L. & Goldshmidt, V. 1995. A new look at the Hebron magnetic anomaly. Israel Journal of Earth Science, 44, 41–49. Rybakov, M., Goldshmit, V. & Rotstein, Y. 1997. New regional gravity and magnetic maps of the Levant. Geophysical Research Letters, 24, 33–36. Rybakov, M., Goldshmit, V., Fleischer, L. & Ben-Gai, Y. 2000. 3-D gravity and magnetic interpretation for the Haifa Bay area (Israel). Journal of Applied Geophysics, 44, 353– 367. Segev, A. 2005. Phanerozoic magmatic activity associated with vertical motions in Israel and adjacent countries. In: Hall, J. K., Krasheninnikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant, Vol. II, the Levantine Basin and Israel, 553– 557. Segev, A. & Eshet, Y. 2003. Significance of Rb/Sr age of Early Permian volcanics, Helez Deep 1A borehole, central Israel. Africa Geoscience Review, 10, 333–345. Sengor, A. M. C. & Yilmaz, Y. 1981. Tethyan evolution of Turkey: a plate tectonic approach. Tectonophysics, 75, 181–241.

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Salhov, S. 1996. Gevim-1 composite log. Isramco Inc. Sawaf, T., Brew, G., Litak, R. & Barazangi, M. 2001. Geologic evolution of the intraplate Palmyride basin and Euphrates fault system, Syria. 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, Paris, 186, 441– 467. Stampfli, G. M. & Borel, G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth and Planetary Science Letters, 196, 17– 33. Stampfli, G. M., Mosar, J., Favre, P. A., Pillevuit, A. & Vannay, J. C. 2001. Permo-Mesozoic evolution of the western Tethyan realm: the Neotethys EastMediterranean Basin connection. 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, Paris, 186, 51–108. Steinitz, G. 1980. Rb –Sr age determinations on basement rocks from Helez Deep 1A well. Geological Survey of Israel Report MM/1/80, 1 –10. Stern, R. J. 1994. Arc assembly and continental collision in the Neoproterozoic E African orogen: implications for the consolidation of Gondwanaland. Annual Review of Earth and Planetary Sciences, 22, 319– 351. Ten Brink, U. 1987. Quantitative basin analysis with applications to the Israeli continental margin. Oil Exploration (Investments) Report, 24. Vidal, N., Alvarez-Marron, J. & Klaeschen, D. 2000. Internal configuration of the Levantine Basin from seismic reflection data (eastern mediterranean). Earth and Planetary Science Letters, 180, 77–89. Walley, C. D. 1998. Some outstanding issues in the geology of Lebanon and their importance in the tectonic evolution of the Levantine region. Tectonophysics, 298, 37–62. Walley, C. H. 2001. The Lebanon passive margin and the evolution of the Levantine Neo-Tethys. 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. Memoires du Museum National d’Histoire Naturelle, Paris, 186, 407 –439. Weber, M., Abu-Ayyash, K. et al. 2004. The crustal structure of the Dead Sea Transform. Geophysical Journal International, 156, 655 –681. Weissbrod, T. 1981. The Paleozoic of Israel and adjacent countries. Geological Survey of Israel, Report MP 600/81, 276 (in Hebrew with English abstract). Weissbrod, T. 2005. The Paleozoic in Israel and environs. In: Hall, J. K., Krasheninnikov, V. A., Hirsch, F., Benjamini, Ch. & Flexer, A. (eds) Geological Framework of the Levant, Vol. II, The Levantine Basin and Israel. Historical Production-Hall, Jerusalem, 283– 316. Wolfart, R. 1981. Lower Paleozoic rocks of the Middle East. In: Holland, C. H. (ed.) Lower Paleozoic of

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Structural setting and tectonic evolution of North Sinai folds, Egypt ADEL R. MOUSTAFA Department of Geology, Faculty of Science, Ain Shams University, Cairo 11566, Egypt (e-mail: [email protected]) Abstract: Detailed study of outcrop and subsurface data of North Sinai indicates the existence of a NE–SW oriented region with very large asymmetric folds lying between the Nile Delta hinge zone and the Sinai hinge belt. The steep southeastern flanks of these folds are often associated with highangle reverse faults. These folds continue northeastward in the Naqb Desert toward the Dead Sea transform. The North Sinai folds represent inversion anticlines formed by inversion of Mesozoic (Jurassic and probably Cretaceous) extensional basins during Late Cretaceous to pre-Miocene times. Mesozoic extension is related to the divergence between Afro-Arabia and Eurasia and opening of the Neotethys whereas inversion is related to the convergent movement between these two plates. The acme of inversion was at Campanian time. The central Sinai hinge belt is a zone of ENE–WSW oriented, right-lateral strike–slip faults that separate the folded area to the north from the tectonically stable area of central and southern Sinai. It responded to the convergent movement between Afro-Arabia and Eurasia by dextral transpression on the faults. Later reactivation of the eastern edges of these faults by drag on the west side of the Dead Sea transform took place in post-Miocene to Recent times.

The Sinai Peninsula lies at the northeastern part of the African plate. It is bounded on the east by the Dead Sea transform, on the west by the Gulf of Suez rift and the Suez Canal, on the north by the Eastern Mediterranean, and on the south by the Red Sea (Fig. 1). Precambrian crystalline basement rocks are exposed in the southernmost part of Sinai forming the Sinai Massif, which reaches a height of 2641 m above sea level. The northern part of these Precambrian basement rocks is unconformably overlain by Phanerozoic sedimentary rocks that have gentle northward tilt and become younger in age toward the north. These Phanerozoic sediments range from Palaeozoic above the Precambrian basement, to Pliocene in the subsurface of the northern coastal area of Sinai (Figs 1, 2 & 3). Cretaceous to Cenozoic (Middle Eocene) carbonate rocks form the surface of the Tih Plateau in central Sinai. A number of large hillocks exist in the northern part of Sinai between the Tih Plateau and the Mediterranean Sea and represent prominent highs within the plains covered mostly by Quaternary alluvium and sand dunes. G. (Gebel, mountain) Yelleg, G. Halal and G. Maghara are the most prominent of these hillocks and are 1090, 890 and 735 m above sea level, respectively. The three hills represent three large anticlines among other smaller anticlines in northern Sinai. Similar folds exist farther to the NE in the Naqb Desert, which, along with those in North Sinai, represent a NE– SW oriented fold system. This fold system is a segment of Krenkel’s (1925) ‘Syrian Arc’ that extends from the Palmyra folds of Syria into the desert west of the Nile, passing through North Sinai.

The present work is intended to throw light on the structural setting and tectonic evolution of the North Sinai folds. This work is based on the results of detailed surface geological mapping at scales of 1:50 000 and 1:20 000 carried out by the author and his graduate students since 1990. In addition, subsurface [borehole and two-dimensional (2D) seismic] data have also been used. North Sinai has been lacking detailed structural studies for a long period of time. Early attempts to study the structures in this area date back to the Early 20th century (Moon & Sadek 1921; Sadek 1928) as well as Shata (1959). The first detailed structural study in Sinai is perhaps that of Bartov et al. 1980. Moustafa & Khalil (1989) carried out detailed photogeological study of the North Sinai folds, which was followed by a program of detailed geological field mapping of the exposed structures. The present work is based on these detailed structural studies. Information about the subsurface of North Sinai includes borehole and seismic data. The latter was only acquired in the late eighties and early nineties as part of the limited hydrocarbon exploration in North Sinai. The onshore area of North Sinai is covered by few 2D seismic sections compared to the northern offshore area, where more seismic data were acquired. The North Sinai folds have NE–SW orientation and include both large (several tens of kilometres long) folds; for example, Gebels Maghara, Halal and Yelleg folds; and small (several kilometres long) folds; for example, Gebels Libni, Minsherah and Falig folds. These folds have been the subjects of several studies (see Moustafa & Khalil 1994 for list).

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 37– 63. DOI: 10.1144/SP341.3 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Simplified geological map of Sinai and the Naqb Desert after Geological Survey of Egypt (1981) and Geological Survey of Israel (2000) with the main tectonic features. Major folds in the area are Maghara (M), Yellleg (Y), Halal (H), Ramon (R), Kurnub (K) and Qatan (Q).

Concerning the geometry of the North Sinai folds, most workers agree that the large folds are highly asymmetrical, with gentle northwestern flanks, dipping at about 5–208NW and steeper southeastern flanks that are vertical or overturned in places. Some workers mapped reverse faults dipping at angles exceeding 458 in some of these folds (e.g. Shata 1959; Al-Far 1966; Bartov et al. 1980; Moustafa & Khalil 1989; Moustafa & Yousif 1990). On the other hand, Abdel Aal et al.

(1992) alleged the existence of low-angle thrusts in North Sinai based on observations of few seismic reflection profiles. All the reverse faults mapped on the surface in North Sinai have a consistent northwestward direction of dip indicating southeastward vergence. Moon & Sadek (1921) and Sadek (1928) mentioned that the North Sinai folds are arranged in distinct NE-oriented parallel lines (axes). Shata (1959) identified a gently folded area in north-central Sinai and a strongly

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Fig. 2. Simplified geological cross section along the Sinai Peninsula. Top Precambrian basement in the subsurface of North Sinai is after Rybakov & Segev (2004) and Abd El-Gawad & Ibraheem (2005). Rock units symbols stand for Palaeozoic (Pz), Mesozoic (Mz), Jurassic (J), Cretaceous (K), Paleocene (Tp) and Eocene (Te) rocks. A and T designate the sense of strike–slip movement on some faults (away from and toward the observer respectively). See Figure 1 for location.

folded area in North Sinai. The two areas are elongated in a NE–SW direction and are separated by a 20-km wide fractured area, which he called the Sinai hinge belt. The North Sinai folds and their associated reverse faults have been explained to be due to pureshear deformation involving tangential compressive stresses (e.g. Moon & Sadek 1921; Shukri 1954; Said 1962; Youssef 1968; Abdel Aal et al. 1992), simple-shear deformation model by a dextral shear couple (Moustafa & Khalil 1989, 1990; Moustafa & Yousif 1990; Moustafa et al. 1991), or a vertical stress model of basement uplifting (Shata 1959). Subsurface data of offshore North Sinai revealed the role of inversion tectonics in forming some structures similar to those in the onshore area (Ayyad & Darwish 1996). Inspired by this model, those authors also proposed the same mechanism for the onshore structures. Most workers agree that folding in northern Sinai took place over a long period of time. Shukri (1954) mentioned that these folds rose from the

bottom of the Cretaceous sea and have been intermittently submerged and emerged in later times up to the Late Oligocene. Shata (1959) proposed that folding acted throughout a long period of time during the Jurassic throughout most of the Early Miocene. Said (1962) stated that folding started as early as the Cenomanian and continued intermittently throughout the latest Cretaceous. He also mentioned that strong folding at the end of the Maastrichtian brought many of the folds above sea level and final uplift of North Sinai took place in post-Eocene time. Bartov et al. (1980) concluded that the Araif El Naqa fold started to develop during the Coniacian and continued developing in the Santonian and up to the Late Campanian. Faulting in this area is evident in two phases during the Senonian and in post-Early Neogene time (post 21 Ma). Sadek (1928) mentioned that the North Sinai folding took place in the Oligocene or Early Miocene time, whereas Farag & Shata (1954) mentioned that folding in G. El Minsherah was after the Early Miocene.

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Fig. 3. Stratigraphic sections based on borehole data and exposed stratigraphic section of G. Maghara showing the changes in thickness of Jurassic rocks across the Sinai hinge belt and of Cenozoic rocks across the Nile Delta hinge zone. Symbols designate: Palaeozoic (PZ), Triassic (Tr), Jurassic (J), Lower Cretaceous (Kl), Aptian– Albian (AA), Cenomanian (Kc), Turonian (Kt), Senonian (Ksn), Paleocene (Tp), Eocene (Te), Oligocene (Tineh, Tn; Qantara, Qt), Miocene (Sidi Salem, SS; Qawasim, QS; Abu Madi, AM) and Pliocene formations (Kafr El Sheikh, KS; El Wastani, WS; Mit Ghamr, MG).

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Stratigraphy The oldest exposed rocks in the study area are Triassic in age and are exposed in the core of G. Araif El Naqa anticline. Jurassic rocks have wider exposures in the area and are exposed in G. Araif El Naqa and G. Minsherah in addition to the main outcrop of G. Maghara. Jurassic rocks have also been penetrated by some boreholes in the study area. Cretaceous outcrops have wider distribution in the area and are exposed mainly in the three major fold structures of Gebels Yelleg, Halal and Maghara, as well as the smaller folds. Senonian and Cenozoic rocks are exposed with gentle attitudes in the topographically low areas lying between the high fold structures, as well as some small tablelands. The youngest exposed sedimentary rocks in North Sinai are of Middle Eocene, Oligocene, Pliocene and Quaternary age, whereas Upper Eocene, most of the Oligocene, Miocene and most of the Pliocene sediments are only found in the subsurface of the northernmost onshore area of Sinai, as well as in the offshore area. Basic igneous intrusives are also exposed in a few parts of the study area, some of which have Early Miocene age (Steinitz et al. 1978) whereas others are not dated (e.g. G. Araif El Naqa area). Similar, nearby basalt outcrops in G. Ramon (Ramon Basalt) are similar to some of those of G. Araif El Naqa and were dated as Early Cretaceous (111 –116 Ma); Geological Survey of Israel (2001). Although the Mesozoic and Cenozoic rocks of northern Sinai show a gradual northward increase in thickness, abrupt changes in thickness of the Mesozoic rocks (especially the Jurassic section) occur in the study area reflecting the presence of

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syn-depositional tectonism during the Mesozoic time. Regional changes in thickness of the different rock units are controlled by the Sinai hinge belt and the Nile Delta hinge zone (Fig. 3). The Sinai hinge belt, which extends in an ENE –WSW direction from the northern tip of the Gulf of Suez toward the Dead Sea Transform, represents the boundary between a tectonically stable (platform) area to the south and a tectonically active area to the north. The tectonically active area included some extensional basins that were actively subsiding during the Mesozoic (mainly Jurassic time). The Nile Delta hinge zone of Schlumberger (1984) or the Nile Delta flexure zone of Sestini (1989) has an arcuate outline and extends from the Bardawil Lake toward the Suez Canal and the central part of the cultivated Nile Delta (Fig. 1). It is a major zone of flexure along which listric normal faults, with a down to the north throw of 5–6 km at the Cretaceous and Middle Eocene levels exist, with abrupt changes in thickness of Cenozoic sediments across this zone (0.5– 1 km to the south versus 5 –7 km to the north), Sestini (1989). Bouguer gravity anomalies indicate a 2– 3 km thick sedimentary section to the south, v. 10 km thick section to the north of the Nile Delta hinge zone (Schlumberger 1995). Table 1 shows the maximum thicknesses of Mesozoic and Cenozoic rocks in the study area, and Figure 3a shows the effect of the Sinai hinge belt on the changes in thickness of the Jurassic rocks and the Nile Delta hinge zone on the changes in thickness of the Cenozoic rocks. Figure 3b also shows that several sub-basins existed in the area, lying north of the Sinai hinge belt as indicated by the changes in thickness of

Table 1. Maximum thicknesses of Mesozoic and Cenozoic rocks in the study area System/series/stage Oligocene Eocene Paleocene Senonian Turonian Cenomanian Lower Cretaceous Jurassic Triassic 1

Kuss & Boukhary 2008. Osman et al. 2000. 3 Farag & Shata 1954. 4 Jenkins 1990. 5 Al-Far 1966. 6 Eicher 1947. 7 Abed et al. 1996. 2

Maximum exposed thickness 1

77 m, G. Risan Aneiza 51þ m, east of G. Halal2 65 m, G. Minsherah3 295 m, G. Minsherah3 213 m, G. Halal2 624 m, G. Halal2 243 m, G. Halal4 250 m, G. Minsherah3 1980þ m, G. Maghara5 203.5þ m6 or 207þ m7, Araif El Naqa

Maximum subsurface thickness 152 m, subsurface of North Sinai 59 m, Nakhl well 326 m, Abu Hamth well 260 m, Nakhl well 370 m, Abu Hamth well 3234 m, Halal-1 well4 914 m, Halal-1 well4

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the Jurassic and/or Cretaceous sediments (cf. the Jurassic sections of G. Maghara and G. Halal as well as the thick Aptian –Albian section in the offshore Mango structure). The following is a brief review of the major stratigraphic units.

Triassic rocks The exposed Triassic section of G. Araif El Naqa was studied by different workers. Bartov et al. (1980) and Jenkins (1990) reported 185þ m thick section whereas Eicher (1947) and Abed et al. (1996) reported larger thicknesses (203.5þ m and 207 m, respectively). The Triassic section of this area includes two main units, a lower clastic unit and an upper carbonate unit. The lower clastic unit is 68 m thick and consists of dark and lightcoloured, fine- to medium-grained sandstone with shaly, marly and silty beds intercalated with a few fossiliferous limestone beds. It also includes two weathered dark igneous sills. The shales and sandstones contain plant imprints and petrified wood fragments. This sandstone unit was deposited in a terrestrial, shallow coastal, to shallow marine environment. The upper carbonate unit is 117 m thick and is made up of shallow marine limestone beds alternating with marls. It also includes an olivine basalt dyke. The upper part of the unit includes some gypsum intercalations.

Jurassic rocks Jurassic rocks are exposed in the core of the G. Maghara anticline (1900þ m; Al-Far 1966), G. Araif El Naqa (141 m), and G. El Minsherah (80þ m). The lowermost Jurassic sediments of G. Maghara (Mashabba Formation, Early Jurassic age) are made up of fluvial sandstones containing large wood fragments that were deposited by northerly flowing braided streams carrying detritus shed off the Arabo-Nubian massif. The Mashabba Formation is followed by interbedded shallow marine carbonates of the Rajabiah Formation and nearshore marine clastics of the Shusha Formation. Lower Jurassic sandstones and shales have also been penetrated by the Nakhl (162 m), Abu Hamth (162 m), Falig-1 (315þ m), and N. Falig-1 (383þ m) wells. A thick clastic sequence in the Ayun Musa-2 (622 m) and Hamra-1 (534 m) wells has been assigned a Middle– Early Jurassic age, while the El Khabra well has an undifferentiated Jurassic section (1430 m) of sandstones, shales and limestones. The Middle Jurassic sediments of G. Maghara are 660 m thick and include a lower carbonate unit (Bir Maghara Formation) and an upper clastic unit (Safa Formation). A Middle Jurassic clasticcarbonate sequence is also exposed at G. Minsherah

and was also penetrated in the Halal-1 (780 m), Falig-1 (310 m), N. Falig-1 (415 m), Katib El Makhazin (502þ m) and Giddi-1 (805þ m) wells. Several coal seams exist in the Lower and Middle Jurassic clastic units (Shusha and Safa Formations, respectively). The Upper Jurassic sediments (Masajid Formation) are dominated by carbonates representing a southerly marine transgression at the end of Bathonian–Callovian times. At G. Maghara, these carbonates are 575 m thick (Al-Far 1966). Shelf carbonates have also been recognized in the Halal-1 (214 m), Falig-1 (530 m), N. Falig-1 (635 m), Ayun Musa-2 (121 m), Hamra-1 (65 m) and Katib El Makhazin (388 m) boreholes. Drilling for hydrocarbon exploration in the northwestern part of G. Yelleg indicated the presence of another Upper Jurassic clastic unit above the Masajid Formation. This unit was named the Gifgafa Formation by geologists working in the Gulf of Suez Petroleum Company (see also Abdel Aal & Lelek 1994). The Gifgafa Formation is a shale section with some sandstone and a few limestone and dolomite streaks. It is 145 m thick in the Falig-1 well and 220 m thick in the N. Falig-1 well

Comments on the changes in thickness of the Jurassic rocks Although the thicknesses of Jurassic rocks penetrated by boreholes and reported in the last section represent true vertical rather than true stratigraphical thicknesses, all of them, perhaps except that of the Halal-1 well, can be considered true stratigraphical thicknesses as these wells occupy crestal positions of folded rocks. The thicknesses presented in the last section indicate that the changes in thickness of the Jurassic rocks in northern Sinai do not reflect a homogeneous northward increase in thickness as one would expect in a normal depositional setting. On a regional scale, Hirsch & Picard (1988) identified two separate ENE –WSW oriented Jurassic basins in the North Sinai–Levant area, namely the Maghara –Halal basin and the Helez – Ramallah trough with Jurassic thicknesses equal to 3000 and 4000 m, respectively. The Mahgara – Halal basin generally corresponds to the area with thick Jurassic rocks in Figure 3. These two basins are separated by a high area containing a thinner Jurassic section. Changes in thickness of the Jurassic rocks in G. Maghara and Halal-1 well indicate the existence of separate sub-basins in northern Sinai like those in the northern Western Desert. The tectonic evolution of the northern Western Desert was previously studied using numerous borehole and seismic data indicating a phase of extensional deformation in the Jurassic time, where several half graben-like basins were formed (Moustafa

NORTH SINAI FOLDS

et al. 1998; Abd El-Aziz et al. 1998). A similar situation might also have existed in northern Sinai. In this regard, G. Halal and G. Maghara would represent two separate Jurassic sub-basins. Table 2 shows a comparison of the thickness of the Jurassic rocks in the two areas indicating that basin opening in G. Maghara and G. Halal probably took place in the Early and Middle Jurassic Epochs, whereas the Late Jurassic Epoch was a period of gradual northward increase in thickness.

Cretaceous rocks Lower Cretaceous rocks. Fluvial-continental Lower Cretaceous sediments unconformably overlie the marine Upper Jurassic rocks as a result of a major eustatic fall in sea level. They belong to the Malha Formation, which has an Early Cretaceous age and have been penetrated in the Abu Hamth (370 m), Nakhl (247 m), Ayum Musa-2 (149 m), Hamra (366 m), Kabrit (103 m), Falig-1 (217þ m), N. Falig-1 (318 m), Katib El Makhazin (32 m) and Mango-2 (242þ m) wells. The exposed section of the Malha Formation is 160 –168 m at G. Yelleg (Jenkins 1990), 243 m thick at G. El Halal (Osman et al. 2000) and 250 m thick at G. El Minsherah (Farag & Shata 1954). The Malha Formation is overlain by an interbedded series of argillaceous clastics and carbonates of Aptian– Albian age. These represent fluvialparalic to shallow shelf facies deposited in a low energy system, with occasional high energy episodes, indicated by the presence of rudists and oolites. The Aptian– Albian sediments are 240 – 382 m thick in the G. Maghara area (Bachmann & Kuss 1998). Thick Aptian–Albian sediments reaching a thickness of 1749 m were encountered in the Mango-1 well in offshore North Sinai. Abu Zeid (2007), based on micropalaeontological work, considered the 300 m thick section overlying the Malha Formation in the northern part of G. Maghara area (G. El Tourkmaniya outcrop) to be of Barremian–Albian age, indicating that Cretaceous transgression in northern Sinai started in the Barremian time. Upper Cretaceous rocks. The Upper Cretaceous rocks in North Sinai are mostly carbonates and

43

form widespread exposures. The Cenomanian section in North Sinai is dominated by a thick sequence of dolomites that were deposited on a broad shallow subtidal shelf. At G. Halal, the Cenomanian Halal Formation is 624 m thick (Osman et al. 2000). According to Bartov & Steinitz (1977), the Cenomanian rocks in North Sinai and the Naqb Desert represent a continuous transgressive event of a relatively long duration without significant regressions. Transgression proceeded from the northwestern part of Sinai that was already submerged during the Aptian or even earlier (Shata 1960; Abu Zeid 2007). The Turonian section in North Sinai includes the Abu Qada and Wata Formations. The Lower Turonian Abu Qada Formation consists of interbedded shales, marls, limestones and sandstones. It is 68 –141 m thick at G. Araif El Naqa and is represented by gypsiferous shales with thick limestone beds (Bartov et al. 1980). The Upper Turonian Wata Formation is a thick carbonate sequence, which forms the prominent dip slopes of the North Sinai anticlines. Its thickness is about 180 m at G. Araif El Naqa (Bartov & Steinitz 1977), 190 m at G. Yelleg (Visser 1941), 132 m at G. El Minsherah (Farag & Shata 1954) and 213 m at G. Halal (Osman et al. 2000). A sequence of massive Turonian limestones with rare shale interbeds was penetrated in the Abu Hamth (254 m), Darag (227 m) and Nakhl (260 m) wells. Deep marine conditions persisted throughout the Cenomanian and Turonian periods in the extreme northern part of Sinai as well as in the offshore area. The Coniacian –Santonian sediments of North Sinai are represented by marine sediments of the Themed Formation (Ziko et al. 1993). The upper part of the Coniacian rocks is truncated by erosion at early Late Coniacian in NE Sinai (Lewy 1975). The Campanian–Maastrichtian sediments are represented by white, soft chalks of the Sudr Formation. Undifferentiated Senonian sediments were penetrated by the Darag (235 m), Nakhl (33 m), Abu Hamth (326 m) and El Khabra (192 m) wells. The exposed Senonian rocks are 265 m thick at G. Yelleg (Visser 1941), 295 m thick at G. El Minsherah (Farag & Shata 1954) and 206 m thick at G. Halal (Osman et al. 2000).

Cenozoic rocks Table 2. Comparison of the thickness of the Jurassic rocks in the G. Maghara outcrop and Halal-1 well Sequence Upper Jurassic Middle Jurassic Lower Jurassic

G. Maghara

Halal-1 well

575 m 660 m 665þ m

214 m 780 m 2240 m

Although the Paleocene sediments (Esna Shale) in Sinai are located in the structural lows between the major anticlinal structures, this does not rule out the possibility that they were once present at the crests of the anticlinal structures and were eroded later. The Esna Shale has a uniform lithology of greenish –grey shale. In the Nakhl and Darag boreholes it is 59 m and 38 m thick, respectively.

44

A. R. MOUSTAFA

To the north, the Esna Shale is exposed on the flanks of G. Maghara (maximum 1 m) as well as in G. El Minsherah (65 m), G. Halal (55 m) and G. Araif El Naqa (50 m). The Eocene sediments crop out in many areas throughout Sinai and have also been penetrated by several wells. The Lower Eocene rocks in Sinai are generally represented by a massive flinty limestone (Thebes limestone) and occupy the broad synclinal lowlands between the large anticlines. To the east of G. Halal, the exposed Middle Eocene rocks are 51 m thick (Osman et al. 2000). No Upper Eocene outcrops have been recognized in North Sinai. However, Upper Eocene rocks were recognized in the shallow subsurface near G. Libni to the east of G. Maghara. Few scattered outcrops of Upper Oligocene shallow marine carbonate rocks (77 m thick) have been reported by Kuss & Boukhary (2008) in G. Rizan Aneiza to the east of G. Maghara. These carbonate rocks rest unconformably above Jurassic and Lower Cretaceous rocks. Basic volcanic igneous rocks of the Late Oligocene –Early Miocene age exist in some parts of northern Sinai, as reported in G. Yelleg and its vicinity. A 40 m thick basic igneous sill is exposed in G. Ikteifa (G. Yelleg area) where it is intruded at the contact between the Maastrichtian chalk and

the Esna Shale. According to Steinitz et al. (1978), this sill is of Early Miocene age (20.5 + 0.7 Ma, K –Ar dating). A NNE oriented discontinuous dyke also exists in the area and extends from G. Ikteifa to the western side of G. El Minsherah. Several small, one-meter thick dykes were also identified at the northeastern part of G. Yelleg by Khalil (1991). A NW–SE oriented basic igneous dyke crops out to the south of G. Maghara (El-Hemma area). Also, a very long east –west oriented basic igneous dyke (Raqabat Naam dyke) exists in central Sinai and has the same age like the G. Ikteifa volcanics (Eyal et al. 1980).

Structural setting The structural pattern of North Sinai is dominated mainly by NE–SW oriented doubly plunging anticlines (Fig. 4). Locally, ENE and east –west oriented faults form boundaries to these folds. Faults of other orientations also exist, but are associated with the anticlinal folds and are thought herein to be fold-related faults. The anticlinal folds of the area are of two obvious sizes, large (tens of kilometres long) and small (several kilometres long). G. Maghara, G. Halal and G. Yelleg folds are the largest three folds in North Sinai and are

Fig. 4. Simplified structural form-line map of North Sinai and the Naqb Desert after Khalil & Moustafa (1994) showing the main structural features of the study area.

NORTH SINAI FOLDS

60, 45 and 43 km long, respectively. Other large folds also exist to the east in the Naqb Desert, for example, Ramon, Kurnub and Qatan anticlines. The small anticlines exist in several parts of the study area, but are dominant in the southern part of the area and are mostly associated with the ENE –WSW and east– west oriented faults. The ENE –WSW and east –west oriented faults exist in two main belts that extend across the study area (Fig. 4). One of these two belts is the Sinai hinge belt, which forms the southern boundary of the area affected by large anticlines. This belt also corresponds to the central Sinai –Negev shear zone of Bartov (1974). The other fault belt generally has east –west orientation and extends from the

(a)

N

45

Gulf of Suez to the northern end of the Gulf of Aqaba and was previously called the themed fault (Moustafa & Khalil 1994). Figure 5 shows the orientations of the different structures in northern Sinai. The folds of the area (both large and small) have a predominant NE– SW orientation with a mean orientation of N458E– S458W (Fig. 5a). The faults of the area (both in the Sinai hinge belt and in the folded area) show a preferred NW to WNW orientation (Fig. 5b). The faults dissecting the large folds have a N558W preferred orientation (Fig. 5c), whereas those in the hinge belt (Fig. 5d ) also include other trends like NNW and ENE. Lengthwise, the ENE –WSW oriented faults are the longest

(b)

N

All faults N = 1197

All folds N = 272

46

165

N

(c)

N

(d)

Sinai hinge belt faults N = 584

Faults dissecting the large folds N = 613

88

95

Fig. 5. Rose diagrams showing the strikes of the different field-measured structures in North Sinai. Number along each circle shows the length of the longest petal in the rose diagram.

46

A. R. MOUSTAFA

faults in the Sinai hinge belt and are the main faults that controlled the deformation of this part of North Sinai.

North Sinai folds G. Maghara folds. The structures of G. Maghara affect an area that is about 60 km long and 13– 17 km wide. They include four main asymmetric NE–SW oriented anticlines (Um Mafruth, Um Asagil, Maghara and Hamayir) separated by two main SE vergent reverse faults (Hamayir and Mizeraa faults), as shown in Figure 6. The Hamayir and Maghara asymmetric anticlines lie on the hanging walls of these two reverse faults (see cross sections in Fig. 6). The Um Asagil asymmetric anticline may lie in the hanging wall of a third reverse fault located south of the G. Maghara complex. To the NE and SW of the Maghara anticline are two narrow ENE –WSW oriented ridges (Um Mafruth and Um Latiya ridges) that extend for 13 and 23 km respectively (Fig. 6). The Hamayir reverse fault dips at an angle ranging from 44 –738NW and has throw equal to 500 m

causing the Middle Jurassic rocks to ride over the Upper Jurassic rocks. The orientation of the eastern portion of this reverse fault changes to north– south with a corresponding change in orientation of its hanging wall anticline. The Mizeraa reverse fault is at least 45 km long and is made up of several segments joined together by strike –slip (tear) faults (Fig. 6). The overall trend of this fault is NE–SW but has some segments oriented NNE and WNW. Sinistral offset characterizes the NNE faults, whereas dextral offset characterizes the WNW faults. The maximum distance of horizontal offset is equal to 2– 2.5 km. The dip angle of the Mizeraa reverse fault reaches a high value of 718. Two large anticlines lie on the hanging wall of the fault. These are the G. Maghara main anticline (where all the section of Jurassic rock is exposed) and the Um Mafruth anticline. The maximum throw of the Mizeraa fault is 1250 m at its middle part. The northwestern flank of the G. Maghara anticline dips at about 158NW whereas its southeastern (forelimb) is narrower and usually vertical to overturned. G. Manzour represents the footwall of the Mizeraa

Fig. 6. Landsat TM image (see Fig. 4 for location) and structural cross sections (from Moustafa 2005) showing the main structures of the G. Maghara area.

NORTH SINAI FOLDS

47

Fig. 7. Field photograph (looking south) showing the asymmetric syncline of G. Manzour that represents the footwall of the Miseraa reverse fault. See Figure 6 for location. Apt, Alb, Kc and Kt designate Aptian, Albian, Cenomanian and Turonian rocks, respectively.

fault and is folded by an asymmetric syncline with a very steep northwestern limb (Fig. 7). The anticlines of the G. Maghara area are pervasively dissected by transverse (NW –SE oriented) normal faults that have relatively small amounts of throw (several tens of metres). These faults are synchronous with the folds and were probably formed due to lengthening of the rocks in a direction normal to the NW–SE shortening, resulting from the folding and reverse faulting. These faults do not dissect the Paleocene– Eocene rocks exposed in the outer edges of the folds. This gives strong evidence that they are of the same age as the folds and can not be attributed to younger extensional deformation (like the Gulf of Suez rifting for instance). Similar observations are also obvious in the G. Yelleg and G. Halal folds. G. Yelleg folds. Like G. Maghara, the G. Yelleg area also has a large NE–SW oriented anticline. This anticline is 45 km long and 18 km wide and affects the exposed Lower Cretaceous to Lower Eocene rocks of the area. It is a large asymmetric fold with a gentle northwestern flank dipping at 5–138NW and a steeper southeastern flank dipping as much as 568SE. The southeastern flank has a monoclinal shape and winding outline (Fig. 8) with segments oriented ENE and NNE within the predominant NE trend. Like G. Maghara, the G. Yelleg anticline is dissected by a large number of NW to WNW oriented (transverse) faults (Fig. 8, inset) that have relatively small amounts of throw in the order of a few tens of metres.

Cretaceous and Cenozoic rocks in the northern part of G. Yelleg area are affected by a number of smaller NE–SW oriented symmetric folds, the most prominent of which are G. Falig, G. Meneidret Abu Quroun and G. Meneidret El Etheili folds (Fig. 8); Moustafa & Gibali (2005) and Moustafa & Khalil (1995). The southernmost part of the G. Yelleg area is bounded by ENE– WSW oriented faults at G. Rishat Saada, G. Rishat Lehman and G. El Minsherah (Fig. 8). These faults belong to the Sinai hinge belt. The G. Yelleg monocline gradually fades away into the ENE –WSW oriented fault of G. Rishat Saada by gradual flattening of its steep southeastern flank. Mesostructures measured in the Upper Cretaceous rocks at the northeastern side of the G. Yelleg area (G. Meneidret El Etheili area) include tectonic stylolites, calcite veins and minor normal faults. These features indicate that s1 was oriented NW–SE and s3 was oriented NE–SW (Fig. 9a). Minor faults affecting Upper Cretaceous rocks of G. Falig also indicate the same orientation of s1 (Fig. 9c). Analysis of the fault– slip data of macrofaults at G. Meneidret El Etheili and G. Falig folds (Fig. 9b, d, respectively) indicate maximum horizontal compressive stress in the NW quadrant and minimum horizontal stress in the NE–SW quadrants. G. Halal folds. The structures of the G. Halal area were mapped in detail by Abd-Allah et al. (2004). The exposed Cretaceous rocks of this area are folded by a very large asymmetric, NE–SW

48

A. R. MOUSTAFA

Fig. 8. Landsat TM image showing the structures of the G. Yelleg area (see Fig. 4 for location). Rose diagram shows the strikes of field-mapped faults of this area.

oriented, doubly plunging anticline that is 43 km long and 14 km wide. This asymmetric anticline has a gentle northwestern flank dipping at about 158NW and a steep southeastern flank that is mostly vertical to overturned (Fig. 10a, c). The steep flank of the G. Halal anticline is expected to be underlain by a NE–SW oriented reverse fault similar to the Mizeraa reverse fault of G. Maghara. Right-stepping en echelon folds in the form of three anticline –syncline pairs affect the southern flank of the G. Halal anticline and give it a characteristic winding outline (Fig. 10a). Long distances separate these three fold pairs and these folds are similar to those affecting the forelimb of the Maghara and Um Mafruth folds. The en echelon folds in these three localities include hanging wall anticlines above en echelon reverse fault segments. The intervening synclines represent transfer zones between these en echelon fault segments (Moustafa 2005). A large number of transverse (NW – SE) oriented normal faults dissect the Cretaceous rocks of G. Halal (Fig. 10b). These faults have steep dip and relatively small throws. Analysis of slip data of these faults indicates lengthening in the NE– SW direction parallel to the fold axis (Fig. 11). Lower Cretaceous sandstones are exposed in the breached core of the G. Halal anticline. The Halal-1

well was drilled in the core of the anticline in 1975 with a total depth of 4311 m subsea. The well penetrated the thickest Jurassic basin, which was inverted later in Late Cretaceous – Early Cenozoic time.

Sinai hinge belt The Sinai hinge belt is a narrow ENE– WSW oriented structural belt forming the boundary between the North Sinai folds and the platform area lying to the south. This belt is 15–20 km wide and is dominated by long segments of ENE – WSW oriented faults that show evidence of rightlateral strike–slip displacement. Detailed field mapping of different parts of the Sinai hinge belt (Moustafa & Yousif 1990; Moustafa & Gibali 2005; Moustafa & Salama 2005) indicate that the ENE –WSW oriented faults in the hinge belt form two sub-belts, northern and southern. In each subbelt, the ENE –WSW oriented faults have a rightstepping en echelon arrangement (Fig. 12). The northern sub-belt extends from the southernmost side of G. Yelleg toward G. El Riash, whereas the southern sub-belt extends from the Mitla Pass to G. Araif El Naqa (Fig. 12). Fault–slip data indicate that these ENE –WSW oriented fault segments show dextral strike–slip displacement (Fig. 13).

N

(a)

(b)

σ3

σ1

N

N 44/356 P

44/356 P

E

W

T 14/252

T 14/252 Tectonic Stylolites

σ1

(c)

S

Calcite Veins Minor Normal Faults

N

N

Meneidret El Etheili Faults (N = 11)

(d)

N

N T 10/017

77/003 T

P 09/140 G. Falig Minor Faults (N = 5)

77/004 T

P 36/280

T 10/017

NORTH SINAI FOLDS

σ3

P 36/280

P 09/140

G. Falig Faults (N = 10)

Fig. 9. (a) Lower hemisphere equal area projection of the poles to mesostructures in G. Meneidret El Etheili (NE part of G. Yelleg) after Moustafa et al. (1991). (b), (c) and (d) show lower hemisphere equal-area projections of field-measured fault planes (great circles) and slickenside striae (small filled circles with arrows) in each left-hand stereogram and analysis of fault– slip data by FaultKin software of Allmendinger (2001) in each right-hand stereogram for Meneidret El Etheili large faults (b), G. Falig minor faults (c) and G. Falig large faults (d). P and T show the orientations of the maximum and minimum principal stress axes. 49

50

A. R. MOUSTAFA

Fig. 10. (a) Landsat TM image of the G. Halal anticline (see Fig. 4 for location). Note the winding trace of the steep southern flank (dotted line). (b) Rose diagram of the strikes of field-mapped faults. (c) Structural cross section across G. Halal fold modified after Abd-Allah et al. (2004).

Some of the ENE –WSW oriented faults in the Sinai hinge belt are characterized by the presence of ENE –WSW oriented anticlines, axially dissected in their middle parts by these faults, for example, G. Rishat Saada (Fig. 14), G. El Minsherah, G. El Riash and G. Kherim folds. Each of the faults dissecting these folds is about 1.5 times the length of the respective fold and it seems that these folds were formed in the middle part of each fault at the early stages of strike– slip movement of the fault. Erosion of the crestal areas of these folds shows more deformation of the older rocks in the axial areas of the folds by tighter folding and steeper

dip of the beds, for example, G. El Minsherah fold (Fig. 15). Some of these folds are dextrally offset by the faults. Thus, in G. Rishat Saada, the southern half of the anticline is dextrally offset for about 2 km relative to the northern half (Fig. 14). Some of the faults enclose pull-apart grabens between their overlapping ends such as the area between the El Minsherah and Rishat Lehman faults (Moustafa & Yousif 1990). Horizontal and gently plunging slickenlines characterize the ENE –WSW oriented faults of the Sinai hinge belt indicating dextral slip (Fig. 13). Bartov et al. (1980) also reported rightlateral strike –slip displacement along the Araif El

NORTH SINAI FOLDS

N

51

N

T

P

N = 30 Fig. 11. Analysis of fault– slip data of the field-measured faults of G. Halal. Symbols are like those in Figure 9. N indicates the number of faults.

Naqa fault that lies in the eastern part of the southern sub-belt. The westernmost side of the southern sub-belt of the Sinai hinge belt is characterized by the existence of right-stepping en echelon doubly

plunging anticlines, rather than an ENE fault segment (Fig. 16 and Moustafa & Khalil 1989). The arrangement of these folds provides evidence for dextral simple shear deformation along the

Fig. 12. A Landsat TM image showing ENE–WSW oriented en echelon faults of the Sinai hinge belt (see Fig. 4 for location). These faults are located in two sub-belts (A & B). Main faults in the northern sub-belt (A) include G. Rishat Saada fault (1), G. Rishat Lehman fault (2), G. El Minsherah fault (3), South Talet El Badan fault (4) and G. El Burqa– G. El Riash fault (5). Main faults in the southern sub-belt (B) include G. El Bruk fault (6), G. Kherim fault (7) and G. Araif El Naqa fault (8). Rose diagram represents strikes of all field-mapped east–west to ENE–WSW oriented faults in the hinge belt.

52

A. R. MOUSTAFA

Fig. 13. Analysis of fault– slip data of the Sinai hinge belt. Symbols are as those in Figure 9.

Fig. 14. Landsat TM image (a) and vertical aerial photograph (b) of G. Rishat Saada fault and the associated fold. See Figure 8 for location.

NORTH SINAI FOLDS

53

Fig. 15. Landsat TM image of G. El Minsherah area showing the Minsherah fault that axially dissects the middle part of the Minsherah anticline (see Fig. 8 for location). Note the small tight folds in the core of the anticline, north of the fault. Structural cross section across the Minsherah fold is after Moustafa & Yousif (1990). A and T (away from and toward the observer, respectively) designate the sense of strike– slip movement on the Minsherah fault.

Sinai hinge belt. Noweir et al. (2006) also confirmed the same style of deformation for this area. The Sinai hinge belt is obvious on the Bouguer anomaly and total magnetic intensity maps of Sinai Peninsula (Egyptian General Petroleum Company 1985) where it shows up like a boundary with a change in geophysical character across it.

Themed fault The tectonically stable area lying south of the Sinai hinge belt is cut by a very long east –west oriented fault zone, the Themed fault. This fault zone

extends for about 200 km from the vicinity of the eastern margin of the Suez rift to the Dead Sea transform (Fig. 1). The themed fault disrupts the monotonously flat-lying rocks of central Sinai by the development of a narrow elongated belt of small doubly plunging tight folds. This fault was rejuvenated along a pre-existing fault marking the southernmost edge of the Early Mesozoic continental margin of the Eastern Mediterranean basin in central Sinai (Moustafa & Khalil 1994). Within its very long tract, the Themed fault truncates rocks ranging in age from Precambrian to Middle Eocene. The eastern and western parts

54

A. R. MOUSTAFA

Fig. 16. Landsat TM image showing the en echelon doubly plunging anticlines affecting the Upper Cretaceous rocks in the Mitla Pass. See Figure 4 for location.

of the fault are very clear, whereas the central part (traversing the Tih Plateau) is less obvious. A basalt dyke of Early Miocene age is intruded along a segment of the Themed fault in the vicinity of G. El Dirsa (see Fig. 4 for location). The Themed fault was mapped in detail by Moustafa & Khalil (1994). The westernmost part of the fault consists of four right-stepping en echelon segments oriented east –west and ENE – WSW. All these segments have steep dips (about 50–908) with slickenside lineations recording dominantly right-lateral strike– slip motion (Fig. 17a). East–west to NE–SW oriented, doubly plunging folds are associated with these fault segments and affect the exposed Cretaceous and Cenozoic rocks. Since a segment of the Themed fault is intruded by an Early Miocene basaltic dyke and the dyke itself does not show any fault-related deformation, Moustafa & Khalil (1994) considered the last movement on the Themed fault to be pre-Miocene in age. The easternmost one third of the Themed fault has east– west orientation where the fault has a steep northward dip that ranges between 628 and 908. Local southward dip of the fault occurs less

often. Nearly horizontal to gently plunging slickenside lineations and associated chatter marks on this segment of the fault indicate right-lateral strike–slip or oblique– slip with a predominant right-lateral strike –slip component (Fig. 17b). In addition, the Rishat El Themed doubly plunging anticline is dextrally offset by a fault for about 300 m (Fig. 17c, d). The en echelon arrangement of the segments of the Themed fault, their associated folds, and the horizontal to gently plunging slickenlines, indicate that the Themed fault is a dextral wrench zone. The Themed fault affects rocks as young as Middle Eocene indicating a post-Middle Eocene slip. The last movement on the fault must be preEarly Miocene, since an Early Miocene basalt dyke was intruded in the weak zone of the fault.

Tectonic evolution The North Sinai folds clearly show the effect of horizontal compressive stress on the deformation of the exposed Mesozoic and Cenozoic rocks. All

NORTH SINAI FOLDS

55

Fig. 17. Lower-hemisphere equal-area projections of the strikes of field-measured faults (great circles) of the western part (a) and eastern part (b) of the Themed fault and their slickenside lineations (small circles) after Moustafa & Khalil (1994). U and D refer to upthrown and downthrown sides respectively. (c) Geological map and (d) vertical aerial photograph of G. Rishat El Themed fold showing its dextral offset by the Themed fault (after Moustafa & Khalil 1994). See Figure 4 for location.

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structures in the three folded areas (Maghara, Yelleg and Halal) indicate shortening in the NW– SE direction and related lengthening in the NE– SW direction. Subsurface data of the study area show the effect of an earlier deformation that proceeded by crustal extension in Mesozoic time, indicating that the exposed folds were formed by basin inversion. Seismic reflection profiles and borehole data show the existence of relatively small extensional basins in the area. A 2D seismic section extending from the G. Um Mafruth anticline (G. Maghara area) to the northeastern nose of the G. Yelleg fold (Fig. 18) clearly shows the structures of this portion of North Sinai. Asymmetric folds associated with reverse faults are clear in the G. Maghara area (northwestern part of the seismic section) followed to the south by a 15-km wide synclinal box fold separating these asymmetric folds from the G. Yelleg anticline. The latter

(southeastern part of the seismic section) is a large box-shaped anticline with two steep flanks on its northwestern and southeastern sides. Both flanks are bounded by faults but the fault bounding the southeastern flank is a blind reverse fault that dissects only the lower part of the Upper Cretaceous section and older rocks. Flattening of the seismic section on the top Masajid Formation (Fig. 19) indicates that all faults had normal slip at Jurassic time and that the Jurassic section lying between the top Rajabiah Formation (intra-Lower Jurassic) and top Masajid Formation (Upper Jurassic) was deposited in two separate extensional sub-basins in G. Maghara and G. Yelleg. These two sub-basins are separated by an inter-basinal high area located at the present-day syncline separating the G. Maghara and G. Yelleg folds. The thickness of the Jurassic sediments in the G. Yelleg sub-basin is smaller than that in the

Fig. 18. Two-dimensional seismic section extending from the eastern part of G. Maghara anticline to the eastern nose of G. Yelleg anticline (see Fig. 1 for location). Symbols designate the following reflectors: top of the Rajabiah Formation (intra-Lower Jurassic) (Jl), top of the Upper Jurassic (Ju), top of the Lower Cretaceous (Kl), tops of the Intra-Upper Cretaceous units (Ku1 and Ku2), top of Santonian (Ku3) and top of Maastrichtian chalk (Ku4).

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Fig. 19. Same seismic section of Figure 18 flattened on the top Jurassic reflector.

G. Maghara sub-basin. The G. Maghara sub-basin includes a southward thickening Jurassic wedge. The thickest part of the Jurassic section in the Yelleg sub-basin is located in its middle part. Small changes in thickness of the Lower and Upper Cretaceous rocks are obvious across some of the faults. Such changes in thickness could either be real or unreal and related to the seismic flattening process. The seismic resolution below the top Rajabiah reflector is not good enough to show any changes in the thickness of the Lower Jurassic rocks. The two Jurassic sub-basins of G. Maghara and G. Yelleg indicate extensional deformation by rifting in North Sinai during the Jurassic. Similar change in thickness of the Jurassic section in G. Maghara and G. Halal areas was also pointed out in a previous section (Table 2). The exposed Jurassic rocks in G. Maghara (1980þ m; Al-Far 1966) represent almost the whole Jurassic system. On the other hand, Halal-1 well that was drilled in the core of G. Halal anticline penetrated a complete Jurassic section, which is 3234 m thick (much thicker than that of G. Maghara). The location of Halal-1 well, relative to the crest of the G. Halal anticline and to its steeply dipping flank (Fig. 10c), indicates that the penetrated Jurassic thickness is close to the true thickness rather than being the vertical thickness of steeply dipping rocks. Because G. Halal lies to the south of G. Maghara, it was expected to

have a thinner Jurassic section assuming northward increase in thickness of the marine Phanerozoic rocks of Egypt. The thick Jurassic section penetrated by the Halal-1 well most probably indicates a separate extensional sub-basin that was subsiding during Jurassic time. Based on these changes in thickness of the Jurassic rocks in the Maghara, Yelleg and Halal areas, as well as the local changes in thickness of the Lower and Upper Cretaceous rocks across the Mizeraa fault, it is proposed herein that three separate extensional sub-basins existed in these areas during the Jurassic and probably during part of the Cretaceous Period. These sub-basins were inverted at a later time as the early normal faults were reactivated by reverse slip and the large anticlines of G. Maghara, Yelleg and Halal are fault-propagation folds associated with the upward propagation of these faults during basin inversion. Detailed field mapping of the G. Maghara area indicates that the NE –SW oriented Mizeraa and Hamayir reverse faults have several segments oriented north–south and WNW– ESE (Moustafa 2005). Slickenside lineations on the north–south and WNW–ESE oriented fault segments show oblique slip where the north–south fault segments show left-lateral and the WNW –ESE fault segments show right-lateral strike– slip components. The horizontal separation on these fault segments is too large to represent the amount of horizontal

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slip on the faults and it is proposed herein that these inverted faults (e.g. the Mizeraa fault) had zigzag shape during the early extensional phase in the Jurassic and Early Cretaceous times. Such fault patterns are common in extensional basins where different rift-parallel segments are linked together by transfer faults (e.g. Colletta et al. 1988; Morley et al. 1990; Ziegler 1992). During the phase of basin inversion, the fault segments oriented at high angle or normal to the direction of compression were reactivated by reverse slip (e.g. the NE–SW fault segments in North Sinai) whereas the transfer faults oriented at an angle to the direction of compression were reactivated by oblique slip. The nature of deformation at the Sinai hinge belt is different from that at the inversion structures of G. Maghara, Yelleg and Halal. Reliable field data support dextral simple shear deformation in the hinge belt along old ENE –WSW oriented faults that separate the inverted basins area of North Sinai from the platform area to the south. ENE – WSW oriented faults in north Egypt are known to be old, probably of Precambrian age. They dissect the exposed Precambrian basement rocks of the northern Eastern Desert, but not the nearby Phanerozoic sediments. Orwig (1982) also mapped faults of this trend in the Precambrian basement in the

subsurface in north Egypt. The Sinai hinge belt faults define the southern boundary of the area affected by Mesozoic extensional deformation. The SE-oriented compression, which led to basin inversion in North Sinai, reactivated these faults by dextral transpression (Fig. 20). The Themed fault is another fault lying further south in the platform area and was also reactivated by dextral slip. Being more oblique to compression, the Themed fault more clearly expresses the effect of strike – slip reactivation. Fault-slip data in Figures 11 and 13 perhaps indicate that faults did not slip under a single stress state. This may be attributed to the fact that all measured faults of each area are included together in one stereogram.

Time of compressional deformation Detailed field investigation indicates that compressional deformation in North Sinai took place in Late Cretaceous –Early Cenozoic times for the following reasons. (1) The en echelon folds of Mitla Pass affect the exposed Cenomanian, Turonian and Lower Senonian rocks. On the other hand, the Campanian –Maastrichtian, Paleocene, and

Fig. 20. Landsat TM image of North Sinai showing the inverted structures of Mahgara, Yelleg, Halal and the Hamayir–Amrar area. The Sinai hinge belt separates these inverted basins from the central Sinai platform area. Heavy arrow represents the direction of maximum compressive stress axis at Late Cretaceous– Cenozoic time.

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(2)

(3)

(4)

Eocene rocks of the Sadr El Heitan Plateau, which lies exactly to the south of these folds, do not seem to be affected by folding. Reworked Lower Cretaceous sandstones deposited within the Campanian –Maastrichtian chalk section of G. Araif El Naqa were reported and explained by Luning et al. (1998) to have been eroded from the crest of the Araif El Naqa fold. This indicates that folding already started before these rocks were deposited. A similar (intra-Late-Senonian) time of folding was reported in the Abu Roash fold (west of Cairo) by Moustafa (1988). An intra-Paleocene syn-tectonic debris flow overlying the Upper Senonian chalk exists on the southern side of the G. Maghara fold (Moustafa 2005). This debris flow contains clasts derived from the Jurassic rocks indicating that intense folding and related uplift took place in pre-Paleocene time. Smaller amounts of debris flow also exist in the Lower Eocene rocks of the same area. The Paleocene and Lower Eocene rocks on the southern flank of the G. Maghara anticline show a gradual upward decrease in dip indicating progressive folding at these times. Folded Lower Eocene rocks on the NE side of G. Falig anticline (NW side of G. Yelleg area),

(5)

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Moustafa & Gibali (2005), clearly indicate continuation of folding in post-Early Eocene time. Folded Middle Eocene rocks in the G. El Dirsa syncline (formed by the dextral slip on the Themed fault) represent the youngest folded rocks within the North Sinai outcrops (Moustafa & Khalil 1994). The lack of younger outcrops in North Sinai prohibits defining the end of the compressive deformation, but the presence of an Early Miocene basalt intrusion along the fault bordering the southern side of the G. El Dirsa syncline led Moustafa & Khalil (1994) to propose that the compressive deformation in North Sinai ended before Early Miocene time.

These observations indicate that the compressive deformation that led to basin inversion in North Sinai started in Late Cretaceous time and ended before Early Miocene time. The acme of this deformation was during the Late Cretaceous and is thought to have led to breaching and erosion along the crests of anticlines and deposition of syntectonic clastics within the Upper Cretaceous chalk of G. Araif El Naqa and in the Paleocene and Lower Eocene rocks on the southern side of the G. Maghara fold. Figure 21 shows an

Fig. 21. Detailed portion of the seismic section in Figure 18 showing thinning of the Upper Maastrichtian chalk (rock unit between the Ku3 and Ku4 reflectors) on the flanks of the G. Maghara and G. Yelleg anticlines.

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intra-Senonian deformation in the broad syncline separating the G. Maghara and G. Yelleg folds, where the Upper Senonian chalk is thicker in the trough of this syncline, compared with its flanks where the chalk was deposited unconformably above the tilted Santonian rocks. Such change in thickness indicates post-Santonian folding in the area. An intra-Upper Senonian angular unconformity was reported on seismic sections of the northern Western Desert of Egypt by Moustafa (2002, 2008). This angular unconformity constrains the acme of Late Cretaceous compressive deformation in north Egypt to be of Campanian age. The Late Cretaceous to pre-Miocene compressional deformation in North Sinai was followed by another deformation leading to dextral offset of the Lower Miocene basalt dyke of G. El Minsherah (120 m of dextral slip reported by Moustafa & Yousif 1990). This post-Early Miocene reactivation of the faults of the Sinai hinge belt is thought to be associated with the slip on the Dead Sea transform (Moustafa & Khalil 1994).

Synthesis The structures of North Sinai are related to the movement of the African plate relative to the nearby plates. Divergent and convergent movements between the African and Eurasian plates account for most of the tectonic deformation of this area. Extensional deformation during the Jurassic and probably part of the Cretaceous is related to the divergent movement between the Afro-Arabian and Eurasian plates that led to opening of the Neotethys in the Eastern Mediterranean area associated with northward drift of some fragments away from northern Africa (Robertson & Dixon 1984). According to Biju-Duval et al. (1979), Argyriadis et al. (1980), Garfunkel & Derin (1984), Dercourt et al. (1986), Mart (1987) and Stampfli et al. (2001), opening started in the Late Triassic –Early Jurassic time. This divergence led to thinning of the continental crust in North Sinai where crustal thickness decreases from 32 km at latitude 288N to about 27 km at the Mediterranean coast (Tealeb 1985; El-Azoni 1992). This divergent movement led to the development of several extensional basins in North Sinai (e.g. Maghara, Yelleg and Halal basins) in the present onshore area. Other extensional basins also exist in the offshore Mediterranean area (e.g. Mango basin, among others) but are not dealt with in the present study. The southern boundary of the extensional basins area is the Sinai hinge belt, which is defined by ENE –WSW oriented faults. Other faults were probably formed at the same time in the platform area lying south of the Sinai hinge belt (e.g. the Themed fault).

Convergence between the African and Eurasian plates in Late Cretaceous –Cenozoic times led to the compressional deformation in North Sinai that caused basin inversion. The direction of the maximum principal stress axis at that time was oriented NW –SE based on the study of mesostructures and/or the analysis of fault–slip data in G. Yelleg (Fig. 9), G. Maghara (Moustafa 2005) and the Sinai hinge belt (Fig. 13). Youssef (1968), Smith (1971), Eyal & Reches (1983) and Letouzey (1986) also reported similar direction of compression. Such convergence started in Late Cretaceous time, reached its acme in Campanian time and continued mildly until the pre-Miocene. Guiraud & Bosworth (1997) and Guiraud (1998) believe that inversion took place during the Santonian and latest Maastrichtian period in north Africa. The acme of the compressive deformation is almost contemporaneous with the collision of Afro-Arabia with Eurasia at the Bitlis suture (Hempton 1985) and the obduction of the Baer – Bassit ophiolites in NW Syria (Delaloye & Wagner 1984). Mild continued compressive deformation till the pre-Miocene time is related to further convergence between the two plates. Compressional deformation in northern Sinai seems to have been relieved during the OligoMiocene time when the Gulf of Suez-ancestral Red Sea rifting took place (Garfunkel & Bartov 1977). Continued deformation on the Sinai hinge belt in post-Early Miocene time is probably associated with the sinistral slip on the Dead Sea transform. The eastern parts of the faults of the Sinai hinge belt were probably dragged by the Dead Sea transform in response to the left-lateral slip in post-Miocene times leading to post-Miocene reactivation of the Sinai Hinge belt (Moustafa & Khalil 1994).

Conclusions Detailed field mapping of North Sinai structures combined with ample subsurface (seismic and borehole) data support the following statements. (1) North Sinai is a province of inverted basins that includes three main inverted structures in G. Maghara, G. Yelleg and G. Halal. The province extends northwards and includes other inverted basins in the Sinai offshore area and eastward to the Dead Sea transform, including other folds in the Naqb Desert. (2) The main asymmetrical folds of the Maghara, Yelleg and Halal areas are fault-propagation folds formed during basin inversion. (3) Neither wrenching nor thin-skinned deformation models previously proposed for the G. Maghara, Yelleg and Halal area could have

NORTH SINAI FOLDS

(4)

(5)

(6)

formed the existing structures. Although the deep geometry of faults in North Sinai is seen on a few seismic sections (e.g. Fig. 18), similarities between the North Sinai structures and those in the northern Western Desert of Egypt (where enough seismic and borehole data prove basement involvement) make it possible to postulate that the North Sinai faults affect the basement and, therefore, the deformation is thick-skinned. The inverted basins province is separated from the tectonically stable area to the south by the Sinai hinge belt. The Themed fault is one of the long faults dissecting the tectonically stable area. Dextral transpressive deformation characterizes the Sinai hinge belt, whereas pure dextral strike –slip deformation characterizes the Themed fault. Three main phases of deformation have affected North Sinai since the Mesozoic time: (a) The earliest (D1) deformation is an extensional deformation of Jurassic (and probably Cretaceous) age related to the divergence between the Afro-Arabian and Eurasian plates and opening of the Neotethys. This deformation formed extensional basins in the G. Maghara, Yelleg and Halal areas, as well as in the North Sinai offshore area and probably in the Naqb Desert. (b) The D2 deformation is compressional and related to the convergence between the African and Eurasian plates. It led to inversion of the extensional basins and dextral strike–slip movement on the Sinai hinge belt and the Themed fault. The acme of this deformation was in the Campanian associated with the collision of AfroArabia with Eurasia at the Bitlis suture and the obduction of the Bassit ophiolites in NW Syria. This deformation continued mildly until the pre-Miocene time. (c) The D3 deformation is post-Miocene and is probably associated with the sinistral strike–slip movement on the Dead Sea transform. This deformation reactivated the Sinai hinge belt and causes the present-day seismicity on faults in the eastern part of the hinge belt.

Fieldwork for this study was partly sponsored by Ain Shams University and the VW Foundation through Joint Research Project between Bremen University (Germany) and Ain Shams University (Egypt). Subsurface data were provided by the Egyptian General Petroleum Corporation. Parts of this paper are based on joint publications with some of my graduate students and colleagues. I specially

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mention the late M. E. Salama (Mansoura University), H. G. Fouda (now with Saudi Aramco), and W. Hashem (Ain Shams University) to whom I am indebted. S. Khalil (Suez Canal University) helped me in the analysis of fault-slip data. Fruitful discussions with J. Kuss and M. Bachmann (Bremen University) were valuable for understanding the stratigraphy of Cretaceous rocks of G. Maghara area. Critical review of the manuscript by M. I. Youssef (Ain Shams University) and constructive comments by Geological Society referees P. Bentham and P. Leturmy are highly appreciated.

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NORTH SINAI FOLDS of Cairo). Middle East Research Center, Ain Shams University, Earth Science Series, 2, 1 –16. Moustafa, A. R. 2002. Structural style and timing of syrian arc deformation in northern Egypt (abstract). American Association of Petroleum Geologists International meeting, Cairo, October 2002. Moustafa, A. R. 2005. Structural architecture of Late Cretaceous–Early Tertiary inverted structures in northern Sinai (Gebel Maghara area): abstract. 43rd Annual Meeting of the Egyptian Geological Society, Cairo, December 2005. Moustafa, A. R. 2008. Mesozoic–Cenozoic basin evolution in the northern Western Desert of Egypt. In: Salem, M., El-Arnauti, A. & Saleh, A. (eds) The Geology of East Libya, 3, 29–46. Moustafa, A. R. & Gibali, H. 2005. Structural Setting and tectonic evolution of Syrian Arc folds at Gebel Yelleg (North Sinai, Egypt), abstract. First International Conference on the Geology of the Tethys, Tethys Geological Society, Cairo University, November 2005. Moustafa, A. R. & Khalil, M. H. 1989. North Sinai structures and tectonic evolution. Middle East Research Center, Ain Shams University, Earth Science Series, 3, 215– 231. Moustafa, A. R. & Khalil, M. H. 1990. Structural characteristics and tectonic evolution of north Sinai fold belts. In: Said, R. (ed.) The Geology of Egypt. Balkema Publishers, Rotterdam, Netherlands, 381– 389. Moustafa, A. R. & Khalil, M. H. 1994. Rejuvenation of the Eastern Mediterranean passive continental margin in northern and central Sinai: new data from the Themed Fault. Geological Magazine, 131, 435– 448. Moustafa, A. R. & Khalil, S. M. 1995. Rejuvenation of the Tethyan passive continental margin of northern Sinai: deformation style and age (Gebel Yelleg area). Tectonophysics, 241, 225–238. Moustafa, A. R. & Salama, M. E. 2005. The Sinai Hinge Zone: a major crustal boundary in northern Sinai), abstract. 43rd Annual Meeting of the Egyptian Geological Society, Cairo, December 2005. Moustafa, A. R. & Yousif, M. S. 1990. Two-stage wrench deformation at Gebel El Minsherah, north Sinai. Middle East Research Center, Ain Shams University, Earth Science Series, 4, 112– 122. Moustafa, A. R., Khawasik, S. M. & Khalil, S. M. 1991. Late Cretaceous – Tertiary rotational deformation in North Sinai (Gebel Yelleg area). Neues Jahrbuch fur Geologie und Palaontologie, 11, 643–653. Moustafa, A. R., El-Badrawy, R. & Gibali, H. 1998. Pervasive E-ENE oriented faults in the northern Egypt and their relationship to Late Cretaceous petroliferous basins in the northern Western Desert. Proceedings of 14th Egyptian General Petroleum Corporation Exploration and Production Conference, Cairo, 1, 51–67. Noweir, M. A., Al Alfy, Z. & Fawwaz, E. M. 2006. Wrench deformation Syrian arc structures: El Hamraa, Umm Busal and Um Horeiba en echelon anticlines, Mitla Pass, west central Sinai, Egypt. Egyptian Journal of Geology, 50, 265– 287. Orwig, E. R. 1982. Tectonic framework of northern Egypt and the Eastern Mediterranean region. Proceedings 6th Egyptian General Petroleum Corporation Exploration Seminar, Cairo, 1, 1 –16.

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Osman, R. A., Ahmed, S. M. & Mahmoud, N. I. 2000. Cretaceous– Lower Tertiary rocks at Gebel Halal area, northern Sinai, Egypt: a Stratigraphy and sedimentary history. Proceedings of 5th International Conference on the Geology of the Arab World, Cairo, 1309– 1332. Robertson, A. H. F. & Dixon, J. E. 1984. Introduction: aspects of the geological evolution of the Eastern Mediterranean. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 1 –74. Rybakov, M. & Segev, A. 2004. Top of the crystalline basement in the Levant. Geochemistry, Geophysics, Geosystems, 5, 1 –8. Sadek, H. 1928. The principal structural features of the Peninsula of Sinai. 14th International Geological Congress, Madrid, 1926, 3, 895–900. Said, R. 1962. The Geology of Egypt. Elsevier Pub. Co., Amsterdam, 377. Schlumberger 1984. Well Evaluation Conference, Chapter 1: Geology of Egypt, 1– 64. Schlumberger 1995. Well Evaluation Conference, Chapter 4: Western Desert, 56–71. Sestini, G. 1989. Nile Delta: a review of depositional environments and geological history. In: Whateley, M. K. G. & Pickering, K. T. (eds) Deltas: Sites and Traps for Fossil Fuels. Geological Society, London, Special Publications, 41, 99–127. Shata, A. 1959. Structural development of the Sinai Peninsula (Egypt). Proceedings of the 20th International Geological Congress, 1956, 225–249. Shata, A. 1960. The geology and geomorphology of El Qusaima area. Bulletin Society Geography d’Egypte, 33, 95– 146. Shukri, N. M. 1954. Remarks on the geological structure of Egypt. Bulletin Society Geography d’Egypte, 27, 65–82. Smith, A. G. 1971. Alpine deformation and the oceanic areas of the Tethys, Mediterranean and Atlantic. Geological Society of America Bulletin, 82, 2039–2070. Stampfli, G., Borel, G., Cavazza, W., Mosar, J. & Ziegler, P. A. 2001. The paleotectonic atlas of the PeriTethyan domain. European Geophysical Society (CD). Steinitz, G., Bartov, Y. & Hunziker, J. C. 1978. K–Ar age determinations of some Miocene– Pliocene basalts in Israel: their significance to the tectonics of the rift valley. Geological Magazine, 115, 329–340. Tealeb, A. 1985. Spectral analysis of the gravimetric bouguer anomaly of Egypt for crustal thickness studies. Helwan Institute for Applied Geophysics Bulletin, 5B, 63–82. Visser, W. A. 1941. Geological Report on Gebel Yelleq (north central Sinai, Egypt). Internal report, Standard Oil Company, Cairo, 21. Youssef, M. I. 1968. Structural pattern of Egypt and its interpretation. American Association of Petroleum Geologists Bulletin, 52, 601– 614. Ziegler, P. A. 1992. North Sea rift system. Tectonophysics, 208, 55– 75. Ziko, A., Darwish, M. & Eweda, S. 1993. Late Cretaceous–Early Tertiary stratigraphy of the Themed area, east central Sinai, Egypt. Neues Jahrbuch fu¨r Geologie und Palaontologie, 3, 135– 149.

Structural setting and tectonic evolution of offshore North Sinai, Egypt M. YOUSEF1*, A. R. MOUSTAFA1 & M. SHANN2 1

Department of Geology, Ain Shams University, Cairo 11566, Egypt

2

BP International Exploration, Chertsey Road, Sunbury on Thames, Middlesex TW16 7LN, UK *Corresponding author (e-mail: [email protected]) Abstract: The offshore area of North Sinai represents the northern extension of the Syrian Arc inversion structures into the southeastern Mediterranean region. Integration of detailed seismic interpretation of key tectonic events in offshore North Sinai and recently acquired gravity and magnetic data reveal structural deformation represented by large buried inversion anticlines that have played an important role in the geological history and hydrocarbon potential of the area. This tectonic inversion took place in the Late Mesozoic and continued slightly during the Cenozoic, and formed NE-trending asymmetrical folds. Three different phases of deformation have been detected in offshore North Sinai: (1) A Jurassic–Early Cretaceous extensional phase, which formed NE trending normal faults bounding asymmetrical half-grabens, (2) Post-Santonian– Middle Miocene positive inversion of these faults and half-grabens and (3) Post–Middle Miocene subsidence. A set of tectonosequences related to the opening and the subsequent convergence of the Tethys was mapped. Each identified tectonosequence has its own unique drive mechanism, geometry, and location with respect to the plate boundary. Recognition of these elements allows illustration of the Tethyan basin evolution of offshore North Sinai through time as well as understanding the tectonic and stratigraphic framework and effective prediction of the petroleum system.

Although the geology of the North Sinai Peninsula has been the subject of study of several researchers, little was written about the offshore area. Being located on the Tethyan margin of the AfricanArabian plate, offshore North Sinai exhibits a sequence of shallow marine and continental sedimentary rocks, which may be up to 12 km thick (Guiraud & Bosworth 1997). Large synsedimentary normal faults, with nearly N108E and N708E average trends, were very active during Permian–Early Cretaceous times (Guiraud & Bosworth 1997). From the Early Senonian, both onshore and offshore domains of this region registered folding and faulting (Ginzburg et al. 1975; Bartov et al. 1980; Jenkins 1990). The compression caused reverse faulting along the old normal faults and the initiation of basin inversion. These events resulted in a well-exposed fold belt of NE– SW oriented, doubly plunging folds referred to as the Syrian Arc fold belt (Krenkel 1925), Figure 1. These folds are of different sizes, including folds tens of kilometres long, several kilometres long, and many others that are less than 2 km in length (Moustafa & Khalil 1989). The area selected for detailed subsurface study in the present study is geographically located in the southeastern part of the Mediterranean Sea, offshore North Sinai (Fig. 2). It covers an area

of about 6630 km2 and is bounded by latitudes 31830 and 318460 N and by longitudes 328410 and 348130 E. The aim of this paper is to address the tectonostratigraphic evolution of offshore North Sinai Basin and to document the phases of rifting and basin inversion, which affected the North Sinai margin of the Tethys from the Jurassic to Middle Miocene.

Regional setting The eastern Mediterranean passive continental margin was formed in the Early Mesozoic, when widespread rifting occurred in the entire Tethys area (Dewey et al. 1973; Garfunkel & Derin 1984; Moustafa & Khalil 1994). This margin was characterized by shallow platform carbonates in the east and by deep-water carbonates on the slope and basin in the west (Bein & Gvirtzman 1977; Druckman et al. 1995). Since the early Late Senonian time, the collision of the African-Arabian plate with the Eurasian plate resulted in the development of the Syrian Arc fold belt (Fig. 1; Neev & Ben-Avraham 1977; Moustafa & Khalil 1989; Eyal 1996; Buchbinder & Zilberman 1997; Garfunkel 1998). The main deformation in the

From: HOMBERG , C. & BACHMANN , M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 65–84. DOI: 10.1144/SP341.4 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Simplified structural form-line map showing the main structural features of North Sinai and al-Naqab Desert after Khalil & Moustafa (1994).

Fig. 2. Location map of the study area showing the used 2D seismic sections and boreholes.

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Syrian Arc fold belt continued into the post-Early Miocene (Moustafa & Yousif 1990; Moustafa et al. 1991; Tibor et al. 1992; Druckman et al. 1995; Eyal 1996). During the Early Miocene the Syrian Arc began to emerge on the Levant margin (Buchbinder & Zilberman 1997). The shelf area underwent localized tectonical uplift and became intermittently emergent (Buchbinder & Zilberman 1997). At the end of the Miocene, the eastern Mediterranean passive continental margin underwent extensive erosion and evaporite deposition (Gvirtzman & Buchbinder 1978; Druckman et al. 1995), a pattern common to most of the Mediterranean basin (Hsu¨ et al. 1978). Thick evaporites (.2 km) were deposited on the seafloor, while the continental margin was affected by erosion producing deeply incised valleys (Cita & Ryan 1978; Garfunkel & Almagor 1987; Abdel Aal et al. 2001). The Pliocene to recent sediments in large areas of the eastern Mediterranean continental margin and adjacent basin are affected by thin-skinned deformation (Bertoni & Cartwright 2006) owing to salt mobilization and shelf loading that caused the collapse and landward tilt of these deposits above the Messinian evaporites (Tibor et al. 1992).

Data and methodology The offshore subsurface data used for this study comprise about 6444 km (3995 miles) of twodimensional (2D) seismic reflection sections with maximum recorded two-way time of 5 to 6 s, 25 m shot point interval, 48-fold geophone coverage, and a line spacing of 400 m, in addition to bore-hole data (wire-line logs and biostratigraphic data) of 11 wells, and gravity and magnetic data. Figures 2 and 3 illustrate the database and the correlation between seismic lines and wells. Fault and horizon interpretations were carried out on a workstation using Landmark’s SeisWorks software. A seismic tectonostratigraphic approach was followed in the present study. Regionally persistent seismic horizons were mapped across the study area. After regionally consistent seismic tectonostratigraphic framework was constructed, biostratigraphic data were used to calibrate the seismic interpretation. Regional events were dated based on ties to apparent depositional hiatuses (condensed sections, erosional unconformities, periods of nondeposition, etc.) interpreted from biostratigraphic well control. Critical well ties were achieved using check shot surveys and synthetic seismograms. There was generally good correlation between the seismic tectonostratigraphy and the biostratigraphic interpretations. Integration of these interpretations

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resulted in a consistent and robust tectonostratigraphic framework.

Tectonosequences Seismic units bounded by regional unconformities are tectonosequences that represent deposition during tectonically controlled phases of basin development. The tectonosequence boundaries reflect major changes in regional tectonic settings (Hubbard 1988) and they are referred to as TB1 to TB5 from oldest to youngest (Figs 3 & 4). The tectonosequences are referred to as TS1 to TS5 (Table 1 & Fig. 5) from oldest to youngest. The geometry of the tectonosequences is illustrated by 2D seismic sections. These sections are located in Figure 2 and are either oriented along the strike of the basin or orthogonal to the major tectonic elements. These data allied to biostratigraphical information from wells, provide important evidence for the timing of deformation in offshore North Sinai. Five tectonosequences interpreted on seismic sections are summarized below (from oldest to youngest); Figure 6.

Tectonosequence 1 (Upper Jurassic– Santonian) Most of the rocks of tectonosequence 1 represents a syn-rift package that comprises the initial clastic, basin-fill and shows significant growth into opposed, fault bounded depocentres. They were deposited during the development of Mesozoic ENE oriented rift basins in North Sinai as well as in the other parts of north Egypt (Orwig 1982; Moustafa et al. 1998), owing to opening of the Neotethys between North Africa-Arabia and Eurasia (Biju-Duval et al. 1979; Argyriadis et al. 1980). The top of tectonosequence 1 is truncated by erosion in the Mango area (Fig. 5), but the boundary becomes a disconformity in the western part of the area of study. Tectonosequence 1 consists predominantly of clastic-dominated successions of the Cretaceous age unconformably overlying the shallow marine Upper Jurassic carbonates as a result of the sharp drop in global sea level in the Late Tithonian (Guiraud 1998). The basin geometry of the tectonosequence 1, illustrated by the isochron map (Fig. 7), shows several sub-basins bounded by extensional faults.

Tectonosequence 2 (Campanian – Maastrichtian) Deposition of tectonosequence 2 locally continued away from the crests of the evolving anticlines leading to onlap of tectonosequence 2 sediments

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Fig. 3. Correlation panel between the main lithostratigraphic units on well data and the interpreted seismic packages. TB1– TB5 refer to tectonosequence boundaries 1 –5.

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Fig. 4. Offshore North Sinai basin cartoon, showing key surfaces mapped. TB1–TB5 refer to tectonosequence boundaries 1– 5.

on older rocks in the flanks. It is thin (,200 ms TWT) over the crests of the anticlines and gradually thickens (.450 ms TWT) in the synclines (Figs 5 & 8). This syn-compressional tectonosequence consists predominantly of shelf and slope mudstones. The onset of Tethys convergence between the Eurasian and African plates in postSantonian time was characterized by positive inversion of the previously formed NE trending extensional basins in offshore North Sinai (Yousef et al. 2006). This inversion led to the development of a series of NE–SW trending folds (Fig. 9a).

Tectonosequence 3 (Palaeocene – Upper Oligocene) Tectonosequence 3 is the second syn-compression tectonosequence identified in offshore North Sinai and is also associated with the folding event, but with a magnitude less than that in tectonosequence 2. It is characterized by onlap and truncation relationships at its margins owing to the active deformation of the antiforms, giving an indication of

the continuation of the post-Santonian basin inversion until the Late Oligocene. Folding and faulting are not as extensive as those of TB2 and TB3 but only gentle folding is obvious (Fig. 9b). Tectonosequence 3 represents deposition in shelf and slope environments. It is also thin (,200 ms TWT) over the crests of the anticlines and gradually thickens (.450 ms TWT) in the synclines (Figs 5 & 8).

Tectonosequence 4 (Miocene) Tectonosequence 4 defines a northward-thickening sedimentary wedge, which thins locally over the crests of the anticlines, delineated by the TB4 and TB5 at its base and top, respectively (Fig. 5). It is almost missing along the present-day shoreline in the south, but thicken to .800 ms in the northern part of the study area (Fig. 5). Tectonosequence 4 comprises open marine deposits, as well as an evaporite sequence deposited during the Late Miocene (Messinian) drying up of the Mediterranean (Hsu¨ et al. 1973). This evaporite sequence consists predominantly of anhydrite and gypsum with

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Table 1. Summary of the principal features of the tectonostratigraphic units in offshore North Sinai Tectonosequence

Age

TWT thickness (ms)

TS5

Pliocene–Pleistocene

TS4

Miocene

400– 600

TS3

Palaeocene –Late Oligocene

700– 1000

TS2 TS1

Campanian –Maastrichtian Late Jurassic –Santonian

local halite deposits that irregularly cover marine sediments of the Middle to Late Miocene age. This is linked to the final convergence of the African and Eurasian plates at the end of the Miocene, where the Late Messinian records the closing of the Strait of Gibraltar and the subsequent evaporation of the Mediterranean Basin (Dolson et al. 2001).

Tectonosequence 5 (Pliocene – Pleistocene) Tectonosequence 5 is the youngest tectonosequence in offshore North Sinai and is represented by a northward-thickening wedge above the TB5. It increases in thickness from nearly 200 ms along the present-day shoreline in the south to nearly 2500 ms along the northern edge of the study area in the north (Figs 5 & 8). Tectonosequence 5 is characterized as a package of relatively undeformed, horizontal to gently inclined parallel reflectors showing moderate to high amplitude and remarkable lateral continuity (Fig. 5). It exhibits a wedgelike geometry and displays extensional faulting with detachment at or near its base. This Neogene extension was gravity driven with general extension directions to the NE. It comprises mainly Nile-derived sediments (marl, shale, and sandstone) prograding into the Mediterranean Sea.

Evidence for basin inversion in offshore North Sinai Tectonic structures The continental margin of offshore North Sinai is a unique example of an apparently simply structured continental margin that, when studied carefully,

0 – 2200

0 – 500 1500– 2000

Lithology Shale and siltstone with sandstone intercalations Shale, mudstone and siltstone with intercalations of sandstone and minor limestone interbeds Evaporites of restricted realm Dark grey to dark green, locally pyritic or glauconitic shale Shallow shelf limestone with few dolomite streaks Shallow shelf carbonates Dark coloured shale, occasionally highly calcareous, with interbeds of sand, sandstone and minor streaks of dolomitic limestone

displays complex structural patterns owing to resurgent and superimposed tectonic activity from the Early Mesozoic to Recent. Detailed study of seismic data interpreted with recently acquired gravity and magnetic data, revealed structural deformation represented by large buried inversion anticlines that played an important role in the geological history and the hydrocarbon potential of the area. The interpretation of different tectonics structures in the area is given in the following paragraphs. Mango structure. The Mango structure is a 24-km long, ENE-oriented (Fig. 10d), doubly plunging anticline with a relatively steep northwestern flank and breached crest. Three wells have been drilled to test the potential of the Lower Cretaceous sandstones and Upper Cretaceous to Eocene carbonates in this structure. The Lower Cretaceous succession is conformable, with the overlying and underlying sequences downdip in the southern flank and the severe erosion at the crestal part of the structure was not observed downdip in the southern flank. The stratigraphic succession, between the Top Jurassic and Top Santonian represented by tectonosequence 1, shows an increase in thickness towards the main bounding fault indicating normal slip on that fault during Jurassic– Santonian time. On the other hand, tectonosequences 2 and 3 are progressively onlapping the northward uprising tectonosequence 1. The east-northeasterly orientation of the Mango doubly plunging anticline, as with other inverted structures of offshore North Sinai, resulted from inversion-related folding over a deep-seated ENEoriented fault. Minor NW-oriented normal faults (Fig. 10b) on the northeastern side of the doubly plunging anticline indicate lengthening in the NE– SW direction, as elsewhere in offshore and

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Fig. 5. 2D seismic section through the Mango structure, offshore North Sinai (see location in Fig. 2). Fault-controlled subsidence in the offshore North Sinai basin resulted in growth of the Upper Jurassic– Santonian TS1 towards the Mango structure main bounding Fault. Later inversion formed the asymmetrical anticline affecting the TS1– TS4. TB1–TB5 refer to tectonosequence boundaries 1 –5.

onshore North Sinai (Moustafa & Khalil 1989; Abd-Allah et al. 2004). Goliath structure. The Goliath structure is a 20-km long, NNE-oriented (Fig. 10d), doubly plunging anticline with a relatively steep NW flank on the northeastern side of the area of study (Fig. 9a). Goliath-1 well was drilled on the crest of the structure and indicated a major stratigraphic break between the Coniacian– Maastrichtian carbonates and the underlying Lower Albian clastics of tectonosequence 1 (Fig. 11a). The Goliath structure

is bounded on the north by a NNE-oriented fault and, like the Mango structure, it represents an example of inverted Late Jurassic– Early Cretaceous half-graben. North Sinai 21-1 structure. The North Sinai 21-1 structure is a 20-km long, ENE-oriented (Fig. 10d), doubly plunging anticline in the central part of the area of study (Fig. 9a). It has a relatively steep northwestern flank. The North Sinai 21-1 well was drilled on the crest of the structure. Updip thickening of tectonosequence 1 (Fig. 8) was

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side of the area of study (Fig. 9a). Ziv-1 well was drilled on the crest of the structure and indicated an increase in thickness of tectonosequence 1 towards the main bounding fault (Fig. 8). The Ziv structure is bounded on the south by a NNE-oriented high-angle reverse fault representing an example of inverted Late Jurassic–Early Cretaceous half graben. This reverse fault has the same orientations of the normal fault affecting the Top Jurassic surface (TB1) (Fig. 10b). Similar onshore observations were reported by Shata (1959), Al-Far (1966), Bartov et al. (1980), Jenkins (1990) and Moustafa (2010).

Fig. 6. Tectonostratigraphy of offshore North Sinai basin.

observed in this structure in addition to an extensive erosional truncation giving indication of an early extensional phase during the Late Jurassic to the Early Cretaceous, followed by a compressional phase starting from the post-Santonian time. On the other hand, the overlying tectonosequences 2 and 3 show a rather progressive onlapping on the crest of the elevated structure. Ziv structure. The Ziv structure is an 8-km long, NNE-oriented (Fig. 10d), doubly plunging anticline with a relatively steep SE flank on the northeastern

Tineh structure. The Tineh structure is a 7-km long, NE-oriented (Fig. 10d), doubly plunging anticline. Three wells have been drilled to explore and evaluate the hydrocarbon potential of Oligocene clastic sequence in this structure. The Tineh-1 well, drilled on the crest of the structure, encountered significant oil indications in some sand reservoirs, interbedded with shales, within the Late Oligocene section. Numerous grabens and half grabens formed during the rifting stage and block tilting recognized in Israel and Sinai (Bartov et al. 1980), as well as in northern Egypt (Keely & Wallis 1991) characterize the Top Jurassic (TB1), Figures 5 and 8. Rose diagram of the mapped faults (Fig. 10a) indicates that they are oriented N30–60E. A few normal faults (Fig. 12) are affecting the Top Miocene (TB5), which acts as a detachment surface for such faults to the north of the study area. A rose diagram of the mapped faults at the Top Miocene surface (Fig. 10c) indicates that they are oriented N60 –70E. TB5 is a major erosional unconformity in the Late Miocene (Fig. 11a, b), which marks a change in the reflectivity pattern (Fig. 5) related to the transition from the shale of tectonosequence 5 (Pliocene–Pleistocene) to the Messinian evaporite complex of the uppermost part of tectonosequence 4 (Miocene). This change is interpreted to be associated with the presence of Messinian evaporite-bearing series (Vidal et al. 2000). This event is proposed herein to be a tectonically driven event related to late-stage convergence of the African plate against Europe causing the closure of the Strait of Gibraltar and evaporation of the Mediterranean Sea (Halbouty & El Baz 1992).

Structural evolution of offshore North Sinai The North Sinai fold belt (part of the so-called Syrian Arc system) shows examples of inversion structures represented by asymmetrical, doubly plunging anticlines that have thick syn-extension strata in the core. These inversion anticlines occur

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Fig. 7. Time-thickness (isochron) map of offshore North Sinai basin for tectonosequences 1. Contour interval is of 200 ms.

onshore (Fig. 1) and also extend offshore as indicated in the present study (Figs 5, 8 & 11b). In order to simulate the structural evolution of the inverted structures in offshore North Sinai, ‘seismic flattening’ was done by Landmark’s SeisWorks/2D software. This operation allows simulating the appearance of the strata below the flattened horizon at the time that horizon was

deposited. Seismic displays before and after the flatten option have been activated for a Goliathinverted structure in the northeastern part of the area of study, as will be summarized in the following sections. Goliath-inverted structure. The Goliath structure is a large northeasterly trending anticline (Fig. 9a)

Fig. 8. 2D seismic section through Ziv and North Sinai 21-1 structures, offshore North Sinai, see Figure 2 for location. TB1–TB5 refer to tectonosequence boundaries 1 –5.

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Fig. 9. Time-structure maps of offshore North Sinai basin for tectonosequence boundaries 2 and 4 (a and b). Contour interval is 200 ms.

and has been interpreted as an inversion-related asymmetrical fold. Because of the onlapping of the Campanian–Maastrichtian sediments onto the structure, the main episode of compression is believed to have occurred in post-Santonian time. By flattening the top Jurassic horizon (TB 1) (Fig. 13a), a northwestward thickening Upper Jurassic section is clear towards a southeasterly dipping fault, in the form of a half graben, indicating an extensional phase during that period. This NE– SW trending fault is bounding the Goliath structure on the NW (Fig. 9a). Another stage of extension took place during the Neocomian time. This is evident from the flattening of the Top Neocomian horizon (Fig. 13b). More

syn-depositional tilting and thickening of the Upper Jurassic and Neocomian sections are obvious towards the same main bounding fault. Seismic data show that the Early Mesozoic extensional phase, started from the Late Jurassic (and may be earlier) and continued until the Late Santonian time. This is also clear from flattening at the top Santonian horizon (Fig. 13c). Extension, followed by contraction (Basin Inversion), is evident from flattening of the top Cretaceous horizon (TB 3) (Fig. 13d). As a result of the switch in tectonic mode from extension to compression, the Goliath half graben becomes a positive structural feature, with a significant erosional truncation surface between the Albian

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Fig. 10. Rose diagrams showing the trends of subsurface structural elements in the offshore North Sinai basin.

clastics and the Campanian–Maastrichtian carbonates (Fig. 5). The normal fault bounding the NW side of the Goliath basin was reactivated as a high angle reverse fault with a switch in depocentres from growth into the fault to growth away from the fault, to the SE of the inversion anticline. The syn-inversion Campanian– Maastrichtian sediments are progressively onlapping the flanks of the elevated structure, attaining a minimum thickness over the crest of the elevated structure with gradual increase in thickness over the flanks (Fig. 13d ). Figure 13e shows continuation of the compressive

event until the Oligocene time, leading to a continuous uplift of the hanging wall against the footwall along the main fault. Flattening at the top Miocene (TB 5) is shown in Figure 13f. Owing to the reduced connection of the Mediterranean basin with the world oceans during that time, an unconformity that irregularly cuts down into the underlying successions resulted due to the onset of the Messinian Salinity Crisis. Flattening such surface would only reflect the magnitude of erosion and its inflected topography on the underlying successions (Ayyad 1997). An indication of

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Fig. 11. Seismic sections through the Goliath structure (a) and the south Tineh area (b), see Figure 2 for location. The Top Neocomian is not shown in the footwall of the fault of section (a) owing to bad seismic resolution. TB1– TB5 refer to tectonosequence boundaries 1 –5.

Early Miocene gentle folding is evident from southern Tineh area (Fig. 11b). Timing of deformation. Most of the mapped structures in offshore North Sinai are located within five NE-oriented inverted structures, whereas the

intervening areas are undeformed or only slightly deformed. The five inverted structures are controlled by major deep-seated reverse faults. Three phases of structural deformation affected the Mesozoic –Cenozoic succession in the study area and led to the development of the mapped

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Fig. 12. 2D seismic section showing growth extensional faults affecting tectonosequences 4 and 5, see Figure 2 for location.

structures. These are Late Jurassic to Early Cretaceous, post-Santonian to Middle Miocene, and Neogene. Dating of the post-Santonian deformation is based on the presence of a stratigraphic gap and progressive onlapping between the Campanian– Maastrichtian carbonates and older rocks in the Goliath-1 well. Additional data from the Mango structure constrain the age of this phase of deformation. At that locality, the progressive onlap within tectonosequence 2 onto growing anticlines marks the climax of the basin inversion in northern Egypt. Basin inversion continued during and after the deposition of the Paleocene to Langhian (Middle Miocene) rocks, but its effect can only be seen in close vicinity of the basin-bounding faults. The third phase of deformation is contemporaneous with Neogene regional subsidence of the shelf margin offshore North Sinai. Jurassic–Early Cretaceous rifting. The earliest phase of deformation is represented by rifting in the area of the eastern Mediterranean Basin. It is mostly considered to be of Late Triassic age (Robertson 1998; Guiraud & Bosworth 1999; Garfunkel 2004). In the Western Desert and North Sinai, owing possibly to the lack of information about the Triassic, rare data suggest that significant extension and associated basin subsidence began during Jurassic time (Moustafa et al. 1998). Tectonic deformation was registered along most of northern and Central Africa basins, around the

Jurassic –Cretaceous transition times (Guiraud et al. 2005). Frequent uplift and block-tilting, sometimes accompanied by slight folding, occurred along the northern Egyptian margin and the Levant (Guiraud 1998; Le Roy et al. 1998; Coward & Ries 2003). These deformations, underlined by hiatuses in the series and unconformities, represent the distant effects of stronger tectonic activity that occurred in southeastern Europe, referred to as the ‘Cimmerian’ or ‘Berriasian’ orogenic event (Nikishin et al. 2001; Stampfli et al. 2001). It led to the development of ENE-oriented basins bounded by major normal faults of the same orientation and they have half graben geometry with a northward tilt toward the boundary faults. Detailed seismic interpretation of key tectonic events in the present study has indicated that the Jurassic –Cretaceous rifting took place in several pulses: † Late Jurassic; † Neocomian; † Aptian –Santonian. Post-Santonian–Middle Miocene basin inversion. Basin inversion in the offshore North Sinai started in the post-Santonian time and proceeded by the reverse reactivation of the NE-trending extensional faults (Yousef et al. 2006). These are the same old faults that had normal slip during the early rifting phase. The unconformity detected at the contact between the tectonosequences 1 and 2

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Fig. 13. 2D seismic sections through the Goliath structure (refer to Fig. 11a), flattened on (a) top Jurassic (TB1), (b) top Neocomian and (c) top Santonian (TB2).

marks the start of basin inversion in the area of study. In most of the study area, it was concluded that basin inversion continued during and after the deposition of the Palaeocene to Oligocene rocks, but its effect can only be seen in close vicinity of the inversion anticlines. Continued folding of the Cenozoic rocks up to the top Langhian (Fig. 11b) indicates that convergence leading to positive inversion in the area continued till that time. It also indicates continuation of closure of the Neotethys till that time. In the northern Western Desert, basin inversion started in the Late Cretaceous and probably proceeded by way of oblique-slip reactivation of old NE-oriented basin-bounding faults (Sultan & Halim 1988; Bosworth et al. 1999). Neogene extension. During the Neogene time, post-Middle Miocene sediments were deposited, onlapping the northward-subsiding shelf margin as

a result of a regional subsidence of the North Sinai passive margin. Several NW-oriented normal faults with small throws cut tectonosequences 4 and 5 (Neogene) in the study area (Fig. 12). They were formed in response to gravity driven extension in Miocene and post-Miocene times.

Discussion Basin development and inversion The development of the Mesozoic ENE –NE oriented rift basins in Egypt is probably related to the rifting between North Africa-Arabia and Eurasia (Biju-Duval et al. 1979; Argyriadis et al. 1980; Garfunkel & Derin 1984; Dercourt et al. 1986; Mart 1987) when small continental blocks drifted away from North Africa (Robertson & Dixon 1984) during the opening of the Neotethys and the formation of a passive continental margin

STRUCTURE AND TECTONICS OF OFFSHORE NORTH SINAI, EGYPT

Fig. 13. (Continued) (d) top Cretaceous (TB3), (e) top Oligocene (TB4) and (f) top Miocene (TB5).

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in northern Egypt and eastern Mediterranean (Moustafa & Khalil 1988, 1989; Cohen et al. 1990; May 1991; Keeley 1994; Guiraud 1998; Moustafa et al. 1998; Ayyad et al. 1998). This phase of rifting was accompanied by alkaline volcanic activity that was reported in different parts in Egypt, for example, Gebel Arif en-Naqa and Um Bogma in Central Sinai (Weissbrod 1969; Bartov et al. 1980; Meneisy 1986), south Eastern Desert (Serencsits et al. 1979), and north Western Desert (El Shazly 1977). But in the study area of offshore North Sinai and its vicinity no volcanic activity related to such rifting event has been recorded. The inversion of the Early Mesozoic basins in North Sinai is related to the convergence between Africa-Arabia and Eurasia, during the closure of the Neotethys (Smith 1971; Moustafa & Khalil 1989; Moustafa et al. 1991, 1998; Moustafa & Khalil 1994, 1995; Guiraud & Bosworth 1997). The study of the Atlantic spreading data revealed that Africa-Arabia moved WNW relative to Eurasia during this event (Smith 1971). This was also synchronous with the obduction of the ophiolites along the northern and northeastern margins of Arabia, the onset of separation of India and Madagascar, and the development of the European Alpine Chain (Guiraud & Bosworth 1997).

Fold development and mechanism of basin inversion The type of deformation of the Syrian Arc System, including offshore North Sinai, was attributed by many authors to thin-skinned tectonic deformation (e.g. Chaimov et al. 1990, 1992, 1993; McBride et al. 1990; Al-Saad et al. 1992; Abdel Aal & Lelek 1994; Guiraud & Bosworth 1997; Bosworth et al. 1999). Although Searle (1994) does not accept this type of deformation, he relates the system to disharmonic folding in the Mesozoic sediments above Triassic evaporites with no major regional thrust. This diversity in the interpretation of the origin of the Syrian Arc folds is probably related to the different studied portions of the Syrian Arc System by different investigators, in addition to the little subsurface data on these folds. On the other hand, the origin of the folds within this system was considered by most authors as thrustrelated folds. These folds were interpreted differently from one area to another. They are considered as drape folds over reverse faults (Moustafa & Khalil 1989) or over thrust faults (Abdel Aal & Lelek 1994), fault-bend folds (Chaimov et al. 1992, 1993; Abdel Aal & Lelek 1994; Salel & Se´guert 1994), forced push-up folds (Moustafa & Yousif 1990; Abdel Aal & Lelek 1994; Moustafa et al. 1998), fault-propagation folds (Salel &

Se´guert 1994; Moustafa & Khalil 1995), and detachment folds (Abdel Aal & Lelek 1994; Salel & Se´guert 1994). The term positive inversion refers to a switch in tectonic mode from extension to contraction, whereas the term negative inversion refers to a switch from contraction to extension (Williams et al. 1989). In positive inversion extensional basins are contracted and become regions of positive structural relief (McClay 1995). In these basins, pre-existing extensional faults are reactivated by reverse slip (Cooper & Williams 1989). Reactivation might affect isolated extensional faults within the basin or all major faults. Syninversion sedimentation over reactivated faults produce growth anticlines that are characteristic of positive structural inversion (McClay 1995). Inversion-related folding, in the area of study, is represented by fault-propagation folds with steep forelimbs and gently dipping back limbs. Faultpropagation folding occurs when a propagating fault loses slip and terminates upsection by transferring its shortening to a fold developing at its tip (Mitra 1990). No intra-basin faults have been evolved and this could be due to the pre-existing structures, with localization of strain and fault lengths as inherited by updip propagation from the pre-existing fabric (Walsh et al. 2002; Paton 2006). The folding and faulting are synchronous as indicated by the studies of Jamison (1987) and Mitra (1990) in other study areas. Inversion is dependent on pre-existing basin configuration in the initial subsidence, usually extensional phase, and the resolution of compressional forces in the later shortening phase (Lowell 1995). However, deformational forces responsible for inversion are oriented from 0–908 to the preexisting basin-bounding faults. The zero direction is pure strike– slip that is not favourable in effecting inversion. Clay model experiments (Lowell 1985) show that pure strike–slip superposed on an earlier normal fault fabric is simply resolved as horizontal slip on the old normal fault surfaces with little to no evidence of the compressional component of strike –slip. Compression at 908 to existing structures should be highly effective in inversion (Letouzey et al. 1990). The offshore North Sinai is a prime example of compression-dominant inversion. During the closure of the Tethys, which resulted because of the convergence of the African-Arabian plate and the Eurasian plate (Olivet et al. 1984) in post-Santonian time, NW– SE oriented compression inverted the NE–SW trending Late Jurassic–Early Cretaceous half grabens (Yousef et al. 2006); Figure 9a. This is further evidenced by reactivation of the older normal faults and possibly creation of new thrusts, so that several kilometres of thrust overlap occurred NE of el-Arish City.

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Conclusions Five regional unconformities were observed and mapped on seismic sections in the study area. These tectonosequence boundaries (TB1 to TB5) reflect major changes in regional tectonic settings. Most of the mapped structures in offshore North Sinai are located within five NE-oriented inverted structures, whereas the intervening areas are undeformed or only slightly deformed. The inverted structures are controlled by major deep-seated reverse faults with doubly plunging anticlines in which there is thick syn-extensional strata in the core. Similar inversion anticlines also occur onshore (Moon & Sadek 1921; Sadek 1928; Shata 1959; Bartov et al. 1980; Moustafa & Khalil 1994). The inversion structures, in offshore North Sinai, are characterized by fault-propagation folds with steep frontal limbs and gently dipping back limbs. Five tectonosequences (TS1 to TS5) related to the main deformational events are recognized: (a) Upper Jurassic –Santonian syn-rift tectonosequence (TS1); (b) Upper Santonian–Upper Cretaceous compressional tectonosequence (TS2) deposited contemporaneously in the synforms developed between rising antiforms; (c) Paleocene –Upper Oligocene compressional tectonosequence (TS3) is also associated with the folding event, but with a magnitude less than that of tectonosequence 2; (d) northward thickening Miocene tectonosequence (TS4); (e) wedge-like Pliocene-Pleistocene tectonosequence (TS5), which displays extensional faulting with a detachment near its base. Tectonosequence boundaries (TB1 to TB5) are regional unconformities reflecting major changes in regional tectonic setting. TB1 corresponds to rifting and block tilting that started during the Jurassic or earlier in North Sinai. TB2 corresponds to the end of this rifting phase. TB3 and TB4 correspond to major regional unconformities, denoting a continuous compressional tectonic event in offshore North Sinai till Langhian (Middle Miocene) time. TB5 corresponds to an erosional unconformity in the Late Miocene, culminating with the Messinian crisis at 6.7 Ma. Three phases of structural deformation affected the Mesozoic –Cenozoic succession in offshore North Sinai. Phase 1 is a Jurassic to Early Cretaceous rifting phase, which has been indicated in this study to have taken place in several pulses during the Late Jurassic, Neocomian, and AptianSantonian times. Phase 2 is a post-Santonian to Middle Miocene inversion phase, and Phase 3 represents the Neogene extension phase. BP Egypt is thanked for providing the 2D seismic and well data, hardware and software, and for giving permission to publish this paper. We are grateful to P. Bentham and J. Cotton (BP Egypt), for fruitful discussions regarding

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the tectonic evolution of offshore North Sinai. M. B. Longacre, M. Hakim and T. Bevan from BP Egypt are also gratefully acknowledged. We would like to thank C. Homberg (Pierre and Marie Curie University), D. Paton and an anonymous referee for constructive comments that improved the manuscript.

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C AMERLENGHI , A. (eds) Proceedings of the Ocean Drilling Program Science Research, 160, 723– 782. R OBERTSON , A. H. F. & D IXON , J. E. 1984. Introduction. In: D IXON , J. E. & R OBERTSON , A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, 1– 74. S ADEK , H. 1928. The principal structural features of the Peninsula of Sinai. 14th International Geological Congress Proceedings, Madrid, 3, 895– 900. S ALEL , J. F. & S EGURET , M. 1994. Late Cretaceous to Paleogene thin-skinned tectonics of the Palmyrides belt (Syria). Tectonophysics, 234, 265 –290. S EARLE , M. P. 1994. Structure of the intraplate eastern Palmyride fold belt, Syria. Geological Society of America Bulletin, 106, 1332–1350. S ERENCSITS , C., P AUL , H., R OLAND , K. A., E L R AMLY , M. F. & H USSEIN , A. A. 1979. Alkaline ring complexes in Egypt: Their ages and relationship to tectonic development of the Red Sea. Annals of the Geological Survey of Egypt, 9, 102 –116. S HATA , A. 1959. Structural development of the Sinai Peninsula (Egypt). 20th International Geological Congress, Proceedings, Mexico, 225– 249. S MITH , A. G. 1971. Alpine deformation and the oceanic area of the Tethys, Mediterranean and Atlantic. Geological Society of America Bulletin, 82, 2039–2070. S TAMPFLI , G. M., M OSAR , J., F AVRE , P., P ILLEVUIT , A. & V ANNAY , J. C. 2001. Permo-Mesozoic evolution of the western Tethys realm: the Neo-Tethys East Mediterranean Basin connection. In: Z IEGLER , P. A., C AVAZZA , W., R OBERTSON , A. H. F. & C RASQUIN S OLEAU , S. (eds) Peri-Tethyan Rift/Wrench Basins

and Passive Margins. Me´moires du Muse´um national d’Histoire naturelle de Paris, 186, 51–108. S ULTAN , N. & H ALIM , M. A. 1988. Tectonic framework of northern Western Desert, Egypt and its effect on hydrocarbon accumulations. Egyptian General Petroleum Corporation 9th Petroleum Conference, 2, 1– 22. T IBOR , G., B EN -A VRAHAM , Z., S TECKLER , M. & F LIGELMAN , H. 1992. Late Tertiary subsidence history of the Southern Levant Margin, Eastern Mediterranean Sea, and its implications to the understanding of the Messinian event. Journal of Geophysical Research, 97, 17593– 17614. V IDAL , N., A LVAREZ -M ARRO´ N , J. & K LAESCHEN , D. 2000. Internal configuration of the Levantine Basin from seismic reflection data (eastern Mediterranean). Earth and Planetary Science Letters, 180, 77– 89. W ALSH , J. J., N ICOL , A. & C HILDS , C. 2002. An alternative model for the growth of faults. Journal of Structural Geology, 24, 1669– 1675. W EISSBROD , T. 1969. The Paleozoic of Israel and adjacent countries: Part II, The Paleozoic outcrops in southwestern Sinai and their correlation with those of southern Israel. Bulletin of the Geological Survey of Israel, 48, 1– 32. W ILLIAMS , G. D., P OWELL , C. M. & C OOPER , M. A. 1989. Geometry and kinematics of inversion tectonics. In: C OOPER , M. A. & W ILLIAMS , G. D. (eds) Inversion Tectonics. Geological Society, London, Special Publications, 44, 376. Y OUSEF , M., M OUSTAFA , A. R. & S HANN , M. 2006. The Petroleum Play Systems in the Mango – Tineh area: Offshore North Sinai, Egypt (Abs). The 8th International Conference on the Geology of the Arab World, Cairo University, Cairo, Egypt.

Deep structures and seismic stratigraphy of the Egyptian continental margin from multichannel seismic data LAURENT CAME´RA*, ALESSANDRA RIBODETTI & JEAN MASCLE Geoazur UMR 6526, Observatoire Oce´anologique de Villefranche-sur-Mer, BP48 06235 Villefranche-sur-Mer cedex, France *Corresponding author (e-mail: [email protected]) Abstract: Regional multichannel seismic reflection (MCS) profiles across the Egyptian continental slope, offshore the Nile delta, were recorded during the MEDISIS survey (conducted in 2002 on board the R/V Nadir). The results of this survey allow an interpretation of the overall structure and evolution of this passive continental margin. The MCS data were processed using an amplitude preserving pre-stack depth migration technique, which has the advantage of providing a quantitative, and geometrically correct, image of seismic horizons. Well-defined reflecting events allow the identification of three main seismic units. The upper unit (a 7 km thick) is interpreted as the post-rift sedimentary cover of the margin; it includes an undisturbed Middle Cretaceous to Upper Miocene sedimentary pile, covered by thick Messinian (latest Miocene) salt-rich layers and by Pliocene to Quaternary sediments, locally intensively deformed by gravity tectonics. The underlying intermediate acoustic unit (6 km thick on average) is interpreted as the Mesozoic syn-rift sedimentary cover of the margin; the end of the last rifting event is marked by a strong angular unconformity, tentatively of Aptian age. The lower unit may correspond to the thinned continental crust of Africa (12 km thick on average in the study area) and its pre-rift cover. Its base is identified by strong, discontinuous reflector packages about 23–25 km below sea floor, interpreted as indicative of the Moho.

The southern border of the Levantine basin is classically interpreted as a Mesozoic passive continental margin now, progressively entering a diachronous collision with the Anatolian –Aegean-European plates as a consequence of long-term convergence and subduction between Africa and Eurasia (Fig. 1). The Egyptian continental margin, presently covered by a thick sedimentary wedge including significant Messinian salt deposits, represents a 500 km long segment of this passive margin. Moreover, the Egyptian margin is an area of fairly active investigations conducted by numerous oil and gas companies as a consequence of its potential, and partly proven, giant gas reserves (Dolson et al. 2000, 2005; Abdel Aal et al. 2001; Samuel et al. 2003). For a decade, systematic surveys of the Egyptian margin have been conducted by scientists from Geosciences Azur and other academic laboratories in order to obtain multi-scale geophysical images of this margin segment and to better understand its tectonic and sedimentary structures. These investigations have focused on the deep-sea terrigenous cone built offshore of the Nile delta and have resulted in geophysical and geological datasets that have provided important results on the various active geological processes that shape this young sedimentary pile (Loncke 2003; Loncke et al. 2006).

However, little is known on the deep structures of this margin segment and only a few simplified geological cross sections are available, based on unpublished MCS data (Dolson et al. 2000, 2005; Abdel Aal et al. 2001) and gravity–magnetic modelling (Segev et al. 2006; Longacre et al. 2007). Most of the proposed models imply a 12–15 km thick sedimentary sequence beneath the Nile deep-sea fan resting above a c. 10 km thick crust, the latter being regarded either as thinned continental crust (Segev et al. 2006), and/or shortened continental crust (Abdel Aal et al. 2001), or as old oceanic crust bounded by a transform margin segment (Longacre et al. 2007). Here we report results from the MEDISIS multichannel seismic (MCS) survey, conducted in 2002 by Geosciences Azur. The processed MCS data reveal a thick sedimentary wedge (12 –15 km) beneath the Nile deep sea fan, covering, depending on the area, a 9 –12 km thick continental crust.

Tectonic setting The Egyptian margin lies in a complex geodynamical setting, where three major plates (African, Arabian, Eurasian) and several minor plates (Anatolian, Aegean and probably Sinai) interact, and where most of the so-called neo-Tethyan oceanic crust

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 85– 97. DOI: 10.1144/SP341.5 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Shaded-relief bathymetry of the Egyptian continental margin (from Sardou & Mascle 2003) and locations of the MEDISIS MCS profiles (lines). The Egyptian margin can be divided into four different morpho-structural domains that are the Western, Central, Eastern and Levantine provinces (Loncke 2003). In the lower right corner: tectonic setting of eastern Mediterranean basin (modified from Loncke 2003).

and its margins have been consumed by subduction and partly incorporated in the Alpine and Caucasus mountain chains (Smith 1971; Sengo¨r 1979; Smith & Woodcock 1982; Sengo¨r et al. 1984; Dercourt et al. 1986; Stampfli & Borel 2002). Currently, most of the deformations occurring in the Eastern Mediterranean relate to two main plate tectonic mechanisms (Fig. 1): (1)

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the northward motion of the Africa, and more recently Arabia, with respect to the Eurasian plate, probably ongoing since the Late Cretaceous (Olivet et al. 1982; Argus et al. 1989; Rosenbaum et al. 2002); the recent westward extrusion of the Aegean– Anatolian micro-plate (30 –40 Ma), which is bounded by the East and North Anatolian fault zones (McKenzie 1970, 1972; Taymaz et al. 1990; Jackson et al. 1992; Jackson 1994; Le Pichon et al. 1995; McClusky et al. 2000).

These mechanisms are expressed in: (1)

major strike –slip faulting (the East and North Anatolian and Dead Sea faults);

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ongoing and ending subductions (Hellenic and Cyprus subduction); active extension (Aegean Sea, Red Sea); and various collision zones (Fig. 1).

In this geodynamic environment, the Egyptian continental margin cannot be considered as a simple passive margin segment. Came´ra (2006) has recently proposed an evolution of the Egyptian margin area including six major phases, following Dolson et al. (2000). In Palaeozoic to Early Triassic times, intraplate deformations, mainly documented by later reactivated structures, led to important continental erosion and sedimentation, generally associated with magmatical activity (Guiraud & Bosworth 1999). From Late Triassic to Jurassic times, the still poorly constrained break-up of Pangea resulted in several rifting events, ultimately leading to the opening of the neo-Tethyan ocean (Smith 1971; Sengo¨r 1979; Sengo¨r et al. 1984; Dercourt et al. 1986; Ben-Avraham & Ginzburg 1990; Stampfli & Borel 2002) and to the development of a passive continental margin north of Egypt (Dixon & Robertson 1984; Moustafa et al. 1998). Between Cenomanian and

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Miocene times, NW–SE-oriented compression episodes, the so-called Syrian Arc tectonism (Krenkel 1925; Moustafa et al. 1998; Dolson et al. 2000), took place in the easternmost Levantine domain and resulted in the structural inversion of former extensional neo-Tethyan rifted features (Lu¨ning et al. 1998a, b; Kuss et al. 2000; Hussein & Abd-Allah 2001; Gardosh & Druckmann 2005). In Oligocene times, renewed continental extension (early motion of Arabia away from Africa) resulted in the creation of the Gulf of Suez and, in the Middle Miocene, in the opening of the Red Sea. Geological and seismological data suggest that the Sinai block, located between the Suez and Akaba gulfs, behaved like a micro-plate trapped between the African and Arabian plates (Le Pichon & Gaulier 1988; Mascle et al. 2000). At the end of the Miocene (from 5.9 to 5.3 ma), the entire Mediterranean basin experienced the Messinian salinity crisis (Ryan et al. 1973; Rouchy 1986; Gautier et al. 1994; Clauzon et al. 1996), which resulted in thick deposits of salt and anhydrite throughout the deep Levantine basin, and erosional unconformities and terrigenous deposits on the continental shelf and upper slope. Finally, during Pliocene and to recent times, the Egyptian margin underwent rapid subsidence and the accumulation of a thick Plio-Quaternary sedimentary pile (Tibor & Ben-Avraham 2005) that constitutes the present day Nile deep-sea fan (NDSF). The NDSF is deposited above the mobile Messinian salt layers and is affected by typical thin-skinned processes, such as gravity gliding and gravity spreading, in relation with the underlying Messinian salt layers (Loncke 2003; Loncke et al. 2006).

The MEDISIS MCS survey During the 2002 MEDISIS survey, about 1000 km of MCS reflection profiles were recorded along seven regional lines crossing the Egyptian continental margin (Fig. 1). The profiles were acquired using an array of 10 GI-guns, with a total volume of 53 l every 30 s. The reflected signals were recorded via a 4700 m long, 360 channel digital streamer, with a 12.5 m group spacing, across record lengths of 15 s at a 4 ms sampling rate. The regional lines cut successively across the Levantine, Eastern, Central and Western province of the Nile deep-sea fan (Fig. 1) as distinguished by Loncke (2003). One of the challenges was to image the deep structures beneath the up to 3 km thick salt layers. The length of the streamer does not allow us, in principle, to obtain well constrained velocities greater than 4.5 km/s. A procedure of iterative adjustments of the interval velocity was thus performed during the processing of lines, described in detail below, to obtain a velocity model that was calibrated

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with the nearby expanding spread profile experiments of De Voogd et al. (1992). Together these methods provided good estimates of the velocities and depth migration on all lines.

Data processing The data were processed at Geosciences Azur using two different techniques: (1) a standard processing using the Geovector package to produce post-stack time-migrated profiles for preliminary interpretations; (2) a preserved amplitude pre-stack depth migration method (PSDM, see Fig. 2) to generate depth sections. In addition, one line (MD06, see Fig. 3) was processed at Geomar (Kiel) using the Sirius package to obtain a reference PSDM profile for comparison (Fig. 3). Geovector software was used to prepare the data for PSDM (via pre-processing to preserve the amplitude of the acoustic signal), and reduce the signal/ noise ratio. The pre-processing included: sort data to 6.25 m common depth point (CDP), first pass velocity analysis, amplitude attenuation of noisy traces, band pass filter (2, 5, 140, 150 Hz), multiple attenuation in the frequency-wavenumber (FK) domain, normal move out (NMO) velocity analysis, loose external mute, and inverse NMO correction, and finally transform from CDPs to shot gather. Note that in order to preserve acoustic signal amplitudes, no spherical divergence and amplitude corrections were applied (Fig. 2). To obtain the final depth-migrated images we used a Ray þ Born migration technique also known as PSDM. The software (Thierry et al. 1999) is the ‘acoustic’ version of an original ‘elastic’ method proposed by Jin et al. (1992), who introduced an attractive asymptotic method for inverting seismic reflection data. Since the technique of Jin et al. (1992) was introduced, several applications to two-dimensional (2D) and three-dimensional (3D) datasets have been developed for acoustic modelling (Lambare´ et al. 1992; Thierry et al. 1999; Ribodetti et al. 2000; Operto et al. 2003). This approach is very sensitive to the initial velocity model estimate. If the model is incorrect, the reflectors are mislocated and the amplitudes of the velocity distributions are biased. To obtain a reliable image (correct geometry and correct velocity perturbations of seismic reflectors) we used a simple and efficient method, developed by Alyahya (1989) and Agudelo (2005), to perform a quantitative estimate and correction of the velocity macro-model, through a standard ‘migration –velocity –analysis’ approach. Iso-X panels (or common image gathers (CIGs)) are stored during migration and semblance panels estimated to obtain a local correction function for the

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Fig. 2. Processing steps in the Ray þ Born pre-stack depth migration technique. The data are first pre-processed but amplitudes are preserved. A velocity macro-model estimated from velocity analysis, changed in depth using Dix equation, is then interpolated and smoothed. In this model, an asymptotic ray tracing (Lambare´ et al. 1996) is estimated for the computation of all the ray related parameters. Common angle migration, including velocity analysis and velocity macro-model correction, is performed iteratively until Iso-X panels appear flat and semblances are around 1 for all the main reflectors. Finally, CIGs are stacked to obtain an accurate migrated image.

velocity-macro model. Confidence in the migrated image is achieved when the Iso-X panels are flat and when the corresponding semblance panels are around 1 (Alyahya 1989). The velocity macromodel is then iteratively corrected during migration until the semblance panels remain around 1 for all the main reflectors. When such conditions are satisfied, all the CIGs are stacked to obtain the final migrated image (Fig. 2). We performed three iterative corrections of the velocity macro-model; using this technique we observe between 0 to 5 km (c. total length of the seismic streamer) very flat CIGs, suggesting a minimum error for the velocity

model; as a consequence depth location of seismic events have uncertainties of less than 30m. Between 5 and 10 km, a slight upward bend of reflectors shown on the CIGs indicates an error in depth location of about 100 m. Between 10 to 15 km, the CIG panels reveal clear downward bends of reflectors indicating over-estimated velocities and an error in seismic event location on the order of 500 m. Below 15 km, the CIG technique is no longer useful but the variability of deeper seismic event locations during the three iterations allows an estimate of the error in depth values of 1000 and 1300 m.

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Results and interpretation The different processing techniques (post-stack time migration versus PSDM using the Sirius package or Ray þ Born technique) result in comparable seismic successions on most of the MEDISIS profiles. As an example, the three different techniques applied to profile MD06 (see location on Fig. 1) clearly yield similar images (Fig. 3). † On Figure 3a, which illustrates the result of a post-stack time migration using Geovector, we observe a series of well-layered seismic sequences (7 –8 s-twt thick) resting on an acoustic basement that includes a few discontinuous internal reflectors (up to 11 –12 s-twt ). † On Figure 3b, a PSDM of the same line made at Geomar using the Sirius package (only processed down to 10 s), a similar succession is observed: well layered and continuous reflecting units can be seen down to 15 km, resting on a pattern of strong and discontinuous horizons detected down to 18–20 km (assuming adjusted velocities between 5.5 to 6 km/s). † Finally on Figure 3c, PSDM of the same MD06 profile processed using the Ray þ Born technique shows a similar succession, above

discontinuous and low frequency reflectors at around 25 km depth. Using the seismic characteristics of line MD06, we have distinguished five main seismic sequences on the Egyptian margin (Figs 4 & 5). From top to bottom, these are referred as sequences E to A. The uppermost, sequence E, is a well-layered unit whose thickness varies between 1 to 3 km (Figs 4– 6). It corresponds to the Pliocene and Quaternary sedimentary cover of the margin. Below this recent cover, an almost reflection-free sequence D shows numerous diapir-like deformations and is inferred to correspond to the Messinian evaporite deposits (Figs 4–6). Its thickness varies between a few hundred meters (upslope) and 3 km (at the base of the slope). It is bounded at its base by a series of strong and continuous reflectors that underline an important velocity inversion (average interval velocities are in the order of 4000 m/s within the salt and 2500 m/s just below). Below the Messinian horizons, sequence C comprises welllayered and undisturbed reflectors (Figs 4–6). It is separated by a strong angular unconformity from underlying discontinuous and strongly fractured horizons of sequence B. Finally, below sequence B, discontinuous, low frequency, and locally

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faulted reflectors define sequence A (Fig. 4), interpreted as evidence of a weakly layered continental crust. At its base, sequence A is bounded by a series of discontinuous and very low frequency reflector packages (see Figs 4 & 7) situated at an average 23 –25 km depth, which we interpret as indicative of the Moho. From these data, the seismic stratigraphy of the Egyptian margin can be interpreted in terms of three main tectono-stratigraphic units (Fig. 8). The two uppermost units 1 and 2 are well-layered and are interpreted as the sedimentary cover of the margin. Unit 1 includes sequences E to C, which are all characterized by continuous horizons (locally disrupted by gravity structures). In contrast Unit 2, which includes only sequence B, shows medium to high-amplitude reflectors and is highly faulted. A strong unconformity separates Unit 1 and 2. Below, Unit 3, which corresponds to sequence A, shows low-frequency discontinuous reflectors bounded at their base by a few, very thick, reflector packages. It thus corresponds to the crustal basement. Detailed description of these units is presented below in dedicated sections.

The sedimentary cover From top to bottom, the sedimentary cover of the margin includes sequences E to B.

As already indicated, sequence E is referred to as the recent Pliocene and Quaternary sediment blanket constituting most of the present Nile deepsea fan. This Pliocene –Quaternary cover, which locally reaches thickness in excess of 4500 m, is intensively cut by a dense set of extensional faults resulting from salt tectonics, creating small polygonal grabens on the upper slope, and elongated crestal grabens on the middle slope (Diegel et al. 1995; Loncke 2003; Loncke et al. 2006). At the base of the slope, this Pliocene–Quaternary pile is strongly deformed by folds, reverse faulting and thrusts that also result from salt movements downslope (Cramez & Jackson 2000; Loncke et al. 2006) (Figs 6 & 7). Sequence D corresponds to the Messinian evaporite section. The lower boundary comprises highamplitude reflectors, while its internal seismic character is a transparent seismic facies that includes a few high-amplitude and low-frequency reflectors; chaotic reflectors are also locally observed, particularly near the base of the sequence and in areas of intense deformation. Since its deposition, the Messinian evaporites have been moving downslope and this motion has introduced important salt thickness variations. Such variations are illustrated in Figure 6 along the MD06 profile that extends from the continental slope up to the deep basin. Along the upper slope/shelf area, the Messinian salt layers are almost absent, either because they were not deposited or/and because part of the salt has moved downslope inducing growth fault systems [see (1) on Fig. 6]. On the middle slope, the salt layers are almost in concordance with the upper cover and their thickness reaches up to 1500 m, except in several crestal grabens [see (2) on Fig. 6]. At the base of the slope, the Messinian layers reach thickness of up to 3000 m as a consequence of their stacking, which leads to various and intensively folded features [see (3) on Fig. 6]. Directly below the salt, bounded at its base by strong reflective horizons, the variably thick sequence C is made of several sub-sequences, including a poorly layered upper section and a lower section made of more continuous horizons. On most profiles sequence C is separated from sequence B, either by a strong acoustic contrast or by an angular unconformity along which internal reflectors are truncated below and onlapping above (Fig. 5). No specific evidence of tectonic activity can be recognized within this sequence. Sequence B includes concordant reflectors, of variable continuity, amplitude and frequency content, characterized by sub-parallel configurations and often dislocated and offset by small faults. Its thickness is highly variable, but is estimated to be 6000 m on average (see Figs 4 & 5).

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The identification of an strong unconformity between sequences B and C is an important result of this study. This surface is clearly observed on profile MD06 (Fig. 5) and has been detected on most of the profiles across the margin, even if locally it is less distinct. These observations indicate that this surface has a regional significance. Using results from ocean drilling program (ODP) site 967 (Emeis et al. 1996) on the Eratosthenes seamount, near profile MD06 (Fig. 1), we correlate this unconformity with the top of indurated limestones of Aptian age (119 –115 Ma). According to Whiting (1998), these sediments, which include interbedded breccias, were deposited in a lower bathyal environment (a few hundred meters water depth). We therefore are inclined to interpret the fractured sequence B (see Fig. 8) as a syn-rift sequence, dominantly made of carbonate and deposited in a rather shallow environment. The unconformity, seen at its top, may mark the end of a rifting phase and is interpreted as a break-up unconformity

occurring here in Aptian time. If correct, this hypothesis implies that sequences C to E correspond to the post-rift evolution of the Egyptian margin, initiated in Middle Cretaceous time (Fig. 8).

The acoustic basement Sequence A is interpreted as the basement of the Egyptian continental margin (see Fig. 8). Its seismic facies is characterized by packages of intermittent, sub-horizontal, low-frequency reflectors. Such deep discontinuous reflectors can be identified on most of the profiles. Moreover one of the post-stack time-migrated profiles (MD08, Fig. 7) displays more continuous seismic events around 11 s-twt slightly deepening toward the south. Despite the presence of acoustic noise generated during MCS data acquisition and possibly during processing, we interpret these few scattered strong reflectors, detected at the base of this unit, as potential indications of the Moho. If our

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interpretation is correct, this indicates a Moho depth averaging 23–25 km over most of the Egyptian continental slope and a basement thickness between 10 to 12 km (Fig. 8), assuming velocities between 5.6 to 6.5 km/s as provided by MCS processing. In absence of refraction data to better constrain crustal velocities, we have used for the basement the average velocity model (between 5.6 to 6.5 km/s) deduced from MCS processing. This indicates that the thickness of the basement decreases from south to north; although beneath the upper part of the slope the thickness is around 14 km, it reaches only 10 to 12 km in most of the study area, whereas below Eratosthenes seamount it reaches at least 18 to 20 km (see Fig. 9). We

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Fig. 9. Geological cross sections of the Egyptian continental margin based on the interpretation of MEDISIS MCS data. The Egyptian continental margin is made of a succession of three main seismic units: (1) a thinned continental crust (and its pre-rift cover) (Unit 1); (2) a syn-rift sedimentary section (Unit 2) and (3) a post-rift sedimentary cover of the margin (Unit 3).

DEEP STRUCTURES OF THE EGYPTIAN MARGIN

Conclusions Processing and analyses of MCS reflection data recorded during the 2002 MEDISIS survey allow the recognition of several seismic sequences along the Egyptian continental margin offshore the Nile delta. The tectonic and sedimentary structures imaged in these seismic profiles lead us to interpret these sequences in terms of three main tectonostratigraphic units, in the overall geodynamical context of the Eastern Mediterranean (Fig. 9): † The upper Unit 1 is on the order of 7 km thick and corresponds to the post-rift sedimentary cover of the Egyptian margin. This post-rift unit comprises three main sedimentary sequences that are, from top to bottom: (a) a 2 to 4 km thick Pliocene–Quaternary cover, intensively deformed by gravity tectonics; (b) a variably thick (few hundred meters to 3 km) latest Miocene (Messinian) blanket in which evaporites dominate along the middle to lower slope, whereas clastic deposits prevail along the upper continental slope – the inter-bedded thick Messinian evaporites has induced a strong decoupling (Loncke et al. 2006) between the upper cover and the remaining sedimentary pile and therefore the superficial tectonic features do not reflect any basement structural trends; (c) a Middle Cretaceous to Upper Miocene sedimentary pile (1 to 4 km thick depending on the area). † The middle Unit 2 averages 6 km in thickness and is interpreted as the syn-rift sedimentary cover of the margin. This unit is characterized by strong and low frequency reflectors that are frequently disrupted by faults, in response to either successive rifting events or to a very slow and long rift evolution. Correlation of the seismic reflectors with ODP data near Eratosthenes seamount suggests that Unit 2 may correspond to relatively shallow water carbonates of Jurassic to Aptian age. If this hypothesis is correct, it would imply the last rifting event to have ended sometime in the Early Cretaceous. This event is marked either by a major angular unconformity or by strong seismic contrasts observed at the boundary between Unit 1 and Unit 2. † The lower Unit 3 is interpreted as a stretched continental crust extending over most of the studied area, with the exception of the northwesternmost region where evidence of oceanic crust has been found from expanding spread profile seismic experiments (De Voogd et al. 1992). This continental basement is bounded by scattered, low-frequency reflector packages, which are good candidates for a Moho lying at about 23– 25 km.

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The authors gratefully acknowledge the MEDISIS crew and scientific staff aboard the R/V Nadir. We thank F. Sage and L. Schenini for help in seismic data processing. We thank D. Klaeschen for help in MD06 profile processing performed using Seismic and Sirius at Geomar centre (Kiel). This work was partly supported by GDR Marges. Detailed reviews by G. Netzeband, D. J. Shillington and D. Praeg have greatly improved preliminary versions of this paper.

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Le Pichon, X. & Gaulier, J.-M. 1988. The rotation of Arabia and the Levant fault system. Tectonophysics, 153, 271 –294. Le Pichon, X., Chamot-Roocke, N. & Lallemant, S. 1995. Geodetic determination of the kinematics of central Greece with respect to Europe: implications for Eastern Mediterranean Tectonics. Journal of Geophysical Research, 100, 12 675– 12 690. Loncke, L. 2003. Le delta profond du Nil: structure et e´volution depuis le messinien (Mioce`ne terminal). The`se de doctorat de l’Universite´ Paris VI, Spe´cialite´ Sciences de la Terre, PhD thesis, 184. Loncke, L., Gaullier, V., Mascle, J., Vendeville, B. & Came´ra, L. 2006. The Nile deep-sea fan: an example of interacting sedimentation, salt tectonics, and inherited subsalt paleotopographic features. Marine and Petroleum Geology, 23, 297–315. Longacre, M., Bentham, P., Hanbal, I., Cotton, J. & Edwards, R. 2007. New crustal structure of the Eastern Mediterranean basin: detailed integration and modeling of gravity, magnetic, seismic refraction, and seismic reflection data. EGM 2007 International Workshop Innovation in EM, Grav and Mag Methods: a new Perspective for Exploration Capri, Italy, April 15–18, 2007, 4. Lu¨ning, S., Kuss, J. & Bachmann, M. 1998a. Sedimentary response to basin inversion: Mid Cretaceous– Early Tertiary Pre- to syndeformational deposition at the Areif El Naqa anticline (Sinai, Egypt). Facies, 38, 103– 136. Lu¨ning, S., Marzouk, A. M., Morsi, A. M. & Kuss, J. 1998b. Sequence stratigraphy of the Upper Cretaceous of central-east Sinai, Egypt. Cretaceous Research, 19, 153–196. Mascle, J., Benkhelil, J., Bellaiche, G., Zitter, T., Woodside, J. & Loncke, L. 2000. Marine geologic evidence for a Levantine-Sinai plate, a missing piece of the Mediterranean puzzle. Geology, 228, 779– 782. McClusky, S., Balassanian, S. et al. 2000. GPS constraints on plate kinematics and dynamics in the eastern Mediterranean and Caucasus. Journal of Geophysical Research, 105, 5695– 5719. McKenzie, D. P. 1970. Plate tectonics of the Mediterranean region. Nature, 226, 239–243. McKenzie, D. P. 1972. Active tectonics of the Mediterranean region. Geophysical Journal of the Royal Astronimical Society, 30, 109– 185. Moustafa, A. R., El-Badrawy, R. & Gibali, H. 1998. Pervasive E-ENE oriented faults in northern Egypt and their effect on the development and inversion of prolific sedimentary basins. EGPC 14th Exploration and Production Conference, Cairo, Egypt, 51– 67. Netzeband, G. L., Gohl, K., Hubscher, C. P., BenAvraham, Z., Dehghani, G. A., Gajewski, D. & Liersch, P. 2006. The Levantine Basin – crustal structure and origin. Tectonophysics, 418(3–4), 167– 188. Olivet, J.-L., Bonnin, J., Beuzart, P. & Auzende, J.-M. 1982. Cine´matique des plaques et pale´oge´ographie, une revue. Bulletin de la Socie´te´ Ge´ologique France, 7, 875–892. Operto, S., Lambare´, G., Podvin, P., Thierry, P. & Noble, M. 2003. 3-d rayþborn migration/inversion – part 2: application to the seg/eage overthrust experiment. Geophysics, 68, 1357– 1370.

DEEP STRUCTURES OF THE EGYPTIAN MARGIN Ribodetti, A., Thierry, P., Lambare´, G. & Operto, S. 2000. Improved multiparameter rayþborn migration/ inversion. Society of Exploration Geophysicists, 70, 1032–1035. Rosenbaum, G., Lister, G. S. & Duboz, C. 2002. Relative motions of Africa, Iberia and Europe during Alpine orogeny. Tectonophysics, 359, 117 –129. Rouchy, J.-M. 1986. Les e´vaporites mioce`nes de la Me´diterrane´e et de la mer rouge et leurs enseignements pour l’interpre´tation des grandes accumulations e´vaporitiques d’origine marine. Bulletin de la Socie´te´ Ge´ologique de France, 8, II, n. 3, 511– 520. Ryan, W. B. F., Hsu¨, K. J. et al. 1973. Initial Reports of the Deep Sea Drilling Project, 13, 1447. Samuel, A., Kneller, B., Raslan, S., Sharp, A. & Parsons, C. 2003. Prolific deep-marine slope channels of the Nile delta, Egypt. American Association of Petroleum Geologists Bulletin, 87, 541–560. Sardou, O. & Mascle, J. 2003. Cartographie par sondeur multifaisceaux du delta sous marin profond du Nil et des domaines voisins. Deux cartes (Morphobathyme´trie et mosaiques d’images acoustiques). Special Publication, CIESM, Monaco. Segev, A., Rybakov, M., Lyakhovsky, V., Hofstetter, A., Tibor, G., Goldshmidt, V. & Ben Avraham, Z. 2006. The structure, isostasy and gravity field of the Levant continental margin and the southeast Mediterranean area. Tectonophysics, 425, 137–157. Sengo¨r, A. M. C. 1979. The North Anatolian transform fault: its age, offset and tectonic significance. Journal of the Geological Society of London, 136, 269– 282. Sengo¨r, A. M. C., Yilmaz, Y. & Sungurlu, O. 1984. Tectonics of the Mediterranean Cimmerides: nature

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and evolution of the western termination of Palaeo-Tethys. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean, Geological Society, London, Special Publications, 17, 77–113. Smith, A. G. 1971. Alpine deformation and the oceanic areas of the Tethys, Mediterranean and Atlantic. Bulletin of the Geological Society of America, 82, 2039– 2070. Smith, A. G. & Woodcock, N. H. 1982. Tectonic synthesis of the Alpine-Mediterranean region: a review. In: Berckhemer, H. & Hsu, K. (eds) AlpineMediterranean Geodynamics. AGU, Geodynamics Series, 7, 15–38. Stampfli, G. M. & Borel, G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters, 196, 17– 33. Taymaz, T., Jackson, J. A. & Westaway, R. 1990. Earthquake mechanisms in the Hellenic Trench near Crete. Geophysical Journal International, 102, 695–731. Tibor, G. & Ben-Avraham, Z. 2005. Late Tertiary paleodepth reconstruction of the Levant margin off Israel. Marine Geology, 221, 331– 347. Thierry, P., Operto, S. & Lambare´, G. 1999. Fast 2-d rayþborn migration/inversion in complex media. Geophysics, 64, 162– 181. Whiting, B. M. 1998. Subsidence record of early-stage continental collision, Eratosthenes platform (sites 966 and 967). In: Robertson, A. H. F., Emeis, K.-C., Richter, C. & Camerlenghi, A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 60, 509– 515.

Controls and evolution of facies patterns in the Upper Barremian– Albian Levant Platform in North Sinai and North Israel MARTINA BACHMANN*, JOCHEN KUSS & JENS LEHMANN Department of Geosciences, Bremen University, P.O. Box 330440, 28334 Bremen, Germany *Corresponding author (e-mail: [email protected]) Abstract: The Upper Barremian–Albian Levant Platform was studied in North Sinai and Israel (Galilee and Golan Heights) by bio- and lithostratigraphy, facies analyses, and sequence stratigraphy. Integrating shallow-marine benthic foraminifera (mainly orbitolines), ammonite, and stable isotope data resulted in a detailed stratigraphic chart. Transects across the shallow shelf in both regions are based on facies analysis and form the basis for depositional models. In both transects five platform stages (PS I–V) were identified, which differ significantly in their stratigraphic architecture, mainly controlled by local tectonics, climate and second-order sea-level changes. In North Sinai, a transition from a shallow-shelf that is structured by sub-basins through a homoclinal ramp into a flat toped platform is recognized, while the sections in North Israel show a transition from a homoclinal ramp into a fringing platform. Local normal faults influenced the depositional architecture of the Upper Barremian–Lower Aptian strata in North Sinai and were attributed to syn-rift extensional tectonics. Four second-order sequence boundaries were identified, bounding MidCretaceous Levant depositional sequences. These well-dated second-order sequence boundaries are MCL-1 (Late Barremian), MCL-2 (earliest Late Aptian), MCL-3 (Lower Albian), and MCL-4 (Late Albian). The sea-level history of the Levant Platform reflects the Late Aptian–Albian global long-term transgression, while the second-order sea-level changes show good correlation with those described from the Arabian plate.

The Late Barremian– Albian Levant Platform is an ideal case study for studying platform development in a setting of global long-term sea-level change, climate variation, and geodynamics. The platform was located at the northern rim of the North African/Arabian plate and the southern border of the Tethyan Ocean, respectively (Fig. 1a). Extending from southern Lebanon to northern Egypt the Levant Platform strikes out in a narrow stripe parallel to the recent coastline (Saint-Marc 1974; Braun & Hirsch 1994; Kuss & Bachmann 1996; Rosenfeld et al. 1998). Its central part was studied in Israel (Galilee and Golan Heights) and its southwestern part in North Sinai, both characterized by good exposures (Fig. 1b, c). During the Late Barremian–Albian about 500 m of shallow marine deposits accumulated on the Levant Platform. During the Mesozoic, the region was characterized by extensional and compressive tectonical processes related to the opening and closure of the Neotethys (e.g. Keeley 1994; Hirsch et al. 1995; Stampfli & Borel 2002; and other papers in this volume) forming the main tectonical structures such as the Syrian Arc fold belts and the Dead Sea Transform (DST). However, the analysed interval represents a consistent platform succession deposited since the Late Barremian. This interval has been interpreted to be deposited in a post-extensional setting during mid-Cretaceous

transgression and indicating relative tectonical quiescence before the onset of Late Cretaceous compression (Moustafa & Khalil 1990; Hirsch et al. 1995; Kuss & Bachmann 1996). Most former studies concentrate on small areas and/or selected stratigraphical or palaeontological parameters. Litho- and biostratigraphical concepts of the Lower Cretaceous of Galilee and the Golan Heights are summarized in Rosenfeld & Hirsch (2005) and Bachmann & Hirsch (2006); tectonic concepts are summarized in Flexer et al. (2005) and Gilat (2005). Stratigraphical subdivisions of the Lower Cretaceous of the Sinai, are mainly based on Said (1971) and Bartov & Steinitz (1977). Younger interpretations include facies, stratigraphic and sedimentological data (Aboul Ela et al. 1991; Askalany & Abu-Zeid 1994; Bachmann & Kuss 1998; El-Araby 1999; Steuber & Bachmann 2002; Bachmann et al. 2003). Only few studies document marine Barremian– Lower Aptian sediments in North Sinai (Arkin et al. 1975; Morsi 2006; Abu-Zied 2007, 2008). In the present paper, we summarize various stratigraphical data with respect to a consistent correlation of various lithological units in the area and add new data from the Barremian–Lower Aptian succession of North Sinai. The methodology includes bio- and isotope stratigraphy, the interpretation of sedimentological structures in the field,

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 99– 131. DOI: 10.1144/SP341.6 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Location of the study area. (a) Tectonic map, simplified after Garfunkel (1998) with the location, the two study areas in North Sinai and northern Israel (Galilee and Golan Heights) straddling the DST fault. (b) Aptian–Albian palaeoenvironmental map of the Levant Platform, modified from Rosenfeld et al. (1998) and supplemented by results from North Sinai indicating the extension of the shallow platform facies belts. (c) Gebel (mountains) examined in North Sinai indicated on a satellite map. ‘B’ indicates the tectonically restored position of the Golan Heights section B, when assumed the lateral movement of 100 km as indicated, for example by Garfunkel & Ben-Avraham (1996).

small-scaled geological mapping, and microfacies analysis. Our interpretations are focused on facies reconstructions and long-term, second-order, sealevel changes within a high-resolution stratigraphical frame. Moreover, we estimated the climatic and tectonic influences on the sedimentation. The data allow us to reconstruct and interpret the varying Late Barremian–Albian shelf geometry, as well as the timing and interpretation of the platform development. In this respect syn-depositional tectonical processes were determined, which were previously unknown. These data are interpreted, allowing the detailed timing of the Late Barremian–Albian succession. This results in an evaluation of the facies, long-term local and regional second-order sea-level changes, and climate framework, characterizing the complex sedimentary system of the Levant Platform. Furthermore, the data allow us to the reconstruct and interpret the varying Late Barremian–Albian shelf geometry. In this respect,

the paper is a review. Together with new data, especially from the Barremian–Lower Aptian succession of North Sinai, the dataset allows the interpretation of the platform development. Most former studies concentrate on small areas and/or selected stratigraphical or palaeontological parameters. Litho- and biostratigraphic concepts of the Lower Cretaceous period of Galilee and the Golan Heights are summarized in Rosenfeld & Hirsch (2005) and Bachmann & Hirsch (2006), tectonic concepts are summarized in Flexer et al. (2005) and Gilat (2005). Stratigraphic subdivisions of the Lower Cretaceous of the Sinai are mainly based on Said (1971) and Bartov & Steinitz (1977). Younger interpretations include (Aboul Ela et al. 1991; Askalany & Abu-Zeid 1994; Bachmann & Kuss 1998; El-Araby 1999; Steuber & Bachmann 2002; Bachmann et al. 2003). Only few studies document marine Barremian– Lower Aptian sediments in North Sinai (Arkin et al. 1975; Morsi 2006; Abu-Zied 2007).

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

Geological framework The Levant Platform developed at the northern passive continental margin of the African– Arabian plate close to the plate boundary with the Anatolian plate. The tectonical pattern in the Eastern Mediterranean Levantine region has important influence on the Cretaceous platform development characterized by this position at the triple junction of the African, Arabian, and Anatolian plates (Flexer et al. 2005). The Early Mesozoic opening of the Neotethys generated an east –west striking rift system with subsiding areas and thus created the Levant Basin at the northeastern edge of the African– Arabian plate (Keeley 1994; Hirsch et al. 1995; Garfunkel 1998; Flexer et al. 2005). The shallow-marine Levant Platform formed at the southeastern passive margin of the Levant Basin trending in a narrow strip parallel to the Mediterranean coastline from Lebanon and Syria through Israel to North Sinai, Egypt (Garfunkel 1998). The breakup of the Levant Platform commenced with the Late Coniacian –Palaeogene inversion of extensional faults (Flexer et al. 2005). In northern Israel, the initial ramp geometry of the Early Cretaceous Levant Platform developed into a flat-topped platform in the Early Aptian (Bachmann & Hirsch 2006). In the Albian it is saddled by the formation of fringing rudist-reefs at the shelf break (Sass & Bein 1982; Ross 1992). In North Sinai, the initial Upper Barremian–Lower Aptian platform is not yet analysed. The Upper Aptian– Lower Albian platform geometry is described as a distally steepened ramp (Bachmann & Kuss 1998). The region in between North Sinai and northern Israel is described as a gradual shallow –deep transition (Rosenfeld et al. 1998) known only from the subsurface (Rosenfeld & Hirsch 2005).

The investigated sections and lithostratigraphical concepts In northern Israel/Galilee and the Golan Heights, most outcrops are located along strike–slip faults orientated parallel to the DST fault and in asymmetric Syrian Arc anticlinal structures (Flexer et al. 2005; Gilat 2005). In North Sinai, several anticlinal Syrian Arc structures form good outcrop conditions (Moustafa & Khalil 1990). Owing to the regional differentiated nomenclature, with a high number of units, members and formations, we compile standard sections for the two platform edges composed from several sections in both regions (Fig. 1). Galilee/Golan. We follow the lithostratigraphical subdivisions of the Cretaceous succession in the Golan Heights and Galilee established by Eliezri

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(1965) and modified by Rosenfeld et al. (1995), while additionally distinguishing a new member (Fig. 2). The terrestrial sandstones at the base of the succession are known as the Hatira Formation and occur in the subsurface of northern Israel (Rosenfeld et al. 1998). The overlying marine succession is subdivided into five formations: Nabi Said (ooid-rich limestones), Ein el Assad (limestones and subordinated marlstones), Hidra (sandstones, marlstones, and few limestone beds), Rama (limestones and marlstones), and Yagur (limestones and dolomites) (Fig. 2). The studied succession is 440 m thick; the standard section is based on three sections studied in Galilee (Har Ramin, Rama and Ein Netofa) and one on the Golan Heights (Ein Quniya) (Bachmann & Hirsch 2006), (Fig. 1). Hence outcrops of marine Lower Cretaceous sediments are generally rare in Galilee and the Golan Heights, two sections (Har Ramin/Galilee and Ein Quniya/Golan Heights) comprise the entire 440 m thick Upper Barremian –Albian succession. Weathering profiles, bedding surfaces, sedimentary structures, stratigraphical, and facies interpretation are described in Bachmann & Hirsch (2006). North Sinai. We use the lithostratigraphical subdivision of the Upper Barremian–Albian succession of North Sinai of Said (1971), applied in current geological maps (Fig. 2). The platform succession starts above terrestrial sandstones (Malha Formation) with a limestone, dolomite, and marlstone alternation characterized by detrital influence (lower Rizan Aneiza Formation). Sandstone, marlstone, limestone, and dolomite alternations with deltaic influence above represent the upper Rizan Aneiza Formation; the subsequent succession of limestones and dolomitic marlstones without considerable siliclastic input was related to the Halal Formation. The thickness of individual sections varies greatly, depending on their palaeogeographical position. From south to north, the thickness of the lower part of the Rizan Aneiza Formation varies from 0 to 170 m and from 120 to 250 m for the upper part of the Rizan Aneiza Formation. The lower Halal Formation studied at Gebel Mansour reaches 220 m. The Upper Barremian–Upper Albian succession of North Sinai was studied in five anticlinal structures, comprising the following Gebels: Amrar, Rizan Aneiza, Raghawi, Mansoura, and southeastern Maghara (Fig. 1c). Within this frame, the Raghawi section (A) describes a proximal setting of the lower Rizan Aneiza Formation (Upper Barremian –lower Upper Aptian) (Fig. 2), while the Amrar section (GA) shows a more distal expression of the same interval. One Rizan Aneiza section (D) includes the Lower/Upper Aptian boundary, while the eastern Maghara section

C. pavonia

oa3

FO O. (M.) s.

deshayesi

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waagenoides sarasini giraudi feraudianus satousi vandenheckii

P. cormyi

tuarkyricus

P. infracretacea

weissi

E. charollaisi

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furcata

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melchioris

Ot-Op O.(M.) texana O. (M.) parva

FO O. (M.) t.

Op O. (M.) parva

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Pc-Ol P. cormyi O.(M.) lotzei

FO P. c.

local LO C. d. local FO P. i.

Pl P. lenticularis

Pl-Cd P. lenticularis C. decipiens FO P. l.

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Unit A Halal Fm.

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oa2

Hidra Fm oa1

Unit A Rizan Aneiza Fm.

Hidra Fm P. wienandsi, H. dinarica, P infracretacea P infracretacea, C. sugdeni, A. podolica, H. dinarica T. marsicana E. charollaisi

Unit 0

Ein El Assad Fm

Nabi Said Fm

Malha Fm Rizan Aneiza Fm.

Hatira Fm

Fig. 2. Biostratigraphic subdivision of the Upper Barremian –Albian ranges of the LFBs, timing of the lithostratigraphical formations and members and ranges of the sections. The ranges of the most important benthic foraminifers and the ammonites are compiled from several authors (see text) and arranged in the chronostratigraphic framework of Ogg et al. (2004). Accompanying organisms are taken from Bachmann & Hirsch (2006) and Bachmann et al. (2003). ‘oa’ refers on ostracod assemblages observed in North Sinai (Bachmann et al. 2003).

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nodoso costatum

S

Raghawi

O. (M.) subconcava

N

Amrar Raghawi Mansour

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mammillatum

Range of sections

Formations Sinai

Amrar

Os

nolani

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E. plicatius periadriatic A. dardei E. murgendsis

dentatus

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Golan

Galilee Golan / Ein Quinya section

loricatus

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oa4

jacobi

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Formations Galilee/ Golan Heights

offshore Galilee

lautus

tardefurcate

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Ammonites Accompanying Northern Sinai organisms

Talme Yafe Fm (basin) Carmel reefs

Thierry et al. 1998

LFB (Larger benthic foraminifer biozones) FO / LO Name

Kutatissites bifurcatus Costidiscus recticostatus Heteroceras coulleti Cheloniceras (C.) cornuelianum C. (C.) cf. quadrarium Paradeshayesites cf. grandis Aconeceras nisus C. (Epicheloniceras) tschernyschewi

Substage Middle Lower

Stage Albian

Time (Ma) Series

Larger benthic foraminifers

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European ammonite zones

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(MgE) completes the Lower–Middle Albian succession (upper Rizan Aneiza Formation) to the SE. These sections are summarized and span a 40 km wide transect from the distal Amrar anticline in the NW to the proximal Maghara structure in the SE (Fig. 1c). Eight sections from the upper Rizan Aneiza and lower Halal formations (Upper Aptian –Albian) were presented in former publications (Bachmann & Kuss 1998; Bachmann et al. 2003); three of them are added in this work (Fig. 1c) to illustrate facies and sequence stratigraphy.

Methods For the North Sinai sections, weathering profiles, bedding surfaces, sedimentary structures, and facies characteristics are documented. Sample distances of the limestone intervals are usually less than 1 m and may be higher in the marlstone and sandstonedominated intervals. Stratigraphy and facies were analysed by using field data, thin sections, ammonite findings and washed samples, and stable isotopes analyses required powdered samples. Biostratigraphy. Biostratigraphy is based on benthic foraminifers (mainly orbitolinids) and ammonites. The orbitolinid distribution (taxonomy according to Schroeder 1975) allows the definition of six larger benthic foraminifera biozones (LFBs), of which five are originally defined on the first occurrence and/or last occurrence (FO/LO) of the eponymous species in the Galilee/Golan area (Fig. 2, Bachmann & Hirsch 2006). We compare our biostratigraphic data with those of the northern Tethys (Schroeder & Neumann 1985; Arnaud-Vanneau 1998; Arnaud et al. 1998; Bernaus et al. 2002; Schroeder et al. 2002; Conrad et al. 2004) and the Middle East (Saint-Marc 1974; Simmons & Hart 1987; Simmons 1994; Simmons et al. 2000), and integrate the LFBs with the chronostratigraphy of Hardenbol et al. (1998a, b) and Ogg et al. (2004) also considering calcareous algae following Bachmann & Hirsch (2006). Ammonites from nearly all sections in North Sinai are determined and allow comparison with the Tethyan ammonite zonation (e.g. Rawson et al. 1999; Garcı´a-Monde´jar 2009) and to correlate the LFBs with the ammonite range charts. The stratigraphical framework of the Upper Aptian –Albian succession additionally comprises the ranges of ostracods and rudists and is supported by graphical correlation (Bachmann et al. 2003). Stable isotopes. A total of 69 bulk rock samples were analysed for carbon isotopes from the Upper Barremian–Lower Aptian of section A (Fig. 1c), North Sinai. To avoid diagenetic alteration, all samples were selected from the micritic parts of

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the limestones. The stable isotope composition was measured using a Finigan mass spectrometer at Research Center Ocean Margins (RCOM) Bremen. The results are expressed in the common d-notation in per mille relative to PDB (PeeDee belemnite) standard. Facies analysis. The thin-section analysis includes the determination of the relative abundance of skeletal and non-skeletal components, matrix composition as well as grain size. We distinguish between abundant, common, and rare occurrence of the components. Microfacies types (MFT) were distinguished according to their texture, matrix, and components, added to data observed already in the field such as lithology, bedding patterns terrigenous input, and abundance of ferruginous ooids and macrofossils. The microfacies types are summarized in facies zones (FZ), which include microfacies types occurring in similar environments. Palaeoenvironment reconstruction and sequence stratigraphical interpretation. Within the stratigraphic framework, the combination of small-scale mapping, log-correlation, and microfacies data (including a review of further data) results in palaeoenvironmental maps of North Sinai to northern Israel for five time slices. Despite the limited number of sections, the general platform geometries could be reconstructed. The stratigraphical correlation of all sections within the two analysed depositional areas allows the interpretation of the platform architecture as a function of accommodation, sediment supply, and production. Lateral and vertical facies changes reflect water-depth and palaeoenvironment variations, which were compared between the Galilee/Golan Heights and Sinai and are interpreted in considering sea-level history and tectonic development. Second-order sequences are reflected by lateral and vertical facies distribution patterns. Sequence stratigraphical terminology as originally developed for thirdorder sequences (Vail et al. 1991) is used herein for the second-order cycles. Our tectonic interpretations benefit from the models given by Moustafa (2010).

Stratigraphy The stratigraphical subdivision is focused on biostratigraphy supplemented by stratigraphical interpretations of d13-C curves. Biostratigraphy of the Upper Barremian–Lower Aptian from northern Sinai on ammonites and orbitolinids is integrated with carbon isotope measurements, summarized in Figure 2. This chart allows the comparison of the studied shoal-water environments with stratigraphical subdivisions of basinal settings and platform settings. Subdivision of the Upper

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Aptian– Albian is based on orbitolinids, ostracods, rudists and ammonites, data based mainly on Bachmann et al. (2003). The stratigraphical interpretation of the sections is shown in Figure 3a, b.

Benthic organisms The abundant orbitolinids in all sections and their wide range of habitats (lagoonal to deeper platform areas, e.g. Pittet et al. 2002) allow the subdivision of the Upper Barremian–Middle Albian succession into five larger benthic foraminifer biozones (LFBs). From base to top these are: Pl–Cd Palorbitolina lenticularis–Choffatella decipiens, Pl P. lenticularis, Pc–Ol Praeorbitolina cormyi – Orbitolina (Mesorbitolina) lotzei, Op O. (M.) parva, Ot –Op O. (M.) texana– O. (M.) parva in the Golan Heights/Galilee area (Bachmann & Hirsch 2006). The occurrence of Eopalorbitolina charollaisi, which is not younger than Late Barremian (Clavel et al. 1995; Arnaud et al. 1998; Conrad et al. 2004), in North Sinai (Amrar section) and Galilee (Ein Netofa area), confirms the Late Barremian age for the LFB Pl–Cd. This stratigraphical frame is extended by a sixth LFB (Os) occurring in North Sinai, which is characterized by the first occurrence (FO) of O. (M.) subconcava (Fig. 2). LFB (Os) ranges from uppermost Aptian to Albian and coincides with the FO of O. (M.) subconcava in the latest Aptian of the Tethyan realm (Castro et al. 2001; Schroeder & Neumann 1985) and in the Middle East (Simmons et al. 2000). The upper boundary of LFB (Os) is placed at the FO of Orbitolina (C.) corbarica in North Sinai (Bachmann et al. 2003), coinciding with the Albian/Cenomanian boundary. The orbitolinid-distribution of the Upper Barremian–Lower Aptian sections in Sinai is presented in Figure 3. Ranges of the LFBs for Golan Heights –Galilee and the younger sediments of Sinai are shown in Figs 3 and 4.

Ammonites Ammonites were sampled in North Sinai, at Gebel Raghawi (section A) and three sections 3 km east and west of section A (R1, K and JJ, Fig. 1c), which can be directly correlated by characteristic marker beds. Additionally, a few ammonites were sampled in the Rizan Aneiza area. The Upper Barremian –Lower Aptian of these sections contains rare heteromorph ammonites, including heteroceratids. One of these, Kutatissites bifurcatus (section R), is an index species for the Upper Barremian, but it ranges into the Lower Aptian (Aguado & Company 1997). Heteroceras coulleti, occurring at Gebel Raghawi, is known from the Upper Barremian of southern France (Delanoy 1997; Delanoy & Ebbo 2000). This

species occurs first in the uppermost Hemihoplites feraudianus Zone and ranges up to the lower Imerites giraudi Zone (middle Late Barremian, Fig. 2) (Delanoy 1997). Cheloniceras (Cheloniceras) and deshayesitids are the most common ammonites of the Lower Aptian in North Sinai. Cheloniceras (C.) cornuelianum and Cheloniceras (C.) cf. quadrarium (section JJ, Raghawi,) are widely distributed taxa. C. (C.) cornuelianum is known from the Boreal of southern England (Casey 1962) and from Spain, Colombia, Texas, Georgia and Turkmenistan. C. (C.) quadrarium is known from southern England and Colombia (Kotetishvili 1970; Young 1974; Lillo Bevia 1975; Sharikadze et al. 2004). These cheloniceratids indicate the Deshayesites deshayesi and Tropaeum bowerbanki zones of the Boreal zonation [sensu Casey et al. (1998), the T. bowerbanki Zone, corresponding to the D. furcata zone in the present paper, late Early Aptian, Fig. 2]. In the same interval Pseudohaploceras matheroni occurs, that is of limited use for biostratigraphy. The base of the Upper Aptian is characterized by an Aconeceras nisus found at northern Raghawi (sections R and A), a species that is also common in Russia and northern Germany (Kemper 1995; Mikhailova & Baraboshkin 2002). A. nisus is accompanied by Cheloniceras (Epicheloniceras) tschernyschewi, which is described from the Boreal in the equivalent of the Cheloniceras (Epicheloniceras) subnodosocostatum zone (early Late Aptian, Fig. 2).

Stable isotopes Carbon isotope values vary between – 1.0‰ and 3.5‰ (Fig. 3). While the Upper Barremian – lowermost Aptian values display a gradually increase from 0.5 to 3.0‰ (interrupted by a small drop only), the upper part of the Lower Aptian succession is characterized by a distinct drop (to – 1‰), a subsequent increase (to 3.5‰), and a drop (to 1‰) of the d13C values until the Lower/Upper Aptian boundary (Fig. 3). Although, the measured data show general lighter values than those from pelagic sections (e.g. Menegatti et al. 1998; Luciani et al. 2001; Weissert & Erba 2004) or shallow platform sections (e.g. Strasser et al. 2001; Wissler et al. 2003) in the northern Tethys, the general trend of our data is similar. The absolute values fit well to measurements from the southern Tethys (e.g. Oman: Immenhauser et al. 2005; Tunisia: Heldt et al. 2008) or the Pacific (Jenkyns & Wilson 1999; Price 2003). Menegatti et al. (1998) divided the Cismon carbon isotope stratigraphy curve for the Upper Barremian and the Aptian in eight segments (C1 –C8), widely used for subdividing Barremian –Aptian carbon isotope curves (e.g. de Gea et al. 2003;

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Fig. 3. Sections A (Gebel Raghawi), D (Rizan Aneiza), and GA (Gebel Amrar) reflect a transect across the Upper Barremian– Lower Aptian shallow shelf in North Sinai. The sections include lithology, d-13C isotope bulk rock values, the occurrence of biostratigraphic marker fossils, and the facies interpretation.

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Fig. 4. Correlation of the North Sinai and Galilee–Golan Heights succession. A North Sinai standard section and the Raghawi standard section was chosen, for Galilee–Golan Heights a composite section was drawn from the sections Har Ramin and Ein Quniya. Indicated are the biostratigraphic interpretation (based on the LFBs), the depositional environment, the interpretation of the second-order sequences, the main changes in detrital input, and ferruginous ooid content, as well as for the Galilee–Golan Heights area the main environmental interpretation and the position of the palaeocoastline.

Heimhofer et al. 2004; Immenhauser et al. 2005; Renard et al. 2005; Heldt et al. 2008). The segments are dated by planktic foraminifera and calcareous nanoplankton zones (Menegatti et al. 1998; de Gea et al. 2003; Fo¨llmi et al. 2006). The sections studied can be correlated with the segments as follows: the Upper Barremian interval, characterized by bulk carbon isotope values between 20.5‰ and 3.0‰, may correlate with segments C1 and C2. The strong decrease to values, around –0.1‰, and

a subsequent increase to 1.0‰, reflect the intervals C3–C4. A relative stable segment may correlate with C5, widely interpreted as reflecting the Oceanic Anoxic Event (OAE) 1a (Menegatti et al. 1998; Luciani et al. 2006). The following increasing carbon isotope values in Cismon (C6) are reflected by an increase to a value around 3.5‰. C7 is indicated again by stable carbon isotope values. The subsequent strong decrease in carbon isotope values reflects the change from segment C7 to C8.

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The Barremian/Aptian boundary in North Sinai coincides with a negative carbon isotope peak between C1 and C2 at 115 m (Raghawi section A, Fig. 3a), which correlates with an interval around the Barremian/Aptian boundary of Moullade et al. (1998) and Godet et al. (2006). Fo¨llmi et al. (2006) interpreted this peak as marking the Barremian/Aptian boundary. The equivalent of an OAE 1a is located at 160 to 175 m in the Raghawi section, which is in the Middle L. cabri/deshayesi zone according to Ogg (2004). The Lower/Upper Aptian boundary is suggested in a position just below the C7 –C8 interval boundary around 200 m in the same section (Fig. 3a).

Stratigraphical interpretation and correlation of the sections All data on stratigraphy are summarized in the regional correlation for the southern Levant Platform (Fig. 4). The first marine sediments are dated as Late Barremian by Kutatissites bifurcatus and Heteroceras coulleti at the lower part of Raghawi sections A, K and R in North Sinai and by Eopalorbitolina charollaizi in the Amrar section (GA, North Sinai) and Ein Netofa section (Galilee). The occurrence of P. lenticularis and C. decipiens in all sections indicate LFB Pl –Cd. The Barremian/Aptian boundary is characterized by the FO of the Lower Aptian index species Triploporella marsicana (Masse 1998) in Galilee/Golan Heights and is correlated with an interval above the last occurrence of H. couleti and K. bifurcatus in North Sinai. The boundary is possibly located at the negative carbon isotope peak at 115 m (Raghawi section A). This interpretation coincides widely with that of Abu-Zied (2007), based on small benthic and planktic foraminifers for the Raghawi area. The occurrence of P. lenticularis, C. decipiens, P. cormyi, P. wienandsi and O. (M.) lotzei allows the subdivision of the Lower Aptian into the LFBs Pl and Pc–Ol in both areas (Figs 3 & 4). While LFB Pl comprises the basal Lower Aptian, LFB Pc–Ol widely coincides with the negative isotope interval marking the OAE 1a in the upper Lower Aptian. In North Sinai, the Lower/Upper Aptian boundary is indicated by the FO of A. nisus [C. (E.) subnodosocostatum ammonite zone] and a decrease of d13C values (Raghawi, section A, Fig. 3). This is 35 m below the boundary defined by Abu-Zied (2007) on the base of planktic foraminifers and suggested by Morsi (2006) on the base of ostracods. The upper part of the basal Upper Aptian C. (E.) subnodosocostatum ammonite zone furthermore comprises the FO of O. (M.) parva (e.g. Rizan Aneiza sections D and RN) and C. (E.) tschernyschewi (Fig. 3). In the Raghawi area, LFB Op–Ot is indicated

only a few metres above the FO of O. (M.) parva (Fig. 3) indicating a late FO of O. (M.) parva in comparison to other studies (Lower/Upper Aptian boundary according to Arnaud-Vanneau 1998). In Galilee/Golan Heights, the Lower/Upper Aptian boundary is suggested to be 40 m below the FO of O. (M.) parva, according sequence stratigraphical and lithological correlations with North Sinai. The correlation of the Upper Aptian –Albian succession between both areas is more difficult because of the rare biostratigraphical markers in the Galilee– Golan Heights region, while the North Sinai succession is quite well dated. The Aptian/ Albian boundary in North Sinai is marked by two distinct ostracod assemblages and confirmed by graphic correlation (Bachmann et al. 2003). This boundary was interpreted to be within the Limestone Member (lower Rama formation) in the Golan Heights (Bachmann & Hirsch 2006). Based on the correlation of second-order sequences between Sinai and the Golan Heights (presented in this paper, Fig. 4) the boundary is expected to be located lower in the section, around the boundary Hidra/Rama Formation, which is conformed to ostracod data from Rosenfeld et al. (1995). In North Sinai, the Early/Middle Albian boundary is placed above the FO of Eoradiolites liratus (occurring since the latest Early Albian, Steuber & Bachmann 2002), by graphic correlation and the FO of Desmoceras (D.) latidorsatum (Bachmann et al. 2003), which first appears in the Late Albian (Gale et al. 1996). The Middle/Late Albian boundary is characterized by the LO of O. (M.) parva and the FO of N. simplex (Bachmann et al. 2003). Biostratigraphical data on the upper part of the succession in Golan–Galilee are generally rare and is based on ostracod assemblages or rare findings of the ammonite Knemiceras in Galilee (Rosenberg 1960; Rosenfeld et al. 1995).

Platform development Facies zones: platform environments The classification of facies zones (FZs) and the interpretation of the depositional environments are summarized in Figure 5, following the classification given by Bachmann & Hirsch (2006, Galilee – Golan Heights) and Bachmann & Kuss (1998, Sinai). The FZs range from deeper open platform to terrestrial and siliciclastic-influenced restricted platform. Nearly every FZ contains more than one facies type reflecting similar environments of deposition. Facies interpretations involve the comparison with other Barremian– Albian Tethyan platforms (Arnaud-Vanneau & Arnaud 1990; Masse 1993; Masse et al. 1995; Husinec 2001; Bernaus

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Fig. 5. Facies classification and interpretation on the base of components, composition and sedimentary structures.

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Fig. 5. (Continued)

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Fig. 5. (Continued)

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et al. 2003; Masse et al. 2003), and particularly with those in the southern Tethys (Pittet et al. 2002; Van Buchem et al. 2002; Immenhauser et al. 2005). Most FZs occur in both study areas and at different stratigraphical levels. Their distribution is indicated for three Upper Barremian– Aptian sections: section A (Fig. 6e & f) (supplemented by data from the sections K and R), sections GA and D (Figs 1c, 3a, b & 6d). To investigate the lateral facies variations in North Sinai during the Upper Aptian– Lower Albian (Fig. 7) sections D and MgE are added to data by Bachmann & Kuss (1998) and Bachmann et al. (2003). For the Galilee –Golan Heights the facies description summarizes data presented in Bachmann & Hirsch (2006). Two standard sections illustrate the facies variations between the platform settings at Sinai and Galilee– Golan Heights (Fig. 4). General trends for terrigenous input, formation of ferruginous ooids, main carbonatic components, and grain diversity are added. Lateral variations of FZs are recognized when correlating the different sections and allow an interpretation of the platform geometry for both areas studied during several time slices. Vertical variations indicate the changes of the platform environment associated with varying external and internal parameters (e.g. climate, sea-level change, sediment accumulation). The distribution of FZs characterizes major lithological, facies, and environment changes, leading to a subdivision of five PS (PS I to V) within the evolution of the Levant Platform. These PS are decoupled from the definition of formations. Facies development is described separately for both study areas, to focus on the regional geometric patterns.

The North Sinai platform stages PS I. PS I comprises the first Lower Cretaceous marine sediments and is characterized by limestones, marlstones, and siliciclastics deposited in varying depositional environments reaching from terrestrial to deeper marine. PS I comprises the lowermost Rizan Aneiza Formation (Upper Barremian– basal Lower Aptian, LBFs Pl –Cd, and lower Pl). An alternation of marl, siltstone, dolomite, and limestone belonging to various FZs (FZ 1 to 8, Fig. 5) in the northern sections A (Raghawi, 85 m thickness) and GA (Amrar, 95 m thickness, Figs 3 & 7) is interfingering with terrestrial sandstones (Malha Formation, FZ 9) 5 km south of Raghawi (Mansoura area – Fig. 1c). At section A (Figs 1c & 3a) PS I is dominated by protected lagoonal environments, characterized by bioclastic wacke and packstones, with benthic foraminifers, shell and echinoderm debris and

oncoids (FZ 6), partly rich in orbitolinids and cyanophyceans. These limestones alternate with marlstones, silty marlstones, and siltstones containing few autochthonous benthic foraminifers, ostracods, small solitary coral colonies, echinoderm debris, and further bioclasts (FZ 6-2) that are indicating a varying terrestrial input. Several levels contain ferruginous ooids. In the lower part, intercalated peritidal to coastal sandstones and siltstones (FZ 8a) indicate two regressive events. In the same interval, ammonite-bearing marlstones indicate a sporadic influence of open-marine conditions. Among those ammonites are several phylloceratids (e.g. Euphylloceras aff. inflatum), which are interpreted as reflecting deeper-marine conditions (Westermann 1996). Intercalations of oolithic, intraclastic and bioclastic grainstones (FZ 2) mark the repeated influence of shallow subtidal, high-energy conditions. Upward increasing, intercalated limestones rich in oncoids (FZ 3, mainly MFT 3a) reflect increasing open platform shallow subtidal conditions. In the uppermost part of PS I frequent intercalated dolomites (early diagenetic dolomite formation), are suggested as representing short events of emergence. The upper boundary of PS I is settled at the point of increasing carbonate content and marked by the last occurrence of the ferruginous ooids and the first occurrence of dasycladacean-rich open lagoons (FZ 4) above. At section GA (Figs 1c & 3a) the succession is dominated by oncoid-rich limestones of FZ 3, showing a trend to more open lagoonal/platform conditions. Siliciclastic input is less frequent and of finer grain size, while ferruginous ooids occur only rarely. Deeper platform environments (FZ 1), comparable to the ammonite bearing beds at section A, occur only in the upper part of PS I at section GA. About 5 km towards the south, terrestrial sediments characterize PS I and indicate that the coastline was located north of the Gebel Mansour (Fig. 8a, f). The proximal position of section A is reflected by intercalated coastal sediments and high contents of clay and ferruginous ooids. Further north, at the distal section GA, the openmarine facies lacks coastal influence. However, depositional environments are partly shallower than in the proximal SW, suggesting deposition on a submarine high, separating the protected area at section A from open-marine environment. The open-marine deeper environments are expected to be in a close position to the study area because of the occurrence of phylloceratids. PS II. During PS II open platform environments in the north are interfingering with terrestrial sediments in the south.

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

The main part of the Lower Aptian sediments and the first metres of the lowermost Upper Aptian (lower LFB Op-zone/lower C. (E.) subnodosocostatum zone), belong to PS II. Marine limestones, marlstones, siltstones, and sandstones of PS II crop out at section GA (90 m thickness, Fig. 3b) and at section A (110 m thickness, Figs 3a & 6e, f), while at Gebel Mansour further to the south, only terrestrial siliciclastics occur. PS II starts with shallow-subtidal openmarine environments containing diverse bioclasts and dasycladaceans (section A, FZs 3 and 4). Intercalated near-shore sandstones (FZ 8) and a remarkable red-brownish dolomite horizon (early diagenetic dolomite formation, Fig. 6e) are suggested as representing events of regression and emergence. Above, marlstones containing some ammonites (FZ 1) are overlain by limestones rich in calcareous algae (FZ 4), limestones with diverse biota and planktic foraminifers (FZ 1a) and limestones characterized by high amounts of orbitolinids (FZ 1b). About 40 m of marlstones with corals at the base and ammonites and planktic foraminifers (FZ 1) are marked by upward decreasing carbonate content. About 4 m of intercalated grainstones with ferruginous ooids (FZ 2), indicate a short event of shoaling. Maximal 10 m of marly claystones of FZ 1 form the top of this interval. The upper boundary of PS II is formed by an erosional unconformity, which cuts into the upper marly claystone at section A and into the ferruginous-ooid interval below E of the section A. At Gebel Amrar limestones deposited in high-energy shoal environments dominate the lower part of PS II, while silty marls of open-marine and lagoonal environments (FZ 1 and 6-2) occur in the upper part. Generally PS II is characterized by subtidal open-platform environments with upward increasing influence of open-marine conditions. The coarse grained terrestrial input is strongly reduced in the upper part and ferruginous ooids are missing until a short shoaling interval, allowing the deposition of near coast sediments with ferruginous ooids, at Gebel Raghawi. Its southward proximal regional extension is very similar to PS I. Terrestrial sediments crop out in the entire southern region and indicate a widely unchanged position of the coastline between the northern Raghawi and Mansour area (Fig. 1). The only exception is one marly bed with corals, which was observed in a single locality at southern Gebel Maghara close to section A (Fig. 7). Because of the erosive nature of PS III above, a now eroded uppermost part of PS II, ranging further to the south, cannot be excluded. The deposition of high-energy shoals at section GA with deeper sediments behind makes the existence of a barrier in that area highly probable.

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PS III. Sandstones deposited in terrestrial and coastal environments characterize PS III. This interval is attributed to the lowermost Upper Aptian (subnodocostatum zone; Fig. 2). At section A (most proximal section) the interval consists of only 10 m of caolinitic quartzose sandstones (Fig. 6e), which are interpreted as terrestrial deposits (FZ 9). The erosional base of the succession cuts into different levels of the underlying succession. At section D (Fig. 1c) 40 m of coastal siliciclastic sediments (FZ 8a) mark that interval, which are two times interrupted by lagoonal and shoal originated carbonates (FZ 4 and 2). At section GA the interval is formed of 10 m of tidal flat dolomites (Fig. 3b). Compared to PS I and II, the coastline moved to the north and was now located between sections A and GA, close to sections RN and D (Fig. 1). Section A comprises the terrestrial realm, while sections RN and D indicate a near shore position. The more distal section GA was also partly emerged, but not affected by siliciclastic deposition, speaking for a local high/swell (Figs 7 & 8a). PS IV. Shallow subtidal marine limestones, deltaic siliciclastics, and a southward transgression mark this platform stage. PS IV comprises the upper Rizan Aneiza Formation, which was deposited during the Late Aptian until the end of the Middle Albian (Bachmann et al. 2003). Marine sediments of this interval were documented from sections D (Fig. 3a) and RN (240 m, Fig. 7), section R (225 m, Fig. 4), sections ME and M (190 m, Figs 7 & 6h), and section MgE (140 m, Fig. 7). In the northern and central distal sections, the marine sedimentation starts in the upper C. (E.) subnodocostatum zone (sections C, RN and R), while in the proximal southeastern area the first marine sediments are dated as Lower Albian (sections M, ME; Bachmann et al. 2003; and section MgE, Fig. 7). At the northern section RN marlstone –limestone alternations of the shallowsubtidal FZ 3 to protected lagoon FZ 6 with some intercalated small-scaled rudist biostromes (FZ 2-2), bioclastic to oolithic shoals (FZ 2), and some meters of open-marine sediments (FZ 1) prevail. The Raghawi– Mansour area was characterized by a delta system with sandstones, siltstones, and marlstones of FZ 8, prograding repeatedly into the area studied and alternating with shallow subtidal, openmarine (FZ 3), and lagoonal sediments (FZs 4, 6 and 7; Bachmann & Kuss, 1998) and furthermore high energy shoals (FZ 2, Fig. 5). The amount and frequency of fine and coarse grained siliciclastic intercalations increase to the SE until they dominate the succession at section MgE (Fig. 7). Up to 4 m thick rudist biostromes (FZ 2-2) with

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Fig. 6. Facies and bedding-patterns of the Upper Barremian– Albian strata. (a) Nabi Said Fm. The Har Ramin section, Galilee, comprises the entire Lower Aptian– Middle Albian succession. (b) Ein El Assad Fm. In the Ein Netofa section (Galilee), the cross-bedded limestones of the Upper Barremian Nabi Said Formation are well developed. (c) Hidra Fm. Well-bedded limestones mark the Ein El Assad Formation in the Har Ramin section. (d) Rama Fm. The Rizan Aneiza

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Fig. 7. Correlation of the sections in North Sinai summarizing new and known sections (Bachmann et al. 2003). The sections are correlated along marker horizons, sequence boundaries, and the biostratigraphic frame. Platform stages are indicated by different colours. The location of the sections is shown on the small map.

lateral extensions of tens of metres are intercalated in the succession at sections RN, D, A and ME (Figs 7 & 6h). At section A, an up to 4 m-thick coral biostrome (FZ 2-2a) is intercalated in the lower part of PS IV, while the entire succession is marked by high amounts of ferruginous ooids, which decrease in frequency and amounts distally (section RN, Fig. 7). Intercalated emergence

horizons are common and marked by rhizolithes or ferruginous crusts that can be correlated over large areas in the region (Bachmann & Kuss 1998). The facies evolution of PS IV reflects third-order sealevel changes (Bachmann & Kuss 1998). All the described sections of PS IV follow a proximal–distal transect across the shallow shelf (Fig. 8). A retrogradation of facies belts and

Fig. 6. (Continued) Formation at its type-locality is characterized by marly and sandy limestones (section D). (e) Lower part of the Raghawi Section A: coastal sandstones marlstones, and upward increasing dolomites and limestone mark the Upper Barremian and Lower Aptian platform stages PS I and PS II (Rizan Aneiza Formation, Gebel Raghawi). The arrow indicates a prominent dolomitic horizon 150 m above the base of the section, which is interpreted as emergence surface. (f) At northern Gebel Raghawi the entire Barremian– Middle Albian succession crops out and is bounded in its upper part by a major reverse fault. The arrow indicates the terrestrial sandstones of PS II shortly above the Lower– Upper Aptian boundary. The succession above the sandstone shows the Upper–Middle Albian PS IV marked by increasing limestone content (Upper Rizan Aneiza Fm). (g) The figure comprises nodular limestones, marlstone, and the sandstone (arrow) of the uppermost Lower Aptian to lowermost Upper Apian. (h) The Upper Aptian to Lower Cenomanian succession at Gebel Mansoura is characterized by deltaic sediments in its lower part and pure limestones with huge rudist biostromes (arrow) in the upper part.

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Fig. 8. Schematic transect through the North Sinai shallow shelf displays the Late Barremian– Aptian evolution and involves the factors extension along normal faults and sea level change (a –e). Palaeoenvironmental maps indicate the lateral extension of facies zones (f–h). (a, b) PS I and PS II are characterized by increasing sea level. (f) The southern coastline was located at the Raghawi–Mansoura normal fault, with a possible short-termed southward transgression

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

transgression onto former terrestrial areas is obvious (Fig. 8). The coastline was located between Gebels Raghawi and Mansour during Late Aptian times, between Mansour and Maghara East around the Aptian/Albian boundary, south of Maghara area during Early and early Middle Albian times, and more than 50 km further south during the late Middle Albian (Bachmann & Kuss 1998). This stepwise retrogradation is also indicated by a southward decreasing thickness, which is mainly owing to a later onset of sedimentation (Fig. 8). PS V. PS V is characterized by relatively uniform deposition of limestones in the entire region. PS V comprises the lower part of the Halal Formation (Late Albian –Early Cenomanian, Bachmann et al. 2003). At Rizan Aneiza and Raghawi only the lower part of PS V crop out (section RN: 40 m, section R: 145 m, Fig. 4), but at Mansour (section M, Fig. 6h) and Maghara SE (section MgE) 200 m of sediment occur (Fig. 7). The base of PS V is characterized by an up to 20 m thick rudist biostrome (FZ 2-2, sections RN, R, M, Figs 4 & 7), which is interfingering southward (section MgE) with lagoonal limestones (FZ 6a). In all localities, a sudden stop of the siliciclastic input occurs below this bed. Overlying this, limestone– marlstone alternations of FZs 5, 6 and 7 prevail. Marly lagoonal environments characterize PS V with high amounts of gastropods and orbitolinids (FZ 6-2) and the repeated formation of small and larger rudist biostromes (FZ 2-2) of protected lagoons, as well as stromatolithic tidal flats (FZ 7) above. Intercalated oolithic and bioclastic shoals (FZ 2) indicate high-energy facies. Open-marine subtidal facies are rare, while deeper platform facies types are missing. Repeated emergence horizons characterized by rhizolithes, tepee structures, dolomitic, and in the lowermost part by carstification are intercalated and reflect sequence boundaries of the third-order sequences (Bachmann & Kuss 1998). The increasing abundance of dolomitized tidal flat deposits (FZ 7) in the upper part of PS V indicates a general shoaling of the depositional area. During PS V, the main position of the coastline was located clearly south of the investigated area (Bachmann & Kuss 1998). Decrease in siliciclastic input and short-termed sea-level lowstands produced emergence in contrast to the prograding delta wedges of former periods. Thickness of the

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sediments is clearly higher in the respective distal sections. The lateral distribution of FZs indicates a very similar, shallow, protected marine environment covering large parts of the North Sinai platform (Fig. 7).

The Galilee –Golan Heights platform stages The platform stages PS I –V are well documented at Galilee –Golan Heights and allow the correlation of four sections (Figs 1b & 9) with North Sinai. These sections represent a transect across the shallow shelf of the northeastern Levant platform located on both sides of the DST fault. The marine succession starts above terrestrial sandstones of the Hatira Formation. PS I. In northern Israel/the Golan Heights PS I comprises the first marine Lower Cretaceous sediments similar to North Sinai. Limestones deposited in shallow subtidal environments predominate. PS I correspond with the Nabi Said Formation, which is of Late Barremian to basal Early Aptian age (LBFs Pl– Cd and lower Pl) (Fig. 4). In central Galilee, they are well exposed in the Ein Netofa section, but at the Golan Heights only a few metres of basal Lower Aptian occur (Bachmann & Hirsch 2006). The Ein Netofa section is characterized by about 32 m of cross-bedded oolithic and bioclastic grainstones, partly enriched in ferruginous ooids and quartz (FZ 2). They increasingly alternate upward with bioclastic packstones (FZ 3) and wackestones reflecting protected environments (FZ 4, 5 and 6). At the Golan Heights 10 m of sediment were deposited in protected lagoons with tidal flats (FZ 4 and 7). During the Barremian, an open-platform highenergy facies belt (Ein Netofa section) was interfingering to the east and SE with terrestrial sediments (Hatira Formation) in the region between Galilee and Golan (Hirsch 1996). The abundance of ferruginous ooids and coarse grained quartz indicates the vicinity of the proximity to the coastal area, which retrograded during the Late Barremian– earliest Aptian to an area east of the Golan Heights. A slightly inclining ramp geometry is documented by a gradient from high- and low-energy open subtidal (FZ 2 and 3) facies occurring in the Ein Netofa section to protected subtidal environments (FZ 3 and 4) at Har Ramin, and intertidal facies-types (FZ 7) at the Golan Heights.

Fig. 8. (Continued) at the end of PS II. (c) A significant drop in sea level caused emergence of wide areas and erosion of the former relief. (d) During PS IV, a homoclinal ramp characterizes the depositional architecture. A delta system developed in the proximal areas, interfingering with lagoonal and shallow ramp sediments. (g) The strike of the facies belts changes from SW– NE to WSW–ENE. (h) During Late Albian drown by rising sea-level, a shallow platform without major changes developed.

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Fig. 9. Six models explaining the development of the depositional architecture of the Levant Platform in North Sinai and Galilee–Golan Heights area.

PS II. Limestones upward alternating with marlstones characterize the marine sediments of PS II. PS II comprises the Ein El Assad Formation and the lower part of the Hidra Formation, corresponding to the Lower Aptian (upper LBFs Pl and Pc–Ol, Fig. 2) – lowermost Upper Aptian. At the boundary between PS I and PS II terrigenous input was interrupted and the lower part of PS II (Ein El Assad Formation) forms a prominent landmark of 55 m-thick pure limestones in Galilee and the Golan Heights. The upper part (lower Hidra Formation) crops out in the Golan Heights and is known from the subsurface only in the Galilee (Rosenfeld et al. 1998). Bioclastic packstones and wackestones of the Lower Ein El Assad Formation belong mainly to the low-energy open and protected lagoon FZs 3, 5 and 6 with high amounts of calcareous green algae in some layers (FZ 4). Intercalated are a few beds deposited under high-energy (FZ 2), restricted or tidal flat conditions (FZ 7). In the upper part of the Ein El Assad Formation increasing

open marine conditions are indicated by the occurrence orbitolinids-rich and planktic foraminiferscontaining sediments of FZ 1 in all sections. Relatively uniform depositional conditions prevailed in the entire area with an only small gradient to more protected environments from west to east characterizing a homogenous shallow-platform environment reaching to an area west of the Golan area. Decreasing siliciclastic input of quartz and ferruginous ooids indicate greater distance of the shoreline. The upper part of the Ein El Assad Formation (LFB Pc –Ol) is characterized by a significant synchronous deepening event establishing deeper marine environments over the entire Galilee – Golan Heights region (Fig. 4). The onset of the Hidra Formation marks increasing terrigenous input. The siltstones, marlstone and limestone alternations were deposited in open-marine subtidal (FZ 3, 4) to protected environments (FZ 6 and 8). In the higher part of PS II shoaling and progradation of the coast line occur (Fig. 4).

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

PS III. Siliciclastic sediments characterize PS III in northern Israel. PS III comprises 5 m of the middle part of the Hidra Formation, marked by fluvial sandstones of FZ 9 that can be not dated exactly. The underlying PS II extends at least into the late Early Aptian and the overlaying PS VI is clearly related to the lower Upper Aptian. Thus, PS III is not older than late Early Aptian and not younger than early Late Aptian. PS III represents an important emergence horizon. A terrestrial facies can be proved from the Golan Heights, documenting the progradation of the coast line. Rosenfeld et al. (1998) described similar siliciclastic sediments of the Hidra Formation from the subsurface of the Galilee area, however without emergence. They describe a brackish facies from the Middle Hidra Formation indicating at least a near coast position for that region. PS IV. PS IV comprises about an 150 m thick succession of marine marlstones and limestones. It corresponds to the Upper Aptian to Middle Albian (LBFs Op and Ot, possibly Os) Rama Formation and is documented from the Golan Heights (Ein Quniya section) and from Galilee (Har Ramin section, Fig. 6a). The 45 m-thick limestone member forming the base of the lower Rama Formation (35 m at the Golan Heights) is characterized by often nodular bedded limestone deposited in the various FZs, reaching from openmarine deeper platform (FZ 1; only at Har Ramin) to tidal flat sediments (FZ 7) with a maximum occurrence of the open-marine shallow platform environments (FZ 3). Terrestrial input is low and rarely ferruginous ooids occur. The upper part of the Rama Formation is composed of limestone – marlstone alternations, mainly deposited in protected environments (FZs 5, 6 and 7) in Galilee and the Golan. Open lagoonal environments (FZ 3) occur only at Har Ramin. Slightly higher siliciclastic input of sand and marlstones occurs in the Lower Albian. During deposition of the Rama Formation, the marine facies belt reaches again to an area significantly west of the Golan area. However, siliciclastic input of quartz and ferruginous ooids is still obvious. The facies indicates that a low-energy lagoon dominated the entire Formation, reaching from Golan Heights to Galilee, and limited to the west by fringing rudist bioherms at the Carmel area in the western Galilee (described by Sass & Bein 1982), and to the east by terrestrial deposits in Jordan. PS V. PS V includes dolomites of the Yagur Formation and was analysed only at Har Ramin (Galilee), where pure dolomites reflect protected lagoon (FZ 6) and tidal flat deposition (FZ 7). According to Rosenfeld et al. (1998) a late Middle

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Albian can be assumed. Similar sediments are described from cores in Galilee, from the Golan (Hirsch 1996), and from southern Israel (Rosenfeld et al. 1998; Rosenfeld & Hirsch 2005) indicating an enlarged shallow platform covering Israel.

Structural control on the Late Barremian – Albian deposits of the Levant Platform The North Sinai record. Deposition of Early Cretaceous strata took place between Late Jurassic extensional and Late Cretaceous compressional tectonics. While the Late Jurassic extension resulted in predominantly NW dipping extensional normal faults (Moustafa & Khalil 1990; Garfunkel 1998), the Late Cretaceous compressional stage resulted in three fold ranges in North Sinai, owing to the inversion of the former extensional faults. The interpretation of the Late Barremian –Albian sedimentary record and the arrangement of facies belts in the palaeogeographical maps allow the reconstruction of extensional movements along these faults during Early Cretaceous times. Facies zones are generally orientated in WSW– ENE trending belts parallel to these major extensional faults (Figs 8 & 9). The prominent anticlinal structure of Gebel Maghara lies at the northernmost fold range, in between east –NE elongated belts of right-stepping en-echelon folds, interpreted as representing the strike –slip rejuvenation of deep-seated earlier extensional faults (Moustafa & Khalil 1990). Those reverse faults were observed at Gebel Amrar and Gebel Raghawi (Fig. 8g; compare Moustafa 2010), where abrupt facies and thickness changes of Early Cretaceous sediments across these faults indicate their importance during earlier extensional stages and might have been active until the early Late Aptian. The southern SW –NE trending fault (Raghawi –Mansour fault) separates the Raghawi and the Mansour areas with its lateral elongation further NE, at Rizan Aneiza. The Amrar section (Fig. 8f ) is located at a second extensional SW –NE trending fault, again rejuvenated during the Upper Cretaceous compression (Amrar fold belt sensu, Moustafa & Khalil 1990). Two similar sets of faults with vertical movement were also described by Moustafa (2010). The following sedimentological and stratigraphical observations suggest a structural control on Early Cretaceous deposition along normal faults of North Sinai. (1) The Upper Barremian to lower Upper Aptian (PS I –II) sections at the neighbouring Gebels Raghawi and Mansour (located c. 5 km apart) show significantly different sedimentary environments (Fig. 7). In the Mansour area, sedimentation took place under terrestrial

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conditions, while more than 240 m of marly and limy sediments were deposited in marine environments at Gebel Raghawi. To interpret this record we either have to assume a primary, very steep, declining coast (with vertical differences of more than 200 m along a 5 km horizontal distance), or more probable, an active normal fault that resulted in higher subsidence rates in the northern areas. At the end of PS II (latest Early Aptian and earliest Late Aptian) water-depth at the Gebel Raghawi area increased significantly in contrast to still terrestrial environments at Gebel Mansour, and may thus reflect different subsidence owing to syndepositional tectonic activities along a normal fault. (Alternative interpretations of deposition and later erosion of marine sediments at Gebel Mansour seem unlikely because of missing reworking horizons). Our facies data hint on a SE – NW striking normal fault active during the Late Barremian–earliest Late Aptian (and reactivated in Late Cretaceous times), which caused the different depositional environments (Fig. 8f ). At Gebel Amrar, further to the NW, shallow marine Upper Barremian– lower Upper Aptian deposits indicate a swell, which may be correlated with another active normal fault belonging to the Amrar fold belt (Moustafa & Khalil 1990).

During PS III (early Late Aptian) terrestrial sedimentation occurred in the entire northern Sinai, which was accompanied by erosion in the southern sections (Fig. 8c). During PS IV (latest Aptian – Middle Albian) stepwise encroaching of the marine sediments from north–NW to south–SE indicates a slightly dipping ramp (Fig. 8d ). Uniform sediments at Raghawi and Mansour characterize the next stage (PS V), with some beds traceable for long distances (Fig. 8e). This means that the Raghawi –Mansour fault was active only until the end of the Early Aptian, possibly until the earliest Late Aptian. During the younger stages (PS III) erosion and deposition reduced the former elevation gradient along the normal faults, until more homogenous depositional environments became obvious during PS IV (Fig. 8). The sedimentological and stratigraphical data suggest extensional tectonical activity during the Barremian–Early Aptian, which is younger than described before (Garfunkel 1998; Moustafa & Khalil 1990; Moustafa et al. 1990). Only since the Late Aptian, subsequent to tectonical activity, a stepwise retrogradation took place and the facies development was controlled only by global sealevel change and supraregional tectonics.

The Galilee– Golan Heights structural development. The northern Israel Barremian –Albian structural development is much simpler as in North Sinai. During the Late Barremian, flooding of the terrestrial area created a slightly inclining ramp structure. Owing to different depositional rates, a landward-thinning wedge of sediments (described by Rosenfeld et al. 1998) filled the available accommodation space until a shallow, uniform platform developed during the Early Aptian (Bachmann & Hirsch 2006). Distally steepening of the ramp and a shelf break is suggested for the Carmel area, where rudist bioherms fringe the shelf-break and are closely allied to deeper marine sediments during the Albian period and deeper marine sediments were deposited nearby (Bein 1971; Sass & Bein 1982). This is confirmed by seismic profiles (Garfunkel 1998). In summary, we suggest a depositional regime, which was mainly controlled by sedimentation rates, sediment production, and sealevel changes without evident tectonic influence. Facies models for the Galilee –Golan Heights area include a post-depositional, Oligocene–recent sinistral strike–slip movement along the Jordan– Gulf of Aqaba line (Fig. 1b, e.g. Garfunkel & Ben-Avraham 1996; Flexer et al. 2005; Mart et al. 2005). The rate of sinistral displacement is controversially discussed, reaching from 10 km (Mart et al. 2005) to 110 km (Garfunkel & Ben-Avraham 1996; Gilat 2005). Our sections show no evidence for strong lateral displacement since displacement was parallel to facies belt boundaries.

Second-order sequences Second-order sequences were already discussed for the Galilee –Golan area and for the upper part of the North Sinai succession (Bachmann et al. 2003; Bachmann & Hirsch 2006). Comparing both platform settings allow us to recognize second-order sequences that have controlled a large part of the Levant Platform. We adopt three second-order sequences, formerly described from the Galilee –Golan Heights area and prefixed MC (mid-Cretaceous) EL (eastern Levant) and numbered (e.g. MCEL-1). For the larger extension of the analysed region, we transfer the prefix into MC (mid-Cretaceous) L (Levant) and add another sequence (Fig. 10).

MCL-1 In both areas, terrestrial sediments were flooded in the Late Barremian. For both areas, we document initial deepening and increasing accommodation during the Late Barremian –earliest Aptian, and at Galilee –Golan landward retrogradational patterns characterize the transgressive systems tract (TST).

Levant platform

AlSin8

K100 (101) SB MCL 4

SB 103

M. inflatum

H. dentatus

L. tardefurcata

Al 3 Al 2 Al 1

H. jacobi

Ap6

N. nolani

Late

Aptian Early

D. deshayesi D. weissi

AlSin1

TST

TST

SB 109.5

SB MCL 3

HST

HST

K90 (111) Late Aptian unconformity

ApSin3

SB 113

Op ApSin2

TST

TST K80 (117)

Op O.(.M.) parva

Pc-Ol

P. waagenoides C. sarasini I. giraudi H. feraudianus G. sartousi A. vandenheckii

ApSin1 Ap4

P. cormyi O.(.M.) lotzei

Pl

D. oglanlensis

Late

AlSin2

SB 107

P. melchioris

D. furcata

Barremian

K100 (106)

HST

SB

O.(.M.) parva O.(.M.) texana

E. subnodosocostatum

125

Al 4

D. mammillatum

115

120

Os O.(.M.) subconcava

Al 6 Al 5

AlSin7 AlSin6 AlSin5 AlSin4

Pl-Cd P. lenticularis C. decipiens

ApEl4 Ap3

ApEl1

Ap2 Ap1

BaEl2

Barr6

LST

SB MCL 2

HST

LST

(120.1) K70 (122.5)

HST

SB (125)

BaEl1

TST

Emergence

SB

III

E. loricatus

HST

Al 7

TST

SB (127)

OAE1a

K60 (126)

I

Middle Early

Albian

Early Cretaceous

110

Al 8

E. lautus

IV

105

Levant platform

arid

arid

humid

Europe Price, 1999

humid

Arabian Platform

AlSin9

Tectonics

fault movement/extensional tectonics (this work) Moustafa and Khalil, 1990

Late

Al 9

Israel

terrigenous input and ferruginous ooids orbitolinid beds

Al 10

Sinai

karstification at 3rd-order SBs tepees, stromatolithes

Al 11

Sinai Israel

Sharland Bachmann Bachmann modified from et al. 2004 et al. 2003 et al. 2003 Bachmann & Hirsch 2006 recalibrated by Bachmann & this work Haq & AlHirsch 2006 Qahtani 2005

V

Hardenbol et al.1998

Climate

Sinai: ramp-platform Israel: fringed platform

Global or Tethyal

121

Onlap curve, crisis & platform-types

Sinai: deltaic ramp Israel: carbonate platform

Bachmann & Hirsch 2006, 2010

nd

2 -order sequences

II

Gradstein et al. 2004

100

Sea level rd

3 -order sequences

Sinai: sub-basin Sinai: sub-basin Israel: siliciclastic ramp Israel: platform

Ammonites

Stage

Substage

Epoch

Chronostratigraphy

larger benthic foraminifer zones

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

Fig. 10. Chronostratigraphic chart summarizing the stratigraphic and sequence stratigraphic interpretation, as well as the climate and tectonic data, in comparison with data on the Arabian Plate and the regions around the Tethys.

These retrogradation patterns were retarded in North Sinai, when the coastline reached the Raghawi area by the activity of the Raghawi – Mansour fault (Figs 7 & 8), while westward retrogradation crossed the Golan Heights without barrier (Fig. 4). The TST, characterized by high siliciclastic input and ferruginous ooid, was deposited under more or less open marine locally high-energy conditions (Fig. 4). A major change to low-energy inner platform conditions in the basal Lower Aptian marks the maximum-flooding surface (MFS) at the Galilee –Golan Heights. The early highstand systems tract (HST) comprises limestones deposited under lagoonal to open-marine conditions and the late HST more protected environments with increasing siliciclastic input (Fig. 4). No former southward transgression is documented and the creation of accommodation space is clearly reduced. In North Sinai, a steady coast line and a decrease in terrigenous input characterize the HST with a depositional environment clearly changing to open marine, deeper-shelf environments. However, no reduction of accommodation is documented, before the late HST, because of subsidence owing to local fault activity. In both areas an exceptional

third-order deepening event is observed superimposing the second-order HST, which largely correlates with the platform drowning observed at the Oceanic Anoxic Event 1a (Weissert et al. 1998; Bachmann & Hirsch 2006).

MCL-2 The sequence boundary (SB) of the MCL-2 is the most prominent emergence horizon in the entire succession, indicated by terrestrial sandstones which eroded into the underlying sediments (Sinai only, Fig. 4). For North Sinai the age of the SB can be dated as being earliest Late Aptian (Figs 3 & 4). In Israel it is recognized only in the Golan Heights, where its dating is unsure. Similarly pronounced is the transgressive surface (TS), when retrogradation of the facies belts is marked by the renewed marine sedimentation in both areas. During the TST, both areas are marked by increasing accommodation, retrograde submergence of the platform, and siliciclastic input (reduced in the late TST) in dominantly protected environments. Slightly higher accumulation at the Golan Heights, compared to North Sinai, may result from earlier

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onset of the marine deposition. While at North Sinai the TS is dated as early Late Aptian [C. (E.) subnodosocostatum zone], the age is not confirmed yet for the Golan Heights. The basal Upper Aptian MFS is documented in the Golan area by the deepest environment in the study area, indicated by a sudden reduction in terrigenous input and followed by HST carbonates. The mfs in the North Sinai is interpreted as lying at a surface of re-establishment of carbonate dominated sediments (Raghawi and Rizan Aneiza sections), which coincides with a decrease of the southward transgression, indicated by interfingering of high-energy shoals at Raghawi with sediments of the proximal delta at Gebel Mansour. The depositional rate of MCL-2 is slightly reduced compared to MCL-1, while the observed transgression to the south, especially at North Sinai, is much greater (Fig. 8g). However, observed depositional environments, which are generally characterized by shallower water-depth during MCL-2 indicates a general lower accommodation increase than before.

MCL-3 The Lower Albian SB MCL-3 is clearly defined at the Golan Heights (Ein Quniya section) and at Galilee (Har Ramin section) by a facies change from open marine to restricted environments (Fig. 4). At North Sinai, it is interpreted as lying below a distinct zone of delta progradation. The LST is indicated by terrigenous delta-influenced sediments in the northern distal sections and by coarse-grained sand dominated delta sediments in the proximal section M. The TSs are marked in both areas by deepening of the environment and by a clear trend to high-energy setting in the individual sections indicating increasing accommodation (Fig. 4). In North Sinai, prominent transgression during the TST results in submergence of the southern and eastern part of the Maghara area (section MgE, Fig. 8h). At Galilee the MFS is indicated by a significant change from open marine to protected environments under constant accommodation. At North Sinai, a change from consequent transgression to a succession with higher-frequency facies changes may result from reduced accommodation and mark the late Early Albian MFS. The HST in both areas is characterized by repeated siliciclastic input, but also by the development of rudist biostromes (North Sinai) or rudist debris fearing sediments (Galilee –Golan Heights).

MCL-4 The SB MCL-4 is marked by emergence in North Sinai and by facies change at Galilee (Fig. 4). The

Late Albian age confirmed at North Sinai may correlate with Galilee, where an accurate dating of the SB is not possible. In both areas, the LST contains few beds of intertidal dolomite. While only few intercalations of protected and open platform sediments mark the TST at Galilee, a succession with shoals, protected lagoons, and rudist biostromes dominate the TST at North Sinai. In the latter area the maximum flooding is characterized by the establishment of tidal flat sediments on large parts of the platform in the uppermost Upper Aptian. The uppermost part of the sequence is present in the Mansour north sections only; here the subsequent SB is of Early Cenomanian age.

Discussion Platform development – geometry and facies development as a consequence of tectonics, siliciclastic input, production, and sea-level change. The five platform-stages (I –V) reflect different depositional environments, accumulation patterns, and platform geometries (Fig. 9). We observe significant variations in one or more of the following parameters: tectonical activity, siliciclastic input (driven by climate and tectonics), carbonate production, and second-order sea-level. To chart the depositional models for the Levant Platform stages (Fig. 9) published subsurface data were incorporated, spanning the region between Galilee – Golan Heights and North Sinai. PS I: Late Barremian–earliest Early Aptian. In both areas siliciclastic influenced carbonate sediments characterize the near shore environments, rich in ferruginous ooids and quartz (Fig. 9a). Marly sediments, with upward increasing carbonate content, are described from the central Israel hinge belt. A shallow water depositional environment was defined on the base of ostracod assemblages, with intercalated fresh-water signals, observed in central and northern Israel (Rosenfeld et al. 1998). A strongly increasing sea level (TST MCL-1) resulted in transgression. Different subsidence, basic geometric regimes, and tectonical influence resulted in different depositional regimes in Sinai and Israel. At North Sinai, low-energy sub-basins developed between active normal faults, which caused enhanced subsidence rates, a stable coast line, and phases of higher siliciclastic input owing to higher-frequency sea-level changes. Sedimentological and tectonical data additionally indicate the existence of a swell in the northernmost part of the Sinai (Gebel Ambra region, Fig. 8) causing a lowenergy depositional regime south of the swell with shallow and deeper lagoonal environments interrupted by sporadical coastal progradation. Those

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

features result in small-scaled variation of the sedimentation rates, with high sediment accumulation in front of the normal fault at Gebel Raghawi and low in the area south. We suggest a shelf break situated only a few kilometres north of the studied area, possibly linked with normal faults further north. This fits with deeper marine marls characterizing the offshore area only 20 km north of the studied sections at Rizan Aneiza (Martinotti 1993) and the occurrence of phylloceratid ammonites in the Lower Aptian sediments probably suggesting that deeper marine conditions were present in an area close to the studied ramp sediments (Lehmann et al. 2007). However, an odd taphonomic history has been suggested for this group of ammonites, with surfacing and floating of shells first that is followed by resinking by cameral puncture, and therefore a substantiation of an autochthonous origin can be complex (Maeda & Seilacher 1996). In northern Israel– Galilee, sediments form an eastward thinning wedge (Rosenfeld et al. 1998; Rosenfeld & Hirsch 2005), with more restricted environments in the eastern part (Fig. 9a). There is no sign of active faulting and sea-level rise result in gradual flooding of the former terrestrial sediments, with a coastline moving from the Galilee to east of the Golan Heights (Bachmann & Hirsch 2006). At Galilee a thick succession of high-energy grainstones indicate increasing accommodation. Gradual transgression and the landward increasing protection of the environments suggest a shallow ramp geometry with a huge high-energy facies belt. During the entire succession a continuous input of detrital components and ferruginous ooids indicates a humid weathering regime, confirmed by the intercalations of freshwater ostracod assemblages (Rosenfeld et al. 1998). PS II: Early Aptian –earliest Late Albian. During the Early Aptian (HST MCL-1) both platform areas pass through significant changes; coarsegrained siliciclastic input is reduced, and thus carbonate dominates the submerged part of the platform (Fig. 9b). In the Galilee –Golan Heights and North Sinai, low-energy environments rich in calcareous algae and bioclasts occur. In the Galilee –Golan Heights, only slight lateral facies variations in large parts of the proximal platform indicate a large extension of the low-energy openplatform facies belt with similar water depths. Thus, flat-topped platform geometry was suggested to form homogenous shallow platform environments at the beginning of PS II. The seaward margin of this platform is not exposed in the studied area and rudist fringing reefs found in the Carmel area (30 km west of Galilee sections) were documented from younger successions only (Bein 1971; Sass & Bein 1982).

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In North Sinai no change of the depositional geometry occurs compared to PS I. The Mansour– Raghawi fault and the Gebel Amrar swell controlled the depositional processes within the small subbasin. In between Sinai and Galilee the subsurface descriptions indicate a transitional zone, without a major shelf break (Rosenfeld et al. 1998). Deepening in the late Early Aptian starts with a succession of orbitolinid beds, followed by openmarine, deeper platform, marly sediments in both areas (Fig. 9c). This deepening causes a shorttermed southward transgression exceeding the Mansour –Raghawi fault until the southwestern Maghara area and may reach into southern Negev (southern Israel), where short-termed marine transgressions is known from the Lower Aptian (Deshayesites deshayesi and D. forbesi ammonite zones) dated by fossils and radiometric ages of an overlying basalt (Gvirtzman et al. 1996). PS III: Earliest Late Aptian. This short-termed platform stage marks a major break in the North Sinai depositional geometry, but is characterized by shorttermed emergence in the Galilee –Golan Heights only (Fig. 9d ). At North Sinai, the fault controlled differential subsidence patterns terminated and the sub-basins were filled with near shore and erosional deposits. Moreover, terrestrial (Raghawi), marine sandstones (Rizan Aneiza), and dolomites (Amrar) correlate with a major second-order lowstand. Regression is also marked by a northward movement of the coastline to an area in between Raghawi and Amrar (Fig. 9d ). Erosional unconformities observed at Gebel Raghawi are possibly owing to denudation of the relief around the normal fault (Fig. 8) and may have shaped the initial stage of a low dipping ramp geometry (observed in the subsequent platform stages). Fluvial sandstones deposited at the Golan Heights indicate a major regression in the northeastern part of the study area. However, erosion was not observed and the geometry persisted homogenous. Altogether PS III presents a stage characterized by a widely emerged platform with distinct regression. PS IV: Late Aptian–Middle Albian. The Late Aptian –Middle Albian PS IV comprises secondorder sequence MCL-2 and the LST and TST of MCL-3, in both areas characterized by temporarily detrital influenced carbonate production. The platform was widely submerged and a gradual transgression marked the entire area (Fig. 9e). At North Sinai, the transgression submerged the Rizan Aneiza, Amrar, Raghawi, and Mansour regions during Late Aptian and the southeastern part of the Maghara area during the Lower Albian (Fig. 7). First marine sediments occur 20 km south of Maghara at the end of PS IV (Bachmann et al.

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2003). A broad delta system developed, which strongly influenced the depositional processes at Mansour and Maghara SE area interfingering with marine marl and limestones around Gebel Raghawi, while Rizan Aneiza was only slightly influenced by siliciclastic deltaics (Fig. 8g). A north dipping ramp geometry is indicated by increasing thickness of open marine sediments with rudist patch reef, oolithic shoals, and lagoonal sediments towards the north. This gradient was considerably developed at the end of PS IV, when mid-ramp environments at Rizan Aneiza gradually passed into lagoons at Raghawi. The dipping direction changed from NW during PS I– III to N (NNW) during PS IV, triggered by the deltaic input from the southern direction. The ramp can be subdivided in three main facies belts: a southern deltaic depositional facies (Mansour and Maghara SE), a shallow protected subtidal facies (Raghawi), and an open subtidal facies with rudist patch reefs in the north (Rizan Aneiza, Fig. 8g). Those facies belts prograded and retrograded following higher frequency sea-level changes. The extensive siliciclastic input in the delta realm caused higher sedimentation rates in the proximal parts of ramp. Only in the upper PS IV did the depocentre move towards north– NW, when the proximal platform was emerged during the third-order LSTs of the early second-order HST of MCL-3 and sedimentation rates increased in the more distal parts of the platform. At Galilee –Golan Heights the entire platform was submerged since the Late Aptian. Marine sedimentation in the adjacent Jordan is not older than latest Albian (e.g. Schulze et al. 2004), which indicates that the coastline was close to the Golan Heights. Rudist fringing reefs were established at the shelf hinge (Bein 1971; Sass & Bein 1982). Ross (1992) pointed out that the back reef gravel may have formed a major hydrologic barrier, while Bein & Weiler (1976) documented the lower slope nature of muddy sediments deposited in close vicinity to the fringing reefs at the Carmel area. Homogenous flat-topped platform geometry and a strong decrease in hydrologic energy in the proximal platform areas resulted in a shallow lagoonal inner platform environment that varied little over a large area. Sedimentation rates were very similar in the inner platform area, comprising Galilee and the Golan Heights. In the area in between Sinai and northern Israel, a facies established similar to the northern region (Rosenfeld & Hirsch 2005). PS V: Late Albian. PS V is characterized by gradual change of the platform geometry at Sinai and constant filling of the accommodation at Galilee and Golan Heights (Fig. 9f ). PS V comprises the last

herein studied second-order sequence. In both areas sediments of the protected inner platform area with strong reduced siliciclastic amounts are prevailing. Crucial for the development of North Sinai was the dislocation of the rising sea-level. The sedimentation rates increased in the more distal sections resulting in filling of distal accommodation. A very homogenous flat-topped platform developed. During the Late Albian this platform extended southwardly to an area 100 km south of Maghara (Bachmann & Kuss 1998). Within this large extended platform, a succession of rudist biostromes, alternating with protected lagoonal and algal-laminate dominated tidal flats, were deposited resulting from higher frequency sea-level changes. Deposition of grainstones and oolithic facies was restricted to short intervals at Gebel Raghawi, indicating the low-energy dominance during that interval and suggesting a shelf break in an area further north. In Galilee, dolomites representing tidal flats and lagoon prevail, which are also described from the Golan Heights and from Judea in between Galilee and North Sinai (Rosenfeld & Hirsch 2005). The Upper Albian distribution of facies belts indicates a large extended homogenous, very shallow, platform without major changes in facies reaching from northern Israel and the Golan Heights to North Sinai. A continuously increasing sea-level resulted in high accumulation rates and southward transgression. However, platform growth kept up with the sea-level rise and shallow-water, and intertidal deposition continued.

External controlling factors As external factors controlling the sedimentation patterns on the carbonate platform we consider climate and global sea-level changes. Imprints of those changes should occur at both platform areas studied, independent from different tectonical regimes, and are comparable to similar changes observed on other Tethyan platforms (Fig. 10). Second-order sea-level change and its influence on the formation of the platform stages. Second-order sea-level changes highly influenced the Levant platform development by controlling the accommodation and the centres of deposition. During the Late Barremian–Early Aptian a second-order TST created the accommodation to start up the marine sedimentation. Reduced accommodation during second-order HST resulted in distal dislocation of the depocentre and thus controlled the replacement of the ramp by flat topped platform architecture in Galilee –Golan Heights. Erosion during the early Late Aptian LST (SB MCL-2) influenced the ramp

BARREMIAN –ALBIAN LEVANT PLATFORM EVOLUTION

geometry in North Sinai, while reduced accommodation during MCL-3 HST and MCL-4 LST initially induced the distal dislocation of the depocentre and the change from ramp to platform architecture in North Sinai. A comparison of global sequences with that of the Arabian plate origin is shown in Figure 10. Data are constrained by the stratigraphical frame slightly clouded by the use of different time scales in the several studies. The southern Arabian plate exhibit similar features concerning the widespread continental siliciclastic sediments flooded during the Late Barremian, and marine sedimentation during Aptian –Albian times (Haq & Al-Qahtani 2005; Sharland et al. 2001, Fig. 10). The early Late Aptian sequence boundary (MCL-2) coincides with prominent sequence boundaries observed at various localities of the Arabian plate (Van Buchem et al. 2002; Gre´selle & Pittet 2005; Haq & Al-Qahtani 2005). A second match with the Arabian plate feature is observed for the Late Aptian SB MCL-4, while SB MCL-3 is younger than that observed on the southern Arabian plate (Gre´selle & Pittet 2005; Haq & Al-Qahtani 2005). Some further second-order SBs observed on the southern Arabian plate by Gre´selle & Pittet (2005) or Haq & Al-Qahtani (2005) coincides with SBs regarded as a higher frequency in origin in North Sinai, because they are not clearly indicated at Galilee –Golan Heights. Important differences to the southern Arabian plate were observed in the Upper Aptian succession. While a Late Aptian unconformity is documented from several platforms (Van Buchem et al. 2002; Haq & Al-Qahtani 2005), continuous marine sedimentation took place in the Levant Platform at the northern edge of the Arabian plate. Comparing the sea-level changes with patterns supposed to be global, similarities are much less obvious. However, the transgression characterizing the Upper Aptian– Albian Levant succession coincides with the long-term global transgression pointed out by Hardenbol et al. (1998), while the lower part of the succession falls into a global sealevel minimum. This agrees with transgressions of smaller amplitude during that time interval on the Levant Platform. Additionally SB MCL-2 possibly correlates with Ap 4 defined by Hardenbol et al. (1998) and with a few time-equivalent exceptional SBs in the northern Tethyan realm, like the Alpine mountains (e.g. Roter Sattel, Strasser et al. 2001), while other areas are characterized by transgression or high stand during the same time. Altogether, the sea-level history of the Levant Platform reflects the Late Aptian– Albian global long-term transgression and the second-order sealevel changes are highly correlated with that one observed on the Arabian plate.

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Climate. The composition of the siliciclastic input and the formation of ferruginous ooids allow the interpretation of general climate conditions in the study area (Fig. 10). The Upper Barremian– Albian succession comprises two intervals with coarse grained siliciclastic input and occurrence of ferruginous ooids, one reaches from the Late Barremian to the earliest Aptian (PS I) and the second comprises the Late Aptian until the mid-Albian (PS IV). Within these parts of the succession marly facies dominates and carbonatic samples often show a marly matrix. These intervals were interpreted as indicating humid conditions (Bachmann & Hirsch 2006; Bachmann & Kuss 1998) in concordance with Mu¨cke (2000), who demonstrated that lateritic weathering in the hinterland was producing the protoliths for Upper Cretaceous oolithic ironstones. Those interpretations are underlined by the common occurrence of plant remains in the succession of PS I at Gebel Raghawi and by the formation of the pronounced delta system and carstification horizons in North Sinai during PS IV. Orbitolinid-rich facies types, associated with abundant calcareous algae and echinoderms and argillaceous muddy limestones, are common in PS IV. Those facies types are typically interpreted as reflecting mesotrophic conditions (Van Buchem et al. 2001; Pittet et al. 2002), which may reflect nutrient input owing to weathering and humid conditions. Two intervals are characterized by reduced siliciclastic input and the absence of the ferruginous ooids (Early Aptian/PS II and the Late Albian/PS V). Such a reduced siliciclastic input can result from reduced sediment supply owing to less intensive weathering or from an increasing distance of the delivery area. For PS II, the steady coast line between Mansour and Raghawi during lower PS II argues for a climate origin of the reduction of detrital input. The Late Albian (PS V) of reduced siliciclastic input is accompanied by an increase of rudist and miliolid dominated pure limestones in North Sinai. Those facies types are commonly interpreted as reflecting oligotrophic conditions (Van Buchem et al. 2002), which fit well with the observed transgressional, less humid system. Changing third-order low stand features accompany humid and less humid phases. LSTs are characterized by detrital input during the humid phases and by emergence and dolomitization during the less humid periods (Bachmann & Kuss 1998). Our data fit with the general ideas of Price (1999) and Hillga¨rtner et al. (2003) for the Oman, who indicated a small trend to humidity in the Late Barremian –Early Aptian. However, the position of the Levant Platform near the equator may have triggered humidity in northeastern Africa, while more

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arid conditions influence the northern Tethyan margin during the Barremian–earliest Aptian. Tectonics. Tectonical activity played an important role, especially during the initial phase of the platform development in North Sinai (Fig. 10). Our data suggest that during the Late Barremian–Early Aptian, extensional normal faults influenced the North Sinai environment and result in the formation of sub-basins, while at the same time a slightly dipping ramp, without fault segmentation, is recorded in northern Israel. The Late Aptian – Albian distribution of the facies belts in North Sinai indicates a uniform deepening in a NW direction and confirms that sedimentation geometry was orientated parallel to SW –NE striking normal faults pointed out by Moustafa & Khalil (1990). Pre-deposition extensional tectonics creating normal faults are described from the Levant margin offshore Israel as well as from northern Sinai until the mid-Jurassic (Moustafa & Khalil 1990; Garfunkel 2004; Gardosh & Druckman 2006). In North Sinai, a set of extensional basins was created from which the Maghara area represents the northernmost point that was located onshore (Moustafa 2010). Extensional tectonics at the North African –Arabian continental margin are interpreted as reflecting divergent movement between the Afro-Arabian and the Eurasian plate, which led to the Mesozoic to Middle Jurassic opening of the Neotethys (Stampfli et al. 2001). Our data indicate that extensional faults were active until the late Early Aptian in North Sinai and may represent a Lower Cretaceous syn-rift stage according Guiraud et al. (2005), who observed continental rifts along the African –Arabian Tethyan margin in the Late Berriasian –earliest Aptian.

Conclusions During the Late Barremian–Albian, the southern Levant Platform was studied in the two regions: northern Israel (Galilee –Golan Heights) and North Sinai. Facies, stratigraphy and stable isotopes of several sections were studied, to reconstruct transects across the shallow shelf. In conclusion we establish depositional models and discuss the controlling factors for the shallow water deposition. We combined biostratigraphy on the base of benthic foraminifers, mainly orbitolinids, and ammonites with stable isotope data, which allow us to date the shallow water strata and all controlling factors in detail. The depositional architecture was controlled by local tectonics, climate, and second-order sea-level changes affecting the sedimentation patterns. Four secondorder sequence boundaries were identified in the Levant area (MCL-1– MCL-4). They partly

correlate with those observed on the Arabian plate (Haq & Al-Qahtani 2005; Sharland et al. 2001), suggesting a regional control. Until the late Early Aptian, the North Sinai was influenced by normal fault development, while the Galilee area– Golan Heights exhibit continuous sedimentation without tectonical influence. Both regional transects reveal five platform stages (PS I –V) that differs with respect to platform architecture, siliciclastic input, and response to sea-level changes. PS I (Late Barremian–earliest Aptian). The Upper Barremian marine sediments transgraded on terrestrial deposits. In northern Israel a homogenous ramp existed, while North Sinai was subdivided in SW –NE striking sub-basins, marginally bounded by active normal faults. Open marine high-energy sedimentation and continuous transgression characterized the Galilee– Golan Heights area during the TST of MCL-1, while near coast siliciclastic and protected lagoonal carbonates alternate with deeper marine sediments in North Sinai. PS II (Early Aptian –earliest Late Aptian). During the HST of MCL-1, continuous filling of the accommodation resulted in shallow marine protected facies belts marking flat-topped platform architecture at northern Israel. Owing to higher subsidence around normal faults, sub-basin development kept on in North Sinai. The sedimentation was characterized by protected– partly deeper– subtidal environment within these sub-basins, separated from eachother by shallower swells. Significant reduction of detrital input in both areas may result from changing weathering regimes. PS III (early Late Aptian). During the LST of MCL-2 the tectonical activity terminated resulting in reorganization of the North Sinai platform. The former fault-controlled sub-basins became inactive and were covered by a shallow ramp architecture that controlled the depositional processes of the western Levant Platform. Emergence was evidenced from North Sinai to northern Israel indicating an extended platform system. PS IV (Late Aptian– Middle Albian). A stable platform architecture, marked by a homoclinal ramp in North Sinai and a flat-topped platform in Galilee –Golan Heights, was sandwiched between transgressive surface MCL-2 and the end of TST of MCL-3. In the North Sinai a broad delta system with high siliciclastic input was deposited, interfingering in the north with carbonate ramp deposits. High accumulation rates on parts of the ramp resulted in slight changes of the dipping direction. Simultaneously,

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a less siliciclastic input and inner platform sediments are characterized in the Galilee –Golan Heights region. PS V (Late Albian and younger). During HST of MCL-3 and TST of MCL-4 pure limestones and dolomites were accommodated in North Sinai, during the ongoing transgression a flat topped platform was established here as well as in northern Israel. Our data indicate that extensional faults controlled sedimentation until the late Early Aptian, which is significantly younger than observed before (Moustafa & Khalil 1990; Garfunkel 2004; Gardosh & Druckman 2006). They may represent a syn-rift stage according to Guiraud et al. (2005). Concerning the climate history, we point out that the larger input of detrital sediment, including ferruginous ooids, indicates an increased humidity owing to an accelerated continental weathering, particularly in the Late Barremian –Early Aptian. We are indebted to F. Hirsch (Jerusalem) and R. Stein (Bremerhaven) for discussions on the Lower Cretaceous of Israel, respectively geochemical data. We are especially thankful for J. Thielemann (Bremen) and M. Heldt (Hannover) for assistance in the field, many successful discussions, and providing valuable data of their theses. A. R. Moustafa (Cairo) helped with discussing the tectonical background of the working-areas. Technical support by M. Krogmann (Bremen) is appreciated for preparing a part of the ammonite fauna. We thank the two reviewers M. Simmons and R. Scott for helpful and instructive discussions and comments. We appreciate financial support by the German Science Foundation and the VW foundation.

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The mid-Cretaceous carbonate system of northern Israel: facies evolution, tectono-sedimentary configuration and global control on the central Levant margin of the Arabian Plate RAN FRANK1*, BINYAMIN BUCHBINDER2 & CHAIM BENJAMINI1 1

Department of Geological and Environmental Sciences, Ben – Gurion University of the Negev, P.O. Box 653, Beer-Sheva 84105, Israel

2

Geological Survey of Israel, 30 Malkhe Yisrael Street, Jerusalem 95501, Israel *Corresponding author (e-mail: [email protected]) Abstract: This study deals with the sedimentary evolution, tectonic configuration and global imprints of a Cenomanian –Turonian carbonate system located in northern Israel, on the central part of the Levant margin of the north Arabian Plate. Detailed sampling of field sections, mesoscopic features, petrography and microfacies form the database for this study. Facies units are integrated into high- and low-order cycles that comprise a sequence stratigraphic model. Two palaeo-highs, separated by a subsiding trough, all striking east– NE, govern the pattern of carbonate deposition in northern Israel. An additional subsiding region extended northward into Lebanon. Eustatic and palaeoenvironmental imprints are represented by earliest Cenomanian subaerial exposure; Early Cenomanian maximum flooding and oxygenation of hypoxic sea-floor; Middle Cenomanian high-stand progradation followed by forced regression and mass transport; Middle Cenomanian subaerial exposure; Late Cenomanian eutrophication during sea-level rise; Late Cenomanian subaerial exposure; latest Cenomanian–Turonian eutrophication and gradual development of the OAE-2 (oceanic anoxic event). A Late Cenomanian eustatic rise was locally masked by uplift and subaerial exposure. We conclude that the tectono-sedimentary regime of northern Israel represents an east– NE branch-off of the depositional strike from the north– south striking Levant margin, and that the carbonate system of this region was strongly influenced by eustasy and palaeoceanographic trends of the Tethys.

Mid-Cretaceous carbonate rocks are spectacularly exposed in the Galilee and Carmel regions of northern Israel, part of the central Levantine passive margin of the Arabian Plate (Fig. 1). Previously, sequence stratigraphic and palaeoenvironmental evolution of mid-Cretaceous carbonate systems were described for parts of the southern and eastern Levant margin (Bachmann & Kuss 1998; Bauer et al. 2003; Schulze et al. 2003, 2004, 2005). Sedimentary configuration for the Galilee and Carmel of northern Israel was previously studied by Freund (1965), Kafri (1972, 1991), Bein (1974) and Sass & Bein (1982), but a comprehensive sequence stratigraphic evolution for northern Israel was not described. Therefore, the integration of this region into the tectono-sedimentary framework of the northern Arabian plate is lacking. Previous work on the carbonate successions of northern Israel were related separately to the Carmel and Galilee regions (Fig. 1). Bein (1976, 1977), Bein & Weiler (1976), Sass & Bein (1982) and Segev & Sass (2006) considered rudist accumulations in the Carmel as the key to understanding

the configuration of the carbonate system in this region. The Carmel (Fig. 1b) was considered as a transitional zone, where shallow-marine carbonate platform sediments passed westwards via a rim of rudist barrier-reefs into deeper water laminites, turbidites and contourites. NE of the Carmel, in the Galilee, this type of east to west (proximal-distal) facies transition was not described. Freund (1965) and Kafri (1972, 1991) considered the Galilee to be a shallow carbonate platform transected by intrashelf basins fringed in places by rudist-reefs. Our comprehensive study of the Cenomanian– Turonian succession of northern Israel is based on bed-by-bed sampling, mesoscopic observations in the field, and petrography and microfacies of these rocks. Using the sequence stratigraphic paradigm, stage-by-stage evolution of this region was reconstructed, emphasizing the genesis of the carbonate system, tectonic processes, and the roles of eustasy and global palaeoenvironments. Our results fill a knowledge gap regarding the evolution and configuration of the Arabian plate margin in this area by supporting a markedly different model of

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 133–169. DOI: 10.1144/SP341.7 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Location of the study area within the Arabian plate, with distribution of mid-Cretaceous sedimentary units and major tectono-sedimentary features in this region (modified after Ziegler 2001). The ‘Levantine hinge-belt’ along the present coastline of Israel and Lebanon (thick bold line) is a narrow north–south striking zone across which east-to-west facies changes occur, from shallow water carbonates (Judea, Ajlun groups) to the east to deeper water carbonates (Talme Yafe Group) in the west. (b) Details of the study area in northern Israel with location of columnar sections and other relevant outcrops. (c) Graphic legend for columnar sections presented in Figures 6–10, 17 and 18.

THE CARBONATE SYSTEM OF NORTHERN ISRAEL

tectono-sedimentary evolution for northern Israel than presented in previous works, emphasizing global palaeoceanographic control including sealevel change, nutrient levels, oxygenation of the water mass, and shelf edge configuration. Specific goals of this study were (1) to establish facies types describing the depositional environments in the mid-Cretaceous of northern Israel, (2) to integrate facies types into genetic units and to determine cyclic patterns and their hierarchical organization, (3) to document proximal-to-distal facies-thickness changes and the geometries of genetic-stratigraphic units at system-tract level, (4) to establish a local tectonic framework for these trends, (5) to uncover effects of Tethyan and global palaeoceanographic/palaeoecological events and eustatic fluctuations on the succession and (6) to present the mid-Cretaceous depositional and structural configuration of the Arabian plate margin as expressed in northern Israel.

Geological setting Tectonic setting of the Levant margin in the Cretaceous The Levant margin of the Neotethys (Fig. 1a) was shaped by Late Permian, Middle to Late Triassic and Early Jurassic extensional tectonics. Rifting led to separation of the Tauride and Eratosthenes blocks from the Arabo-Nubian Platform, forming the Levantine Basin of the eastern Mediterranean (Bein & Gvirtzman 1977; Garfunkel & Derin 1984; Garfunkel 1998; Robertson 1998). Normal faulting associated with this rifting determined the

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depositional strike along the Levant margin, constituting the eastern margin of the Levantine Basin. The north–south trend exercised control on facies transitions along a narrow depositional belt, termed the ‘Levantine hinge-line’ or ‘belt’ (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977). Cessation of faulting in the Middle Jurassic, lengthy tectonic quiescence, and slow passivemargin type subsidence prevailed along the Levant margin from the Mid-Jurassic to the mid-Cretaceous. During this period, east –west proximal-to-distal facies transitions were recorded in the hinge-belt, normal to the north–south striking shelf-margin. The mid-Cretaceous ‘Levantine hinge-belt’ is described from the southern coastal plain of Israel extending to the Carmel region (Bein 1971, 1974; Gvirtzman & Klang 1972; Bein & Weiler 1976; Bein & Gvirtzman 1977; Sass & Bein 1982). Northwards in the Galilee and southern Lebanon, the directional trend and associated facies transitions of the ‘Levantine hinge-belt’ were not reported. However, Walley (1998) described a similar hingebelt trend (Fig. 1) along the western Lebanon flexure in northern Lebanon, based on interpretation of data from Saint-Marc (1974) and others. He considered Galilee and southern Lebanon to be underlain by a southwestern extension of the Palmyride basin of Syria, representing an interruption of the north– south trend.

Cenomanian – Turonian lithostratigraphy in northern Israel Figure 2 represents the lithostratigraphy currently used for the Upper Judea Group carbonates

Fig. 2. Lithostratigraphic scheme for Cenomanian and Turonian mapped units in the Galilee and Carmel. Based mainly on the 1:200 000 geological map of Sneh et al. (1998) with some new observations. Not to scale.

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(Cenomanian–Turonian) of northern Israel. It incorporates schemes used by Picard & Kashai (1958), Freund (1959), Kashai (1966), Kafri (1972, 1991), Sneh et al. (1998), Sneh (2002) and Segev & Sass (2006) with some local modifications. There has been a tendency to use different lithostratigraphic concepts for Galilee and Carmel, and the schemes are presented here separately. In Galilee, limestones of the Deir Hanna Formation occur at the base of the Upper Judea Group, overlying Upper Albian Yagur or Kamon dolomites (Fig. 2). The Deir Hanna Formation consists of well bedded or laminar chalks and limestones with chert nodules or bands, sometimes dolomitized near the top. Sakhnin Formation dolomites directly overlie the Deir Hanna Formation in most of Galilee. The Sakhnin dolomites become wellbedded upwards, or are sometimes highly brecciated (e.g. in the regions of Adamit, Peqi’in, Deir El-Assad, Fig. 1). In some localities, breccias of the Sakhnin Formation pass laterally into dolomitized laminites of the Upper Deir Hanna Formation (e.g. Adamit, Betzet, Fig. 1). The Sakhnin Formation was not mapped in the western Galilee (e.g. Picard & Kashai 1958; Freund 1965), where bioturbated limestones of the Yanuch Formation (Freund 1959) unconformably overlie chalky laminites of the Deir Hanna Formation. Kafri (1972) showed that the bioturbated limestones of the Yanuch Formation in western Galilee pass into chalks and bedded limestones toward the NW Galilee (e.g. Hila and Hurfesh regions, Fig. 1). Typical lithologies, microfacies and cyclic patterns of the Yanuch Formation also occur in NE Galilee (e.g. Dishon, Manara sections, Fig. 1). From western to central Galilee, the Yanuch Formation is overlain by ammonite-bearing marls of the Lower Turonian Yirka Formation. In other parts of Galilee, a limestone complex mapped as the Bina Formation (Shadmon 1959) directly overlies either the Yanuch or Sakhnin Formations. In the Carmel, the Yagur Formation dolomites of the Late Albian age (Fig. 2) form the lower part of the Judea Group. The Yagur Formation is overlain by the Isfiyye Formation, consisting of well-bedded chalks with chert nodules and bands, and in some places with pyroclastics interbedded at the base (e.g. Sass 1980; Segev & Sass 2006). The Isfiyye Formation is overlain by bioclastic limestones of the Beit-Oren Formation. The Beit-Oren limestones are overlain by chalks with some pyroclastics of the Arkan Formation (Segev et al. 2002; mapped as Khureibe and Junediyye Formations in older studies). The Arkan Formation is overlain by dolomites and limestones of the Muhraqa Formation. This transition is unconformable (Bein 1974; Lipson-Benitah et al. 1997), in places accompanied by pyroclastics (Segev & Sass 2006). Above the

Muhraqa Formation are Lower Turonian ammonitebearing marls, chalks and limestones of the Daliyya Formation in the central-northern Carmel (Freund & Raab 1969). Limestones of the Bina Formation form the top of the succession. In the region of southern Carmel, the Cenomanian succession is only partly exposed and is mostly dolomitic. At the base a dolomitized chert-bearing equivalent of the Isfiyye Formation is overlain by bioturbated dolomites of the Zikhron Formation, becoming well bedded toward the top. Dolomites of the Zikhron Formation are considered by Segev & Sass (2006) as the lateral continuation of the Arkan chalk complex to the north. The Zikhron dolomites are overlain by a pyroclastic bed (Segev & Sass 2006) followed by massive dolomites and limestones mapped as Sakhnin and Bina Formations respectively. These lithostratigraphic units, for the most part, are useful for depiction on geological maps. In order to determine the sedimentary evolution of this region, however, it was necessary to deconstruct them into their basic genetic components, and to reconstruct their cyclic patterns, systems tracts and sequences, using biostratigraphic and other chronostratigraphic controls.

Material and methods The database for this study was acquired by bed-by-bed sampling and analysis at outcrop, mesoscopic (hand sample) and petrographic scales. Thin sections were examined for non-skeletal and skeletal grains, cements and early diagenetic history. Limestone classification follows Dunham (1962) and Embry & Klovan (1971). Key micro- and macro-fossils were determined by experts. Beds in each columnar section were classified by their sedimentological features into facies units representing palaeoenvironments, broadly following the subdivisions of Read (1985) and Burchette & Wright (1992). In each section vertical cyclic pattern and cyclic hierarchies were analysed. The sequence-stratigraphic approach implemented here is based on the ‘four system-tract model’ (Hunt & Tucker 1992; Helland-Hansen & Gjelberg 1994) modified for carbonate systems (Hunt & Tucker 1993). Lowstand systems tract (LST), transgressive systems tract (TST), highstand systems tract (HST) and forced regressive (FRST) systems tract are recognized, with bounding surfaces that include (a) a basal surface of forced regression (BSFR) at the base of the FRST; (b) a sequenceboundary (SB) unconformity at the top of the proximal part of the HST, and basinwards, overlying the FRST formed during relative sea-level fall. Proximally, the SB unconformity may coalesce with the younger BSFR, while distally the SB may be expressed by a correlative conformity; (c) a

THE CARBONATE SYSTEM OF NORTHERN ISRAEL

transgressive surface at the base of the TST; (d) a maximum-flooding surface [in fact, maximum flooding interval, (MFI)] at the top of the TST. The concept of a type-3 SB (drowning unconformity) introduced by Schlager (1999) was adopted for short-lived subaerial exposures topped by transgressive deposits, without intervening LST deposits. A time-boundary definition of some of the system tracts employed the transgressive– regressive sequence stratigraphic method of Embry (2002), with SB and maximum-flooding surfaces as main bounding surfaces defining a regressive system tract (RST). Time-correlative bounding surfaces were established using published biostratigraphy (mostly Freund & Raab 1969; Lewy & Raab 1978; LipsonBenitah et al. 1997), radiometric age determinations (Segev et al. 2002; Segev & Sass 2006), and some personal communication with experts, as explained where presented.

Results: facies types and cyclic patterns Sedimentary facies types and their environments of deposition Types of carbonate facies occurring in the Cenomanian–Turonian of northern Israel are summarized in Table 1. In total, 20 facies types are described (marked FT1– FT20 in Table 1 and Fig. 3). Column 2 in Table 1 provides details of sedimentology, bedding and faunal components. These sedimentary facies reflect a variety of depositional environments ranging from autochthonous basinal deposits and condensed surfaces, to supratidal and subaerially exposed sea-floor, as well as environments characterized by submarine transport. Column 3 in Table 1 and Figure 3 shows the palaeoenvironments indicated by these features, following the carbonate ramp subdivision into basin, outer ramp, mid-ramp and inner ramp of Burchette & Wright (1992). Column 4 lists the lithostratigraphic units within which these facies types were found, while column 5 presents their position within the sequence stratigraphic scheme. Figures 4 and 5 depict some of these features in the field and under the microscope.

Cycles and cyclic hierarchy in the Cenomanian –Turonian succession of northern Israel Sedimentary cycles in northern Israel were classified into two orders according to the following criteria: high-order cycles that cannot be further divided into smaller cycles, and therefore termed undividable cycles (UC); and low-order cycles that

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are composed of high-order cycles, therefore termed composite cycles (CC). High-order cycles are arbitrarily numbered UC-1 –12. They are graphically presented and explained in Figures 6 and 7. UC are mostly decimetres to a few metres in thickness, but may in some cases be more than ten metres thick. UC are combinations of different facies types (Table 1, Fig. 3). UC-1 –6 (Fig. 6), are shallowing-upward cycles and UC-7 –12 are deepening-upward cycles (Fig. 7). UC-1 –UC-6, and UC-7, 11 and 12 are combined into low-order CC, but deepening-upward cycles UC-8 –10 (Fig. 7) are independent. The low-order composite cycles CC-1–7, are graphically simplified in Figure 8. Their thicknesses ranges between a few metres to a few tens of metres. † Type-1 low-order CC-1 (Figs 8a & 9) begins with basinal laminites (in some parts dolomitized), passes upwards into bioturbated fossiliferous mid-ramp deposits, and is topped by stacked UC-1 peritidal cycles. This is a shallowing-upward, progradational cycle, with the mid-ramp facies commonly rich in macrofossils and skeletal debris, reflecting high skeletal production on the mid-ramp (facies type-12 in Table 1). This high production results in filling the accommodation space to near sea level, as reflected in development of UC-1 peritidal cycles at the top. † Type-2 low-order CC-2 (Figs 8b & 9) is composed of stacked UC-3 high-order cycles (Fig. 6). Each UC-3 cycle may begin with a basal hiatal shell-concentration, then passes upwards into pelagic basinal or outer-ramp facies, rich in pithonella calcisphaeres, phosphatic grains and bioeroded bioclasts, and is topped by cross-bedded or bioturbated shoreface grainstones. The resulting CC-2 cycle indicates pulsed progradation/aggradation and filling of the accommodation space to fair-weather wave base only. In comparison with the CC-1 described above, CC-2 reflects a lower rate of skeletal carbonate production. † Type-3 low-order CC-3 (Figs 8c & 9) is composed of lower shoreface UC-4 cycles, passing upwards into upper shoreface UC-5 cycles, and ending with peri-tidal UC-1 cycles. In-situ skeletal growth in CC-3 is limited, but accommodation space was filled to near sea level, to a great extent by transported peloids. The composition and fabric of the peloids indicate origination in a predominantly shallow inner-ramp setting. Thus, progradation and accommodation filling were maintained by grain supply from the inner-ramp to the shoreface. † Type-4 low-order CC-4 (Figs 8d & 9) begins with stacked peritidal UC-1 cycles with rare

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Table 1. Facies types of the Cenomanian –Turonian succession of northern Israel 1. Facies type FT-1

FT-2 FT-3

FT-5 FT-6 FT-7

FT-8

Platy or laminated chalk, indurated chalk, or marl, bioclastic/ peloidal calcisiltite, or micrite mudstone. Chert as nodules or bands. Fish debris, echinoids, pelagic crinoids. Planktonic, esp. heterohelicid foraminifera, pithonellid calcispheres. Pyrite, glauconite. Phosphatic grains, especially in Yanuch/ Yirka Fm. Decimeter-thick calcarenite, calcisiltite beds. Cross-laminated, graded to massive, rippled, planar. Occasionally dolomitized. Cm- to m-scale breccias, embedded in laminated or homogeneous dolomite, or chalk matrix. Internally deformed (fragmented, brecciated, folded, faulted). Dolomitized matrix. Shear features at base: boudins, folds, shear foliation, d-microstructures. Thick (.100 m) well-stratified, thin-bedded/laminated calcarenitic clinoforms dipping 228 – 358. Beds to 30 cm thick. Well sorted, mud-free, mostly non-graded. Bioturbated or thin-bedded fine-grained packstones. Pelagic microfossils: pithonellid calcispheres, planktonic foraminifera, pelagic crinoids. Phosphatic grains; mm-sized Fe concretions. One or more limestone beds within pelagic beds. Coarse bioclastic or poorly washed peloidal grainstone. Pelagic microfossils, crinoids, echinoderm debris in matrix. Bioerosion. Quartz, phosphatic grains. Chalcedony cement in borings. Densely packed bivalve shell-beds embedded in basinal chalks. Gryphaeids, Pycnodonte vesiculosa in Carmel. Shells bioeroded, some fragmented. Matrix calcisiltite and pelagic microfossils. NE Galilee: flat elongated orbitolinids; serpulids; some large bivalves. Rudists shell beds in the Bina (Kishk Fm).

3. Environment

4. Lithostratigraphic units

5. Sequence stratigraphic position

Hypoxic basinal or outer-ramp deposits. Yirka Fm – hypoxia, nutrient rich.

Deir Hanna, Yanuch, Yirka, Isfiyye, Arkan

Ce TST-1; Low Ce RST-1; Ce TST-2; Tu TST

High – low density Turbidites.

Sakhnin; mid Muhraqa. Sakhnin, Bina, Muhraqa Sakhnin

Ce FRST-1; Ce FRST-2 Ce FRST-1; Ce FRST-2 Ce FRST-1

Grain flow (Beds ,5 cm); Sandy debrites (Beds .5 cm).

Yanuch (W Galilee)

Ce FRST-2

Pelagic basinal (pelagites).

Muhraqa, Yanuch, Bina

Ce HST-2; Ce TST-2

Fragmental lag concentration. Winnowing of basinal hypoxic sea-floor.

Deir Hanna, Yirka

Top Ce TST-1; Tu TST (Max.-flooding).

Lag concentration. Colonization of hypoxic sea-floor. Low sed. rate.

Within upper Deir Hanna; part of Beit Oren (Carmel).

Top Ce TST-1; Tu TST ;

Bina Fm – mid-ramp rudist conc. during storm decline.

Within Bina (Kishk) Fm of western Galilee.

Tu HST.

Debrites. Translational sliding.

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FT-4

2. Description

Thin reddish-black crusts, mm thickness, paving irregular bedding planes and exposure surfaces. Lag conglomerate.

FT-10

Bioturbated coarse-grained pack- or wackestone; chert nodules. Bioclasts: rudists, oysters, gastropods, echinoderms, fish. Pithonellid calcispheres, planktonic, rare benthonic foraminifera; pelagic crinoids. Phosphatic grains. Bioeroded bioclasts, occasional chalcedony cement. Laminated, thin-bedded or well-stratified, very fine-grained mudstones or wackestones, sterile to slightly bioclastic. Rotaliid foraminifera (especially gavelinellids); fecal pellets; rare planktonic foraminifera and fragments of echinoderms (Bina Fm). In South Galilee: dolomitized with some shear folds and synsedimentary breccias. Dolomitized bioturbated pack-, float- or wackestones. Complete and disarticulated rudists, oysters, gastropods, echinoderm debris; poriferan spicules; pithonellid calcispheres. Shallow-water benthic foraminifera (Nezazzatids, Cuneolina, Dicyclina) upwards. Hummocky Grainstones with peloids, quartz cross-stratification. silt. Algal, molluscan, echinoid SE-dipping bioclasts. Shallow-water benthic foraminifera (especially clinoforms. miliolids, Cuneolina). Massive, cross-bedded or planar grain to rudstones. Rudist and oyster bioclasts, calc. algae. Abundant mud-peloids and lithoclastic grains; some ooids (Yanuch, Bina). Isopachous rim cement; blocky spar; poikilotopic calcite. Bioturbated mudstones and wackestones with shallow-water benthic foraminifera: Nezazzatids, alveolinids, Cuneolina, Dicyclina. Rarely, small gastropods, rudist and echinoderm debris.

FT-11

FT-12

FT-13 FT-14 FT-15

FT-16

Can be found in the basin and on the platform. ferruginous mineralization of submarine omission surfaces; some reworking forming lag cgl. Outer ramp. Nutrient-rich surface water. Reduced sedimentation rate.

Sakhnin, NE Galilee; Yanuch, West Galilee; Muhraqa, Carmel.

Base UC-1; base Ce TST-2; base Pelech seq., base Tu TST

Muhraqa; Yanuch; Bina.

Ce HST-2; Ce TST-2

Hypoxic mid- or outer ramp. Hypoxic lagoonal (e.g. Betzet) Mass-transport features in South Galilee (Mt. Kedimim).

Sakhnin (Mt. Kedumim); Bina

Ce RST-1; Tu TST

Mid-ramp. Increased skeletal production.

Lower Sakhnin (Galilee), Zikhron (Carmel), Bina (NE Galilee)

Mid- e HST-1

Lower shoreface.

Bina; Muhraqa

Tu HST

Offshore bars. Upper shoreface.

Bina

Tu HST

Upper shoreface.

Yanuch, Bina

Ce HST-2

Low energy, subtidal lagoon. Intensive dolomitization beneath supratidal facies (FT-19).

Sakhnin, Bina, Daliyya

Ce HST-1; Tu TST-2; Ce HST-2

THE CARBONATE SYSTEM OF NORTHERN ISRAEL

FT-9

(Continued)

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Table 1. Continued 1. Facies type FT-17

FT-19

FT-20

Bioturbated limestone, decimeter to 2 m beds. Acteonella sp., chondrodonts, radiolitid, hippuritid rudists, poriferan, echinoderm debris; shallow-water benthic foraminifera. Platy or massive foraminiferal packstones to poorly washed grainstones. Oncoids, microbial lumps and films, micritized bioclasts. Dolomite, laminated auto-micrite, peloids. Laminae irregular to wavy; coating free surfaces and bioclasts. Fenestrae, birds-eyes, fragmentation, flat-pebble conglomerates; dissolution cracks; Fe stained. a) Thin horizon with spheroidal calcitic, silicic or partly-silicified vadose pisoids, pore-filling speleothems (flowstones), empty or infilled dissolution vugs, concentric pendant calcites. b) Irregular rugged pitted surface covered by highly-weathered marl layer. c) Thin reddish crust (millimetre-thick) with pedogenic pisoliths, circumgranular cracks and alveolar septal fabrics.

3. Environment

4. Lithostratigraphic units

5. Sequence stratigraphic position

Open-marine shallow subtidal.

Top Bina

Tu HST

Semi-restricted lagoon. Slow carbonate sed. rate. Condensed.

Bina

Base Ce TST-2

Supra- and inter-tidal flats with microbial laminites.

Sakhnin and Bina

Ce HST-1; Ce HST-2; Tu HST

Sub-aerial exposure facies: a- Calcrete; b- Karst (Lapies); c- Sub-aerially exposed surface.

a) Top Sakhnin (E Galilee). b, c) Top Yanuch

SBs:Ce SB-2. Ce SB-3. Ce SB-4.

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FT-18

2. Description

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Fig. 3. Schematic model for position of the major Cenomanian –Turonian facies types (facies types 1 –20, Table 1) on two carbonate ramp profiles. Note that facies type-4 includes the range from incipient slides on the inner ramp to disconnected blocks downslope. Transgressive surfaces (facies type-9) are expressed both in the basin and on the inner ramp. Dashed line represents different types of basinal condensed sections (facies types- 7, 8 and 9). Temporal distribution is presented in Table 1; spatial distribution is discussed at length in the text.

†

†

†

†

macrobenthos and passes upwards into openmarine inner-ramp UC-2 cycles in which the sub-tidal mud substrate is colonized by rudists, large chondrodont bivalves, opisthobranch gastropods (Acteonella) and sponges. CC-4 is a peritidal shallowing-upward progradational cycle. Type-5 low-order CC-5 (Figs 8e & 9) is constructed of stacked UC-6 cycles, each beginning with well-laminated mudstones and topped by massive shallow subtidal wackestones. The laminated mudstones of CC-5 were deposited on a stagnant hypoxic sea-floor and the massive beds reflects oxic bioturbated mid-ramp facies. CC-5 represents fluctuations in the oxygen level on the mid-ramp and not necessarily expansion or contraction of accommodation space. Type-6 low-order CC-6 (Figs 8f & 9) is composed of stacked UC-7 cycles with platy or laminated marl or chalk at the base (turbidites, ammonites) and skeletal hiatal concentrations at the top. CC-6 is a deepening-upwards basinal cycle. Type-7 low-order CC-7 (Figs 8g & 9) is composed of stacked UC-11 and/or UC-12 deepening-upward cycles. CC-7 reflects retrogradation of hypoxic laminites. Type-8 low-order CC-8 (Fig. 8h) is composed of UC-8 cycles, each commencing with

shallow-lagoonal facies, deepens upwards into shoreface grainstones, and ends with pelagic outer ramp facies. CC-8 is a low-order deepening-upward cycle.

Sequence stratigraphic subdivision and interpretation of systems tracts The correlations shown in Figures 9 and 10 are based on common cyclic trends found in the columnar sections, identification and extension of bounding surfaces, and biostratigraphic and radiometric age controls (mostly from Freund & Raab 1969; Lewy & Raab 1978; Lipson-Benitah et al. 1997; Segev et al. 2002). These correlations show that the Upper Judea Group consists of three Cenomanian sequences and one Turonian sequence, summarized in Figure 11. Systems tracts and bounding surfaces are presented with discussion of their facies composition and reconstruction of palaeoenvironmental configuration.

Sequence 1: Early – Middle Cenomanian Albian/Cenomanian sequence boundary (Alb/Ce SB-1). The boundary between the Yagur dolomites, considered as Albian in age (Lewy & Raab 1978), and the overlying indurated chalks of the Deir

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Fig. 4. (a) Laminites (facies type 1) of the Ce TST-1. Lower Deir Hanna Formation, Kziv east section, eastern Galilee. (b) Bored and bioeroded bioclasts in poorly washed packstones (facies type-7) of the Early Cenomanian maximum-flooding bed (Ce MFI-1). Borings filled by chalcedonic cement. Upper Deir Hanna Formation, Betzet section, NW Galilee. (c) Poorly washed peloidal grainstones (facies type-7) of the Early Cenomanian

THE CARBONATE SYSTEM OF NORTHERN ISRAEL

Hanna/Isfiyye Formations (Galilee/Carmel respectively; Figs 2 & 9) is a discontinuity surface, described previously by Karcz (1959), Folkman (1969), Bein (1974), Kafri (1986) and LipsonBenitah et al. (1997). Folkman (1969) and Kafri (1986) described pisolitic calcretes, monomictic conglomerates, oxidized Fe crusts, quartz grains and silicification and dedolomitization phenomena, and concluded that this discontinuity surface represents subaerial exposure. Gardosh et al. (2006) associated this subaerial exposure event with sea-level fall and a submarine canyon incision offshore. Ammonites and planktonic foraminifera (Avnimelech 1965; Lewy & Raab 1978; LipsonBenitah et al. 1997) from chalks somewhat above the Albian/Cenomanian SB in the Carmel region give an Early Cenomanian age; thus this surface is close to the Albian –Cenomanian transition. The Early Cenomanian transgressive system tract (Ce TST-1). The Ce TST-1 and the overlying Early Cenomanian maximum-flooding interval (Ce MFI-1) are in the lower –middle part of the Deir-Hanna Formation of Galilee, and in the Isfiyye and Beit-Oren Formations of Carmel (Figs 9 & 11). Ammonites and planktonic foraminifera suggest that the age of the TST is Early Cenomanian (Fig. 11). The Cenomanian TST-1 is composed of wellbedded to laminated, fine-grained, mostly calcisiltitic bioclastic debris (mudstone), with some pelagic microfossils, pyrite, and glauconite grains (facies type-1; Fig. 4a). The fine-grained bioclastic debris of this facies type, originated from the remote proximal environment, and the pelagic microfossils from the overlying water column. The laminated texture with pyrite and glauconite suggests that deposition took place in proximity to the hypoxic basin floor, presumably where the oxygen minimum zone (OMZ) impinged on the slope. This system is composed of one high-order (UC-7) deepening-upward cycle, bounded at the base by the Albian/Cenomanian SB, and at the top by the Early Cenomanian maximum-flooding interval. Lateral thickness variations of Ce TST-1 across northern Israel, from the Manara section in the NE to the Isfiyye section in the SW, are

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simplified and scaled in Figure 12a. The Ce TST-1 is relatively thin and condensed in NE Galilee (14 m at Manara), thickens toward NW Galilee (50 m at Betzet), and reaches a maximum in the Carmel region (up to 120 m; Bein 1974). Thicker sections of more basinal facies occur towards Carmel in the SW. Early Cenomanian maximum-flooding interval (Ce MFI-1). The second correlated surface, the Ce MFI-1 (Figs 9 & 11), consists of a limestone bed intercalated within basinal chalks of the Deir Hanna Formation in Galilee, and within the Isfiyye-Arkan chalk complex in Carmel. This bed is a few centimetres to a few metres in thickness, overlying basinal thin-beds or laminated beds of the Ce TST-1. Orbitolina sefini occurs in this bed in NE Galilee (identified by M. Simmons, 2007, pers. comm.), and Early Cenomanian ammonites and planktonic foraminifera occur below and above this bed in Carmel, suggesting that it lies within the Early Cenomanian M. mantelli Zone (Fig. 11). In NE Galilee (Manara section) the Ce MFI-1 occurs as a limestone bed rich in large and flat orbitolinids, serpulids and large bivalves (facies type-8; Fig. 4e). In NW and central Galilee (e.g. Kziv-east, Betzet sections; Fig. 9) this bed appears as skeletal or peloidal hiatal concentrations with bioeroded bioclasts and pelagic microfossils (Fig. 4b, c; facies type-7) or as a shell-concentration (Kziv-east section; Fig. 4d; facies type-8). In Carmel and at Mt. Kedumim near Nazerath, the Ce MFI-1 is identified as a highly bioturbated bed or cemented basinal shell bed (‘Beit-Oren’ limestones in Carmel) capped by glauconites (Fig. 9) (Weiler 1968; LipsonBenitah et al. 1997). The matrix of the Ce MFI-1 contains pelagic micro-elements. Bioturbation, colonization by macrobenthos, and formation of lag concentrations by winnowing suggest that the Ce MFI-1 represents oxygenation of the Ce TST-1 submarine hypoxic bottom, accompanied by decreased sedimentation. Facies configuration of the Ce MFI-1 (Fig. 12) is based on the subdivision of this interval into three zones: the horizon of flat orbitolinid of NE Galilee represents the proximal facies zone of the Ce MFI-1

Fig. 4. (Continued) maximum-flooding bed (Ce MFI-1). Lower Deir Hanna Formation, Kziv east section, eastern Galilee. (d) Bivalve shell-bed (facies type-8) of the Early Cenomanian maximum-flooding bed (Ce MFI-1). Lower Deir Hanna Formation, Kziv east section, eastern Galilee. (e) Flat orbitolinids, serpulids and bioclastic debris (facies type-8) in the Early Cenomanian maximum-flooding bed of NE Galilee. Deir Hanna Formation, Manara section, NE Galilee. (f) Distal debrites (facies type-3) between basinal Deir Hanna laminites and Sakhnin Formation dolomites at Sheikh-Danun, western Galilee. Clasts are dolomitized. Matrix is composed of dolomitized Deir Hanna laminites. Breccias of this type appear locally at the base of the Sakhnin Formation in western Galilee. Lower Sakhnin Formation, Sheikh-Danun section, western Galilee. (g) Debrite breccias (facies type-3). Clasts are of peritidal origin, matrix is a textureless dolomite. Sakhnin Formation, Betzet section, NW Galilee. Hammer for scale (circled).

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Fig. 5. (a) Northern slope of Betzet Valley, NW Galilee, near Betzet section. Albian Yagur Formation at base. Above are basinal laminites (facies type-1, indurated chalks) of the Deir-Hanna Formation passing upwards into laminites and bioturbated limestones (facies types-1, 10) of the Yanuch Formation. Two lenticular dolomite bodies (Channels 1 and 2) composed of dolomitized graded grainstones (a.1, a.2), interpreted as turbitidic channel fill

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Fig. 6. Shallowing-upward high-order UC units found in the Cenomanian –Turonian succession of northern Israel. UC-6 is a cyclic unit controlled by repeated bottom stagnation on the mid-ramp zone (sterile laminites) and does not necessarily reflect expansion or contraction of accommodation space.

Fig. 5. (Continued) (facies type-2). They are incised into Deir-Hanna laminites. (b) Well-laminated graded, planar, and ripple cross-stratified dolomitized calcisilts (facies type-2) of the Sakhnin Formation, interpreted as the turbiditic Ta– Td Bouma (1962) divisions. Pen for scale. Sakhnin Formation, Gat section, western Galilee. (c) Synsedimentary foliated (yellowish) shear-zone lithology (facies type-4). Sakhnin Formation, Betzet section region, NW Galilee. (d) Photomicrograph of the shear-zone lithology (facies type-4): d– texture – mudstone fragment in foliated fine-grained dolomite. Shape indicates elongation and rotation owing to sub-horizontal shear. Sakhnin Formation, Betzet section region, NW Galilee. (e) Bedding-parallel shear planes. Shear-zone lithology such as that seen in (c) is found on these bedding planes. Sakhnin Formation, Betzet section region, NW Galilee. (f) Boudins and planar foliation in shear-zone lithology, Sakhnin Formation, Betzet section region, NW Galilee. (g) Steeply inclined slump fold in upper slope laminites (facies type-4), Sakhnin Formation, Mt. Kedumin section near Nazerath, southern Galilee. (h) Pendant calcitic/silicic cement of the Middle Cenomanian sequence boundary (Ce SB-2). Sakhnin Formation, Manara section, NE Galilee. (i) Deformed pisoid at the Ce SB-2. Interpreted as a vadose/calcrete pisoid (facies type-20). Sakhnin Formation, Manara section, NE Galilee. (j) Disolved vadose pisoids (facies type-20) at the Ce SB-2. Sakhnin Formation, Manara section, NE Galilee.

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Fig. 7. Deepening-upward high-order UC found in the Cenomanian– Turonian succession of northern Israel. Early Turonian cycles overlying latest Cenomanian sequence boundary are the UC-7 of the Yirka Formation, the UC-8 of the Daliyya Formation, the UC-9 of the Bina Formation, the UC-11 of the Bina Formation and the UC-12 of the Daliyya and Bina Formations. They are assigned here to OAE-2.

(Fig. 12a). According to Vilas et al. (1995) and Hottinger (1997), flat orbitolinids are more common than conical forms under conditions of water cloudiness, low illumination, and increased water depth. Simmons et al. (2000) and Di Lucia et al. (2007) showed that flat orbitolinids characterize transgressive systems tracts, with the flattest forms found around maximum-flooding surfaces. Thus, flat orbitolinids at MFI-1 at Manara section indicate increased nutrients and phytoplankton at the

surface, causing turbidity and low light levels on the sea floor. Other authors also attributed flat orbitolinids to increased water cloudiness, either owing to terrigenous influx or eutrophic conditions in transgressive contexts (e.g. Immenhauser et al. 1999; Pittet et al. 2002; Al Juboury et al. 2006). The intermediate facies belt of shell-beds and bioclastic fragmental concentrations in NW and central Galilee (Fig. 12b) represents weak bottom-currents and winnowing of fine laminated

THE CARBONATE SYSTEM OF NORTHERN ISRAEL

Fig. 8. Low-order CC-1-8 occurring in northern Israel. They are composed of high-order UC cycles. Schematic and simplified presentation: the exact number of cycles and their thicknesses are shown in Figure 9. Inferred ramp profiles are indicated for CC-1 and CC-2.

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148 R. FRANK ET AL. Fig. 9. Sequence stratigraphic correlation of cyclic units in the Cenomanian–Turonian succession of northern Israel, from NE Galilee to the Carmel region. See Figure 1 for locations of columnar sections. Detailed explanation in text. The Pelech sequence (Seq. 3) is locally exposed in western Galilee (Pelech and Hamra Valley sections). The Middle Cenomanian lowstand system tract (Ce LST), represented in the Sheikh-Danun section, is not expressed in this correlation scheme, but is clearly expressed in the correlation scheme of Figure 10. Large parts of the sequences and system tracts are composed of two cyclic ranks, high-order UC and low-order CC.

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Fig. 10. SW to NE sequence-stratigraphic correlation across the Galilee. Ce RST-1 (HST-1 and FRST-1) at Dishon is proximal, composed of CC-1 peritidal cycles; towards the SW it is much more distal and composed of basinal laminites and mass-transport deposits. CC-2 cycles build the LST in Sheikh-Danun and Kziv-west sections.

sediments of the TST-1, leaving fragmentalbioclastic or peloidal lag concentrations. In the Kziv-east section, the oxic sea-floor was subsequently colonized by bivalves and bioeroding organisms (e.g. Fig. 4d). The lower slope, or basinal facies zone, of the Ce MFI-1 occurs in the Nazareth Hills and Carmel. It is characterized by bioturbated and biodegraded macrobenthic shell material indicating oxygenation of the sea-floor, decelerated sedimentation, and probably also increased flux of nutrients (cf. Kidwell 1986; Hallock 1988). Glauconites at the top of the Ce MFI-1 were formed in this distal zone as hypoxia resumed (cf. Porrenga 1967; Carson & Crowley 1993). The Lower –Middle Cenomanian regressive system tract (Ce RST-1). The Early–Middle Cenomanian system tract corresponds in Galilee to the Upper Deir Hanna limestones and the mid-ramp/peritidal facies of the Sakhnin dolomites, and in the Carmel region to the Zikhron and Arkan Formations. The Ce RST-1 is composed of two parts, the midCenomanian HST (Ce HST-1) and the midCenomanian forced-regressive system tract (Ce FRST-1) (Figs 9, 10 & 11). These systems tracts are combined into a ‘regressive systems tract’ sensu Embry (2002) (Ce RST-1) for the purpose of time-boundary definition and palaeogeographic interpretation, but each part will be also addressed separately. The Ce RST-1 is bounded below by the Ce MFI-1 and above by the mid-Cenomanian

SB (Ce SB-2; Fig. 11), a termination most probably placed within the mid-Cenomanian A. rhotomagense zone. Its age is bracketed by Early Cenomanian Orbitolina sefini in the underlying Ce MFI-1, and by mid-Cenomanian radiometric and ammonite data from overlying beds (Fig. 11). Radiometric dating of pyroclastic rocks above the Ce SB-2 in southern Carmel gives midCenomanian date of 95.4 + 0.5 Ma (Segev et al. 2002; top Zikhron Formation). Ammonites of the A. jukesbrownei zone (Kennedy & Jolkicˇev 2004; Fig. 11) occur in chalks from above the Ce SB-2 in Galilee (Calycoceras sp. identified by Z. Lewy, pers. comm. 2007, and Protacanthoceras sp. in Freund 1958). The mid-Cenomanian HST (Ce HST-1) forms the lower part of the Ce RST-1 (Fig. 11). In most of Galilee and southern Carmel, Ce HST-1 is a progradational type-1 low-order shallowing-upward cycle (CC-1), 60– 200 m in thickness. The vertical architecture of this cycle and the arrangement of its facies components on the ramp are shown in Figure 8a. The mid-ramp zone of this cycle (facies type-12 in Table 1) is highly productive, providing lime mud downslope to the outer ramp/basin by wave and current action and by suspension, and also upslope to the inner ramp. A sharp boundary within CC-1 separating lower basinal laminites (facies type-1) from overlying bioturbated midramp facies (Fig. 8a) reflects the transition from the zone of impingement of the OMZ to the mixing zone. This boundary represents a marked

150 R. FRANK ET AL. Fig. 11. Correlation of sequences and systems tracts with chronostratigraphic controls (cf. Gradstein et al. 2004) and event chronology across northern Israel and adjacent regions. T1 to T7 are local ammonite zones of the Upper Cenomanian– Turonian succession in Israel (Freund & Raab 1969). Lithostratigraphy follows Sneh et al. (1998) for Galilee and Segev & Sass (2006) for the Carmel region. Biostratigraphic control based on planktonic foraminifera biostratigraphy in Carmel (Lipson-Benitah et al. 1997), ammonite biostratigraphy in Carmel (Kashai 1966; Lewy & Raab 1978) and Galilee (Freund 1958; Glikson 1966; Freund & Raab 1969), rudist biostratigrapy in Carmel and Galilee (Freund 1965; Buchbinder et al. 2000), identification of orbitolinids (M. Simmons 2007, pers. comm.), and other benthic foraminifera. Radiometric dating of volcanic rocks in Carmel are by Segev et al. 2002 [Dating of pyroclastics overlying Yagur Formation (97.1 + 1.7 Ma) omitted owing to wide error-bar]. Comparisons of bounding surfaces are to the northern Negev (Lewy 1990; Buchbinder et al. 2000), west-central Jordan (Schulze et al. 2003), Sinai (Bauer et al. 2003), Oman Platform (Philip et al. 1995; Van Buchem et al. 1996, 2002) and parts of the Arabian plate (Sharland et al. 2001).

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Fig. 12. (a) SW to NE facies-thickness changes in the Ce TST-1 across northern Israel. Note system-tract thickening and MFI-1 deepening toward Carmel. (b) Palaeogeographic reconstruction of Early Cenomanian maximum-flooding interval.

ecological transition between non-productive sea-floor in the OMZ to highly productive mid-ramp in the mixing zone, a transition that could generate distally steepened ramp morphology. The mid-Cenomanian forced-regression system tract (Ce FRST-1) is the upper part of Ce RST-1 (Figs 10 & 11). Its base corresponds to the BSFR (sensu Hunt & Tucker 1992, 1993), upon which mass-transport deposits were emplaced (mapped as Sakhnin Formation dolomites). The upper boundary corresponds to the mid-Cenomanian SB (Ce SB-2) at the transition to overlying lowstand or transgressive deposits (Fig. 10). Mass-transport features of the Ce FRST-1 include debrite breccias (Fig. 4f, g; facies type-3), channelized and nonchannelized turbidites (Fig. 5a, b; facies type-2), and shear-planes and translational slides (Fig. 5c, g; facies type-4). Facies, geometries and sedimentary configuration of the Ce RST-1 (HST-1 and FRST-1) are presented in Figure 13. The type-1 low-order cycle of the Ce HST-1 (CC-1 in Fig. 8a) is present in most of Galilee and in the southern Carmel region (peritidal zone in Fig. 13a). Blocks and skeletal grains derived from the ramp were redeposited as forced-regressive debris-flow breccias, deformed and sheared slabs, and channelized and nonchannelized turbidites in western Galilee (facies types 2, 3 and 4). Toward NE Galilee the peritidal ramp passed distally into a productive mid-ramp zone (Fig. 13a, b). In southern Carmel, the equivalent mid-Cenomanian peritidal ramp of the Zikhron Formation passed toward the north into

mid-Cenomanian basinal deposits of the Arkan Formation (Figs 9 & 13c; see also Segev & Sass 2006). This transition is associated with ramp disintegration and formation of synsedimentary breccias. These south to north facies changes in Carmel are reflected in the transition from the Hotem –Carmel section to the Rakit–Beit Oren section (Figs 9 & 13c). A NE–SW transverse in the Ce RST-1 across Galilee is presented in Figure 13d. The peritidal system of Galilee passes gradually toward the Mt. Kedumim section at the south into thicker and deeper mid-outer-ramp deposits with some slumps and debrites (facies type-3). Farther to the south, this facies shallows once more, as represented by the peritidal facies of the southern Carmel region (Zikhron Formation of the Hotem –Carmel section). A more distal expression of this south – north transition is shown in Figure 13c. In summary, facies distribution in the Ce RST-1 indicates peritidal palaeo-highs prevailing in most of Galilee and in southern Carmel. The slope of the shallow ramp of Galilee faced north, west and south. The slope of the shallow ramp of southern Carmel faced north (and west; Bein 1974). These two palaeo-highs were separated by a deeper zone of increased subsidence that extended roughly east –west from central Carmel to southern Galilee. The mid-Cenomanian sequence boundary (Ce SB-2). The mid-Cenomanian SB separates the Ce RST-1 from the overlying TST or LST (Figs 9, 10 & 11). Pyroclastic rocks above this SB in southern Carmel (top Zikhron Formation; HC section in

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Fig. 13. Ce RST-1. (a) Platform-basin configuration of Ce RST-1 with locations of cross sections shown in (b), (c) and, (d). Ce HST-1 forms autochthonous mid/peritidal ramp zones in Galilee and southern Carmel. The peritidal zone of Galilee passed toward northern Galilee into a deeper mid-ramp zone. Mass-transport deposits of the Ce FRST-1 were deposited on the slope. (b) Facies-thickness changes across the A–A0 –A00 line. Note ramp deepening toward the north (A0 –A) and mass-transport complex emplaced toward the west (A0 – A00 ). (c) Facies-thickness changes across line C–D. Progradational ramp system of the southern Carmel (CC1 cycle) deepens toward the north. (d) Facies-thickness changes across line A– B. The peritidal shallow-water ramp system of Galilee passes into thicker and deeper facies in southern Galilee, and further to the south into shallow-water ramp system of southern Carmel.

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Fig. 9) were dated to 95.4 + 0.5 Ma (Segev et al. 2002), corresponding to the Middle Cenomanian A. rhotomagense ammonite zone. The ammonite Calycoceras sp., recovered from above the Ce SB-2 at the Manara section, supports this age also in Galilee (Figs 9 & 11). Diagenetic events connected with the SB affected the upper surface of the Ce HST-1. On this surface at the Manara section is a thin reddishblack ferruginous pavement, and some special diagenetic micro-features penetrate a few centimetres beneath the ferruginous crust. These features include sphaeroidal calcitic/silicic or partly silicified vadose pisoids, empty or infilled dissolution vugs, pore-filling karstic cements and concentric pendant calcites (Fig. 5h –j; facies type-20). These are vadose diagenetic features indicating subaerial exposure of the Ce HST-1 ramp. Teepee structures reported by Bogoch et al. (1994) from the Sakhnin Formation/Bina Formation transition in eastern Galilee (Kadarim region, Fig. 1) are additional evidence for subaerial exposure of the proximal ramp. The thin reddish-black mineralized surface that paves this zone of vadose vugs and cements, is interpreted as a transgressive feature (facies type-9) of the following TST. Rapid sea-level rise at the TST phase, and flooding of the SB led to sediment starvation, condensation, and mineralization. The mineralized crust sealed open cavities and inhibited sea-water percolation. This interpretation is supported by Longman (1980) who proposed that preservation of vadose cements requires rapid transgression or subsidence. Ce HST-1 (CC-1 cycle) is absent in western Galilee and in central-northern Carmel (e.g. Yanuch, Rakit/Beit-Oren sections; Fig. 9). The Ce SB-2 is represented in these regions by truncation of bedded/laminated chalks of the Deir Hanna and Arkan Formations, but with no evidence for subaerial exposure. Bein (1974) showed this truncation at the top of the chalk complex of the Carmel region (top Arkan Formation) and a coeval truncation surface is found in western Galilee at the transition from the Deir Hanna chalk laminites to the bioturbated limestones of the Yanuch Formation (Yanuch section, Fig. 9). These truncations represent submarine erosion of basinal deposits on the distal outer slopes of the western Galilee and Carmel regions and represent the distal part of the Ce SB-2 (Fig. 13a, c).

Sequence 2: late –Middle to Late Cenomanian The Middle Cenomanian lowstand system tract (Ce LST). The Middle Cenomanian lowstand system tract is shown in Figure 10. Breccias of the Ce FRST-1 (facies type-3 in Table 1) and Ce SB-2 features in

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western Galilee (Sakhnin Formation of the SheikhDanun and Kziv-west sections, Fig. 10) are overlain by a number of subtidal UC-3 shallowing-upward cycles, forming a single CC-2 low-order cycle (Fig. 10). Such shallowing-upward cycles, directly overlying a SB, reflect progradation at the initial stage of relative sea-level rise, characterizing a lowstand systems tract (cf. Hunt & Tucker 1993). The late –Middle to Late Cenomanian transgressive system tract (Ce TST-2) and the LateCenomanian maximum-flooding interval (Ce MFI-2). The Ce TST-2 is composed of limestones and chalks bounded at the base by the Ce SB-2 or transgressive surface and at the top by a maximumflooding interval (Ce MFI-2). It corresponds to bioturbated limestones of the lower Yanuch Formation of western Galilee (Yanuch section), to limestones of the lower Bina Formation of NE Galilee (Dishon section), and to bedded chalks at Manara (part of the Yanuch Formation, but considered Deir Hanna Formation by Kafri 1991; Sneh & Weinberger 2003) (Figs 9, 10 & 11). Ammonites of the late Middle Cenomanian Acanthoceras jukesbrownei zone were recovered from the lower part of this systems tract in northern Galilee (Fig. 11) but proximal sections are more poorly constrained: benthic foraminifera from the proximal part of the Ce TST-2 at the Dishon section (Fig. 9) belong to the Pseudorhapidionina dubia total range zone (Fig. 11), spanning the Middle to Late Cenomanian (Aguilera-Franco 2001, 2005). Thus, the correlation of proximal to distal sections is based on similar facies patterns and the comparison of bounding surfaces, and supported by low-resolution biostratigraphic correlation. In both proximal sections (e.g. Dishon) and distal sections (e.g. Manara), the Ce TST-2 begins with relatively low rates of carbonate deposition, or submarine omission. Slow deposition in proximal settings is expressed by the growth of heavily micritized grains, microbial lumps, coated grains and oncoids (base Ce TST-2 at Dishon section; facies type-18). Submarine omission in distal settings is expressed by the transgressive ferruginous pavement of the Ce SB-2 at Manara (facies type-9; see above). The facies configuration of the Ce TST-2 is shown in Figure 14. The Ce TST-2 is thin and proximal at Betzet and Dishon sections. At Betzet it is composed of mid-ramp UC-6 redox cycles and at Dishon it is composed of lagoonal to outer-ramp UC-8 deepening upward cycle. The Ce TST-2 becomes much more distal toward the Carmel region in the SW, and also toward northern Galilee. In the Carmel region, it is represented by relatively condensed fine grained basinal pelagites (facies type-6 at Oren-Valley section). In northern

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Fig. 14. (a) SW–NE cross-section along line A –B from central Carmel to the Manara section, NE Galilee, showing facies-thickness changes in the Ce RST-1 and overlying Ce TST-2. Note thickness variations (in metres) of the two superimposed system tracts. Ce TST-2 is relatively condensed and proximal in Galilee and more distal and deeper in Carmel and NE Galilee. Also note rapid thickening of the Ce TST-2 toward the north that could be related to downfaulting. (b) Sedimentary/structural configuration of northern Israel in the Ce TST-2. The Galilee palaeo-high is bounded by a basinal chalky facies belt in the north and by outer-ramp to basinal facies in the south; boundary to the basin in the north may correspond to a normal fault.

Galilee (Kedesh and Manara sections) it is relatively thick (65 –70 m) and composed of laminated calcisiltite chalks (facies type-1). The rapid thickening of the Ce TST-2 toward northern Galilee, from the Dishon section towards the Kedesh and Manara sections (Fig. 9), may reflect downfaulting to the north (Fig. 14). A deep basinal facies belt extended from the Kedesh and Manara sections in the east toward the Sarach Valley and Adamit in the west (Fig. 14b). The maximum-flooding interval terminating the Ce TST-2 (Ce MFI-2) was recognized on the Galilee palaeo-high, where it forms well-laminated pelagites (facies type-6) at the termination of the deepening-upward cycles UC-8 and UC-10 (Fig. 7) of the Dishon and Yanuch sections.

Late Cenomanian regressive system-tract (Ce RST-2). Included in the composite Ce RST-2 are the Late Cenomanian highstand and the Late Cenomanian forced-regressive systems tracts (Ce HST-2 and Ce FRST-2; Fig. 9, 11). The Ce RST-2 is placed in the Late Cenomanian (Fig. 11), as it is bounded at the base by the Late Cenomanian maximumflooding interval (Ce MFI-2) and at the top by the Late Cenomanian SB (Ce SB-3 of western Galilee; Fig. 11). The Late Cenomanian highstand systems tract (Ce HST-2) forms the lower part of the Ce RST-2. It consists of parts of the Yanuch and Bina Formations in the Galilee, and the lower part of the Muhraqa Formation at Oren Valley section in the western Carmel region (Figs 9 & 11). Latest

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Middle Cenomanian ammonites of the previous TST, and the occurrence of the benthic foraminifera Cisalveolina fallax in the upper part of the Ce HST-2 at the Kishor section, place this system tract in Late to latest Cenomanian. In western Galilee the Ce HST-2 is topped by a basal surface of forced regression (Ce BSFR-2; Fig. 9). The Ce HST-2 is composed of a type-2 loworder cycle (CC-2) forming an aggradational or poorly progradational homoclinal ramp (Fig. 8b). In central and NW Galilee the Ce HST-2 is thin and shallow, composed of peritidal UC-1 cycles (e.g. at the Betzet section). This peritidal zone passes into thicker and deeper UC-3 subtidal cycles (outer ramp to shoreface) toward the south (Yanuch and Beit Ha’Emek sections) and NE (Dishon section). These thickness-facies changes show that central Galilee was peritidal and elevated, while to both the south and north, the ramp was deeper and more section accumulated. Further toward the south, in Carmel (Oren Valley section; Fig. 9), this system is composed of deep basinal fine-grained bioclastic pelagic ooze in the Lower Muhraqa Formation (facies type-6). The Late Cenomanian forced-regressive system tract (Ce FRST-2) terminating the Ce RST-2 in the regions of western Galilee and Carmel (Figs 9 & 11) is corresponding to parts of the Yanuch and Muhraqa Formations. In western Galilee, spectacular calcarenitic clinoform units (facies type-5) occurs from the Kishor section to the Shagor canyon, 5 km to the south (Figs 1, 15a & 16a). The clino-beds composing this unit dip 228 –358 to the west –SW, and the clinoform unit thickens rapidly toward the south, from 0 m at Beit-Ha’Emek section, 35 m at Kishor section, to more than 100 m in Hamra section (Figs 9 & 15a). The base of the clinoform body is a sharp, erosive contact, considered a basal surface of forced regression (BSFR-2 in Figs 9 & 15a). The top of the clinoform unit is truncated by a subaerial unconformity (Fig. 16b, c) onlapped by latest Cenomanian limestones (Fig. 17). In the Carmel region c. 40 km to the south –SW, the Ce FRST-2 is composed of 25– 30 m of thin-bedded graded and rippled calcarenites and calcisiltites in the middle part of the Muhraqa Formation (Figs 15b, 16d, e). The unconformity-bounded clinoform unit of western Galilee, with subaerial erosion at its top, is a forced-regression feature (cf. Plint 1988; Hunt & Tucker 1993; Posamentier & Morris 2000). The rapid thickening of the clinoform unit toward the south, from the Beit-Ha’Emek section to the Hamra section, is explained by downfaulting. Accommodation space created on the hanging walls was filled by steeply dipping calcarenites (Fig. 15a). The calcarenitic grains originated from shoreface erosion of HST-2 deposits (UC-3 cycles)

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on the nearby shelf to the north, in the region of the Beit-Ha’Emek and Yanuch sections. The calcarenites were swept off-shelf toward the south –SW across a fault scarps developed on the south –SW facing Galilean slope (Fig. 15a, b). The toe of slope or basin can be found c. 40 km to the south – SW in distal Carmel, where thin-bedded calcarenites and calcisiltites were deposited as a turbidite sheet composed of partial Bouma (1962) sequences within the Muhraqa Formation (Figs 15b, 16d, e).

Sequence 3: Late Cenomanian SB (Ce SB-3) and Pelech sequence In the Hamra Valley of western Galilee, near the village of Pelech (PL section in Fig. 1), a 35 m section of well-bedded limestones overlies the Ce RST-2. This succession is bounded at the base and top by SBs, thus forming an independent sequence, termed the Pelech sequence (Fig. 17). The basal discontinuity is the Late Cenomanian SB (Ce SB-3), a ferruginous marl bed penetrated by decimetre-scale iron concentrations interpreted as root casts. The upper discontinuity is the latest Cenomanian SB (Ce SB-4), an irregular surface with karst features (Fig. 16c). The Pelech sequence is composed of fine-grained pelagic packstones, skeletal wackestones, or sterile mudstones. The detailed lateral relations between the Pelech sequence and adjacent facies units are not fully presented here, but this locally exposed sequence onlaps to the west over the forced-regressive clinoform unit of Ce FRST-2 (Fig. 17). The benthic foraminifera Cisalveolina fallax from the underlying HST, and caprinid rudist fragments from the Pelech sequence, suggest a latest Cenomanian age, probably corresponding to the M. geslinianum ammonite zone (cf. biostratigraphic scheme of Aguilera-Franco et al. 2001). The Pelech sequence is composed of finegrained pelagic outer-ramp packstones and midramp wackestones (facies types 10 and 17; Table 1). The SB at its base reflects sea-level fall at the end of the Ce RST-2, and the Pelech sequence represents the subsequent rise. Most of the Pelech sequence was removed by erosion during the latest Cenomanian exposure of the latest Cenomanian SB (Ce SB-4). The latest Cenomanian sequence boundary in northern Israel (Ce SB-4). The latest Cenomanian SB occurs between Cenomanian and Lower Turonian strata in northern Israel. In western Galilee, the clinoform unit (Ce RST-2), and the Pelech sequence are truncated by a reddish, encrusted surface with pedogenic pisoliths, circumgranular cracks, alveolar septal fabrics and karst features (facies type-9; Figs 16b, e & 17). In the Carmel

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Fig. 15. Ce RST-2 and Tu TST. (a) facies-thickness changes in the Ce FRST-2 (Ce HST-2 and Ce FRST-2) and Tu TST showing a SW– NE cross-section along line A– B from the Oren Valley section of western Carmel to the Dishon section of NE Galilee. Note thickness variations (in metres) of the superposed system tracts. All systems tracts show that part of Galilee is proximal and condensed with respect to a thicker and deeper facies to the NE and SW. The Ce FRST-2 of western Galilee consists of thick (35–100 m) and steep (22–358) calcarenitic clinoform unit rapidly thickening to the south. Thickness variations suggest that this calcarenitic clinoform body reflects a syndepositional response to block downfaulting toward the south (see text for details). (b, c) Sedimentary/structural configurations in the Ce RST-2 and Tu TST.

region, the Cenomanian–Turonian transition lies within the Muhraqa Formation complex, which has individual caprinid rudists in its lower part and individual hippuritids in the upper part (Buchbinder et al. 2000). In the Muhraqa Formation of the Isfiyye

section of the Carmel (Fig. 1) a ferruginous crust (facies type-9) overlies Upper Cenomanian forcedregressive debrite breccias (Fig. 18). In western Galilee, subaerial exposure accompanied by pedogenesis and karstification (facies

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Fig. 16. (a) Well-stratified clinoform unit at Kishor dipping 20–258 to the SW. Clinobeds downlap on erosion surface (BSFR). Underlying beds are tilted 10– 128 to the NE. Yanuch Formation, Kishor section, western Galilee. (b) Palaeo-calcrete fabric from top Mt. Gamal clinoform unit. Left: Dense micrite with a complex network of cracks forming irregular micrite nodules. Right: heterogeneous fabric with sub-spherical concentric pisoids (rectangle). Top Yanuch Formation, Mt. Gamal section, western Galilee. (c) Highly irregular karst surface at the top of the Yanuch Formation interpreted as ancient lapie´s. Above is highly weathered soft horizon, either basal Yirka Formation or a palaeosol. Top Yanuch Formation, Hamra section, western Galilee. (d) Well-stratified or thin-bedded limestones of the Muhraqa Formation. These beds are composed of peloidal-bioclastic grainstones that are commonly graded. Muhraqa Formation, Oren Valley section, western Carmel. (e) Turbidite from the middle Muhraqa Formation. Sample is graded at the base (Ta), passed upwards to sub-planar lamination (Tb) and is rippled at top (Tc). Muhraqa Formation, Oren Valley section, western Carmel.

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Fig. 17. The Pelech sequence of the Hamra Valley, western Galilee (region of Hamra Valley section in Fig. 1) as seen by correlation between two adjacent sections, the Pelech (PL) section in the east and the Hamra (HM) section in the west. For location of these sections see Figure 1. The Pelech sequence, recognized only in western Galilee, is composed of well-stratified wackestones and packstones of outer- to mid-ramp origin. It is bounded by discontinuities and onlaps the forced-regressive clinoform unit of the Hamra Valley from the east. Note faulted clinoform unit and amalgamation of the Late Cenomanian SB (Ce SB-3) with the younger latest Cenomanian SB (Ce SB-4).

type-20, Table 1) affected Upper Cenomanian strata prior to Early Turonian transgression (Fig. 17). On the other hand, Upper Cenomanian strata in the Isfiyye section of the Carmel region lack subaerial exposure features and the thin ferruginous crust is interpreted as submarine omission surface at the beginning of the subsequent Early Turonian transgression. Therefore, at the end of the Cenomanian, relative sea-level fall resulted in exposure in Galilee while the Carmel region remained submerged. Buchbinder et al. (2000) placed the latest Cenomanian SB in northern Israel between the Upper Cenomanian ‘Yanuch and Lower-Muhraqa’ Formations and the Turonian ‘Yirka and UpperMuhraqa’ Formations (Fig. 2). A prolonged hiatus spanning 1.25 ma was indicated, extending from the upper part of the Late Cenomanian N. vibrayeanus zone, until the middle part of the Early Turonian W. coloradoense zone. In the present study, the benthic foraminifera Cisalveolina fallax was recovered from below this SB in western Galilee (Fig. 11) and ammonites of the Turonian C. securiforme Zone (T4 of Freund & Raab 1969) occur c. 20 m above it. A C. fallax zone was defined by Saint-Marc (1974) as spanning the latest Cenomanian to Early Turonian. Therefore, the hiatus in this region may not be as extensive as the previous estimation.

Sequence 4: Early and Middle Turonian Turonian transgressive system tract (Tu TST). The Tu TST (Figs 9 & 11) is bounded at the base by the latest Cenomanian SB (Ce SB-4) and at the top by a maximum-flooding interval. In the Carmel region it corresponds to part of the Upper Muhraqa Formation limestones and the Daliyya Formation, and in Galilee to the uppermost Yanuch Formation, Yirka Formation and part of the Bina Formation. The Tu TST mostly consists of deep basinal facies, turbidites, hiatal concentration (sensu Kidwell 1986), and hypoxic laminites, but shoreface deposits are present as well (facies types 1, 2, 7, 11 and 15 in Table 1). In the Carmel region, the Tu TST begins with deepening upwards cycles from the lower shoreface to the basin (UC-12 cycles forming a single CC-7 cycle, Fig. 8). Towards southern Carmel these cycles pass into deepening-upward cycles with lagoonal aspects (UC-8 cycles forming one CC-8 cycle, Fig. 18). This facies transition indicates that in the Lower Turonian shallow ramp conditions prevailed in the southern Carmel region and deeper outerramp to basinal conditions prevailed toward the north. In western Galilee, the Tu TST similarly begins with basinal marls (facies type-1) of the Yirka Formation (e.g. at Hamra section, Fig. 9)

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Fig. 18. Correlation between the Rakefet section (RKF) and the Isfiyye section (ISF) located in the Carmel region (for location of these sections see Fig. 1). Datum is the latest Cenomanian SB (Ce SB-4). Deepening-upward CC-8 cycles are noted in the Rakefet section of southern Carmel and CC-7 cycles are noted in the Isfiyye section at the north. Note that abundant shallow-water microfauna is present at the CC-8 cycle at Rakefet, compared to deep water pelagic microfauna (especially planktonic foraminifera and calcisphaeres) in the CC-7 cycle at Isfiyye.

and with retrogradational shoreface grainstones passing upward to turbidite/pelagic marl farther to the north (facies types 2 and 15). In other parts of Galilee the Tu TST corresponds to hypoxic midto outer-ramp deepening-upward cycles in the Bina Formation (UC-9, UC-11 and UC-12). Geometry and facies of the Tu TST are shown in Figure 15a, c. In central and NW Galilee the Tu TST is condensed (mid- to outer-ramp UC-9 cycle). At the Dishon section in the NE, it is a much thicker deepening-upward cycle (55 m) composed mainly

of mid-outer-ramp laminites (CC-7; facies type-11). At the Yanuch section in the SW the Tu TST is composed of a 30 m thick deepening-upward cycle (30 m) (CC-6), and at the Hamra section to the SW it is composed of 80 m of basinal marls of the Yirka Formation (facies type11). Composite facies configuration of the Tu TST is shown in Figure 15c. The thin condensed cycle of central and NW Galilee reflects condensation on a palaeohigh. This zone was bounded to the north and south by deeper and more rapidly subsiding

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regions. Outer-ramp laminites of NE Galilee onlapped the Galilean palaeo-high from the north, and basinal marls onlapped the steep clinoformic slope of western Galilee from the south.

Discussion Tectono-sedimentary framework of the central Levant margin, northern Israel The mid-Cretaceous ‘Levantine hinge-belt’ (see above; Fig. 1) extended from the southern coastalplain of Israel northwards to the Carmel region (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977), and also formed a coast-parallel, north– south striking facies belt in northern Lebanon (Walley 1998). Galilee and southern Lebanon were considered by Walley (1998) as an interruption of this trend, caused by the extension of the Palmyride basin to the SW. The data and sequence stratigraphic interpretations presented here support the overall palinspastic model presented by Walley (1998). We have shown that a large part of Galilee was structurally elevated and bounded to the south and north by deeper, more rapidly subsiding regions. This tectono-sedimentary configuration is consistent throughout the studied succession, expressed in system tracts Ce TST-1 (Fig. 12), Ce RST-1 (Figs 13 & 14), Ce TST-2 (Fig. 14), Ce RST-2 (Fig. 15) and Tu TST (Fig. 15). There is also some evidence that the southern Carmel region was elevated with respect to a deeper, more subsident zone in central-northern Carmel, a trend expressed in system tracts Ce RST-1 and Tu TST. Data from the Ce RST-1 at Mt. Kedumin and southern Carmel (Figs 9 & 13) and data from the Tu TST in southern Carmel (Figs 9, 15c & 18) suggest that a subsiding trough, also trending east– NE, separated the shallow and more elevated southern Carmel region from a likewise shallow and more elevated Galilee. This trough extended from central-northern Carmel toward the southern part of the Galilee. Thus, the Cenomanian –Turonian sedimentary configuration in northern Israel is characterized by facies transitions indicating that the north– south striking hinge-belt (depositional strike) of the Levant margin (Fig. 1) shifted to the east –NE, northwards of the southern Carmel region (Hotem Carmel section, Fig. 1). Structural control on facies and systems-tract geometry extends beyond Galilee. Data of Saint-Marc (1972, 1974) from Mt. Sannine in central Lebanon show that the Cenomanian succession further thickens towards central Lebanon (Fig. 19). This thickening represents the northward continuation of the subsidence trend recorded here for northernmost Galilee.

Fig. 19. SSW–NNE cross-section from central Carmel, via Galilee, towards Mt. Sannine, central Lebanon, showing thickness variations of the Cenomanian succession. Note significant thickening from Galilee toward Mt. Saninne at the north. Data from Mt. Saninne are from Saint-Marc (1972, 1974).

The east –NE structural trend characterizing northern Israel in the Cenomanian and Turonian stages corresponds to other large-scale trends in the Levant. The SW –NE Palmyride trend in Syria was active, at least since the Triassic and possibly even Late Palaeozoic period (Ponikarov et al. 1967; Brew et al. 2001). Similarly, Ferry et al. (2007) showed that depositional strikes represented in the Lower Aptian Jezzine Formation and Upper Albian Niha Formation of Lebanon trend east –NE. This trend is also sub-parallel to depositional strikes described from mid-Cretaceous carbonate platforms in the Levant such as those studied by Carter & Gillcrist (1994) from the northern margin of the Arabian Plate (SE Turkey), by Schulze et al. (2005) from west-central Jordan, and by Bachmann & Kuss (1998) and Bauer et al. (2003) from northern Sinai. These depositional trends have not been described from the mid-Cretaceous of Israel. The integration of all these trends into a complete framework of the northern Arabian margin

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may be premature. Clearly, a comprehensive compilation of all thickness and facies data is necessary to complete the tectono-sedimentary configuration of the Levant margin in the Mesozoic.

Controlling mechanisms and correlation across the Arabian Plate and beyond Regional significance of the Albian–Cenomanian SB of northern Israel. The Alb/Ce SB-1 of northern Israel was described as a subaerial unconformity surface (Folkman 1969; Kafri 1986). The Alb/Ce SB-1 in central and southern Israel corresponds to a discontinuity characterized by a shell-bed of Pycnodonte vesiculosa (Lewy & Weissbrod 1993; Braun & Hirsch 1994) but evidence for subaerial exposure was not reported. The Albian– Cenomanian boundary in central Jordan (Abed 1984) was described as a subaerial erosion surface paved by conglomerates, and in Lebanon (Ferry et al. 2007), as an emergent surface. Approximately coeval hiatuses and relative sea-level fall have been recorded globally, for example, in SW England (Simmons et al. 1991), the Anglo-Paris basin, and as far as Crimea, Kazakhstan, Turkmenistan and Iran (Gale et al. 1996), as well as from the North American western interior basin (Gro¨ke et al. 1998). Lower part of Cenomanian sequence-1: eustatic and environmental controls. A plate-wide earliest Cenomanian maximum-flooding surface K120 was described by Sharland et al. (2001) at 98 Ma. It was recognized in SE Turkey, Lebanon, eastern Syria, Iraq, Iran, Kuwait, Saudi Arabia, Qatar, UAE, Oman and Yemen. In the Arabian Plate, the K120 interval largely corresponds to organic-rich bituminous shales, or carbonate mudstones and wackestones. A section consisting of organic-rich shales from the base of the Natih-E member in Oman was chosen as a ‘reference section’ of this maximum-flooding event. The Early Cenomanian maximum-flooding episode of northern Israel (Ce MFI-1) correlates with K120 (Fig. 11). This is indicated by the presence of Orbitolina sefini in this interval at the Manara section and occurrences of Early Cenomanian ammonites below and above it in the Carmel region (Fig. 11). However, the Ce MFI-1 has a different sedimentological expression than the K120. It corresponds to a sharp facies transition from stagnant basinal laminites (Ce TST-1) into bioturbated –bioeroded shell-beds, bioclasticpeloidal hiatal-concentrations, glauconite-rich horizons and flat-orbitolinid limestones. This facies transition represents weak current-winnowing and submarine bottom reworking, oxygenation of hypoxic sea-floor, and possible elevation of nutrient levels in the water body. A similar timeequivalent shell bed occurs within the well-stratified

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basinal succession of the Ein-Yorkeam Formation in northern Negev, southern Israel (Figs 1 & 11). In north-central Jordan, Schulze et al. (2003) correlated the K120 to the surface separating Lower Cenomanian marls of the Naur-A member from shallower rudist buildups and shallowing upward cycles of the Naur-B member. Other Early Cenomanian maximum-flooding episodes from Europe (Caus et al. 1997; Wilmsen 2008) are somewhat younger than the K120 of the Arabian plate but may provide some explanations for the sedimentological differences between the shelly bioclastic Ce MFI-1 of northern Israel and the bituminous K120 described from the Arabian plate. In the Sopeira Formation of SE Spain, Caus et al. (1997) showed basinal deposits capped by a maximum-flooding bed near the transition from the Early Cenomanian to the Middle Cenomanian. This is also an Early Cenomanian maximumflooding event with nearly identical facies to Galilee, for example, a proximal facies with abundant orbitolinids. Another maximum-flooding event was described by Wilmsen (2008) in the Early Cenomanian mid-M. dixoni zone of northern Germany. It corresponds to inoceramid shell-beds (Schloenbachia/Inoceramus virgatus bioevent) widely distributed in NW Europe. Despite uncertain timing or diachrony, Lower Cenomanian shell-beds of Israel and Europe were similarly affected by mechanisms of bottom reworking near or below storm wave base. Therefore, while Sharland et al. (2001) attributed the K120 event in the Arabian plate to eustatic rise and local subsidence, the Cenomanian maximum-flooding intervals in Israel and Europe record an additional palaeo-environmental input. End of Cenomanian sequence-1: eustatic and palaeoenvironmental controls. The Lower to Middle Cenomanian highstand system tract of northern Israel (Ce HST-1) represents a progradational ramp system. In most of Galilee and southern Carmel it is a type-1 low-order cycle (CC-1; Fig. 8a) characterized by complete filling of the accommodation space to peritidal depths as a consequence of two synchronously operating factors: (a) the development of an effective, mollusc-dominated carbonate-production zone on the mid-ramp that supplied carbonate mud both upslope to the innerramp zone and downslope to the basinal areas (Fig. 8a, facies type-12 in Table 1), and (b) reduction of the accommodation space owing to eustatic fall. The development of a mollusc-dominated carbonate-production zone on the mid-ramp indicates mesotrophic conditions favoured by rudists and other molluscs. Termination of this highly productive carbonate ramp was first by gravity collapse represented by the Middle Cenomanian FRST, and

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subsequently by ramp emergence forming the Middle Cenomanian SB. The ramp system of this stage represents an Early to Middle Cenomanian regressive event in northern Israel. This Early to Middle Cenomanian regressive event of northern Israel can be identified in the northern Negev of southern Israel (Fig. 11). Lewy (1990) described an Early to Middle Cenomanian regressive event (Zafit Formation), bounded at the top by subaerial exposure surface (Lewy & Avni 1988) and overlain by transgressive deposits of the Avnon Formation. The age and facies stacking pattern of the Zafit Formation are similar to those of the CC-1 cycle of northern Israel (Fig. 8a). An Early to Middle Cenomanian regressive event was also recorded from other locations in the Arabian Platform (Fig. 11) and Europe. In west central Jordan, Schulze et al. (2003, 2004) considered the rudist-bearing shallowing-upward cycles of the Naur-D member as a regressive HST. The Naur-D member is capped by a bored ferruginous hardground (SB CeJo2) overlain by the drowning succession of the Fuheis Formation. Philip et al. (1995) recognized an Lower to Middle Cenomanian shallowing-upward HST pattern in the Natih-E member of the Natih Formation in the Oman Platform. Van Buchem et al. (2002) assigned this interval in the Natih Formation to sequence-1 of the Cenomanian, terminated by a Middle Cenomanian exposure surface (type-1 SB) 10 m below top Natih-E member, and Gre´laud et al. (2006) associated this boundary to eustatic fall, incisions, and deposition of forced-regressive wedges. Caus et al. (1997) considered the Middle Cenomanian part of the Santa-Fe limestones of NE Spain as a progradational system redeposited as breccias in the basin and topped by a subaerial type-1 SB at the Middle/Late Cenomanian boundary. In the deeper marine environments of the Anglo –Paris basin, Robaszynski et al. (1998) described a major fall in sea-level in the late Early Cenomanian, in the later part of the M. dixoni zone, followed by a strong transgression in the earliest A. rhotomagense zone, somewhat earlier than the late Middle Cenomanian transgressive episode recorded in northern Israel. Middle Cenomanian hiati, omission, condensation, and deep erosion were reported across western Europe, England, and Ireland and the Caucasus, as well as from the western interior seaway of America (e.g. Baraboshkin et al. 1998; Hancock 2003; Tro¨ger 2003). These studies suggest that the pattern of sedimentation in the late Early to Middle Cenomanian in marginal marine environments of the Arabian plate and parts of Europe is regressive –progradational, terminating by subaerial or subaquatic omission. In deeper environments sedimentation is associated with condensation and hiati, and redeposition of shallow-water deposits

(northern Israel, Spain and England). Hancock & Kaufman (1979), Van Buchen et al. (2002) and Hancock (2003) associated this Middle Cenomanian regressive event with an eustatic sea-level fall. The beginning of sequence-2: eustatic and palaeoenvironmental controls. Sequence-2 of northern Israel begins with transgression in the late Middle Cenomanian (Ce TST-2). This interval is represented by deposition of pelagites in basinal regions and by starvation, non-deposition, and development of a proximal deepening-upward cycle on the palaeo-high of Galilee. This transgressive event can be correlated across Israel and the Arabian plate (Fig. 11). In southern Israel (Fig. 1), transgressive outer-ramp nodular limestones of the Avnon Formation appear above the Middle Cenomanian SB of the top Zafit Formation, indicating platform drowning (Wald 2004; Fig. 11). As in Galilee, the transgressive succession of the northern Negev terminates with a maximum-flooding bed composed of laminated pelagites (facies type-6) with abundant small calcisphaeres, planktonic foraminifera, and other pelagic fossils. The last occurrence of orbitolinid foraminifera is slightly above this maximum-flooding bed, suggesting that the Ce MFI-2 of the northern Negev is Middle Cenomanian in age as in Galilee. In west-central Jordan, the equivalent drowning succession above the CeJo2 SB starts with dark bituminous marls and limestones of the Middle Cenomanian Fuheis Formation (Schulze et al. 2003). This transgressive event may correspond with the Ce TST-2 of northern Israel but the overlying progradational phase (Karak limestone) was not identified in northern Israel. In the Oman Platform, the Middle Cenomanian transgression is recognized above shallowwater lowstand deposits at the Upper Natih-E member with a maximum-flooding within the Lower Natih-D Member (Sharland et al. 2001; Van Buchem et al. 2002). This maximum-flooding in Oman corresponds to the calcareous shale unit of the K130 MFS of the Arabian plate and approximately corresponds to the laminated pelagites of the Ce MFI-2 of northern Israel. The Middle Cenomanian drowning succession in northern Israel represents the beginning of a lengthy episode of eutrophication associated to low skeletal productivity, spanning the entire later Cenomanian and Early Turonian (cf. Buchbinder et al. 2000). Eutrophication is especially expressed by blooms of pithonellid calcisphaeres. Cretaceous pithonellids are planktonic, unicellular, thermophilic and opportunistic dinoflagellates typical of surface waters of the continental margins in the neritic –pelagic transitional zone (c. 80–300 m water depth). They thrive under stressed conditions of elevated salinity, temperature and saturation with

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respect to CaCO3 (Dias-Brito 2000). Other studies connect mid-Cretaceous pithonellid-rich deposits to nutrient-rich surface waters above coastal upwelling (Banner 1972; Jarvis et al. 1988; Wendler et al. 2002). In central Jordan eutrophication controlled coeval facies of sequence-2, and Schulze et al. (2004) invoked both sea-level rise and deposition of dysoxic bituminous successions in local basins to explain the Jordanian succession. Van Buchem et al. (2002) preferred increased fresh-water, nutrient and clay injection to explain coeval water cloudiness in the Oman platform. Latter part of sequence-2: eustatic, palaeoenvironmental, and tectonic controls. Late Cenomanian changes in northern Israel are expressed in the Ce RST-2 and the overlying Pelech sequence (Figs 9 & 11). Depositional processes were affected both by sea-level fluctuations and localized tectonic movements. Late Cenomanian aggradation began with the establishment of a CC-2 cycle reflecting the establishment of a homoclinal ramp (Ce HST-2; Figs 8b & 11). Sea-level fall and normal faulting in Galilee led to the transformation of this ramp into an open, non-rimmed shelf, with a steep clinoform-bearing slope facing south– SW, and with toe-of-slope turbidites distally in the Carmel region (Ce FRST-2; Fig 15). This shelf-slope system was exposed for the first time during the Late Cenomanian as evidenced by the Late Cenomanian SB (Ce SB-3; Fig. 11). Shortly afterwards, the exposed shelf-slope system of western Galilee was drowned below the storm-wave base, forming the drowning succession of the Pelech sequence (Figs 11 & 17), partly removed by erosion in a second phase of emergence in the latest Cenomanian (Ce SB-4). The Late Cenomanian SB (Ce SB-3) is coeval with subaerial exposure recorded by Voigt et al. (2006) in the M. geslinianum zone of NW Europe, correlated by them to other surfaces in the Anglo– Paris basin, NE Germany, Crimea, Ukraine, Mangyshlak (Kazahstan) and SE India. As in Galilee, this surface was subsequently submerged, but a second exposure surface, as found in northern Israel (Ce SB-4; Figs 11 & 17) was not reported. Equivalents of the Late Cenomanian SB, the Pelech sequence, and the latest Cenomanian SB of northern Israel are recorded from Wadi Feiran of Sinai. In this region, Kassab & Obaidalla (2001) show a basal Late Cenomanian erosional surface truncating part of the M. geslinianum ammonite zone, overlain by a 7.5 m thick succession, and then with a second hiatus at the Cenomanian– Turonian boundary. The Late Cenomanian SB may correspond to the Ce Sin-6 reported by Bauer et al. (2003) from Sinai, while their Ce Sin-7 may

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correspond to the overlying latest Cenomanian SB. The Late Cenomanian SB may also correspond to the slightly younger Ce Jo4 SB reported by Schulze et al. (2003) from Jordan and to the Ce Up SB at the top of the Tamar Formation of the Negev (Buchbinder et al. 2000), both within the Late Cenomanian N. vibrayeanus zone (Fig. 11). The termination of the Upper Cenomanian homoclinal carbonate ramp system of Galilee (Ce HST-2) is unique, as drowning (by the Pelech sequence) was preceded by forced-regression and subaerial exposure, and followed by a second subaerial exposure event (latest Cenomanian SB). Regression and platform exposure at the end of the Cenomanian in Israel and Sinai were suggested by Flexer et al. (1986), Lewy & Avni (1988) and Bauer et al. (2003) on the basis of Fe-crusts and borings, but robust indications for subaerial exposure such as karst, calcrete, palaeosols, or fluvio-deltaic deposits were not reported. For example, the top of the uppermost Cenomanian in the Negev and central Israel is characterized by ferruginous burrowed or brecciated horizons that are more likely to be indicative of subaquatic omission (see recent example of Heck et al. 2007, ancient example of Immenhauser et al. 2000). Furthermore, in many marginal carbonate systems around the Tethys subaerial exposure of Upper Cenomanian platforms was not recognized and platform termination resulted from drowning owing to sealevel rise, or eutrophication and anoxia (Jenkyns 1991; Philip & Airaud-Crumiere 1991; Gusˇic & Jelaska 1993; Philip et al. 1995; Caus et al. 1997; Drzewiecki & Simo 1997; Schlager 1999; Buchbinder et al. 2000; Scott et al. 2000; Bauer et al. 2001, 2003; Schulze et al. 2004). The first Late Cenomanian subaerial exposure event (Ce SB-3, Fig. 11) was influenced by a eustatic fall at the M. geslinianum zone (Flexer et al. 1986; Voigt et al. 2006), and also by local faulting in Galilee (Fig. 15). On the other hand, the younger subaerial latest Cenomanian SB (Ce SB-4, Fig. 11) seems to result from a local tectonic uplift of Galilee and does not have equivalents in Europe. Late Cenomanian tectonic movements associated to local uplifts were also reported from North Sinai by Bauer et al. (2003), and localized subaerial unconformities attributed to local tectonic events were reported from the Cenomanian on the eastern Arabian plate and the northern margins of the Tethys (Van Buchem et al. 1996 and Scott et al. 2000 – Oman; Borgomano 2000 – Apulian platform; Wilmsen 2000 – Cantabrian platform; Mouty et al. 2003 – NE Palmyrides). Manifestation of oceanic anoxic event-2 (OAE-2; ‘Bonarelli’ Event) across the latest Cenomanian sequence boundary: palaeoenvironmental and eustatic control. The latest Cenomanian SB of

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Galilee represents both a subaerial unconformity and transgressive surface onlapped by Lower Turonian deepening-upward cycles UC-7, UC-8, UC-9, UC-11 and UC-12 (Figs 7 & 8). These Early Turonian cycles represent initiation of drowning caused by sea-level rise and onlap of deep-water facies onto the Galilee palaeo-high (Tu TST in Fig. 15). Above the latest Cenomanian SB, the Lower Turonian Daliyya marls of Carmel contain heterohelicid foraminifera, dinoflagellates, evidence for high vegetative productivity, high concentrations of total organic carbon (1.06–2.02%), and pyrite (Honigstein et al. 1989). The Daliyya marls were considered by Honigstein et al. (1989) to have formed under ecological stress and hypoxia, and were related to the second oceanic anoxic event of the Tethys (OAE-2; Arthur et al. 1987). The correlative marls in the CC-6 cycles of the Yirka Formation, onlapping Galilee from the south (Fig. 15), contain similar features indicating increased nutrients, and hypoxia (facies type-1). The well-laminated and thin-bedded CC-7 and UC-9 cycles of the Bina Formation (facies type-11), onlapping Galilee palaeo-high from the north (Fig. 15) represent hypoxia and stagnation of Lower Turonian sea-floor deposits as well. These

well-stratified, well-laminated mudstones and wackestones are characterized by sparse but consistent occurrences of rotaliid foraminifera, especially gavelinellids (facies type-11 in Table 1, Fig. 3). Rotaliid foraminifera began to dominate oxygenpoor bottoms from the Albian, and small benthic foraminifera including gavelinellids, were used as deep-water hypoxic markers of the Cenomanian – Turonian OAE in the Tarfaya basin, Morocco (Gebhardt et al. 2004). This well-laminated rotaliidbearing facies in the Bina Formation (facies type-11) therefore also reflects hypoxia related to the OAE-2. The succession of events above the Middle Cenomanian SB (Fig. 20) shows that Lower Turonian hypoxia related to the OAE-2 of the Tethys was at the culmination of a lengthy trajectory of eutrophication (Fig. 20). In the Ce TST-2, mass occurrences of pithonellid calcisphaeres, phosphatic grains, and condensation indicate increased nutrients in the water body. During the subsequent Ce HST-2 the carbonate system failed to keep up with the rising sea level as a result of low auto-production of skeletal carbonate (as reflected in the UC-3 cycles), hence development of a homoclinal ramp (Fig. 8b). The Late Cenomanian SB reflects final

Fig. 20. Palaeoenvironmental evolution toward the Early Turonian OAE-2 climax, as reflected in the Upper Cenomanian– Lower Turonian sections of Galilee. The Upper Cenomanian carbonate system preceding the Middle Cenomanian SB shows gradual decrease in skeletal production, increase in eutrophication and deterioration to significant hypoxia in the Early Turonian.

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deterioration of this non-productive homoclinal ramp at a time of Late Cenomanian emergence. The subsequent Pelech sequence, depleted in skeletal grainstones but rich with pithonellid calcisphaeres, reflects further eutrophication-induced reduction of skeletal production leading to drowning of the system. This sequence was uplifted and subaerially exposed within the latest Cenomanian, but the overlying Lower Turonian ammonite marls and laminites reflect culmination of eutrophication and hypoxia of the OAE-2. This chain of events (Fig. 20) tracks the gradual reduction in production of skeletal carbonate through the Late Cenomanian owing to increased eutrophication.

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Summary (1)

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Faciestypesofthe Cenomanian– Turoniansuccession of northern Israel are highly variable, and indicate environments ranging from basinal, outer, mid- and inner ramp, to peritidal and subaerially exposed. Facies units are organized into two orders of sedimentary cycles. High-order cycles may be shallowingor deepening-upward, and are usually up to a few metres in thickness. Most high-order cycles are stacked into low-order cycles. Low-order cycles reflect progradational, aggradational, or retrogradational trends and homoclinal and distally steepened ramp profiles. The cycles construct three Cenomanian and one Turonian sequence. The origin of major Cenomanian–Turonian depositional events that can be correlated across the Arabian plate and Europe is eustatic. The Late Cenomanian eustatic rise was, to some extent, decoupled and masked by local tectonism and uplift in northern Israel. Sedimentary overprints of bottom winnowing, oxygenation of hypoxic sediments, eutrophication, or hypoxia reflect palaeoceanographic influence on some of the depositional events. Eustatic and/or palaeoenvironmental imprints were revealed for the first Cenomanian subaerial exposure; Early Cenomanian maximum flooding and oxygenation of hypoxic sea-floor; Middle Cenomanian highstand progradation followed by forced-regression and masstransport toward the basin; Middle Cenomanian subaerial exposure; increased Late Cenomanian eutrophication throughout sea-level rise; Late Cenomanian subaerial exposure; latest Cenomanian –Turonian eutrophication, and anoxia of the OAE-2 (‘Bonarelli’) event. Two episodes of carbonate ramp demise were recorded: (a) an abrupt deterioration of productive prograding ramp in the

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Middle Cenomanian by sea-level fall, gravity collapse, subaerial exposure, and transgression and (b) deterioration of a non-productive aggradational homoclinal ramp by sea-level fall, faulting, subaerial exposure and drowning in the Late Cenomanian. Throughout the Late Cenomanian, the carbonate system of Galilee experienced a continuous decrease in skeletal grains production owing to increasing eutrophication. These conditions culminated in the Early Turonian by extreme hypoxia related to the OAE-2. Facies-thickness trends of Cenomanian – Turonian systems-tracts show that southern Carmel and much of Galilee were elevated during the Cenomanian and Turonian stages. Galilee and southern Carmel palaeo-highs were separated by a subsiding trough extending fromcentral–northernCarmeltowardsouthern Galilee. The Galilee palaeo-high was bounded at the north by a deeper zone of rapid subsidence, extending across northernmost Galilee and into Lebanon. This sedimentary configuration represents a shift of the north– south depositional strike of the mid-Cretaceous Levantine Hinge Belt toward the east –NE. This trend corresponds to other large-scale tectono-depositional features in the Mesozoic Levant. Late Cenomanian normal faults and a latest Cenomanian SB reflect tectonic activity and uplift of Galilee to above sea level near the end of the Cenomanian. Late Cenomanian faulting was associated with transformation of a Late Cenomanian homoclinal ramp into a steep margin. Locally, calcarenitic clinoform unit-formed steep slope beyond this margin, facing to the south–SW. A concomitant, turbiditic toe-of-slope developed distally in Carmel to the south.

This research was done in the framework of a PhD thesis of the senior author at Ben-Gurion University of the Negev, funded by the Geological Survey of Israel and the Earth Sciences Administration, Israel Ministry of National Infrastructures. The authors thank Z. Lewy (Geological Survey of Israel) for identifying macrofossils and M. Simmons (Neftex) for identifying orbitolinid foraminifers. The authors especially acknowledge the insightful and constructive contribution of the reviewers and editors. Y. Raphael (GSI) and D. Kusashvili (BGU) provided logistic assistance in the field and laboratory.

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Orbital time scale, intra-platform basin correlation, carbon isotope stratigraphy and sea-level history of the Cenomanian – Turonian Eastern Levant platform, Jordan JENS E. WENDLER1,2*, JENS LEHMANN1 & JOCHEN KUSS1 1

Department of Geosciences, Bremen University, P.O. Box 330440, 28334 Bremen, Germany 2

Present address: Smithsonian Institution National Museum of Natural History, 10th & Constitution NW, Washington, DC 20560-0121, USA *Corresponding author (e-mail: [email protected]) Abstract: Two Cenomanian– Turonian boundary (CTBE) sections (KB3 and GM3) of the Karak– Silla intra-platform basin of the Eastern Levant carbonate platform, Jordan, are correlated based on high-resolution calcimetry. KB3 contains black shales with over 7 wt% total organic carbon (TOC). GM3 was deposited at shallower water depth and reveals four conspicuous gypsum beds used for sea-level reconstruction. Spectral analysis of carbonate content and TOC reveals forcing, mainly by the 100 ka cycle of Earth’s orbit eccentricity. Whole rock stable carbon isotope data show a conspicuous positive d13C excursion representing the Oceanic Anoxic Event 2 (OAE2). The carbon isotope records of KB3 and GM3 correspond well with the cycles in the d13C record of the global stratotype (GSSP) at Pueblo (USA). The GSSP orbital timescale, thus, can be applied to the Jordan record. Furthermore, all stable isotope events defined in the English chalk reference record are recognized in Jordan. Our orbital model for the Jordan sequence-stratigraphical framework reveals approximately 1.2 (þ0.2) Ma duration of a third-order sequence, proposed to represent one cycle of the long obliquity (1.2 Ma). This longterm period is superimposed on three fourth-order fluctuations of 400 ka length (long eccentricity; fourth-order sea-level fluctuations), each of which comprises four carbonate cycles (100 ka eccentricity; fifth-order sea-level fluctuations). Demise of the Levant platform occurred during the phase of decreasing d13C values after OAE2 in the interval between the Cenomanian –Turonian (C– T) boundary and the end of the Early Turonian.

The Levant carbonate platform deposits of central Jordan represent a textbook-like shallow-marine platform setting subdivided into intra-platform basins during Cenomanian–Turonian (C– T) times (Kuss et al. 2003). The morphological structuring by these basins induced lithologically highly variable successions, especially during black shale deposition that preferably occurred in deeper subbasins. A correlation of these different successions is possible by means of high-resolution calcimetry and stable carbon isotope stratigraphy. In this paper we present such a correlation that enables the study of three research aspects: a) linking different palaeoenvironments over the C –T boundary interval; b) refining the sequence-stratigraphic model by constructing an orbital time scale; c) correlating the successions to the global record (exemplified for the well-dated C– T boundary interval) in order to eliminate ambiguities in the local stratigraphy. Furthermore, our data support the global picture and estimates of duration of the C –T boundary interval that includes the global Oceanic Anoxic Event 2 (OAE2).

The duration of OAE2 has been the matter of many studies in recent years resulting in a range of values from 320 to 960 ka (Obradovitch 1993; Sageman et al. 2006) partly caused by considering different intervals. An orbital timescale for the Pueblo stratotype section was given by Sageman et al. (2006) providing a precise time measure for the period investigated in our study. In addition, a comprehensive study of stable carbon isotope records in Europe was given by Jarvis et al. (2006) providing a detailed set of isotope events suitable for global correlation. Recently, an orbital model was also presented for the Wunstorf (Germany) section (Voigt et al. 2008). This record correlates well with the present results and can be used to support the new hypothesis put forward in the present paper regarding the orbital trigger of third-order sequences.

Geological setting The Eastern Levant carbonate platform in Jordan is characterized by 300–400 m thick successions of

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 171–186. DOI: 10.1144/SP341.8 0305-8719/10/$15.00 # The Geological Society of London 2010.

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nodular limestone, massive limestone and laminated limestone with intercalated clay, marl and gypsum beds that were deposited during the Cenomanian and Turonian. According to Schulze et al. (2004), carbonate platform demise occurred during the Late Cenomanian and Early Turonian when the deposition of marl and clay became dominant. The sections GM3 (Ghawr Al Mazar) and KB3 (Kuthrubbah) are outcrop sections, about 30 km apart, in the Wadi system cutting east– west into an extended plateau area east of the Dead Sea (Fig. 1). Palaeogeographically the NW-deepening Levant carbonate platform extended over the passive margin of the Arabo– Nubian shield during C –T times. The sections studied comprise the Upper Cenomanian Hummar Formation and the Upper Cenomanian to Middle Turonian Shueib Formation, which are part of the Ajlun Group (Powell 1989). They represent deposits of an intra-platform basin (Karak-Silla basin) at c. 100 km distance from the palaeocoastline of the Arabian Shield. Schulze et al. (2003) report a range of facies representing supratidal to shallow subtidal deposits, based on the analysis of a diverse shallow-water benthic association including calcareous algae, rudists, larger benthic foraminifera, oysters and ostracodes of brackish to hypersaline environments (see Morsi & Wendler 2010). The intra-platform basin was connected to the open marine environment and only temporarily experienced restricted conditions (formation of evaporites) during regressions. GM3 had a marginal position, while KB3 represents deeper parts of the Karak –Silla intra-platform basin. The described facies indicate prolonged phases of low activity of the platform carbonate factory from the Late Cenomanian to Early Turonian. Despite the general carbonate platform setting, the profiles investigated here show only short periods of normal carbonate production, while most parts of the section consist of marls and clays. So, the material represents a depositional system with relatively high siliciclastic input.

Material and methods For this paper we focus on the interval section metre 47–83 of section GM3 (Ghawr Al Mazar: 318150 3400 N; 358350 4100 E) representing the OAE2. It is part of a mid-Cenomanian to Lower Turonian section published in separate publications (see Morsi & Wendler 2010). The section develops from green clays and marls into a unit of platy, bituminuous limestone beds, followed by brown marly clays, grey marls and limestone at the top. A total of 155 samples were collected from the GM3 section interval 47– 83 m at sample spacing of 10 –25 cm.

Section KB3 (Kuthrubbah: 318090 1300 N; 358360 0600 E) is 25 m thick and comprises an alternation of black shales and platy, bituminuous limestone beds, and shows a massive limestone at the top. An interval of about 2.5 m below this topmost limestone could not be sampled owing to poor outcrop conditions. 117 samples were taken with 10 –20 cm sample spacing (section metre 0 –16 m) and 30 –50 cm (above 16 m). Bulk samples were crushed with an agate mortar. The measurements of the carbonate content and the total organic carbon (TOC) were performed at the Alfred Wegener Institute Bremerhaven, Germany, using a LECO CS-125 carbon –sulphur determinator. For total carbon a LECO CNS-2000 was used. Stable carbon isotopes were measured on bulk carbonate at the isotope laboratory of Bremen University, Germany, using a Finigan MAT 251 mass spectrometer. The results are reported relative to the V-PDB standard. Thin sections were prepared from limestone samples for microfacies analysis. Standard smear slides were used for determination of coccolith assemblages.

Results Biostratigraphy The integrated biostratigraphic framework of Schulze et al. (2003) using nannofossils and ammonites, supported partly by larger benthic foraminifera, forms the stratigraphic basis (Fig. 2). The sections comprise nannoplankton zones CC10 and CC11. For details on index species see Schulze et al. (2003). A revision of the nannofossil content of all correlated sections previously studied and new analyses resulted in a repositioning of the zone boundary (Fig. 2) based on earlier appearances of the Turonian index fossil Quadrum gartneri already in the section part formerly placed in the Late Cenomanian by Schulze et al. (2003), taken parallel to the new GM3 section. The ammonite occurrences in the C –T of Jordan (Schulze et al. 2004) provide a more detailed biostratigraphy that can be correlated with ammonite zone schemes of Southern Europe (Hardenbol et al. 1998), Israel (Lewy 1989, 1990), and the Middle East in general (Lewy & Raab 1976). In conjunction with the isotope record presented here it enables a good time control. The Hummar Formation (section metre 47– 56 m, Fig. 2) contains abundant Neolobites vibrayeanus, encountered also in the present material at section metre 47.5 m. Wiese & Schulze (2005) stated that the stratigraphic range of N. vibrayeanus in Jordan is not yet clear, but as far as biostratigraphical control is given, the species range correlates approximately with the early Late

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Fig. 1. (a) Location of sections GM3 and KB3 in Jordan. (b) Investigated section part GM3 showing lithological markers. (c) Overview of the northern slope of Wadi Al Karak; frame marks the section enlarged in (b).

174 J. E. WENDLER ET AL. Fig. 2. Biostratigraphy, stable carbon isotope, carbonate and TOC content data of sections GM3 and KB3. Carbonate data connected by correlation lines; lines are extended toward the isotope curves at key points to show correspondence of d13C records. TOC, main correlated TOC maxima. Black shales in GM3 ¼ grey shaded intervals. Timescale in the carbonate panel of GM3 assumes 100 ka (eccentricity) duration of the numbered cycles, zero point according to Pueblo (see Fig. 3). Dolo., dolomite unit; TS/FS, transgressive/ flooding surface; mfs, maximum flooding surface (according to Schulze et al. 2003). Vertical bar marks OAE2 as indicated by Sageman et al. (2006). Stratigraphy panel includes the referred lithological units of GM3. m1, m3 – Ammonite marker beds. Positive d13C excursions within OAE2: a, b, c; negative d13C excursions: N1, N2.

CENOMANIAN– TURONIAN ORBITAL TIMESCALE FOR THE LEVANT

Cenomanian Calycoceras guerangeri zone (Lehmann in Wiese & Schulze 2005). Concerning the first occurrence of N. vibrayeanus, this is contrary to Schulze et al. (2003, fig. 7; 2004, fig. 3) who correlated it in Jordan with the middle of the Acanthoceras rhotomagense zone of the northern European zonation. N. vibrayeanus is accompanied by further ammonites in Jordan, Proeucalycoceras haugi, Pseudocalycoceras harpax and Turrilites acutus. This assemblage can be assumed to represent Middle to Late Cenomanian (pers. comm. Lewy in Schulze et al. 2004). P. haugi is a species of the early Late Cenomanian (Kennedy & Juignet 1994). P. harpax is also early Late Cenomanian (Kennedy & Juignet 1994), though the delineation between the type material from India and that from the Near East still needs clarification. T. acutus mainly occurs in the Middle Cenomanian, and can range into the Late Cenomanian (Juignet & Kennedy 1976). To summarize, the first appearance of N. vibrayeanus in Jordan occurs already in the Middle Cenomanian and the last occurrence is correlated with the top of the European Calycoceras guerangeri zone. Section part 56 –64 m contains the platy limestone beds described as ‘ammonite marker bed 1’ by Schulze et al. (2003, 2004) who reported Vascoceras cauvini, Metoicoceras geslinianum and Burroceras transitorium, which indicate Near East ammonite zone T1 of Lewy & Raab (1976) correlating with the European M. geslinianum and N. juddii zones. The record of B. transitorium needs further confirmation, since this species is hitherto recorded from New Mexico, Arizona and Brazil only (Kirkland 1996; Gale et al. 2005). In the middle part of this interval ammonites are abundant at section metre 60.5 m in section GM3 and at 3.2 m in section KB3. Although most of the specimens are poorly preserved, besides Puzosia sp. the largest part can be possibly attributed to very feebly ribbed Watinoceras spp., namely W. guentheri, W. hesslandi, W. inerme and W. semicostatum. All of these are previously known from western Morocco and western Afrika (Reyment 1955, 1957; Collignon 1967). This group of species might be conspecific following Wright & Kennedy (1981), if so W. hesslandi Reyment (1955) has priority. The total range of these smooth Watinoceras spp. is poorly known, but the genus is widely accepted as appearing not before the Lower Turonian of the modern substage definition following the discussion of Collignon’s (1967) data by (Lehmann & Herbig 2009). The first and possibly most common occurrence of these forms can be placed into the basal Turonian W. moremani zone (Wiedmann & Kuhnt 1996), just above the latest Cenomanian Neocardioceras juddii zone.

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Ammonite zones T2-4 correlate to the basal CC11 and can be only roughly determined by scarce findings of Choffaticeras pavillieri and Ch. quaasi in section part 67–74 m. The limestone in the top of the section represents the Wala Limestone Member, which was described as ‘ammonite marker bed 3’ by Schulze et al. (2003, 2004). It is characterized at the base by an Early Turonian ammonite assemblage with Vascoceras durandi, Thomasites rollandi, Choffaticeras luciae, Ch. quaasi and Fagesia lenticularis, indicative of zones T5 –6a. The first three species are typical for the Pseudaspidoceras flexuosum to Thomasites rollandi zones in the middle part of the Lower Turonian in Tunisia, the detailed range of Ch. quaasi and F. lenticularis, the latter is a somewhat obscure species following Chancellor et al. (1994) in the Early Turonian. Transition into zone 6b (M. nodosoides/C. woollgari zone boundary, base Middle Turonian) was placed by Schulze et al. (2003, 2004) in the Middle Wala Limestone.

Lithological evolution at the different intra-platform basin sections The GM3 section can be subdivided into six lithological units (Fig. 2). Overlying the platform limestone called Karak Limestone, the analysed profile starts at profile metre 47 with a 12 m succession of greenish marls and clays (lower 4.5 m not exposed) with interbedded nodular limestone beds (green clay unit). Occasionally, gypsum beds and crosscutting diagenetic gypsum veins are intercalated at section metre 49 –50 m. At 52.30 m a thin (c. 15 cm), iron-rich black clay layer marks a first event of increased TOC accumulation (Fig. 2). This bed is very conspicuous in the field and contains a 3 cm thick, very dense, limestone bed. Above this black shale, the succession of greenish clays continues to 54.8 m. A 2.5 m thick unit of dolomite and ankeritized platform carbonate (dolomite unit) follows with a layer of strongly distorted gypsum-carbonate breccia with signs of reworking at the base, possibly connected to a minor hiatus (TS in Fig. 2). It marks an abrupt change in lithoand bio-facies. The section continues into a 7.5 m thick alternation of brownish marl and bituminuous, platy, partly laminated limestone beds (platy limestone unit). The third bed of this unit is an oysterlimestone typical of the platform carbonate facies of the Cenomanian investigated here. The two basal and the upper two bituminuous limestone beds are calcisphere (calcareous dinoflagellate cysts Pithonella) packstones showing bioturbation (Fig. 2). The other beds, in contrast, are laminated. Thin sections reveal a strong alteration of the limestone. The dark brown marl at the base of the platy limestone unit forms a second black shale (TOC 1).

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This interval is followed by 10 m of monotonous brown clays/marls (brown marl unit). The lower half of this unit can be considered a third black shale. The next unit is a 6 m thick interval of greenish-grey marls (grey marl unit) that contain TOC peak 2 representing the fourth black shale of the section. At section metre 80, a 2 metre thick green to grey, gypsum-rich marl is present. It is followed by a 1 m thick interval of red marls at the base of the following limestone sequence (socalled Wala Limestone). This red marl represents a marine red bed, which potentially is a marker horizon throughout the Levant platform (Wendler et al. 2009a). The grey, bioclastic Wala Limestone is approximately 14 m thick and represents platform-type carbonate deposits rich in oysters. Section KB3 is essentially a succession of black shale with intercalated limestones. Section interval 0–5 m displays six equally-spaced, dm-thick, limestone beds, light-beige in colour, alternating with dark-grey marl. Section interval 5 –16 m is dominated by dark grey to greenish marl with regular intercalations of dm-thick harder beds. At section metre 16.1 a limestone bed, comparable to the beds in the lower part of the section, is present. From section metre 16.2– 24 of the section greenishgrey marl were deposited. The upper part of this unit is barely exposed and strongly weathered and distorted. Thus sampling ends at section metre 20. The Wala Limestone member in the top of the section is dislocated by some metres owing to rock sliding.

Isotope record Whole rock stable carbon isotopes are in the range of –3 to þ4‰ (Fig. 2). High-frequency fluctuations (metre-scale) are evident and a superimposed trend (2–3 m bundles) forms conspicuous negative and positive excursion phases. A broad positive excursion above mean value is present in both sections. The carbon isotope excursion (CIE) of GM3 and KB3 (Fig. 2) shows the following shape (section metres are given for GM3): † pre-excursion background around 1‰ (with a decreasing trend) in section GM3 part 47– 51 m; † pre-excursion strong negative excursion to – 4‰ (section GM3 part 51 –54.5 m, double peak: negative peak 1: N1; negative peak 2: N2); † first build up (peak a) – slow increase to 1.7‰; † a significant decrease to 0‰ at 57.5 m (the so-called ‘trough’); † second build up (peak b) – maximum around 3‰, † after peak b, a period of significant fluctuation down to 0‰ and below occurs (around 60 m); † a third build up (peak c) starts around section metre 63 and is part of a plateau of d13C values c. 2.5‰ from 62 to c. 70.5m;

† the post excursion interval 70.5– 83 m (slow decrease to mean value), which contains one significant positive excursion of above 3‰ at 74.5 m, and a strong negative excursion to –2.8‰ at 83 m. This negative excursion represents the lowest value since the pre-excursion negative spike.

Calcimetry and gypsum marker beds In section GM3 carbonate content is between 10 and 90%, and the record shows a distinctive cyclic pattern consisting of high-frequency fluctuations bundled into sets (Fig. 2). These bundles are numbered starting at the base of the isotope excursion. Section part 47–55 m below the isotope excursion can be subdivided into three bundles (– 1 to –3) of about 3 m thickness. Abundant disperse gypsum (19%) is present in the clay and marl in the lower and upper parts of bundle –2 (gypsum bed 1). Gypsum content otherwise is generally below 1%. The platy limestone facies (section part 57 –64.5 m) shows high-frequency cyclicity in the carbonate content values that range between 50 –90%. A conspicuous minimum of 11% carbonate related to elevated gypsum content occurs at bundle boundary 2/3 (gypsum bed 2). From 64.5 m to 80 m the cyclicity continues to show bundling, but bundle thickness increases. At the base of bundle 7 a reddish bed reflects increased contents of iron oxides and gypsum (gypsum bed 3). A bed with a substantial minimum of ,10% carbonate content at 80 m exhibits the highest gypsum content of the section and marks the boundary between bundles 10 and 11 (gypsum bed 4, 69% gypsum). The four so-called gypsum beds are clays and marls with disperse gypsum rather than massive sulphate beds. The most conspicuous one, gypsum bed 4, potentially is connected to a minor hiatus. TOC is mostly around 0.5 wt% throughout the section (Fig. 2). With the onset of the platy limestone facies related to isotope peak b a 2 m thick peak in TOC (1.5–3%) can be observed (TOC1). Throughout the platy limestone and brown marl units TOC remains around 0.5% with short peaks between 1– 2%. Another major peak in TOC straddling about 4 m in thickness (1.5–3% TOC) occurs in the brown marl unit in the carbonate bundles 9 and basal 10 (TOC2). In the KB3 section carbonate values are between 30 and 95% (Fig. 2). Carbonate content also shows high-frequency cyclicity with superimposed longer-period bundles. The carbonate-minima of these bundles are correlated with the GM3 record. The carbonate cycles of both sections are positively correlated with the cyclic stable carbon isotope excursion.

CENOMANIAN– TURONIAN ORBITAL TIMESCALE FOR THE LEVANT

TOC values abruptly rise from a background level of ,1% at the base of the section to a first peak of .7% at 2.5 m (TOC1). Above that level the mean value is constantly decreasing from 4 to 2% and sharply increases to a second major peak of 7% at section metre 18 m to form a 3.5 m thick double peak with maxima of 7% (TOC2). Between these two main peaks TOC values remain high from 5.5 m to 12.5 m with peaks .5%. Similar to the carbonate content, high-frequency cyclicity characterizes the TOC record.

Isotope and calcimetry correlation: spectral analysis The correlation of carbonate cycles (bundles) between GM3 and KB3 is corroborated by the comparable shape of the d13C records and the position of major TOC peaks. Thus, section KB3 comprises cycles 1–10 of GM3. Owing to the high sampling resolution of KB3 the carbonate and TOC records of this section are suitable for spectral analysis in order to determine the type of cyclicity represented. We estimate mean accumulation rates using possible ages given for the stratigraphic interval studied (based on Ogg et al. 2004; Sageman et al. 2006) assuming an equal length of the carbon isotope excursion all over the globe. These theoretical accumulation rates are applied to the spectra of carbonate and TOC in order to search for the best fit to the Milankovitch frequency band. According to the Ogg et al. (2004) timescale the duration from the base of the Middle Cenomanian to the basal Middle Turonian is about 95.7 to 92.15 ¼ 3.55 Ma. This interval is represented by c. 80 m in the GM section (see Schulze et al. 2004 for entire section) resulting in 22.5 m/Ma mean accumulation rate. With respect to the isotope excursion the time span between the first increase (peak a) and the end of the plateau is in both time scales about 870– 1000 ka (vertical bar in Figs 2, 3 & 4) resulting in mean accumulation rates of 16– 19 m/Ma (KB3) and 20–23 m/Ma (GM3). This interval contains eight cycles in both sections. Since cycle thicknesses in GM3 are highly variable (Fig. 2) the range of estimated accumulation rates is larger: cycles 4 –6 are 2 m thick while all other cycles are 3–3.8 m thick. Thus, accumulation rates around 30 m/Ma dominate. Power spectra were generated using the Lomb – Scargle algorithm. The TOC spectrum of KB3 (Fig. 5a) shows a strongest power signal at 7.2 m, followed by peaks in power at a 16 m long period (close to section thickness of 20 m: low significance), 2.83 m, 1.65 m, 0.34–0.41 m (only two samples per cycle: low significance). A less significant signal occurs at 0.88 m. Assuming 18 m/Ma mean accumulation rate the 7.2 m cycle would be

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400 ka long eccentricity, 2.83 m represents a period of c. 200 ka not known from the Milankovitch frequency band. The cycle at 1.65–1.84 m is in the range of short eccentricity, and 0.34 –0.41 m would represent precession. The peaks at 0.66 m and 0.88 m could reflect the obliquity cycle. The spectrum of carbonate content (Fig. 5b) of the most symmetrical, highest resolution part of KB3 (3–11 m) shows two prominent peaks of 1.6 –2.65 m cycles. Power is also recorded at a cycle of c. 7 m thickness which is, however, about the thickness of the measured section. Two small signals occur at 0.34 m and 0.36 m. Cycle and bed thickness is lower in this part of the sections (Fig. 2). Thus, using 17 m/Ma proposed accumulation rate for this particular interval (derived from high-resolution correlation of Fig. 4) the two strongest peaks are related to 156 and 97 ka, the 7 m cycle spans 400 ka, and the 0.34 m cycle would comprise 20 ka. The cycle ratio 1.6:0.34 m is 4.7:1. Figure 5c shows the power spectrum of carbonate content of section GM3. Three peaks occur in the frequency band of short eccentricity (2.65 m, 3.3 m, 4.1 m), and strong peaks are present in both the obliquity (1.29 m) and precession (0.71 m and 0.59 m) frequency bands. In Figure 5d the section part 67 –79.8 m is analysed. We find the strong 3.42 m cycle (cycles 8–11 in Fig. 2) corresponding to a duration of 107 ka when 32 m/Ma accumulation rate is assumed. This latter analysis suggests that accumulation rates of GM3 are higher than those of KB3, especially in the Lower Turonian part of GM3.

Discussion High-resolution correlation with Pueblo: an orbital timescale for the Levant In Figure 3 we attempt a correlation of the Jordan GM3 isotope record with the Pueblo GSSP section record of Sageman et al. (2006). We can assume a reasonably complete stratigraphic record based on the lithological properties of the GM3 section, which indicates only two potential positions of minor hiati. Those are, first, the base of the dolomite unit where reworking is indicated, and second, the gypsum bed 4 where a substantial sulphate content suggests enhanced evaporitic conditions during deposition linked to temporary emersion (Wendler et al. 2009a). The 100 ka cycles in the carbonate record positively correlate with the d13C curve but are more detailed. Thus, we use the clearly defined cycles of the carbonate record and correlate excursions of decreased carbonate content (negative d13C excursions) with negative excursions in the Pueblo

178 J. E. WENDLER ET AL. Fig. 3. Correlation of the GM3 section and Pueblo (Colorado) GSSP; Pueblo orbital time scale from Sageman et al. (2006); correlation lines use minima in the isotope/carbonate curves. 100 ka-cycle numbering (carbonate curve) indicates Jordan fifth-order cycles; 100 ka-cycle numbers have been suggested accordingly for the Pueblo d13Ccarb record between the base of the CIE and the C–T boundary in concert with the Pueblo orbital model. GM3 isotope record: Isotope events from Jarvis et al. (2006) are labelled according to Figure 4: M, Monument; h, Holywell; c1; c2; L, Lulworth; lpm, late plateau minimum (new event introduced here); the respective numbering according to Voigt et al. (2006) is TU1 to TU5. Jordan orbital model: stippled curve, 1.2 Ma long-obliquity cycle (third-order sea-level cycle); solid curve, 400 ka long-eccentricity (fourth-order sea-level cycle). In the stratigraphy column, sequence S5 is grey-shaded; 5a, 5b and 5c are the fourth-order sub-sequences according to orbital model; SB ¼ sequence boundary.

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Fig. 4. Correlation of d13C records Jordan and Pueblo (Sageman et al. 2006) with the English chalk reference curve (Jarvis et al. 2006). SB ¼ sequence boundary. Positions of marker beds of Plenus Marl (England) and Bridge Creek Limestone (Pueblo) are given to aid lithological comparison with Jordan (platy limestone unit grey shaded). Left panel: qualitative representation of platform extent redrawn from Philip & Airaud Crumiere (1991).

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7 cycles = >30 m/Ma 3 cycles = <20 m/Ma

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0.7

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cycle ratios 1.84:0.66 = 2.8---eccentricity:obliquity 1.84:0.4 = 4.6---eccentricity:precession 1.84:0.35 = 5.3

0.8 0.7

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GM3 carbonate 67 m – 79.8 m, cycles 7–10 thickness : c. 13 m est. acc. rate : 32 m/Ma cycle ratios 3.42:1.77 = 1.9---eccentricity:obliquity? 3.42:0.71 = 4.8---eccentricity:precession

0.7 0.6

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1.77 m (55 ka) 1.03 m (32 ka)

0.71 m (22 ka)

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.1 1.2 1.3 1.4

Frequency (1/m)

cycle ratios 3.3:1.29 = 2.6 --- eccentricity:obliquity 3.3:0.71 = 4.6 --- eccentricity:precession 3.3:0.59 = 5.6

Fig. 5. Power spectra of TOC and carbonate values. (a) KB3 TOC record assuming mean accumulation rate of 18 m/Ma. (b) KB3 carbonate record of section part 3 –11 m (Late Cenomanian). (c) GM3 carbonate record. 26 m/Ma mean accumulation rate for this section, dominant cycle thickness .3 m (.30 m/Ma). Match with Milankovitch frequencies occurs at 32 m/Ma accumulation rate. (d) Lower Turonian section part of GM3.

d13C record (Fig. 3). The spectral analysis strongly supports the idea that, such as in Pueblo, orbital forcing dominated by the 100 ka eccentricity caused the cycles in the Jordan record. Correlation starts at the zero-level of the Sageman et al. (2006) orbital time scale, which we interpret to be corresponding to the base of peak a of Jordan based on biostratigraphy and interpretation of the isotope record. Further levels to tie the records are the peak positions b, c in both Jordan and Pueblo. 100 ka cycle lengths of the Pueblo section are determined by the orbital time scale panel in Figure 3. Comparing the number of cycles in the interval from the base of the CIE to the

Cenomanian –Turonian boundary (CTB) unfolds a cycle-by-cycle correspondence between the two records. Hence, the five cycles determine a duration of 500 ka for this interval in both records (Fig. 3). This is one cycle longer compared to only four cycles detected by Voigt et al. (2008) for the same interval in the Wunstorf (Germany) section. Similarly, slight differences to the Pueblo and Wunstorf records occur in the Early Turonian interval: We have six 100 ka cycles between the CTB and the Lulworth isotope event in Jordan; in Pueblo (800 ka) two more 100 ka cycle were calculated; in Wunstorf a length of 6.25 cycles (100 ka) is seen for that same period of time.

CENOMANIAN– TURONIAN ORBITAL TIMESCALE FOR THE LEVANT

Global carbon isotope events Figure 4 is based on the correlation in Figure 3 and integrates the English chalk reference record of Jarvis et al. (2006) into this correlation in order to identify the isotope events of the latter for the Jordan record. The correlation of peaks a, b and c of the OAE2 between England and Pueblo is followed the one given by Jarvis et al. (2006) who used the Pratt (1985) isotope curve for Pueblo. Thus, Figure 4 integrates the detailed stratigraphies of the Plenus Marls (England), Bridge Creek Limestone (Pueblo) and the platy limestone (Jordan), part of which contain the main isotope peaks in the respective regions. The platy limestone unit, together with the dolomite unit below, can be considered the time-equivalent of the Plenus Marl. The facies change, from clays to a conspicuous dolomitic limestone bed, occurs at the base of the isotope excursion. This facies change is interpreted as a transgressive surface (TS) owing to its litho-facies properties and marks the onset of isotope peak a. It can be correlated using stable carbon isotopes with bed 63 of the Pueblo section, and relates to the base of the Plenus Marls. The base of Bed 63 of Pueblo is a TS. Deepening above the mentioned facies change in GM3 is also indicated by bio-facies properties of the platy limestone unit representing a substantial change from a brackish fauna below to an assemblage consisting of normal marine microfaunal elements (compare Morsi & Wendler 2010). Based on the correlation in Figure 4, this onset of a marl– limestone alternation corresponds to Pueblo bed 67 and Plenus Marl bed 3 both covering the ‘a’ peak of the isotope excursion and indicating progressive onlap. The continuation of cyclic limestone bed deposition above isotope peak a in Jordan (see also Fig. 3) represents progressive flooding. It is followed by the oyster bed (Fig. 2), which appears to be an equivalent of the ‘oyster packstone’ of Caus et al. (1997), which represents this transgression in the Pyrenees platform of NE Spain. The latter can be related to a major transgressive pulse throughout Europe (e.g. comprehensive overview in Voigt et al. 2006). A discrepancy concerning the correlative beds within the CIE peak-interval ‘a’ occurs in the literature as some studies correlate Plenus Marl bed 3 with bed 63 of Pueblo (Keller et al. 2004; Voigt et al. 2006), which has implications for the timing of the initial transgression. The correlation of the Lulworth peak in the top of the three sections is based on the assumption that the nodosoides/woollgari (resp. T6a/T6b) ammonite zone boundary is synchronous in the three sections as corroborated for Pueblo and England by Kennedy & Cobban (1991) and for the

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Middle East by Lewy & Raab (1976) and Schulze et al. (2003). Within this correlation scheme a peak by peak correlation is possible for further events. The good fit approached by this procedure (Fig. 4) implies high-resolution correspondence of the three records and the absence of considerable hiatuses in any of them. Besides the positive d13C excursions, the double-peaked negative excursion N1 and N2 merits attention. It correlates in timing and shape with a recently described strong negative d13C excursion preceding OAE2 in carbonate platform deposits from Mexico (Elrick et al. 2009) suggesting a common, perhaps global, cause.

Sea level across the C– T boundary interval: implications of isotope record for sequence stratigraphy Our sedimentological analysis combined with isotope data implies the following model of sealevel variation (Fig. 3). The four gypsum beds mark periods of evaporitic deposition during times of low sea level. The basal and top one of these gypsum-rich marl beds are the two most conspicuous ones. Thus, we relate them to major third-order lowstands. The lower one relates to sequence boundary CeJo4 of Schulze et al. (2004) [corresponding to the global Ce5 of Haq et al. (1987) and Haq & Al-Qahtani (2005)]; the upper one near the Early–Mid Turonian boundary is correlated with the globally recognized sequence boundary Tu1 (Fig. 6). CeJo4 marking the base of the fifth Cenomanian sequence (S5) of Jordan was defined by Schulze et al. (2004) based on a hardground at the top of the Hummar limestone, which is present in more distal sections while not developed in the study area, where this stratigraphic level is represented by lowstand deposits (green clay unit, Fig. 2) followed by a conspicuous surface marked by the black shale at section metre 52.3 (CeJo4). Regarding the south-western Levant platform, a sequence boundary corresponding to the global sequence boundary Tu1 in the Early Turonian is present in the Sinai area (CeSin7 of Bauer et al. 2003), however, its exact stratigraphic position cannot be determined there owing to a major hiatus. Because of its global significance, this major sequence boundary needs to be newly defined for Jordan. It should be numbered TuJo1 (Fig. 6). Being located near the Early–MidTuronian boundary this third-order sequence boundary TuJo1 is, in contrast to Schulze et al. (2004), positioned below instead of being above the Wala Limestone Member. This controversy shows the importance of the distinction of fourthorder and third-order sequences in order to

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J. E. WENDLER ET AL. Global eustatic Haq et al. 1987 rise

Sinai Bauer et al. 2003

Tunisia Robaszynski et al. 1993

rise

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Ce5

CeJo4

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CeJo3 CeJo2

Ce4

CeJo1

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Ce3/4 Ce3

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Fig. 6. Overview of sequences at various positions at the Levant platform (Robaszynski et al. 1993; Hardenbol & Robaszynski 1998; Bauer et al. 2003; Schulze et al. 2004) and global sequences (Gradstein et al. 2004). Note differences in sequence boundaries of Jordan sequence S5 resulting from distinction of fourth-order and third-order sequences (grey bar).

define co-relatable sequence boundaries. Besides different-scale Milankovitch cycles, as shown in the present case, tectonically controlled local sea-level fluctuations may lead to a mismatch in the inter-plate sequence-stratigraphic correlation approaches (Strasser et al. 2000) because they may, independently from the global eustatic control, pronounce sedimentary surfaces. We interpret the sequence boundary in the top of the Wala limestone member, formerly assigned to TuJo1 by Schulze et al. (2004), to correspond to a fourth-order sea-level fluctuation and propose the following third-order sequence architecture. There are four carbonate cycles of 100 ka duration between two gypsum beds (Fig. 3). Hence the sea-level fluctuations between the gypsum rich marl beds follow a 400 ka period that can be assigned to the long eccentricity. This cycle is confirmed by spectral analysis (Fig. 5) although this kind of analysis must be interpreted with care because of the shortness of the analysed section. Consequently, the two major sequence boundaries comprising an interval of 12 carbonate cycles of length 100 ka are 1.2 Ma apart. Hence this is the total duration of third-order sequence S5 of Jordan (Figs 3 & 6). It is composed of three cycles of 400 ka length that are interpreted to represent fourth-order sea-level fluctuations. The 100 ka carbonate cycles apparently were forced by sea-level variations too, as they may show flooding surfaces

(FS in Fig. 3) at the base. Therefore, we can interpret short eccentricity to have forced fifth-order sea-level changes. The carbon isotope record corresponds to this succession of sequences (Fig. 3) – the main negative isotope excursion occurring with a phase difference of about one fifth-order cycle-length (100 ka) after the gypsum beds (i.e. gypsum bed 1 followed by negative spike N1; gypsum bed 4 followed by Lulworth event). Likewise, the two minor minima in the d13C record, the interval between d13C peaks b and c, and the late plateau minimum (lpm), occur above the fourth-order cycle boundaries, gypsum bed 2 and 3 respectively. Thus, these minima are apparently linked to lowest sea level. Major positive d13C excursions correspondingly are positioned between the gypsum beds during transgression and highstand: a and b in sequence 5A; c in sequence 5B; Holywell, c1, c2 in sequence 5C (Fig. 3). Concluding, based on our cycle analysis and the orbital model for third- to fifth-order sea-level cycles in Jordan we can support the orbital model of Sageman et al. (2006). Hence, the Jordan record underpins the notion of Sageman et al. (2006) that the Early Turonian comprises about 800 ka, that is, 640 ka less than in the GTS 2004. Regarding the time interval from Monument to Lulworth (approximately corresponding to one third-order sea-level cycle), both the Jordan and

CENOMANIAN– TURONIAN ORBITAL TIMESCALE FOR THE LEVANT

the Pueblo records reveal a duration of 1.2 to c. 1.4 Ma, respectively. The recent spectral-analysisbased time scale for the Wunstorf record (Voigt et al. 2008) reveals c. 12 cycles of 100 ka duration (c. 1.2 Ma) for the same period of time.

1.2 Ma long obliquity cycle and third-order sea-level change There is a long period cycle in Earth’s obliquity of c. 1.2 Ma length, which has been interpreted to drive glacio-eustatically controlled third-order sequences of the Pliocene–Pleistocene (Lourens & Hilgen 1997), and glacio-eustatic sea-level fluctuations with effects on equatorial bio-productivity in the Oligocene (Wade & Pa¨like 2004). The 1.2 Ma obliquity cycle was also found to be responsible for Middle Miocene global cooling (Abels et al. 2005). Thus, we put forward the hypothesis that the 1.2 Ma obliquity cycle had an influence on third-order sea-level variations in the Cretaceous as well. Cycles of similar duration also occur in the Early–Middle Turonian of the Alps (Wendler et al. 2009b) stratigraphically continuing the pattern presented here. Wilmsen (2003) considers fourth-order sea-level cycles to be 400 ka eccentricity forced, which consequently relates the third-order sequences to a longer-term cyclicity. This is in accordance with our results. A relation between the main carbon isotope excursions in the English Chalk and eustatic sea-level fluctuations has been demonstrated by Jarvis et al. (2002, 2006). A critical point, which is frequently mentioned in this context, is the assumption that these sea-level changes require some ice-volume control, which for the Cretaceous is still contentious (e.g. Miller et al. 2005).

The timing and possible cause of platform demise of the Levant carbonate platform It has been stated that OAE2 was coeval with the widespread demise of carbonate platforms in the Tethyan Realm (Masse & Philip 1981; Schulze et al. 2004; Voigt et al. 2006). Carbonate platform demise associated with OAE’s has also been recorded for other stratigraphic levels, such as the Early Cretaceous (e.g. Fo¨llmi et al. 1994; Weissert et al. 1998). Our study shows that demise of the platform follows the onset of OAE2 with a considerable time lag. While on the one hand, reduced activity in carbonate production of the Levante carbonate platform in the Jordan research area (Schulze et al. 2004) started already in mid-Cenomanian times, we note that on the other hand, platform carbonate production resumes to some extend exactly during OAE2. So in the case of the green clay unit, platform

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retreat can be related to sea-level drop and the establishment of brackish environments well before OAE2. The demise of carbonate production here was likely caused by the enhanced delivery of terrigenous siliciclastic material spanning an episode of several 100 ka, which probably was a precursor to trigger OAE2. Then, at the base of the isotope excursion (globally correlated with the base of the M. geslinianium ammonite zone), that is, with the onset of OAE2, we can observe a short-term re-establishment of platform carbonate production (oyster bed, Fig. 2). This level of platform growth, spanning the maximum phase of OAE 2 (CIE peaks a, b) is very noteworthy because it correlates to the conspicuous coral limestones of the Naqb Limestone Member (Powell 1989), which are found in proximal sections of the Arabian block in Jordan along the palaeocoast of the Levant platform. Furthermore, it is coeval with the Late Cenomanian platform extension event of Philip & Airaud Crumiere (1991), which lasted throughout the Late Cenomanian (Fig. 4). So the problem that arises here is that the major peaks in Jordan CIE are in fact correlated with a re-establishment of carbonate platform-type deposits in conjunction with sea-level rise rather than platform demise. Towards the top of this phase, the start of the global calcisphere bioevent of Hart (1991) marks a change in the environmental conditions (note dots in Fig. 2). From this point, that is, above the C– T boundary, platform demise characterized by the absence of platform carbonates follows, that is, with a time discrepancy of about 500 ka after the onset of OAE2. During this phase, spanning the whole Early Turonian, increased cycle thickness was caused by elevated accumulation rates that might indicate higher accommodation space. This can be assumed to correspond to platform drowning. This phase is characterized by Pithonella abundance peaks (Fig. 2) continuing the calcisphere bioevent, and suggesting high productivity (Caus et al. 1997; Wendler et al. 2002a, b; Wendler & Willems 2002; Wilmsen 2003). Since the demise of the platform postdates OAE2, it cannot have played any role in triggering the disturbance in the carbon cycle related to the anoxic event and vice versa – a strong time lag is clearly involved. It is interesting that high bio-productivity apparently is a widespread feature of this phase of carbonate platform retreat (e.g. Hallock & Schlager 1986; Brasier 1995), and in the sections studied here it is particularly well represented by an exceptionally high abundance of chlorophyll-derived pristane and phytane (up to 2 mg g-1 TOC) (Sepu´lveda et al. 2009). It might reflect a response of the biosphere following fertilization during OAE2. Excess phytoplankton production, however, will have diminished the light transmission of surface

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waters, possibly strongly enough for hampering activity of the benthic platform carbonate producers (rudists, corals, larger benthic foraminifera), hence the platform demise. While on the one hand this reduction of the platform growth induced hiatuses in the platform sedimentary record near the C –T boundary elsewhere, for example, Bauer et al. (2002), the sedimentary record of the intra-platform basin studied here on the other hand provides a continuous record. Platform growth resumes with the onset of deposition of the Wala Limestone (at the top of the section) comprising the uppermost Lower Turonian and basal Middle Turonian.

phase of decreasing d13C values after OAE2 spanning the Early Turonian. Biological factors (strong phytoplankton productivity) apparently played a dominant role for the ceasing of platform carbonate producers. A. Masri and the Jordan Geological Survey (NRA) are thanked for providing ideal field campaign logistics. Laboratory facilities at the AWI, Bremerhaven, were made available by R. Stein. This research benefitted from review of an earlier manuscript version by S. Voigt. We thank M. Wilmsen and K. Fo¨llmi for their insightful and constructive reviews. Funding was provided by the Deutsche Forschungsgemeinschaft (grant KU 642/ B20-1).

Conclusions The two sections of the Karak –Silla intra-platform basin in Jordan, GM3 and KB3, are conspicuously cyclic as revealed from carbonate, TOC and isotope data, and supported by a symmetrical lithological architecture with repeated intercalations of gypsum beds as indicators of sea-level decrease. Stable carbon isotope, carbonate and TOC records of these two sections show good correspondence at high-resolution. Carbon isotope curves exhibit a broad positive excursion related to OAE2 and comparable to the global record. Correlation with the Pueblo GSSP section (Sageman et al. 2006) enables the establishment of an orbital timescale for the Jordan sections. Isotope events of the European records can be detected in the Jordan sections, thus supporting the wide significance of these time markers. Based on the constructed timescale a model of orbitally-forced sea-level fluctuation is presented: The Jordan third-order sequence S5 had a duration of 1.2 Ma and is composed of three fourth-order sequences (400 ka). Fifth-order fluctuations were controlled by the 100 ka eccentricity cycle. Gypsum beds indicate the lowstands of the fourthorder cycles. Negative isotope excursions are related to these levels. The Jordan and Pueblo timescales correlate precisely during OAE2 thus giving equal age determination for both regions: Base peak a to top peak c: c. 500 ka. The Early Turonian part from the C –T boundary to the Lulworth event in Jordan consists of six 100 ka eccentricity cycles. For this interval of the section, increased cycle thickness caused by elevated accumulation rates should indicate higher accommodation space during platform drowning during the Early Turonian. The section interval of GM3 representing the lower part of OAE2 contains indications of platform limestone formation. Thus, demise of the Levant platform only occurred later, namely during the

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Philip, J. & Airaud Crumiere, C. 1991. The demise of the rudist-bearing carbonate platforms at the Cenomanian/Turonian boundary; a global control. In: Montaggioni, L. F. & Macintyre, I. G. (eds) Reefs as Recorders of Environmental Changes. Coral Reefs 10; 2. Springer International, Berlin– Heidelberg–New York, International, 115– 125. Powell, J. H. 1989. Stratigraphy and sedimentation of the Phanerozoic rocks in Central and South Jordan. Pt.B: Kurnub, Ajlun and Belqa groups. NRA Geological Bulletin, 11, 130. Pratt, L. M. 1985. Isotopic studies of organic matter and carbonate in rocks of the Greenhorn Marine Cycle. In: Pratt, L. M. (ed.) Fine Grained Deposits and Biofacies of the Cretaceous Western Interior Seaway: Evidence of Cyclic Sedimentary Processes. SEPM Field Trip Guidebook, Tulsa, 4, 38–48. Reyment, R. A. 1955. The Cretaceous ammonoidea of southern Nigeria and the southern Cameroons. Bulletin Geological Survey of Nigeria, 25, 1–112. ¨ ber einige wirbellose Fossilien Reyment, R. A. 1957. U aus Nigerien und Kamerun, Westafrika. Palaeontographica, 109, 41–70. Robaszynski, F., Hardenbol, J. et al. 1993. Sequence stratigraphy in a distal environment: the Cenomanian of the Kallat Senan area. Bull. Centres Recherche Exploration-Production Elf Aquitaine, 17, 395–433. Sageman, B., Meyers, S. R. & Arthur, M. A. 2006. Orbital timescale and new C-isotope record for Cenomanian–Turonian boundary stratotype. Geology, 34, 125– 128. Schulze, F., Lewy, Z., Kuss, J. & Gharaibeh, A. 2003. Cenomanian–Turonian carbonate platform deposits in west central Jordan. International Journal of Earth Sciences, 92, 641–660. Schulze, F., Marzouk, A. M., Bassiouni, M. A. A. & Kuss, J. 2004. The late Albanian– Turonian carbonate platform succession of west-central Jordan: stratigraphy and crises. Cretaceous Research, 25, 709– 737. Sepu´lveda, J., Wendler, J., Leider, A., Kuss, J., Summons, R. E. & Hinrichs, K.-U. 2009. Molecularisotopic evidence of environmental and ecological changes across the Cenomanian– Turonian boundary in the Levant Platform of central Jordan. Organic Geochemistry, 40, 553–568. Strasser, A., Hillga¨rtner, H., Hug, W. & Pittet, B. 2000. Third-order depositional sequences reflecting Milankovitch cyclicity. Terra Nova, 12, 303– 311. Voigt, S., Gale, A. S. & Voigt, T. 2006. Sea-level change, carbon cycling and palaeoclimate during the Late Cenomanian of northwest Europe; an integrated palaeoenvironmental analysis. Cretaceous Research, 27, 836–858. Voigt, S., Erbacher, J., Mutterlose, J., Weiss, W., Westerhold, T., Wiese, F., Willmsen, M. & Wonik, T. 2008. The Cenomanian –Turonian of the Wunstorf section – (North Germany): global stratigraphic

reference section and new orbital time scale for Oceanic Anoxic Event 2. Newsletters on Stratigraphy, 43(1), 65–89. Wade, B. S. & Pa¨like, H. 2004. Oligocene climate dynamics. Paleoceanography, 19, 1 –16. Weissert, H., Lini, A., Foellmi, K. B. & Kuhn, O. 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events; a possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189–203. Wendler, J. & Willems, H. 2002. Distribution pattern of calcareous dinoflagellate cysts across the Cretaceous– Tertiary boundary (Fish Clay, Stevns Klint, Denmark); implications for our understanding of species-selective extinction. In: Koeberl, C. & MacLeod Kenneth, G. (eds) Catastrophic Events and Mass Extinctions; Impacts and Beyond. Geological Society of America (GSA), Boulder, CO. Wendler, J., Graefe, K. U. & Willems, H. 2002a. Palaeoecology of calcareous dinoflagellate cysts in the mid-Cenomanian Boreal Realm; implications for the reconstruction of palaeoceanography of the NW European shelf sea. Cretaceous Research, 23, 213–229. Wendler, J., Graefe, K. U. & Willems, H. 2002b. Reconstruction of mid-Cenomanian orbitally forced palaeoenvironmental changes based on calcareous dinoflagellate cysts. Palaeogeography, Palaeoclimatology, Palaeoecology, 179, 19– 41. Wendler, J., Wendler, I. & Kuss, H. J. 2009a. Early Turonian shallow marine red beds on the Levant carbonate platform (Jordan), Southern Tethys. SEPM Special Publication, 91, 179– 187. Wendler, I., Wendler, J., Neuhuber, S. & Wagreich, M. 2009b. Productivity fluctuations and orbital cyclicity during Early to Middle Turonian development of marine red beds. SEPM Special Publication, 91, 209–221. Wiedmann, J. & Kuhnt, W. 1996. Biostratigraphy of Cenomanian/Turonian organic carbon-rich sediments in the Tarfaya Atlantic Coastal Basin (Morocco). In: Berichte-Reports Geologisch-Pala¨ontologisches Institut der Universita¨t Kiel (Jost Wiedmann Symposium. Cretaceous Stratigraphy, Paleobiology and Paleobiogeography, Tu¨bingen, 7 –10 March 1996, Abstracts), Kiel, 76, 195– 200. Wiese, F. & Schulze, F. 2005. The upper Cenomanian (Cretaceous) ammonite Neolobites vibrayeanus (d’Orbigny, 1841) in the Middle East: taxonomic and palaeoecologic remarks. Cretaceous Research, 26, 930–946. Wilmsen, M. 2003. Sequence stratigraphy and palaeoceanography of the Cenomanian Stage in northern Germany. Cretaceous Research, 24, 525–568. Wright, C. W. & Kennedy, W. J. 1981. The Ammonoidea of the Plenus Marls and the Middle Chalk. Palaeontographical Society Monographs, 560, 134–148.

Biostratigraphy, palaeoecology and palaeogeography of the Middle Cenomanian – Early Turonian Levant Platform in Central Jordan based on ostracods ABDEL-MOHSEN M. MORSI1* & JENS E. WENDLER2,3 1

Geology Department, Faculty of Science, Ain Shams University, 11566 Cairo, Egypt

2

Bremen University, Geosience department, P.O. Box 330440, 28334 Bremen, Germany

3

Present address: Smithsonian Institution National Museum of Natural History, 10th & Constitution NW, Washington, DC 20560-0121, USA *Corresponding author (e-mail: [email protected]) Abstract: Study of a Cenomanian –Turonian sequence, including the oceanic anoxic event 2 (OAE2) in Central Jordan, yielded 22 ostracod species from the Middle– Late Cenomanian interval; no ostracods were found in the Early Turonian. The majority of the taxa have a wide geographical distribution along the southern shores of the Tethys; from Morocco in the west to the Arabian Gulf region in the east. Biogeographical homogeneity of the ostracod associations in North Africa and the Middle East reflects facilitated communication along the whole expanse of the southern Tethys margin during the Cenomanian, and suggests similar living conditions and absence of important geographical barriers that could hinder marine faunal exchange. Biostratigraphically, the investigated fauna revealed five informal ostracod biozones (I to V from older to younger). The recorded assemblages are characterized by ostracod faunas of typical marine shelf setting in biozone I, shelf lagoonal setting with fresh-water influence in biozone II, marine shelf setting with intervals of fresh-water supply in biozones III and IV, and reduced oxygen levels in the interval of biozone V. This sequence of biozones provides palaeontological evidence for the occurrence of an interval of enhanced fresh-water influence in Levant platform lagoons preceeding OAE2. A combined biostratigraphic and chemostratigraphic time scale based on stable carbon isotopes reveals the first appearance of Reticulicosta kenaanensis, previously described as an Early Turonian indicator species already in the Late Cenomanian. Absence of ostracods throughout the Early Turonian indicates environmental conditions adverse to ostracods during most of OAE2 and its aftermath interpreted to reflect strong water column stratification.

Sediments rich in organic matter reflect a significant disturbance in the global carbon cycle related to an oceanic anoxic event (OAE) at the Cenomanian– Turonian (C –T) transition interval (OAE2). Deposits of this global phenomenon are characterized by major biotic changes in microfaunas (e.g. Leckie et al. 2002). Increased deposition of organic carbon in the oceans and shelf seas has been explained by both enhanced bio-productivity and unusual preservation of the total organic carbon (TOC) (Arthur et al. 1987; Kuypers et al. 2002). While the effects of OAEs in pelagic to hemipelagic deposits have been studied intensively, the palaeoenvironmental conditions related to these significant events in the near-shore environments still need to be investigated in more detail. OAE2 was preceeded by a significant sea-level drop, followed by a rapid transgression in conjunction with the onset of a positive d13C isotopic excursion, which reflects a disturbance in the global carbon cycle (e.g. Schlanger et al. 1987; Voigt et al. 2006;

Elrick et al. 2009; Wendler et al. 2010). While being defined as an oceanic event, this anomaly in the life system of Earth also involves the shallow marine and terrestrial environments (e.g. Hasegawa 1997; Davey & Jenkyns 1999). Regarding shallow marine environments, OAE2 is well-developed in North African shelf –sea deposits, including the Levant carbonate platform of the southern rim of the Tethys. The aim of this paper is to study the ostracod record in the Cenomanian–Turonian section of Ghawr Al-Mazar (GM3) in order to enhance the biostratigraphical framework of the study area, and reconstruct the mid-Cenomanian through Early Turonian palaeo-environmental conditions related to OAE2. The ostracod data are being integrated into a biostratigraphical scheme based on calcareous nannoplankton and ammonites. Biostratigraphy is supported here by a high-resolution stable carbon isotope stratigraphy as an independent chemostratigraphical tool. The ostracod assemblages discussed

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 187–210. DOI: 10.1144/SP341.9 0305-8719/10/$15.00 # The Geological Society of London 2010.

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in the present paper represent the prime palaeontological tool for reconstruction of the palaeoecology of the study area because other fossils are scarce and fail to provide continuous records over the section. Previous studies of Cenomanian–Turonian ostracods of Jordan comprise Babinot & Basha (1985), Powell (1989), Basha (1997), and Schulze et al. (2004).

by an ammonite assemblage with Vascoceras durandi, Thomasites rollandi, Fagesia lenticularis, Choffaticeras quaasi and Ch. luciae indicative of zones T5 –6a. Transition into zone 6b (M. nodosoides/C. woolgari zone boundary, base Middle Turonian) was detected by Schulze et al. (2003, 2004) in the middle Wala Limestone.

Material and methods Location and stratigraphic setting The studied section (GM3) is situated at Ghawr Al-Mazar (318150 3400 N; 358350 4100 E) in central Jordan. It crops out in the Wadi system cutting east– west into an extended plateau area east of the Dead Sea (Fig. 1). The Karak Limestone Member is not well-exposed in the studied section. Therefore, a section representing this member at Mujib Dam (MD5), NNE of Ghawr Al-Mazar, was logged into the profile (see Fig. 2). The investigated interval comprises the Middle Cenomanian –Lower Turonian, which incorporates the Fuheis, Hummar, and Shueib Formations (Masri 1963), being subdivisions of the Ajlun Group (Quennel 1951).

Biostratigraphy The multi-biostratigraphical framework of Schulze et al. (2003), using nannofossils and ammonites supported partly by larger benthic foraminifera, forms the base of the stratigraphy presented here (Fig. 2). The section comprises the calcareous nannoplankton zones CC-9 to CC-11. For details on indicative species see Schulze et al. (2003). The ammonite occurrences in the Cenomanian–Turonian of Jordan (Schulze et al. 2004) provide a relatively detailed biostratigraphy correlated with ammonite zone schemes of southern Europe (Hardenbol et al. 1998) and Israel (Lewy 1989, 1990). In conjunction with the isotope record presented here it enables an excellent time control. The Fuheis and Hummar formations (0–56 m) contain abundant Neolobites vibrayeanus indicating a Middle to Late Cenomanian age (Fig. 2). In section part 38– 43 m, it co-occurs with the larger benthic foraminifer Praealveolina cretacea marking the basal Late Cenomanian. Vascoceras cauvini, Metoicoceras geslinianum and Burroceras transitorium indicate ammonite zone T1 for the basal Shueib Formation, correlating with the M. geslinianum and N. juddi zones of Europe. Ammonite zones T2 –4 correlate to the basal CC-11 and can be detected only by scarce and poorly preserved not in situ findings of Choffaticeras pavillieri and Ch. quaasi in section part 67 –74 m. The limestone in the top of the section represents the Wala Limestone Member, which is the ammonite marker bed 3 of Schulze et al. (2003, 2004). It is characterized at the base

A total of 255 samples from the GM3 (Ghawr Al-Mazar) section were collected at 15 –25 cm sample spacing, attaining an 83 metre-thick Middle Cenomanian–Lower Turonian profile. The section exhibits the following lithological units: 0–9 m nodular limestone rich in oysters; 9–18 m dark grey clay with grey nodular limestone beds; 18–40 m (Karak Limestone) laminated limestone alternating with nodular and massive platform limestone and marl beds (recorded at section MD5 (Mujib Dam 5; see Fig. 2; 79 samples); 40–43 m (Hummar Limestone) massive limestone with larger benthic foraminifera; 43–54.8 m green clays and marls; 54.8–57.3 m dolomitic limestone; 57.3–64.8 m platy, bituminuous limestone bed/ brown marly clay alternation; 64.8–74 m brown calcareous clays; 74–80 m grey marls; 80–83 m (Wala Limestone) massive platform limestone. The section presented here (Fig. 2) is a composit with the section part MD5 inserted into section GM3. MD5 represents the most complete record of the Karak Limestone and the Hummar Limestone, which vary considerably in thickness owing to lateral facies shifts. The latter is especially significant for the Hummar Limestone Member, which is replaced in the GM3 section by clays. This lateral facies change is plausible from field observation, mapping (Powell 1989) and the repetitive isotope record at that position (Fig. 2). A gap in the ostracod record in the MD5 section part is due to the fact that this limestone succession is unsuitable for disaggregational processing and isolation of ostracod tests. Abundant ostracods can be observed in thin sections of this material; however, a determination of these specimens on a species level is impossible. 100 g of sample were disaggregated by repeated freezing and thawing in saturated sodium sulphate (Glauber’s salt) solution, and washed through sieves at mesh sizes of 20, 63, 100 and 630 mm. The sieved fractions were cleaned by applying 0.5–2 min ultrasonic treatment with parallel microscopic controlling to avoid fossil specimen destruction. Additional cleaning with

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REWOQUAT-tenside was applied when high amounts of clay were present. For intervals with very high amounts of fine fraction up to 500 g of sample were prepared in a second processing in order to obtain more corse fraction and micropalaeontological specimens therin. The 100 –630 mm size range was then dry-sieved into fractions 100– 250, 250 –400 and .400 mm for systematic assessment of ostracods. Systematic subsamples of the 100 –250 mm fraction were prepared using a micro splitter. Ostracods were picked at a binocular microscope at 50 magnification. All samples were analysed, but only the samples indicated in Figure 2

yielded ostracods. For taxonomic purpose and documentation, selected specimens of the ostracod faunas were analysed and photographed using a Cam Scan SEM at Bremen University, Germany. Samples are stored at the Geology Department, Faculty of Science, Ain Shams University (Cairo, Egypt). Reference numbers (JC-01 to JC-57) are given only to the illustrated specimens. Stable carbon isotopes of bulk carbonate (complete sample set) were measured at the isotope laboratory of Bremen University, Germany, using a Finigan MAT 251 mass spectrometer. The results are reported relative to the PDB-standard.

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Fig. 2. Middle Cenomanian –Early Turonian section (GM3) showing formations, members, calcareous nannoplankton zones, ammonite zones, and stable carbon isotope record. English chalk reference curve from Jarvis et al. (2006), note change in scale (carbon isotopes and age) below OAE2 for enlargement of the Middle– Late Cenomanian part of the record. Indicated isotope events and intervals mark long-term inflection points. For details in section interval 47– 83 m (OAE2) see Wendler et al. (2010). Sample numbers indicate samples with ostracods. Shaded lithological intervals are black shales. Right panel: ostracod assemblages.

Isotope stratigraphy Stable carbon isotopes of bulk carbonate are in the range of –4 to þ4‰. Compared to deeper marine records the present data reveal a relatively noisy signal with an enhanced amplitude range. The record contains some significant negative excursions, the most significant one occurring just

below OAE2. Major positive excursions of the record are distinctive and can be used by correlation for a chemostratigraphic subdivision of the Middle Cenomanian through Early Turonian. Correlation of stable carbon isotope events recorded in the Jordan data with the English Chalk reference record of Jarvis et al. (2006) is given in Figure 2. From the bottom to the top of the investigated

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

section, these d13C-excursions can be correlated to the major global isotope events mid-Cenomanian event 1 (MCE1), Jukes-Brown Event, and OAE2, as well as some minor peaks (Fig. 2). The record is embedded in the multi-biostratigraphical framework given above, and supported by the detailed analysis of the OAE2 (see Wendler et al. 2010). The timescale of Ogg et al. (2004) that was applied to the English Chalk reference d13C-record and the orbital timescale established for the GSSP stratotype (Sageman et al. 2006) can be transferred by this correlation to the Jordan record for age control. The isotope record is used here mostly for stratigraphical control. An in-depth discussion, being beyond the scope of this paper, is given in a separate publication (Wendler et al. 2010).

Systematic descriptions The studied fauna yielded 22 ostracod species (Fig. 3) that are assigned to 16 genera belonging to 10 families. The classification has been made following Horne et al. (2002). Species synonymies are abbreviated; synonymy lists include the first naming of species and the names other than those utilized herein. Complete records of the species are given in Figure 4. Morphological and taxonomic comments are given wherever necessary. Subclass PODOCOPA Mu¨ller, 1894 Order PLATYCOPIDA Sars, 1866 Suborder PLATYCOPINA Sars, 1866 Superfamily CYTHERELLOIDEA Sars, 1866 Family CYTHERELLIDAE Sars, 1866 Genus Cytherella Jones, 1849 Cytherella aegyptiensis Colin & El Dakkak, 1975 Figure 5, figs 1–3 1974 Cytherella gr. ovata (Roemer); Rosenfeld & Raab, p. 3, pl. 1, figs 3– 5. 1975 Cytherella aegyptiensis, Colin & El Dakkak, p. 50, pl. 1, figs 2– 3. 1991 Cytherella cf. eosulcata, Colin; Shahin (1991) p. 133, pl. 1, figs 3–4. 1994 Cytherella ahmadiensis, Al-Abdul-Razzaq; Shahin et al. (1994) p. 36, pl. 1, figs 1 –2. Material: 384 specimens. Dimensions: Length: 0.76– 0.82 mm; height: 0.51 – 0.54 mm; width: 0.37– 0.39 mm. Stratigraphical and geographical distribution: This species is known in the Cenomanian of Egypt (Colin & El Dakkak 1975; Shahin et al. 1994; Morsi & Bauer 2001; Szczechura et al. 1991) and Morocco (Andreu-Boussut 1991), Cenomanian –Turonian of Israel (Rosenfeld & Raab 1974) and Middle Cenomanian of Jordan (Schulze et al. 2004).

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Occurrence in the studied section: Middle– Upper Cenomanian. Cytherella dhalalensis Morsi & Bauer, 2001 Figure 5, figs 4 –5 1991 Cytherella sp. 5, Andreu-Boussut, p. 450, pl. 5, figs 5– 10. 2001 Cytherella dhalalensis, Morsi & Bauer, p. 383, pl. 1, figs 3–5. Material: 20 specimens. Dimensions: Length: 0.63 –0.68 mm; height: 0.38– 0.42 mm; width: 0.29 –0.31 mm. Stratigraphical and geographical distribution: Middle Cenomanian of Morocco (Andreu-Boussut 1991) and Upper Cenomanian of Egypt (Morsi & Bauer 2001) and Jordan (present section). Cytherella cf. gambiensis Apostolescu, 1963 Figure 5, figs 6 –7 cf. 1963 Cytherella gambiensis, Apostolescu, p. 1680, pl. 1, figs 1 –3. 2002 Cytherella cf. gambiensis, Apostolescu; Bassiouni, p. 91, pl. 22, figs 13 –15. Material: 28 specimens. Dimensions: Length: 0.63 –0.69 mm; height: 0.42– 0.51 mm; width: 0.30 –0.31 mm. Stratigraphical and geographical distribution: This species was recorded in the Lower Turonian of Egypt (Bassiouni 2002) and Middle–Upper Cenomanian of Jordan (Schulze et al. 2004). Occurrence in the studied section: Middle Cenomanian. Cytherella sp. Figure 5, figs 8 –9 Material: 11 specimens. Dimensions: Length: 0.64 mm; height 0.41 mm; width: 0.31 mm. Description: Carapace subrectangular in lateral outline. Dorsal margin of right valve slightly convex; dorsal margin of left valve straight. Ventral margin slightly concave. End margins rounded, smoothly joined to anteriorly converging longitudinal margins. Right valve larger than left valve; maximum overlap at mid-dorsum and midventrum. Maximum length central; maximum height behind middle, at posterior one third of length. Lateral surface smooth, with pronounced mediodorsal shallow sulcus. Anterior margin occupied by a marginal rim. In dorsal view, carapace elongated ovate with greatest width at posterior one third of length. Sexual dimorphism pronounced, males are thinner than females. Remarks: Cytherella dhalalensis Morsi & Bauer (2001), which originally comes from the Late

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Fig. 3. Stratigraphical distribution of ostracods in the studied section.

Portugal Jugoslavia

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Fig. 4. Geographical distribution of ostracod taxa. Vranc., Vranconian; Apt., Atian; Cenom., Cenomanian; Turon., Turonian; l., lower; m., middle; u., upper.

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Fig. 5. 1– 3: Cytherella aegyptiensis Colin & El Dakkak (1975). Hummar Formation, sample GM3-109, 1, JC-01, L 0.78 mm, LVC; 2, JC-02, L 0.77 mm, RVC; 3, JC-03, W 0.39 mm, DVC; 4– 5: Cytherella dhalalensis Morsi & Bauer (2001). Hummar Formation, sample GM3-114, 4, JC-04, W 0.29 mm, DVC; 5, JC-05, L 0.63 mm, LVC; 6– 7: Cytherella cf. gambiensis Apostolescu (1963). Fuheis Formation, sample GM3-31, 6, JC-06, L 0.63 mm, LVC; 7, JC-07, W 0.30 mm, DVC; 8– 9: Cytherella sp. Shueib Formation, sample GM3-152, 8, JC-08, W 0.31 mm, DVC; 9, JC-09, L 0.64 mm, LVC; 10: Bairdia youssefi Bassiouni (2002). Fuheis Formation, sample GM3-05, JC-10, L 0.88 mm, RVC; 11: Bairdia sp. Shueib Formation, sample GM3-162, JC-11, L 0.83 mm, RVV; 12–13: Bythocypris sp. 1

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Cenomanian of Egypt, is similar to the present species. However, it deviates by its more symmetrical posterior margin and having the dorsal margin joined to pthe osterior margin at posterior one fifth of length; in the present species, the posterior margin is asymmetrical and the dorsal margin joins the posterior margin at posterior one third of length. Cytherella sp. 11 Andreu-Boussut (1991) from the Late Cenomanian of Morocco also resembles the present species. However, it differs in having a straight ventral margin, unlike the present species which has a concave ventral margin. Stratigraphical and geographical distribution: Upper Cenomanian. Order PODOCOPIDA Mu¨ller, 1894 SuborderBAIRDIOCOPINA Sars, 1968 Superfamily BAIRDIOIDEA Sars, 1968 Family BAIRDIIDAE Sars, 1888 Genus Bairdia McCoy, 1844 Bairdia youssefi Bassiouni, 2002 Figure 5, fig. 10 1991 Bairdia sp. 1, Andreu-Boussut, p. 470, pl. 11, figs 1–6. 2002 Bairdia youssefi, Bassiouni, p. 15, pl. 1, figs 1–4. Material: 8 specimens. Dimensions: Length: 0.88 mm; height: 0.51 mm; width: 0.47 mm. Stratigraphical and geographical distribution: This species was previously recorded from the Lower Cenomanian of Egypt (Bassiouni 2002) and Middle Cenomanian of Morocco (Andreu-Boussut 1991). Occurrence in the studied section: Middle Cenomanian. Bairdia sp. Figure 5, fig. 11 Material: 8 specimens. Dimensions: Length: 0.83 mm; height: 0.50 mm. Remarks: The present species is represented only by separate valves. They differ from Bairdia youssefi Bassiouni (2002) in having a broader anterior margin and shorter caudal process.

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Stratigraphical and geographical distribution: Upper Cenomanian. Family BYTHOCYPRIDIDAE Maddocks, 1969 Genus Bythocypris Brady, 1880 Bythocypris sp.1, Rosenfeld & Raab, 1974 Figure 5, figs 12 –13 1974 Bythocypris sp.1, Rosenfeld & Raab, p. 6, pl. 1, figs 17–18. Material: 5 specimens. Dimensions: Length: 0.85 mm; height: 0.45 mm; width: 0.38 mm. Stratigraphical and geographical distribution: This species was recorded from the Upper Cenomanian of Israel (Rosenfeld & Raab 1974) and Egypt (Morsi & Bauer 2001). Occurrence in the studied section: Middle Cenomanian. Suborder CYPRIDOCOPINA Jones, 1901 Superfamily CYPRIDOIDEA Baird, 1845 Family CANDONIDAE Kaufmann, 1900 Subfamily PARACYPRIDINAE Sars, 1923 Genus Paracypris Sars, 1923 Paracypris dubertreti, Damotte & Saint-Marc, 1972 Figure 5, figs 14 –15 1972 Paracypris dubertreti, Damotte & Saint-Marc, p. 276, pl. 1, fig. 1. 1974 Paracypris acutocaudata, Rosenfeld; in Rosenfeld & Raab, p. 8, pl. 1, figs 22– 24. 1991 Paracypris cf. dubertreti, Damotte & Saint-Marc; Andreu-Boussut, p. 485, pl. 18, fig. 9. 1999 Paracypris acuta (Cornuel); Ismail, p. 309, pl. 3, figs 14–15. Material: 14 specimens. Dimensions: Length: 0.76 –0.78 mm; height: 0.30– 0.31 mm. Remark: complete synonymy and arguments for placing P. acutocaudata Rosenfeld as a junior synonym for P. dubertreti are found in Bassiouni (2002). Stratigraphical and geographical distribution: This species was first described from the Middle and

Fig. 5. (Continued) Rosenfeld & Raab (1974). Fuheis Formation, sample GM3-05, 12, JC-012, L 0.85 mm, RVC; 13, JC-13, W 0.38 mm, DVC; 14–15: Paracypris dubertreti Damotte & Saint-Marc (1972). 14, Fuheis Formation, sample GM3-31, JC-14, L 0.76 mm, LVC; 15, Hummar Formation, sample GM3-100, JC-15, L 0.78 mm, RVC; 16–17: Paracypris mdaourensis Bassoullet & Damotte (1969). Shueib Formation, sample GM3-162, 16, JC-16, L 0.85 mm, RVC; 17, JC-17, L 0.85 mm, LVC; 18: Monoceratina? trituberculata Rosenfeld (1974). Hummar Formation, sample GM3-111, JC-18, L 0.53 mm, LVC; 19– 21: Dolocytheridea atlasica Bassoulet & Damotte (1969). Hummar Formation, sample GM3-110, 19, JC-19, female, W 0.25 mm, DVC; 20, JC-20, female, L 0.50 mm, RVC; 21, JC-21, male, L 0.52 mm, RVC. Abbreviations: RVC, right view carapace; LVC, left view carapace; DVC, dorsal view carapace; LVV, left view valve; RVV, right view valve; L, length; W, width.

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Upper Cenomanian of Lebanon (Damotte & SaintMarc 1972), and subsequently from the Aptian to Upper Cenomanian of Israel (Rosenfeld & Raab 1974, 1984), Middle Cenomanian of Morroco (Andreu-Boussut 1991), Cenomanian –Lower Turonian of Algeria (Viviere 1985, Majoran 1989) and Aptian –Albian and Cenomanian of Egypt (Boukhary et al. 1977; Shahin et al. 1994; Ismail 1999; Morsi & Bauer 2001; Hewaidy & Morsi 2001; Bassiouni 2002). Occurrence in the studied section: Upper Cenomanian.

1979); Middle and Upper Cenomanian of Morocco (Andreu-Boussut 1991); Cenomanian of Tunisia (Bismuth et al. 1981a; Ben Youssef 1980; Gargouri-Razgallah 1983) and Upper Cenomanian of Egypt (Shahin et al. 1994; Morsi & Bauer 2001; Bassiouni 2002). Occurrence in the studied section: Upper Cenomanian.

Paracypris mdaourensis Bassoullet & Damotte, 1969 Figure 5, figs 16–17 1969 Paracypris mdaourensis, Bassoullet & Damotte, p. 140, pl. 1, fig. 10a –d.

Dolocytheridea atlasica Bassoullet & Damotte 1969 Figure 5, figs 19 –21

Material: 8 specimens. Dimensions: Length: 0.85 mm; height: 0.32 mm; width: 0.31 mm. Stratigraphical and geographical distribution: The present species was recorded in the Lower Turonian of Algeria (Bassoullet & Damotte 1969; Viviere 1985). It was also reported from the Vraconian to Upper Cenomanian of Morocco (Andreu-Boussut 1991), Lower Cenomanian to Lower Turonian of Israel (Rosenfeld & Raab 1974; Lipson-Benitah et al. 1985), Albian –Turonian of Egypt (Shahin 1991; Ismail 1999; Morsi & Bauer 2001; Bassiouni 2002) and Middle Cenomanian of Jordan (Schulze et al. 2004). Occurrence in the studied section: Upper Cenomanian. Suborder CYTHEROCOPINA Jones, 1901 Superfamily CYTHEROIDEA Baird, 1850 Family BYTHOCYTHERIDAE Sars, 1926 Genus Monoceratina Roth, 1928 Monoceratina? trituberculata Rosenfeld, 1974 Figure 5, fig. 18 1974 Monoceratina? Trituberculata, Rosenfeld, in Rosenfeld & Raab, p. 11, pl. 2, figs 10 –11; pl. 4, fig. 6. 1979 Exophthalmocythere? Bituberculata, AlAbdul-Razzaq; Grosdidier, pl. 9, figs 52a–d. 1985 Perissocytheridea? Trituberculata, (Rosenfeld); Viviere, p. 150, pl. 5, figs 1–3. Material: 2 specimens. Dimensions: Length: 0.53 mm; height: 0.31 mm. Stratigraphical and geographical distribution: This species is widely known from the Upper Cenomanian of Israel (Rosenfeld & Raab 1974), Algeria (Viviere 1985); Cenomanian of Gabon (Grosdidier

Family CYTHERIDEIDAE Sars, 1925 Subfamily CYTHERIDEINAE Sars, 1925 Genus Dolocytheridea Triebel, 1938

1969 Dolocytheridea atlasica, Bassoullet & Damotte, p. 139, pl. 2, fig. 9a–d. 1973 Dolocytheridea cf. atlasica, Bassoullet & Damotte; Grosdidier (1973), pl. 3, fig. 22. 1985 Dolocytheridea aff. Atlasica, Bassoullet & Damotte; Viviere (1985), p. 154, pl. 4, figs 1–2. Material: 782 specimens. Dimensions: Length: 0.50 mm; height: 0.29 mm; width: 0.25 mm (female). Length: 0.52 mm; height: 0.28 mm (male). Stratigraphical and geographical distribution: Dolocytheridea atlasica is widely known from the Lower and Upper Cenomanian of Algeria (Bassoullet & Damotte 1969; Viviere 1985) and Israel (Rosenfeld & Raab 1974), Cenomanian of Tunisia (Ben Youssef 1980), Upper Albian –Cenomanian of Oman (Babinot & Bourdillon de Grissac 1989; Colin et al. 2001), Upper Albian and Cenomanian of Iran (Grosdidier 1973) and (?)Upper Albian – Turonian of Egypt (Shahin et al. 1994; Colin & El Dakkak 1975; Boukhary et al. 1977; Szczechura et al. 1991; Ismail 2001; Morsi & Bauer 2001; Bassiouni 2002). In Jordan, it was recorded in the Middle Cenomanian (Schulze et al. 2004). Occurrence in the studied section: Middle and Upper Cenomanian. Genus Neocyprideis Apostolescu, 1956 Neocyprideis vandenboldi Gerry & Rosenfeld, 1973 Figure 6, figs 2–3 1964 Fabanella?, sp. A, Bold, p. 120, pl. 14, fig. 5a –d. 1973 Neocyprideis vandenboldi, Gerry & Rosenfeld, p. 103, pl. 1, figs 1–9; pl. 2, figs 1 –6. Material: 1114 specimens. Dimensions: Length: 0.81– 0.83 mm; height: 0.47 – 0.50 mm; width: 0.35 –0.40 mm.

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

197

Stratigraphical and geographical distribution: The present species was previously recorded in the Upper Cenomanian and Lower Turonian of Israel (Gerry & Rosenfeld 1973; Rosenfeld & Raab 1974; Lipson-Benitah et al. 1985) and Egypt (Bold 1964; Shahin 1991; Bassiouni 2002). In Jordan, it was reported from the Middle Cenomanian (Schulze et al. 2004). Occurrence in the studied section: Upper Cenomanian.

Remarks: The present species differs from Schuleridea baidarensis (Damotte & Saint-Marc 1972) in having a narrower posterior end and less steeply downward sloping posterior outline. Stratigraphical and geographical distribution: This species was previously recorded from the Albian of Lebanon (Bischoff 1990) and Lower Cenomanian of Israel (Rosenfeld & Raab 1974). Occurrence in the studied section: Upper Cenomanian.

Genus Perissocytheridea Stephenson, 1938

Family CYTHERURIDAE Mu¨ller, 1894 Genus Metacytheropteron Oertli, 1957

Perissocytheridea istriana Babinot, 1988 Figure 6, figs 4–7

Metacytheropteron berbericum Damotte 1969) Figure 6, figs 8 –13

1988 Perissocytheridea istriana; Babinot, p. 8, pl. 2, figs 1–7. 1991 Pterygocythere? sp. 1, Andreu-Boussut, p. 612, pl. 42, figs 12– 16. pars 1991 Perissocytheridea ignota, Szczechura et al., p. 18, pl. 6, figs 8–11; non fig. 7. 1994 Looneyella sohni, Rosenfeld; Shahin et al. (1994), p. 54, pl. 3, figs 11 –12. 1999 Cytheropteron cf. navarroense, Alexander; Ismail (1999), p. 307, pl. 3, figs 2– 3. Material: 27 specimens. Dimensions: Length: 0.59–0.62 mm; height: 0.34 – 0.38 mm; width: 0.33 mm (female). Length: 0.63 mm; height: 0.31 mm; width: 0.31 mm (male). Stratigraphical and geographical distribution: This species was first recorded in the Cenomanian of Prementura, south Istria, Croatia (Babinot 1988). It was also illustrated from the Cenomanian of Egypt (Szczechura et al. 1991; Shahin et al. 1994; Ismail 1999; Morsi & Bauer 2001; Bassiouni 2002) and Upper Cenomanian of Morocco (Andreu-Boussut 1991). Occurrence in the studied section: Upper Cenomanian. Subfamily SCHULERIDEINAE Mandelstam, 1959 Genus Schuleridea Swartz & Swain, 1946 Schuleridea houneensis Bischoff, 1990 Figure 6, fig. 1 pars 1974 Dordoniella? D. baidarensis, Damotte & Saint-Marc; Rosenfeld & Raab (1974), p. 12, pl. 2, fig. 20; non 21–22; non pl. 4, fig. 11. 1990 Schuleridea (Schuleridea) houneensis, Bischoff, p. 109, pl. 9, fig. 121; pl. 10, figs 122 –137. Material: 28 specimens. Dimensions: Length: 0.47 mm; height: 0.30 mm.

(Bassoullet

&

1969 Cytheropteron berbericus, Bassoullet & Damotte, p. 137, pl. 2, fig. 7a– d. 1973 Metacytheropteron parnesi, Sohn; Grosdidier (1973), p. 150, pl. 6, fig. 54a–d. 1974 Metacytheropteron berbericum (Bassoullet & Damotte); Rosenfeld & Raab (1974), p. 12, pl. 2, figs 26 –28; pl. 5, figs 2 –4. 1983 Metacytheropteron pleura Al-Furaih, p. 2, pl. 1, figs 1–2. 1991 Metacytheropteron gr. parnesi Sohn; AndreuBoussut (1991), p. 563, pl. 31, figs 7 –10. 1991 Metacytheropteron cf. berbericus (Bassoullet & Damotte); Szczechura et al., p. 23, pl. 4, fig. 15; pl. 10, fig. 1. Material: 424 specimens. Dimensions: Length: 0.52 –0.57 mm; height: 0.29– 0.34 mm; width: 0.33 mm (females). Length: 0.56 –0.59 mm; height: 0.29–0.32 mm; width: 0.33 mm (males). Stratigraphical and geographical distribution: This species has a wide distribution in North Africa and the Middle East. In Algeria, where it was first described, it was recorded throughout the Cenomanian (Bassoullet & Damotte 1969; Viviere 1985; Majoran 1989). It was also found in the Upper Albian –Cenomanian of Tunisia (Ben Youssef 1980; Bismuth et al. 1981a; Gargouri-Razgallah 1983; Abdallah et al. 1995), Cenomanian of Morocco (Andreu-Boussut 1991), Israel (Rosenfeld & Raab 1974; Majoran 1989), Saudi Arabia (Al-Furaih 1983), Oman (Athersuch 1988, 1994), Iran (Grosdidier 1973) and Egypt, the (?)Albian – Cenomanian of Egypt (Colin & El Dakkak 1975; Boukhary et al. 1977; Shahin 1991; Shahin & Kora 1991; Szczechura et al. 1991; Shahin et al. 1994; Ismail & Soliman 1997; Ismail 1999, 2001; Morsi & Bauer 2001; Bassiouni 2002). In Jordan, it was also reported from the Middle Cenomanian (Schulze et al. 2004). In southern Europe, the

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Fig. 6. 1: Schuleridea houneensis Bischoff (1990). Hummar Formation, sample GM3-112, JC-22, L 0.47 mm, RVC; 2 –3: Neocyprideis vandenboldi Gerry & Rosenfeld (1973). Hummar Formation, sample GM3-112, 2, JC-23, L 0.83 mm, RVC; 3, JC-24, W 0.40 mm, DVC; 4– 7: Perissocytheridea istriana Babinot (1988). Hummar Formation, sample GM3-118, 4-6 female, 4, JC-25, W 0.33 mm, DVC; 5, JC-26, L 0.62 mm, RVC; 6, JC-27, L. 0.59 mm, LVC; 7, male, JC-28, L 0.63 mm, LVC; 8 –13: Metacytheropteron berbericum Bassoullet & Damotte (1969). Hummar Formation, sample GM3-111, 8, 12, 13 male, 9-11, female: 8, JC-29, L 0.58 mm, LVC; 9, JC-30, L 0.52 mm, LVC; 10, JC-31, L 0.55 mm, RVC; 11, JC-32, W 0.33 mm, DVC; 12, JC-33, W 0.33 mm, DVC; 13, JC-34, L 0.56 mm, RVC;

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

present species was recorded in the Cenomanian of the western Portugese Basin (Babinot et al. 1978). Occurrence in the studied section: Middle and Upper Cenomanian. Family KRITHIDAE Mandelstam, 1958 Subfamily KRITHINAE Mandelstam, 1958 Genus Pararithe van den Bold, 1958 Parakrithe andreui Bassiouni, 2002 Pl. 2, fig. 14 ?1978 Pontocyprella? sp. 14, Andreu, p. 275, pl. 39, figs 8, 10 –11. 2001. Parakrithe sp.2, Morsi & Bauer, p. 390, pl. 3, figs 12–13. 2002 Parakrithe andreui Bassiouni, p. 35, pl. 8, figs 4–11. Material: 78 specimens. Remarks: Pontocyprella? sp. 14, Andreu (Andreu 1978) from the Lower Cenomanian of Portugal fits externally well into the present species. However, it is placed with question mark in the synonymy since its internal features were not described. Dimensions: Length: 0.67–0.75 mm; height: 0.34 – 0.35 mm. Stratigraphical and geographical distribution: This species was previously recorded in the Lower and Upper Cenomanian of Egypt (Morsi & Bauer 2001; Bassiouni 2002) and Middle Cenomanian of Jordan (Schulze et al. 2004). Occurrence in the studied section: Middle Cenomanian. Family TRACHYLEBERIDIDAE SylvesterBradley, 1948 Subfamily TRACHYLEBERIDINAE SylvesterBradley, 1948 Genus Cythereis Jones, 1849 Cythereis namousensis Bassoullet & Damotte 1969 Figure 6, figs 15– 20 1969 Cythereis namousensis, Bassoullet & Damotte, p. 134, pl. 1, fig. 3a–d. Material: 256 specimens. Dimensions: Length: 0.66–0.70 mm; height: 0.37 – 0.38 mm; width: 0.36 mm (females). Length: 0.74–0.83 mm; height: 0.38 –0.43 mm; width: 0.33 mm (males).

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Stratigraphical and geographical distribution: The present species is common in the Cenomanian rocks in Algeria (Bassoullet & Damotte 1969; Majoran 1989), Israel (Rosenfeld & Raab 1974; Majoran 1989), Tunisia (Ben Youssef 1980; Bismuth et al. 1981a; Gargouri-Razgallah 1983; Abdallah et al. 1995), Egypt (Boukhary et al. 1977; Shahin et al. 1994; Shahin 1991; Ismail 2001; Morsi & Bauer 2001; Bassiouni 2002) and Jordan (Schulze et al. 2004). Occurrence in the studied section: Middle and Upper Cenomanian. Genus Peloriops Al-Abdul-Razzaq, 1979 Peloriops aegyptiaca Morsi & Bauer, 2001 Figure 7, figs 1 –2 2001 Peloriops aegyptiaca, Morsi & Bauer, p. 394, pl. 5, figs 4, 5, 8. Material: 2 specimens. Dimensions: Length: 0.58–60 mm; height: 0.32 – 0.33 mm. Stratigraphical and geographical distribution: This species was first described from the Upper Cenomanian of Egypt (Morsi & Bauer 2001). Occurrence in the studied section: Middle Cenomanian. Genus Reticulocosta Gru¨ndel, 1974 Reticulocosta kenaanensis (Rosenfeld 1974, in Rosenfeld & Raab 1974) Figure 7, figs 3 –8 1974 Cythereis rawashensis kenaanensis Rosenfeld, n. ssp., in Rosenfeld & Raab, p. 17, pl. 3, figs 23 –25; pl. 6, figs 5 –6. 2002 Reticulicosta kenaanensis (Rosenfeld), Bassiouni, p. 77, pl. 18, figs 9, 10. Material: 328 specimens. Dimensions: Length: 0.81 –0.84 mm; height: 0.44– 0.46 mm; width: 0.35 –0.36 mm (females). Length: 1.00 mm; height: 0.51 mm; width: 0.36 mm (males). Stratigraphical and geographical distribution: This species was previously recorded from levels assigned to the Lower Turonian in Israel (Rosenfeld & Raab 1974), Egypt (Bassiouni 2002) and Jordan (Schulze et al. 2004). Occurrence in the studied section: Upper Cenomanian.

Fig. 6. (Continued) 14: Parakrithe andreui Bassiouni (2002). Fuheis Formation, sample GM3-05, JC-35, L 0.75 mm, RVC; 15– 20: Cythereis namousensis Bassoullet & Damotte (1969). Hummar Formation, sample GM3-111, 15, 16, 19, female, 17, 18, 20, male: 15, JC-36, L 0.66 mm, RVC; 16, JC-37, W 0.36 mm, DVC; 17, JC-38, L 0.83 mm, RVC; 18, JC-39, 0.33 mm, DVC; 19, JC-40, L 0.74 mm, LVC; 20, JC-41, L 0.79 mm, LVC.

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Fig. 7. 1, 2: Peloriops aegyptiaca Morsi & Bauer (2001). Fuheis Formation, 1, sample GM3-42, JC-42, L 0.54 mm, LVC; 2, sample G3-44, JC-43, L 0.56 mm, RVC; 3 –8: Reticulocosta kenaanensis Rosenfeld & Raab (1974). Shueib Formation, sample GM3-167, 3-5, 7, 8, female, 6, male: 3, JC-44, W 0.35 mm, DVC; 4, JC-45, L 0.84 mm, RVC; 5, JC-46, L 0.81 mm, LVC; 6, JC-47, L 1.00 mm, LVC; 7, JC-48, L 0.82 mm, RVC; 8, JC-49, W 0.35 mm, DVC; 9– 11, 13: Veeniacythereis maghrebensis Bassoullet & Damotte (1969). Fuheis Formation, sample GM3-10, 9-11,

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

Genus Veeniacythereis Gru¨ndel, 1973 Veeniacythereis maghrebensis (Bassoullet & Damotte, 1969) Figure 7, figs 9–11, 13 1969 Cythereis maghrebensis, Bassoullet & Damotte, p. 133, pl.1, fig. 2a –e. pars 1974 Veeniacythereis jezzineensis (Bischoff); Rosenfeld & Raab, p. 21, pl. 3, fig. 30, non figs 28–29, 31 –33. pars 1975 Veeniacythereis jezzineensis (Bischoff); Colin & El-Dakkak, p. 56, pl.1, figs 11–12; pl. 2, fig. 2, non pl. 2, fig. 1. pars 1985 Veeniacythereis gr. jezzineensis (Bischoff ); Viviere, p. 185, pl. 11, fig. 7, non pl. 11, figs 5–6, 8–11. 1991 Veeniacythereis jezzineensis (Bischoff ); Szczechura et al. p. 28, pl. 7, figs 4, 9–11; pl. 10, fig. 5. Material: 255 specimens. Dimensions: Length: 0.82–0.84 mm; height: 0.53 – 0.54 mm; width: 0.48 mm (females). Length: 0.87 mm; height: 0.54 mm; width: 0.48 mm (male). Stratigraphical and geographical distribution: This species is widely known from the Upper Cenomanian of Algeria (Bassoullet & Damotte 1969; Viviere 1985), Tunisia (Bismuth et al. 1981a; Gargouri-Razgallah 1983) and Kuwait (Al-AbdulRazzaq & Grosdidier 1981); Cenomanian of Oman (Athersuch 1988, 1994) and Israel (Rosenfeld & Raab 1974), (?) Albian –Cenomanian of Egypt (Colin & El Dakkak 1975; Boukhary et al. 1977; Shahin 1991; Szczechura et al. 1991; Shahin et al. 1994; Ismail & Soliman 1997; Morsi & Bauer 2001; Bassiouni 2002). In Jordan, it was reported from the Middle Cenomanian (Schulze et al. 2004). Occurrence in the studied section: Middle and Upper Cenomanian. Veeniacythereis streblolophata schista Al-AbdulRazzaq & Grosdidier, 1981 Figure 7, figs 12, 14, 15 1973 Cythereis IR C 4, Grosdidier, pl. 8, fig. 66a –d. pars 1975 Veeniacythereis jezzineensis (Bischoff); Colin & El Dakkak, p. 56, pl. 2, fig. 1, non pl. 1, figs 11 –12; non pl. 2, fig. 2.

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pars 1981a Veeniacythereis streblolophata, Al-Abdul-Razzaq; Bismuth et al. (1981a), p. 233, pl. 10, figs 5 –7, non figs 3–4. 1981. Veeniacythereis streblolophata schista, Al-Abdul-Razzaq & Grosdidier, 185, pl. 2, figs 1– 5. pars 1985 Veeniacythereis gr. jezzineensis (Bischoff); Viviere, p. 185, pl. 11, fig. 5, non pl. 11, figs 6– 11. 1991 Veeniacythereis ex. gr. streblolophata Al-Abdul-Razzaq & Grosdidier; Szczechura et al. (1991), p. 29, pl. 7, figs 1–3, 5–8. 1991 Veeniacythereis gr. jezzineensis (Bischoff ); Andreu-Boussut, p. 659, pl. 67, fig. 5. 1991 Veeniacythereis jezzinensis (Bischoff ); Shahin, p. 144, pl. 3, figs 8 –9. 1997 Cythereis cf. canteriolata (Crane); Ismail & Soliman, p. 178, pl. 3, figs 19, 20. 1997 Cythereis gapensis (Alexander); Ismail & Soliman, p. 180, pl. 3, figs 21, 22. 1997 Veeniacythereis jezzinensis (Bischoff ); Ismail & Soliman, p. 182, pl. 3, figs 14, 15. Material: 16 specimens. Dimensions: Length: 0.60 –0.62 mm; height: 0.36– 0.39 mm (females). Length: 0.67 mm; height: 0.38 mm (male). Stratigraphical and geographical distribution: The present subspecies was previously found in the Arabian Gulf region from the Albian – Lower Cenomanian of Iran (Grosdidier 1973), Cenomanian of Oman (Athersuch 1988, 1994) and Upper Cenomanian of Kuwait (Al-AbdulRazzaq & Grosdidier 1981). In North Africa, it was recorded in the Cenomanian of Morroco (Andreu-Boussut 1991) and Tunisia (Bismuth et al. 1981a; Garghouri-Razgallah 1983), Lower Cenomanian of Algeria (Viviere 1985) and (?) Albian –Cenomanian of Egypt (Colin & El Dakkak 1975; Shahin 1991; Szczechura et al. 1991; Shahin et al. 1994; Ismail & Soliman 1997; Morsi & Bauer 2001; Bassiouni 2002). In Jordan, it was reported from the Middle Cenomanian (Schulze et al. 2004). Occurrence in the studied section: Middle and Upper Cenomanian. Subfamily BRACHYCYTHERINAE Puri, 1954 Genus Brachcythere Alexander, 1933

Fig. 7. (Continued) female, 9, JC-50, L 0.83 mm, RVC; 10, JC-51, L 0.84 mm, RVC; 11, JC-52, L 0.82 mmRVC; 13, male, JC-53, L 0.87 mm, RVC; 12, 14, 15: Veeniacythereis streblolophata schista Al-Abdul-Razzaq and Grosdidier (1981). Fuheis Formation, sample GM3-37, 12, female, JC-54, L 0.60 mm, RVC; 14, female, JC-55, L 0.62 mm, LVC; 15, male, JC-56, L 0.67 mm, RVC; 16: Brachycythere gr. sapucariensis Kro¨mmelbein (1964). Shueib Formation, sample GM3-167, JC-57, L 0.71 mm, RVC.

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Brachycythere gr. sapucariensis Kro¨mmelbein, 1964 Figure 7, fig. 16 gr. 1964 Brachycythere sapucariensis, Kro¨mmelbein, p. 490, pl. 44, figs 1–5. 1991 Brachycythere gr. sapucariensis, Kro¨mmelbein; Andreu-Boussut (1991), p. 602, pl. 44, figs 1 –11. Material: A single specimen. Dimensions: Length: 0.71 mm; height: 0.51 mm. Stratigraphical and geographical distribution: Brachycythere sapucariensis s.s. was originally described by Kro¨mmelbein (1964) from the Coniacian –Santonian of NE Brazil. It was subsequently recorded in the interval from the Early Cenomanian–Early Coniacian in Brazil and West Africa (e.g. Kro¨mmelbein 1996; Neufville 1973, 1979; Grosdidier 1979; Andreu-Boussut 1991; Viviers et al. 2000). In North Africa and the Middle East, related forms were found in the Early Turonian– Early Coniacian of Tunisia (Bismuth et al. 1981a, b), Algeria (Viviere 1985), Egypt (Shahin 1991; Shahin et al. 1994) and Oman (Athersuch 1988). Occurrence in the studied section: Upper Cenomanian.

Ostracod biostratigraphy Ostracod biozonations were suggested by Rosenfeld & Raab (1974) in Israel; Bismuth et al. (1981a) in Tunisia; Athersuch (1988, 1994) in the Arabian Gulf area; Hataba & Ammar (1990), Shahin et al. (1994) and Ismail (2001) in Egypt; Andreu (2002) in Morocco; and Damotte (1995) for North Africa. In the present section, ostracods are significant contributors to the microfauna in the Middle and Late Cenomanian (Schulze et al. 2003, 2004), where several horizons yield a variably diversified ostracod fauna (Fig. 3). On the contrary, no ostracods have been retrieved from the Lower Turonian. Most of the elements recorded in the present section were previously utilized by the above mentioned authors to characterize different parts of the Cenomanian. However, some elements were also recorded in other regions from the Albian and some also extend higher into the Turonian (Fig. 4). Based on relative age dating in previous records, the species making up the associations found herein can be differentiated into: Aptian/Albian – Cenomanian taxa, which include species ranging from the Aptian/Albian to the Cenomanian, represented by Bairdia youssefi, Paracypris dubertreti, Schuleridea houneensis and Metacytheropteron berbericum; Cenomanian taxa, consisting of species known only for the Cenomanian, represented by C. dhalalensis, Bythocypris sp. 1 Rosenfeld & Raab, Parakrithe andreui, Monoceratina?

trituberculata, Perissocytheridea istriana, Cythereis namousensis, Peloriops aegyptiaca, Veeniacythereis maghrebensis (questionable in the Albian) and V. streblolophata schista (questionable in the Albian); and Cenomanian –Early Turonian taxa, including species crossing the C– T boundary such as Cytherella aegyptiensis, Cytherella cf. gambiensis, Neocyprideis vandenboldi, Reticulocosta kenaanensis and Brachycythere gr. Sapucariensis, and besides those, Paracypris mdaourensis and Dolocytheridea atlasica, which were recorded since the Albian, extending up to the Lower Turonian. In terms of biostratigraphical zonation, the following informal ostracod biozones I to V are recognizable in the section of Ghawr Al-Mazar (GM3) from base to top (Figs 2 & 3). These biozones are correlated with the different biozones established in other North African and Middle East regions (Fig. 8). The ostracod biozones established in these regions differ, not only in faunal composition, but also in their stratigraphical ranges and could not be followed in the studied section with a satisfactory exactness. More differentiation was needed owing to distinct palaeo-environmental variability of at least local significance for the time interval preceding OAE2. Owing to the different resolution, creation of an additional formal zonation in the present study has been seen inconvenient, and only an informal biozonation with an attempt to correlate with the previously established ostracod biozones, has been made (Fig. 8).

1-Ostracod biozone I This biozone occurs in the lower part of the section incorporating the lower part of the Fuheis Formation, which is assigned to the Middle Cenomanian. It is an assemblage zone characterized by common to abundant occurrence of Bairdia youssefi, Bythocypris sp.1 Rosenfeld & Raab, Parakrithe andreui, Cythereis namousensis, Cytherella aegyptiensis, C. cf. gambiensis, Dolocytheridea atlasica, Metacytheropteron berbericum, Paracypris dubertreti, Peloriops aegyptiaca, Veeniacythereis maghrebensis and V. streblolophata schista. Of these taxa Bairdia youssefi, Bythocypris sp. 1 Rosenfeld & Raab, Cytherella cf. gambiensis, Parakrithe andreui, and Peloriops aegyptiaca are restricted to this zone. The taxa making up this zone are predominantly shelf taxa, reflecting normal shallow marine environmental settings.

2-Ostracod biozone II This ostracod biozone is separated from the underlying biozone by the Karak and Hummar Limestone Members that are unsuitable for ostracod sampling.

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

Fig. 8. Correlation of Cenomanian– Early Turonian ostracod zonal schemes in North Africa and Middle East regions. 203

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This section part yields only a single specimen of Veeniacythereis maghrebensis at the base, and an ostracod fauna in common with the underlying zone 1 represented by Cytherella aegyptiensis, Metacytheropteron berbericum, Dolocytheridea atlasica, Veeniacythereis maghrebensis and Veeniacythereis streblolophata schista. Zone II is located in the lower part of the Hummar Formation and is distinguished from Zone I by the first occurrence of Neocyprideis vandenboldi in the section. This species dominates the assemblage throughout most of Zone II. Cytherella aegyptiensis, Dolocytheridea atlasica, Metacytheropteron berbericum, Perissocytheridea istriana, Cythereis namousensis and Veeniacythereis maghrebensis are also found. The dominance of Neocyprideis vandenboldi and a scarce presence of Perissocytheridea istriana together with marine faunal elements indicate a shelf lagoonal environmental setting with freshwater influence (Bartov et al. 1980; Rosenfeld et al. 1988; Flexer et al. 1989; Bauer et al. 2003).

3-Ostracod biozone III This biozone is distinguished by a significantly abundant occurrence of Dolocytheridea atlasica, Metacytheropteron berbericum, Cythereis namousensis, Veeniacythereis maghrebensis and Cytherella aegyptiensis and represents the peak interval for these taxa. It lies in the middle part of the Hummar Formation. Representatives of Cytherella dhalalensis, Paracypris dubertreti, Schuleridea houneensis and Monoceratina? trituberculata also co-occur. The assemblage making up this zone indicates a shallow marine environmental setting. In the middle of this zone, two samples (GM3-111 and GM3-112) yield abundant Neocyprideis vandenboldi and Perissocytheridea istriana, thus indicating a significant fresh-water influx in this level.

4-Ostracod biozone IV This biozone is characterized by an ostracod assemblage with rare Dolocytheridea atlasica, Metacytheropteron berbericum, and Veeniacythereis maghrebensis that were significantly abundant in the underlying zone. Cytherella aegyptiensis, C. dhalalensis, Paracypris dubertreti, Cythereis namousensis and Veeniacythereis streblolophata schista also co-occur. This biozone is located in the upper part of the Hummar Formation, and is immediately followed by the OAE2 level. It is probable that the scarcity of the marine ostracod fauna in this zone was generally a response to the Late Cenomanian global regression and decreasing oxygen levels related to this event (e.g. Babinot & Crumiere-Airaud 1990). The interval in the middle of the zone is a marly limestone (samples

GM3-117 to 120) which yields Neocyprideis vandenboldi and Perissocytheridea istreana in the assemblage, indicating fresh-water influence in this level.

5-Ostracod biozone V This biozone occurs in the lower part of the Shueib Formation, which is assigned as Late Cenomanian. It is an acme zone characterized by a highly abundant occurrence of Reticulocosta kenaanensis. This taxon, which dominates the assemblage throughout biozone V, occurs together with rare Cytherella sp., Bairdia sp., Paracypris mdaourensis and Brachycythere gr. sapucariensis. The species found in this zone indicate a shallow marine setting. The base of ostracod biozone V corresponds to a latest Cenomanian major transgression at the onset of OAE2 (Wendler et al. 2010), and this zone comprises the lower part of OAE2 in the studied section. At the peak level of the OAE2 interval, just below the Cenomanian– Turonian boundary, and higher up in the Early Turonian part of the section, no ostracod fauna has been found. The Cythereis rawashensis kenaanensis zone of Rosenfeld & Raab (1974) probably represents the continuation of the present R. kenaanensis zone into the Early Turonian. This assumption cannot be proved here owing to the lack of ostracods in the Lower Turonian of the studied section. However, findings of R. kenaanensis in other Central Jordan sections in the Lower Turonian have been reported by Schulze et al. (2003), together with ostracod species belonging to the Cythereis rawashensis kenaanensis zone (UC-6) of Rosenfeld & Raab (1974).

Discussion In a previous study carried out by Schulze et al. (2004) in west-central Jordan, ostracod faunas were recorded in the Lower– Middle Cenomanian interval represented by the Naur Formation, and higher up in the Fuheis Formation. Of the taxa they recorded in this interval, Cytherella aegyptiensis, C. cf. gambiensis, Dolocytheridea atlasica, Metacytheropteron berbericum, Parakrithe andreui, Cythereis namousenis, Veeniacythereis maghrebensis and V. streblolophata schista are in common with the fauna recorded in the present study in the Middle Cenomanian, which is assigned to the ostracod biozone I. A similar assemblage was also recorded from this interval in Jordan by Babinot & Basha (1985) and Powell (1989). Schulze et al. (2004) also recorded ostracods in the uppermost Cenomanian –Lower Turonian interval represented by the upper part of the Shueib Formation at the OAE 2 level. Their Turonian assemblage (ostracod

CENOMANIAN OSTRACODS OF THE LEVANT PLATFORM

assemblage 2 ‘OA2’) is devided from the lower assemblage (OA1) by a lag in ostracod presence similar to our study. The association they found has Reticulocosta kenaanensis in common with the present record, therefore, belonging to the ostracod biozone V. However, if it is assumed that the gap in ostracod occurrence in Schulze et al. (2004) and the one shown here are time-equivalent, then the assemblage ‘OA2’ of these authors would represent the continuation of the R. Kenaanensis zone in the late Early Turonian. In the interval between ostracod biozones I and V, they only recorded Cytherella sp., which already begins in the Lower–Middle Cenomanian. Comparison with ostracod faunas from other North African and Middle East regions reveal similarities in faunal compositions. However, variations are observed in the stratigraphical ranges of the species, most likely owing to environmental variability. These variations resulted in establishing several local zonal schemes in these regions to which bio-correlation of the present record is attempted in Figure 8.

Palaeoecology The studied ostracod assemblages of the Middle Cenomanian and lowermost part of the Upper Cenomanian are dominated by open marine shelf taxa. Marine associations are mainly comprised of different species of Cytherella, Bairdia, Paracypris, Metacytheropteron, (?)Monoceratina, Cythereis, Peloriops and Veeniacythereis, which are widely distributed in the marine Cenomanian sediments along the Southern Tethys. In contrast, Neocyprideis vandenboldi is quite common in the Hummar Formation at 49 m and 51 –52 m, accompanied in sample GM3-118 by Perissocytheridea istriana. Both species have been interpreted as reflecting fresh-water influence (Bartov et al. 1980; Rosenfeld et al. 1988; Flexer et al. 1989; Bauer et al. 2003). Recent species of the genus Perissocytheridea have been associated with sandy mud substrates in a euryhaline environment, favouring the mesohaline zone with large and quick salinity changes (Keyser 1977). The benthic foraminifera assemblage found in samples that yield Neocyprideis vandenboldi are dominated by the genera Haplophragmoides, Ammobaculites, Cyclammina and Trochammina. Haplophragmoides spp. in particular indicate variable salinity and are interpreted as typical inhabitants of brackish marsh environments (e.g. Murray 1991; Fatela et al. 2009). These species are recorded in co-occurrence with marine species, indicating that coastal marine environmental settings, with rapidly varying conditions including repeated periods of significant fresh-water influx, prevailed during the time of this interval that spans a period of about 300 ka prior to OAE2

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(Fig. 2). Alternatively, a temporal restriction of marine-water influx owing to strong sea-level fall could have led to brackish conditions, under a humid climate. The co-occurrence of brackish and fully marine species is suggestive of time-averaged sampling underpinning the rapidity of salinity fluctuations during this phase. The strong negative d13C excursions in this interval (section metre 51 –54) may support both interpretations since, apart from possible late diagenetic effects, it could be interpreted as representing; a) isotopically light CO2, most probably derived from soil formation during sea-level fall (Allan & Matthews 1982), or b) isotopically light marine carbonates precipitated under fresh-water influence (Elrick et al. 2009). A very similar signal has been recently reported by the latter authors from Mexico so it seems that such features in the d13C record of Cenomanian – Turonian shallow-water sections below OAE2 has a global extend. The uppermost ostracod-yielding levels in the section contain a marine fauna dominated by abundant Reticulocosta kenaanensis. Correlation of these levels, with calcareous nannoplankton and ammonite zonations as well as the stable carbon isotope record (Fig. 2, Wendler et al. 2010), dates them as latest Cenomanian. Reticulocosta avoids strong food pulses and is probably better than other genera adapted to longer but still temporary periods of oxygen depletion-dominating during the transition to permanent oxygen deficiency (Gebhardt & Zorn 2008). Therefore, the occurrence of Reticulocosta kenaanensis in high abundance within a low diversity assemblage could represent a response to environmental stress, which resulted from the early phase of oxygen depletion associated with OAE2. The other Late Cenomanian taxa were seemingly unable to tolerate these conditions, therefore became scarce or even extinct at this level. On the other hand, the ability of R. kennanensis to tolerate oxygen deficiency and excessive nutrients, together with a lack of effective competition from other taxa, allowed for the increased abundance of this species compared with other co-occurring taxa. Many of the south Tethyan marine ostracod species and whole genera of the Cenomanian became extinct below the C– T boundary (see, Bassoullet & Damotte 1969; Rosenfeld & Raab 1974; Babinot & Colin 1988; Damotte 1985, 1995; Ismail 1999; Bassiouni 2002). Owing to the magnitude and global extent of the OAE2, it is possible that only the taxa, which were capable of surviving bottom water oxygen deficiency, could persist into the Early Turonian. Reticulocosta kanaanensis, being such a potentially opportunistic species, was reported in North Africa and Middle East areas as an Early Turonian newcomer indicating the base of the Turonian (Rosenfeld & Raab 1974; Bassiouni

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2002; Cythereis rawashensis kenaanensis Assemblage zone of Rosenfeld & Raab 1974). The appearance of this species in the present material is well below the Cenomanian –Turonian boundary (by 0.5 Ma, Fig. 2). This suggests that Reticulocosta kanaanensis probably represents a taxon, which came into existence along with an event of environmental stress caused by anoxic conditions related to the onset of OAE2 in the Late Cenomanian. Therefore, its use as a marker for the stratigraphic boundary is problematic. Absence of any ostracods higher up in the section merits further discussion. Contrary to this lack of ostracods, the presense of well-preserved calcareous remains of benthic and planktic foraminifera, echinoderms, coccoliths, and calcareous dinoflagellates indicates that diagenesis has not induced a solution of carbonate in this interval. Hence, unless the unlikely case of selective ostracod dissolution, the disappearance of ostracods can be interpreted to reflect palaeoenvironmental conditions adverse to ostracods. A biomarker investigation of the studied section parallel to the present study (Sepu´lveda et al. 2009) showed that water coloumn stratification increased to extreme levels throughout OAE2 and after. As a result of this density stratification, bottom waters were highly saline and depleted in oxygen, which could explain the observed absence of ostracods.

Palaeobiogeography During the Cenomanian, the Levant platform area of today’s Jordan was part of the South Tethyan bioprovince. The largest assembly of palaeobiogeographical aspects dealing with the Cenomanian ostracod faunas from this province and a comparison with the Northern Tethys bioprovince was carried out by Babinot & Colin (1988), Andreu (1993), Gebhardt (1999), Gre´koff (1969) and Luger (2003). Significant contributions have also been added by Babinot et al. (1978), Damotte (1985, 1995), Viviere (1985), Athersuch (1988), Andreu (1993), Majoran (1988, 1989) and Morsi & Bauer (2001). Very similar ostracod assemblages, characterized by representatives of Veeniacythereis, Peloriops and Reticulicosta, dominate this province, which extended from the Atlantic coast of Morocco in the west to the Arabian Gulf region in the east. Many of the elements recorded in Jordan are also known from the different North African and Middle East countries (Fig. 4). Cytherella aegyptiensis, Paracypris dubertreti, Paracypris mdaourensis, Monoceratina? trituberculata, Dolocytheridea atlasica, Metacytheropteron berbericum, Cythereis namousensis, Veeniacythereis maghrebensis and V. streblolophata schista show the widest occurrence in North Africa, as well as in the Middle East. C. dhalalensis, C. cf. gambiensis, Bairdia

youssefi, Bythocypris sp. 1 Rosenfeld & Raab, Schuleridea houneensis, Peloriops aegyptiaca and Reticulicosta kenaanensis are also found in Egypt, Israel and Lebanon. The biogeographical homogeneity of the ostracod associations in North Africa and the Middle East reflects communication along the whole expanse of the Southern Tethys margin during the Cenomanian and suggests, as also indicated by Babinot & Colin (1988), Andreu (1993) and Morsi & Bauer (2001), similar living conditions and the absence of important geographical barriers that could hinder marine faunal exchange. North– south exchange with the northern rim of the Tethys is not clearly pronounced. Only a few species represented by Perissocytheridea istriana and Metacytheropteron berbericum have occurrences in common with the northern Tethys province. Perissocytheridea istriana was previously recorded in Portugal (Babinot et al. 1978) and Metacytheropteron berbericum from Yugoslavia (Babinot 1988). Ostracods are mostly benthic organisms lacking pelagic larvae and having poor dispersal properties. Babinot & Colin (1988) put forward the interpretation that a too-deep central Tethys ocean, and also unfavourable currents, could be probable causes of reduced north– south migrations between the northern and southern Tethys rims. Monoceratina? trituberculata and Brachycythere gr. sapucariensis occur together within basins of West Africa, for example, Gabon (Grosdidier 1979). Probable affinities between the southern Tethys bioprovince and the coastal basins of West Africa during the Cenomanian were indicated by many authors (Grosdidier 1979; Viviere 1985; Athersuch 1988; Babinot & Colin 1988). Brachycythere sapucariensis, found as a single specimen in the studied material, was originally described from Brazil (Kro¨mmelbein 1964) and represents one of the taxa reflecting affinities between South America and West Africa during the Cenomanian.

Conclusions Twenty two ostracod species were found in the investigated section. The assemblages can be grouped into five informal ostracod biozones that reflect considerable palaeoecological variations. Most significantly, brackish water species are recorded in co-occurrence with marine species, indicating that coastal marine environmental settings with rapidly fluctuating salinity conditions, including repeated periods of significant fresh-water influx, prevailed during a period of a few 100 ka prior to OAE2. A flourishing of the opportunist Reticulocosta kanaanensis during the major transgression at the base of OAE2 is replaced soon after by the absence of any ostracod throughout the remaining OAE2 and Early Turonian period of time. The

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investigated mid- to Late Cenomanian ostracod association has much in common with ostracod associations in North Africa and other areas of the Middle East reflecting communication along the whole expanse of the southern Tethys margin during the Cenomanian. The authors express their gratitude to H.-J. Kuss, University of Bremen, for providing field assistance and lab facilities. A. Masri and the Jordan Geological Survey (NRA) are thanked for providing ideal field campaign logistics. We thank S. Papenmeier for careful micropalaeo-ontological preparation and sampling. Thanks to H. Mai, University of Bremen, for assistance at the SEM. M. Segel is acknowledged for stable isotope measurement. The manuscript benefitted from the thoughtful reviews by I. Slipper and B. Andreu. J. Wendler acknowledges funding by the Deutsche Forschungsgemeinschaft (grant KU 642/B20-1).

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Mesoscale folds and faults along a flank of a Syrian Arc monocline, discordant to the monocline trend N. JOSEPH-HAI1*, Y. EYAL1 & R. WEINBERGER2 1

Department of Geology and Environmental Sciences, Ben Gurion University, Beer Sheva 84105, Israel

2

Geological Survey of Israel, 30 Malkhe Yisrael Street, Jerusalem 95501, Israel *Corresponding author (e-mail: [email protected]) Abstract: Orientations of folds and small faults were measured in Turonian and Senonian rocks along the western limb of the Ramallah monocline in Israel, one of the structures comprising the Syrian Arc fold belt (SAFB). The minority of the folds, aligned NNE–SSW, are compatible with the WNW– ESE shortening trend of the SAFB, whereas the majority of them, aligned ENE–WSW, are not compatible with this shortening trend. Kinematic analysis of faults’ attitude indicates NNW– SSE shortening and ENE–WSW extension in accordance with the shortening of the majority of folds. Based on the folds trends, scale, and geometry, as well as the associated fault kinematics, we conclude that the folding mechanism is tectonic shortening and not intraformational folding due to landsliding or collapse owing to karst activity as previously postulated. We propose that a minority of the folds, compatible with the major trend of the Ramallah monocline, are parasitic small folds within the SAFB. The majority of the folds, which are not compatible with the SAFB, were formed owing to NNW–SSE shortening that has been associated with Miocene to Recent movement along the Dead Sea Transform.

The Syrian Arc fold belt (SAFB) (Krenkel 1924) consists of three segments, which are characterized by the distinct orientation of their axes (Fig. 1a). The fold axes in the southwestern segment in Sinai (Egypt) and northern Negev (Israel) are aligned NE–SW; the folds in the central segment, comprising the Judea-Samaria and Galilee mountains in Israel and southern Lebanon, are aligned NNE–SSW to almost north –south, and the northeastern segment, the Palmyrides (Syria), are aligned NE –SW. It is widely accepted that these structures formed owing to displacement along reverse faults (e.g. Mimran 1976), that resulted from reactivation of older, Jurassic, normal faults (Freund et al. 1975; Bruner 1991; Walley 1998; and references therein). The central segment of the Syrian Arc comprises the NNE– SSW trending monoclines of Hebron, Ramallah, and Fariah, whose axes are arranged in en-echelon architecture (Shahar 1994). These monoclines are the most prominent structures in the central and northern parts of Israel and form the mountain backbone west of the Jordan River (Fig. 1b). The early folding of the Syrian Arc owing to WNW– ESE shortening started during the Turonian series and terminated at the Miocene series (Letouzey & Tremolieres 1980; Eyal & Reches 1983) or extended from the Turonian series to the Present (Eyal 1996). A later regional shortening, trending NNW –SSE and related to the Dead Sea Transform (DST) is active

from the Miocene series to the Present (Letouzey & Tremolieres 1980; Eyal & Reches 1983; Eyal 1996; Eyal et al. 2001). New road cuts in the Turonian Bina Formation expose many mesoscale folds, faults, and joints in the western margins of the Judea Mountains during the last decade owing to the construction of Highway 6 (Cross Israel Highway). Previous studies in the surrounding areas already described folds, whose axes are oriented perpendicular to the main NNE–SSW trending of the monocline axes (e.g. Bentor & Vroman 1954; Shomroni 1970; Dimant 1971; Livnat 1971; Ilani 1972; Hildebrand 1975). However, to date, there is no coherent explanation to their origin and time relation with the large-scale monoclines. Several options have been suggested to explain the origin of these folds. (a) They are collapse structures that formed owing to karst activity. This karstic activity formed large sub-surface cavities in the carbonate rocks of Bina formation into which the overlying rocks were tilted and folded by collapse. (b) The folds are intraformational structures of the Turonian age that initiated by landslides in the unstable shelf sediments. (c) The folds have a tectonic origin owing to post-Turonian shortening. Each of these fold mechanisms is known in the area and displays a distinctive pattern (Fig. 2). The goals of the present study are to quantify the orientations of these folds (and associated faults) and

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 211–226. DOI: 10.1144/SP341.10 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Generalized tectonic map of Israel showing the SAFB. (b) Location of the study area along Highway 6.

MESOSCALE FOLDS ALONG A SYRIAN ARC MONOCLINE

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Fig. 2. Schematic representation of deformed and sedimentary structures in the Bina Formation. The thickness of the representative stack of layers is about 10 m.

define their deviation from the major structures of the SAFB, and to ascribe an appropriate mechanism for their formation. We also discuss the formation age of these structures and their relation to known tectonic phases. In addition, karst development related to folding will be described.

The study area, along Highway 6 between Ben Shemen interchange to the south and Nakhshonim interchange to the north, comprises gentle dipping strata (38 –108) to the west, which crop out west of the flexural zone of the Ramallah monocline. The stratigraphic section exposed in this area is mainly composed of the Cenomanian –Turonian carbonate rocks of the Judea Group, uncomfortably overlain by the Senonian chalks and cherts of the Menuha and Mishash formations (Mount Scopus Group) or, in some locations, by the Miocene carbonatic – clastic Saqiye Group (Livnat 1971; Hildebrand 1975). Most outcrops in the study area are comprised of the Turonian Bina Formation (Fig. 3), which is divided into three members (Livnat 1971). (a) Dolomitic-limestone member, comprising of pale grey, dense, fine-grained dolomite with yellow patina, forming well-bedded 1 m thick ledges, interbedded with thin marl layers. (b) Bioclastic limestone member, comprising of massive, well-bedded biosparitic limestone. (c) Sublithographic limestone member, comprising of pale grey-yellowish limestone, and contains layers of irregular chert nodules and concretions. Folds within the Bina Formation are known in this area and are the main focus of the present study. In some locations rocks of the Menuha Formation fill cavities and funnels within the Bina formation, indicating that the area was exposed to surface erosion at the end of the Turonian series (Livnat 1971).

in the overlying Mount Scopus Group were measured along road number 444 (Fig. 1b), closely located to Highway 6. Dips of the folded beds, on both limbs of the fold, were measured, and the fold axes were calculated based on a cylindrical best-fit of poles to bedding (i.e. p-diagram). In a few cases, the fold axes were measured directly. Fault attitude and orientations of kinematic indicators (e.g. displaced layers, striae, slickolites and smallscale pull-aparts) were measured at the same outcrops, to obtain the instantaneous strain axes under which the faults were formed and to compare them with the finite strain axes derived from the folds. One to three folds and up to ten faults were measured at some of the outcrops (Tables 1 & 2), depending upon the quality of exposure. A total of 27 folds and 47 faults were measured, with each fold containing any number from 5 to 27 fold beds. We processed the folds data with Stereonet for Windows v. 1.2 program (Allmendinger 2003). A fold is regarded as cylindrical when 90% of the p-poles fall within an angle of 108 from the constructed p-circle, and sub-cylindrical when 90% fall within an angle of 208 from the p-circle (Ramsay & Huber 1987). The distribution of the fold axes is displayed by rose diagrams and the division of the fold population into different groups is based on common, but different, directions of fold axes. A fault’s sense of displacement was determined by the association of displaced layers and striations or slickolites. Fault data was processed with FaultKin for Windows v. 1.2 program (Allmendinger 2001). In this program the P & T are the instantaneous strain axes for a fault, where P, the shortening strain axis, and T, the extensional strain axis, are at 458 to fault plane. The mean shortening and extensional strains were calculated as the Bingham distribution, which is equivalent to an unweighted moment tensor summation (Allmendinger 2006). Maps were produced with Arcmap 8.2 GIS program (2006).

Method

Results

We measured the orientation of folds and faults at 17 outcrops along Highway 6 and one outcrop in Shilat Junction (Israel Coord. 151820/146440), all of them in the Bina Formation. Two additional folds

Folds

Geological setting

The folds are open (i.e. interlimb angle c. 1408) upright, sub-horizontal (Ramsay & Huber 1987;

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Fig. 3. Representative photos of the open folds along Highway 6 (stations from the top): 2A, 2B, 17. Cars provide a scale.

Twiss & Moores 1992), with wavelengths ranging between 25 –100 m and amplitude between 10– 30 m (Fig. 3). The studied localities along the central part of Highway 6 (Table 1) and the associated p-diagrams and fold axes are presented in Figure 4. The trends of the fold axes and geometry (anticline or syncline) are indicated according to the precise location within the outcrop (Fig. 5). Several observations are apparent from examination of the maps and stereograms (Fig. 6). (a) All

the folds are cylindrical or sub-cylindrical according to Ramsay & Huber (1987) criteria. (b) The trend of the major group of folds (15 folds), Group A, is 28/2548, a95 ¼ 9.38. The mean trend of all the folds is similar to that of the major group, implying that the NNW –SSE shortening is the predominant shortening direction. The trend of the minor group of folds (6 folds), Group B, is 18/2078, a95 ¼ 7.78, implying WNW –ESE shortening. (c) Trends of Group A and Group B vary among the stations

MESOSCALE FOLDS ALONG A SYRIAN ARC MONOCLINE

Table 1. Summary of fold orientation from Highway 6 by outcrop Fold axis orientation Outcrop 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18* 19† 20†

Location

Trend [8]

Plunge [8]

145990/159880 145950/159915 146310/159510 146365/159430 146370/153300 146435/153390 146580/154300 147035/157670 147050/158235 146530/159345 146475/159445 146040/159950 146110/159915 146100/159880 144525/162360 145340/160765 146935/157825 146890/156990 146865/156550 146900/156525 146925/156535 146515/154380 145100/152310 146370/153435 151820/146440 145116/153489 144749/152211

033 027 031 053 325 202 257 249 271 204 270 240 236 245 254 256 077 068 089 122 104 063 323 313 204 280 073

3 16 2 1 3 14 9 9 3 7 12 3 11 4 5 3 0 11 2 4 10 5 6 8 4 3 10

Fold axis orientations were determined by the p-diagram method. Location is given in Israeli Cassini Soldner (I.C.S) grid (old grid). *Outcrop 18 is located near the Shilat junction. † Outcrops 19 – 20 are within the Mount Scopus Group and were measured along road number 444.

without consistent spatial order (e.g. compare fold axes in the neighbouring stations 7 and 8). The four uncommon NW-trending folds are all closely located at the most southern part of the study area (Stations 3, 14B, 16–17) and maybe related to local disturbance at that site. Outcrops of the Menuha Formation, composed of massive chalks, are exposed along Highway 6. However, the absence of any prominent bedding or stratigraphic markers prevent dip measurements of these rocks. On the other hand, dip measurement of 40 cm thick folded layer of the Mishash chert exposed in two localities along road 444 (Israel Coord. 145116/153489, 144749/152211) reveals that the fold axes are similar to that of Group A. In one location, near Re’ut (Israel Coord. 152560/143505), we measured a fold in Pliocene rocks whose 2648 trending axis is similar to that of Group A (Fig. 7).

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Faults A total of 47 normal and strike–slip faults were measured in the study area (Table 2), their exposed length varies between 0.5 to 10 m. The entire population of the faults can be divided into three main groups based on their P & T axes kinematic analysis (Fig. 8). (a) 28 faults, mainly normal, with a mean fault plane of 618/2558 were found at six stations. Their linked Bingham instantaneous strain axes are: 1. 168/2538 (extensional strain axis); 2. 28/3448; 3. 738/0828 (shortening strain axis). (b) 9 faults, mainly left-lateral, with a mean fault plane of 878/2118 were found at three stations. Their linked Bingham instantaneous strain axes are: 1. 48/1678 (extensional strain axis); 2. 818/2848; 3. 88/0768 (shortening strain axis). (c) 10 faults, mainly right-lateral, with a mean fault plane of 88/0218 were found at three stations. Their linked Bingham instantaneous strain axes are: 1. 48/2468 (extensional strain axis); 2. 868/ 0958; 3. 28/3368 (shortening strain axis). Several observations are apparent from inspection of Figure 8 and Table 2. (a) The predominant extension, based on Groups A and C, is oriented ENE –WSW; a second direction, based on Group B is NNW–SSE. (b) Fault orientations vary among the stations; for example, in a few stations (e.g. stations 3 and 18) faults belonging to only one group were found, whereas in other stations (e.g. stations 2, 6 and 14) faults belonging to two groups were found. Field observations show that some of the normal fault planes are sub-parallel, restricted to single mechanical layers, and associated with parallel joint planes without any evidence for significant displacement along them. This phenomenon suggests that these faults are actually reactivated joints or faulted joints (Wilkins et al. 2001).

Karst development related to fold hinge zone Field observations indicate that in many of the examined folds the hinge zone is more deformed and characterized by various stages of karst development. In some folds, only a few joints and faults can be found (Fig. 9b), whereas in others intensively fractured hinge zone with vertical, closely spaced joints with rough walls are found (Fig. 9c). Recent karst activity in some folds is expressed by opened joints filled with loose fragments of the surrounding Cretaceous limestones mixed with Terra Rosa (Fig. 9d ). In several outcrops, fold hinge zones are associated with small karstic caves, which are commonly filled in by impure calcite, sometimes in the stalactite form (Fig. 9e). Notably, palaeokarst of the Late Turonian

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Table 2. Summary of faults orientation from Highway 6 by outcrop Fault orientations Outcrop 2

3

5 6

7 13 14

15 18†

Strike [8]

Dip [8]

Stria (plunge/trend) [8]

Sense

336 295 297 292 285 110 305 150 282 305 298 160 150 153 162 155 160 190 068 167 0 168 178 145 160 160 191 005 178 095 155 107 107 0 205 346 190 354 230 095 170 164 170 145 165 350 170

73 89 80 89 83 85 88 75 78 83 85 65 69 70 67 66 70 65 75 74 55 76 76 79 85 63 71 70 67 70 75 90 90 70 90 60 85 88 80 85 70 62 60 63 65 65 60

63/096 9/295 4/297 3/293 2/285 5/285 28/310 15/150 12/283 10/125 15/115 65/250 69/240 70/243 67/252 66/245 70/250 65/280 75/160 70/218 55/090 76/260 76/270 79/235 77/287 55/224 71/281 70/095 67/268 10/265 25/157 7/280 68/120 70/080 90/295 60/076 2/280 10/165 2/240 4/255 70/260 62/254 60/260 63/235 65/255 65/080 55/245

Normal* Left Left Left* Right Right Right Right Right Right Right* Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal* Normal Normal Normal Normal Normal Normal Normal* Left Left Right Normal Normal* Normal Normal Left Left Right Right* Normal Normal Normal Normal Normal Normal Normal*

Fault attitudes appear as measured in the field. *In these faults the sense was assumed according to the fault direction. † Outcrop 18 is located near the Shilat junction.

age were also found (Fig. 10) and is characterized by cavities within the Bina Formation that are filled by the chalks of the Menuha Formation (Mount Scopus Group).

Discussion Previous studies in and around the study area described folds, whose trend are not aligned parallel

MESOSCALE FOLDS ALONG A SYRIAN ARC MONOCLINE

Fig. 4. p-diagrams plotted according to their geographic location along Highway 6. The pole to the great circle, represented by fill square, is the fold axis.

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Fig. 5. (a) Location of folds along the northern part of the study area along Highway 6. The Bina Formation (Kub), Senonian Menuha Formation (Kum) and Neogene to recent conglomerates (NQa, Al) are exposed next to the highway [after Hildebrand-Mittlefehldt (1993) and Yechieli (2007)]. Each fold axis is represented with an arrow, whose head shows the fold plunge. The three different kinds of arrows relates to the different groups of folds. The number at the base of the arrow is the outcrop number (Table 1), whereas the number at the tip of the arrow is the plunge of the fold axis.

MESOSCALE FOLDS ALONG A SYRIAN ARC MONOCLINE

Fig. 5. (Continued) (b) Location of folds along the southern part of Highway 6. See also caption of Figure 5a.

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Fig. 6. Trend of fold axes in the study area projected on lower hemisphere equal area stereographic nets and rose diagrams. Each point on the projection represents a fold axis whose orientation was determined by the p-diagram method (Table 1). The size of the rose diagram sector is 108 and the petals indicate the relative frequency of the points. Upper stereogram-All fold axes, Group A – ENE–WSW trending folds and Group B – NNE–SSW trending folds.

Fig. 7. (a) Fold in Pliocene rocks near Re’ut. (b) Equal area Lower hemisphere p-diagram of the fold.

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Fig. 8. Stereographic projections, equal area lower hemisphere, of the faults in the study area. The arrows on the great circle (i.e. fault plane projection) show the movement direction of the hanging wall. For each group the scatter P & T axes of all fault planes is projected, filled circles represent P axes, hollow squares represent T axes. Large filled squares are the linked Bingham axes, P indicates shortening and T extension. FP represents the fault plane and AP for auxiliary plane. Arrows along the FP indicate the sense of movement. (a) All faults; (b) normal faults; (c) left-lateral faults; and (d) right-lateral faults.

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Fig. 9. Schematic representation of karst development according to different stages of hinge zone evolution in the study area.

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Fig. 10. Palaeokarst of Late Turonian age along Highway 6. The Menuha Formation is filling cavities within the Bina Formation (right photo) and shows irregular contact with the Bina Formation (left photo).

with that of the NNE –SSW orientation of the major Syrian Arc structures in central Israel in general and the Ramallah monocline in particular. Bentor & Vroman (1954) were the first to indicate that the Tulkarm, ’Azzun, Naballa, Lydda and Barfilya noses are small secondary folds perpendicular to the major Ramallah monocline. Later, Dimant (1971), Ilani (1972) and Hildebrand (1975) mentioned undulations, whose approximate width is 1.5–2.5 km and amplitude 20 –80 m, that are perpendicular to the Syrian Arc structures. Additionally, these studies mention intraformational folding within the Bina formation, whose wavelength varies from tens to hundreds of meters (Shomroni 1970; Dimant 1971; Livnat 1971; Ilani 1972; Hildebrand 1975). Based on the scale, it seems that the folds described in this study correspond to these latter folds. However, in this study we could quantify, owing to the new outcrops exposed by the construction of Highway 6, the trend of folds and faults by measuring their attitudes, whereas in the previous works only Ilani (1972) provides a general direction of these folds. The folded section of Turonian rocks consists of multiple layers with different mechanical properties and thickness (commonly 10 –40 cm). It is well known that the dominant wavelength of a structure is a linear function of the layer thickness, and that the competent (stiff ) layers rather than incompetent (soft) layers control the shape and wavelength of the folds (e.g. Ramsay & Huber 1987). It is accepted that the wavelength of the large-scale Syrian arc monoclines is controlled by the displacement along large buried reverse faults at depth (e.g. Mimran 1976). The studied folds, however, are parasitic features on the western limb of the largescale Ramallah monocline. Their relatively short wavelength and different trend may indicate that their formation is not controlled by buried faults, but by a different mechanism. The data presented in this study indicate that a few folds (Group B)

are compatible with the NNE –SSW orientation of the Syrian Arc structures, and may be regarded as parasitic small folds of the Ramallah monocline. However, the majority of folds exposed along the central part of Highway 6 are aligned ENE –WSW (Group A), and therefore not compatible with the major NNE–SSW trend of the SAFB. Hence, the focus of the following discussion will concentrate on these folds. Karstic processes are well known within the Bina Formation at this area (Dimant 1971; Livnat 1971; Ilani 1972; Hildebrand 1975; Frumkin & Fischhendler 2005), and are also described in this study. In a few locations surrounding the study area, the association between synclines and karst activity was observed or suggested (e.g. Buchbinder & Sneh 1984; Frumkin pers. comm. 2007). We acknowledge the fact that the folding of layers might be associated with karsts (Fig. 2), but argue that the folds along Highway 6 are associated with buckling and not with other mechanisms owing to the following reasons. (a) Deformation by collapse of layers into karstic cavities should result in random trends of the fold axes. The orientations of fold axes measured in the study area are not random and are divided into two coherent subpopulations without an overlap of their a-95 confidence circles. On the other hand, the preferred orientation of these fold populations can be attributed to known regional deformation phases. (b) Commonly, collapse of layered blocks into karstic cavities result in disruption of the bedding (Fig. 11). However, in most of the folds in the study area, the limbs are continuous. Intensive fracturing, when it exists, is located only at the very deformed hinge zone. (c) Field examination in the study area (Fig. 10) shows that the effect of karsts on the surrounding beds is minor and very local. Collapse of bedding may result in the formation of small synclines, above or within the karstic cave, however, it cannot explain the formation of the large

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Fig. 11. Disruption of the Turonian Bina Formation owing to the collapse of layered blocks into a karstic cavity, near Modı`in, Isr. Coord. 149946/147102.

anticlines, with wavelengths of about 100 m, as found in the area (Fig. 3). (d) The formation time of the faults that accompany the folds is uncertain. However, the majority of faults were formed under a strain field that is compatible with the ENE– WSW trending of the majority of the folds. We suggest an alternative scenario for karst evolution in the study area in which the karst development follows the formation of folds similar to the work of Jamicic & Novosel (1999). In this scenario, the spacing of joints and faults at the hinge zone is higher than at the fold limbs. Therefore, karst formation, owing to flowing water followed by block collapse, is concentrate at fold hinge zones. Figure 9 displays such a process according to different deformed hinge zones found in the study area. Large undulations in the Bina Formation were explained, based on thickening of Senonian sediments in the synclines and thinning in the anticlines, by post-Turonian–pre-Senonian formation age correlative with the formation of the main Syrian Arc structures (e.g. Dimant 1971; Ilani 1972; Hildebrand 1975). According to these studies, these undulations were formed by a combination of NW– SE shortening and block movement along faults at the basement. However, no direct evidence supporting the existence of faults in the basement, which is essential for this explanation, was given, and in addition, this explanation cannot account for the formation of structures perpendicular to the SAFB at the same time. Bentor & Vroman (1954) suggested that the ‘intense compression caused a tendency of elongation along the main fold axis. An anticline, such as the Central Israel Fold range, built in part by hard limestones and dolomites, showed resistance to this tendency of elongation and, as the only direction of relief is upward, the

elongating anticline will start to crinkle into small secondary folds perpendicular to the main axis’. Nevertheless, in such a mechanism, elongation parallel to fold axis is necessary and this is not usually observed in most of the folds. Regarding the small-scale folding within the Bina formation, Dimant (1971) and Ilani (1972) suggested that these folds were evolved by syn-sedimentary landslides within the Bina formation, and regarded them as intraformational folding. According to this explanation, local folds were created when landslides of soft beds upon more lithified ones occurred owing to tectonic movements that started during the Upper Cenomanian series and continued to the Turonian series. Landslides within the Bina formation were found in the study area (e.g. near Shilat Junction, Israel Coord. 151696/147182) and are characterized by landslide scars and change of bedding dip within a small distance (Fig. 2). The folds measured in the study do not contain scars and the change of the dip is gradual. Furthermore, landslides upon the limb of a large structure, like the Ramallah monocline, is expected to produce small parasitic folds whose axes are parallel to the main fold and not perpendicular as recorded in this study. The tectonic movements mentioned in previous studies seem to be related to the early folding phase with a WNW– ESE plate-related Syrian Arc shortening, which started during the Late Turonian series (Eyal & Reches 1983). This shortening direction can account for only a minor group of measured folds (Group B). The ENE –WSW trending folds comprising the major folds group (Group A) indicate a shortening direction of NNW–SSE, which is perpendicular to the expected trend. Analysis of the strike –slip faults (Group C) also indicates a

MESOSCALE FOLDS ALONG A SYRIAN ARC MONOCLINE

NNW–SSE shortening direction and ENE –WSW extension similar to that of the main group of normal faults (Group A), and to the main joint system, 340 –3458, found in the area (Livnat 1971; Hildebrand 1975) (Fig. 12). Evidence for both the Syrian Arc Stress (SAS) and Dead Sea Stress (DSS), were found by macro- and mesostructures that indicate regional shortening parallel to SHmax, accompanied by simultaneous regional extension normal to SHmax (Eyal & Reches 1983; Eyal 1996; Eyal et al. 2001). The SAS is associated with WNW– ESE shortening and NNE –SSW extension, and the DSS by NNW –SSE shortening and ENE– WSW extension. We suggest that most structures studied in this work may be related to the Miocene to the Present regional NNW–SSE shortening associated with the DST (Eyal & Reches 1983).

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Although this regional stress/strain field is related to the DST, evidence for structures compatible with the DSS was found as far as 200 km on both sides of the transform (Eyal & Reches 1983; Eyal 1996). Therefore finding structures compatible with the DSS in the study area, about 30 km west of the transform, is plausible. To strengthen the argument that the folding in the study area did not resulted from intraformational folding during the deposition of the Bina Formation, similar structures should be also found in younger formations. This would indicate that the deformation is not restricted to the Turonian Bina Formation and resulted from a younger tectonic phase. Indeed, observations in two localities in the study area revealed similar ENE – WSW trending folds, comprised of Senonian rocks (e.g. Mishash Formation), and in one location

Fig. 12. Lower hemisphere equal area stereographic projection of the main groups of mesostructures measured at the study area. (1) Solid circle represents the mean fold axes (n ¼ 15). (2–3) Empty squares represent the mean extension directions based on normal (n ¼ 28) and right-lateral strike– slip (n ¼ 10) faults. (2–3) Solid squares represent the mean shortening directions based on normal (n ¼ 28) and right-lateral strike–slip (n ¼ 10) faults. (4) Empty triangle represents the mean joint direction (Livnat 1971; Hildebrand 1975). P indicates shortening and T extension.

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(Re’ut) even Pliocene rocks seem to be folded. However, a final prove for this argument, needs an additional study on a larger scale.

Conclusions The most prominent fold trend observed in the carbonatic Bina Formation along the central part of Highway 6 in Israel is ENE– WSW. This trend is not compatible with the NNE–SSW trend of the major Syrian Arc structure of the Ramallah monocline. Field observations and kinematic analysis discard the possibility that these folds are intraformational structures or were formed owing to karstic activity. The well-grouped fold trends, folds appearance and existence of similar folds in formation younger than Bina Formation indicate that the folds were formed owing to tectonic shortening. The NNW– SSE shortening and ENE –WSW extension indicated by the faults found in the study area support the suggestion that these folds should be attributed to a younger stress field, such as that associated with the DST. This research was supported by the Samuel Sebba Chair for Structural Geology. We thank the editor, C. Homberg and the reviewers C. Druckman and P. Vergely for their helpful comments and suggestions. We are grateful to A. Sneh, who introduced us to the remarkable collapse of Bina layers into a karstic cavity near Modı`in. We also thank L. Korman, G. Tirosh, U. Makover, and D. Hadar for their help in the field, and to Y. Asher for the figure editing.

References Allmendinger, R. 2001. FaultKin for Windows v. 1.2 program. Allmendinger, R. 2003. Stereonet for Windows v. 1.2 program. Allmendinger, R. 2006. Structural Geology. World Wide Web Address: http://www.geo.cornell.edu/ geology/classes/RWA/GS_326/GEOL326. Bentor, Y. K. & Vroman, A. 1954. A structural contour map of Israel 1:250 000, with remarks on the dynamic interpretation. Geological Survey of Israel Bulletin, 7. Bruner, I. 1991. Investigation of the subsurface in the northern Negev, Israel using seismic reflection techniques. PhD thesis, Tel-Aviv University, 105 (with English abstract). Buchbinder, B. & Sneh, A. 1984. Marine sandstones and terrestrial conglomerates and mudstones of Neogene– Pleistocene age in the Modi’im area: a re-evaluation. Geological Survey of Israel Current Research, 1983–84, 65– 69. Dimant, E. 1971. The geology of Bet Horon region and the western flanks of the Bet El anticline. Geological Survey of Israel Report, MM/101/71, 79 (in Hebrew).

Eyal, Y. 1996. Stress field fluctuations along the Dead Sea Rift since the middle Miocene. Tectonics, 15, 157– 170. Eyal, Y. & Reches, Z. 1983. Tectonic analysis of the Dead Sea rift region since the Late Cretaceous based on mesostructures. Tectonics, 2, 167 –185. Eyal, Y., Gross, M. R., Engelder, T. & Becker, A. 2001. Joint development during fluctuation of regional stress field in southern Israel. Journal of Structural Geology, 23, 279 –296. Freund, R., Goldberg, M., Weissbrod, T., Druckman, Y. & Derin, B. 1975. The Triassic–Jurassic structure of Israel and its relation to the origin of the Eastern Mediterranean. Geological Survey of Israel Bulletin, 65, 26. Frumkin, A. & Fischhendler, I. 2005. Morphometry and distribution of isolated caves as a guide for phreatic and confined paleohydrological conditions. Geomorphology, 67, 457 –471. Hildebrand, N. 1975. The geological map of Israel, 1:50 000. Sheet 8-I: Kefar Sava, explanatory note. Geological Survey of Israel, 23 (in Hebrew). Hildebrand-Mittlefehldt, N. 1993. The geological map of Israel, 1:50 000. Sheet VII-I: Kefar Sava. Geological Survey of Israel. Ilani, S. 1972. The geology of Tulkarm area. Geological Survey of Israel Report, MM/105/72, 71 (in Hebrew). Jamicic, D. & Novosel, T. 1999. The dynamics of tectonic modeling of some caves in the karst region. Geology Croatia, 52/2, 197–202. Krenkel, E. 1924. Der Syrische Bogen. Zentarbladt fur Mineralogie, Geologie, Palaeontologie, 9, 301– 313, 10, 274– 281. Letouzey, J. & Tremolieres, P. 1980. Paleostress around the Mediterranean since the Mesozoic from microtectonic: comparison with Plate tectonic data. Rock Mechanics, 8, 173 –192. Livnat, A. 1971. The geology of the north western slopes of the Judean hills. Geological Survey of Israel Report, MM/102/71, 60 (in Hebrew). Mimran, Y. 1976. Deep-seated faulting and the structures of the northern Negev (Isreal): an analysis and experimental study. Israel Journal of Earth Sciences, 25, 111–126. Ramsay, J. G. & Huber, M. I. 1987. The Techniques of Modern Structural Geology. V. 2: Folds and Fractures. Academic Press, London. Shahar, J. 1994. The Syrian arc system: an overview. Palaeogeography, Palaeoclimatology, Palaeoecology, 112, 125 –142. Shomroni, A. 1970. Structures in Shephelat Lod. Israel National Oil Company Interim G.O.N. – A.S. 51 Report. Twiss, R. J. & Moores, E. M. 1992. Structural Geology. W. H. Freeman & Company, New York. Walley, C. D. 1998. Some outstanding issues in the geology of Lebanon and their importance in the tectonic evolution of the Levantine region. Tectonophysics, 298, 37–62. Wilkins, S. J., Gross, M. R., Wacker, M., Eyal, Y. & Engelder, T. 2001. Faulted joints: kinematic, displacement-length scaling relations and criteria for their identification. Journal of Structural Geology, 23, 315– 327. Yechieli, Y. 2007. Geological map, 1:50 000, Sheet 8-III: Lod. Geological Survey of Israel.

Middle and Upper Jurassic stratigraphy and sedimentary evolution of Lebanon (Levantine margin): palaeoenvironmental and geodynamic implications PIERRE-YVES COLLIN1*, ANNA MANCINELLI2, MAURIZIO CHIOCCHINI3, MUSTAPHA MROUEH4, WALID HAMDAM4 & FAHTI HIGAZI4 1

UPMC University Paris 06, UMR 7193, ISTEP (Institute of Earth Sciences of Paris), 4 place Jussieu, case 117, F-75005, Paris, France 2

Universita` di Camerino, Dep. Scienze della Terra, via Gentile III da Varano, 62032 Camerino, Italy 3

Via Tazio Nuvolari n.15, 00144 Roma, Italy

4

Agronomy Department, Lebanese University, BP13-5368 Chourane, Beirut 1002, Lebanon *Corresponding author (e-mail: [email protected]) Abstract: The Arabian, African and Eurasian plates interact in the Levantine region. Despite numerous studies of the region, many geological issues relating to Mesozoic times remain unresolved. The Lebanon passive margin is a key area for understanding Neo-Tethyan sedimentary history during this period. The Jurassic succession in Lebanon is well exposed and thick (more than 1000 m). It is more or less complete and relatively undeformed. With a few recent exceptions most studies of the area were made in the 1950s and so the sedimentary evolution of the Jurassic is only partly understood. This study provides (1) a new sedimentary and sequence stratigraphic framework, and (2) a new biostratigraphic framework based on benthic foraminifera and calcareous algae. Palaeoenvironmental and geodynamic conclusions are inferred. Jurassic outcrops occur in both the Mount Lebanon and the Anti-Lebanon areas. Here, they were studied essentially in Mount Lebanon. The Jurassic succession can be divided into three parts: (1) the lower part (Kesrouane Formation) is a thick succession of marine limestones or dolomites; (2) the middle part (Bhannes Formation) consists mainly of basaltic eruptive rocks associated with pyroclastic strata; (3) the upper part (Bikfaya Formation) is a succession of marine limestones. During the Bathonian, Callovian, Oxfordian and parts of the Kimmeridgian, a large epicontinental shelf, with very shallow marine environments, extended across Lebanon (Kesrouane Formation). The period was characterized by a stable platform morphology. It was a tectonically quiet period, although intense subsidence allowed the accumulation of a thick sediment package. During the Kimmeridgian, the carbonate platform regime that had dominated Lebanon during the Middle Jurassic came to an abrupt end, as evidenced by a regional unconformity, a regression and block faulting. This rifting phase is associated with a volcanic event (Bhannes Formation) that is recognized from northern to southern Lebanon. During the Lower Jurassic (Kimmeridgian p.p. to Tithonian p.p.) shallow marine carbonate shelf deposits are observed again (Bikfaya Formation), indicating a marine transgression. This last formation exhibits rapid lateral thickness variations, because of active block faulting and erosion, and is overlain by continental sandstones of the basal Cretaceous.

The Levant is a geographical region where the Arabian, African and Eurasian tectonic plates interact (Fig. 1a). Despite numerous studies of the Mesozoic and Cenozoic geology of the region many issues remain unresolved, especially with regard to sedimentology and stratigraphy. Jurassic sedimentary rocks crop out particularly well on Mount Lebanon and the Anti-Lebanon or Mount Hermon (Fig. 1b). The succession occurring in the Mount Lebanon area belongs to

the Middle and Upper Jurassic. The lowermost rocks cropping out on the eastern slopes of the Anti-Lebanon are Lower Jurassic (Vautrin 1934; Renouard 1951; Dubertret 1955). These oldest units are accessible in Syrian territory only and could not be included in this study. The sedimentary units recognized on the western slopes of Mount Hermon are the lateral equivalents of units observed in Mount Lebanon. The entire Jurassic sedimentary succession is particularly thick,

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 227–244. DOI: 10.1144/SP341.11 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. (a) Map of the eastern Mediterranean showing the location of the study area. Simplified plate tectonic setting for the region (modified from Walley 2001). (b) Simplified geological map of Lebanon (modified from Dubertret 1955; Walley 2001). Location of the main Jurassic sections.

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

reaching up to 2000 m. Here, only the top 500 m, corresponding to parts of the Middle Jurassic and the Upper Jurassic, were accessible and studied in the field. Most studies of this part of the Middle East have concentrated on tectonics and geodynamics (e.g. Beydoun 1988; Walley 1998, 2001; Butler & Spencer 1999; Gomez et al. 2001), with stratigraphic attributes of the formations based largely on much older works by Renouard (1951) or Dubertret (1955). Much less work has been done on the detailed stratigraphy of the Jurassic of Lebanon, and what there is relies mostly on stratigraphic data published by Saint-Marc (1980), Dubertret (1955), and Noujaim Clark & Boudagher-Fadel (2001, 2004). These studies are supplemented by more regional studies (e.g. Kuss 1990). Detailed sedimentological studies of the Jurassic succession of the Middle East are scarce and most were made to assess Lebanon’s petroleum potential (Nader et al. 2003, 2004; Nader & Swennen 2004). Consequently, no study since Dubertret (1955) has sought to synthesize both the stratigraphy and sedimentology of this part of the Levantine margin in an attempt to clarify its palaeogeographical evolution and to address its geodynamic context. Accordingly the updating and development of a new stratigraphic and sedimentary framework for this part of the Arabian plate is an important objective in understanding the history of the Neo-Tethys in Jurassic times.

Geological and regional setting The Mesozoic carbonate platforms of Lebanon were deposited in the Palmyrides intra-plate basin that extended from Egypt to Iraq via Israel, Lebanon, and Syria (Walley 1998). In Lebanon, Middle and Upper Jurassic rocks crop out in the Mount Lebanon and the Anti-Lebanon ranges. Upper Triassic and Lower Jurassic sedimentary rocks have been reported only in the Anti-Lebanon of southwestern Syria (Mouty & Zaninetti 1998). These Dogger and Malm units have been studied in this work mostly in the central and northern part of Mount Lebanon (Fig. 1b). Outcrops further south are intensely tectonized because of the proximity of the Yammouneh Fault and so no uninterrupted successions of any significant thickness occur. Although the Jurassic sedimentary succession exposed in the Anti-Lebanon displays significant differences to that one of the Mount Lebanon (Dubertret 1955, 1975), it has only been partly revised here, since outcrops in Syria were not visited. The Middle and Upper Jurassic sedimentary succession can be divided into three major units (Dubertret 1955, 1975; Walley 2001; pers. obs.) as listed below.

(1)

(2)

(3)

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A unit nearly 1000 m thick (Kesrouane Formation), composed at its base of a thick dolostone unit (Chouane Member), which was only studied in this work as it generally crops out only in the cores of anticlines. It is overlain by a 300 m thick unit (Fig. 2a) consisting of monotonously stacked, massive, decimetres to metres thick fine-grained limestone beds (Nahr Ibrahim Member; Fig. 2b – d), the main subject of this study. A unit up to 100 m thick (Fig. 2e), made up of basalt flows (Fig. 2f) and volcano-sedimentary beds (Fig. 3g) interspersed with palaeosols (Bhannes Formation) and rare marine limestone beds (Fig. 3h). A massive limestone unit up to 30 m thick (Fig. 2e), consisting of various fine- to coarsegrained sediments (Bikfaya Formation).

The Jurassic is capped by a more or less erosional discontinuity surface, occasionally with incised valleys (Fig. 3i). Above the surface, fluviatile sandstones (Gre`s de Base Formation, or Chouf Sandstone Formation of some authors; Figs 2 & 3j) form the base of the Cretaceous succession.

Objects and methods In the course of two fieldtrips in 2003 and 2004 the Jurassic of Lebanon was explored from North to South, in order to (1) make an inventory of potential study sites and check the validity of geological maps, and (2) test whether exposure information (e.g. site access or formation thickness) was still valid. Nine reference sections were selected as a result of this work and investigated (Fig. 1b): the thickest, exposing successions of 300 to 400 m thick (Bechare section, Kleiat section, Tarchich section, Safsafe section), cover the topmost part of the Kesrouane Formation, the Bhannes Formation and the Bikfaya Formation (Fig. 4), and are capped by the Gre`s de Base or locally by the Salima Formation, which are the earliest Cretaceous sediments. The other sections were several tens of metres thick (Lacloucq section, Nahr Ibrahim section, Faitroun section, Beskinta section, Bakaa section) and expose the top of the Kesrouane Formation, the Bhannes Formation and the Bikfaya Formation (Fig. 4). These latter sections were studied mainly in order to date the volcanic episode as precisely as possible and to document the lateral variability of these formations. The thickest Jurassic succession of northern Lebanon (Val D’Enfer) was not visited during these fieldtrips for safety reasons; only the Safsafe section was studied there. In the field, the beds were measured one by one and their lithology, texture, and macro- and microfossil content was described, to obtain the most

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Fig. 2. (a) Panorama of the Kleiat section. The Kesrouane Formation (Nahr Ibrahim Member) crops out from the valley bottom to the cliff top. The detailed log was made at the valley bottom and along the road. (b) Bechare section. Top of the Kesrouane Formation (Nahr Ibrahim Member; around the 170–250 m points of the log in Fig. 7). (c) Detail of limestone (wackestones and packstones) stacking pattern, with planar, horizontal and parallel bases and caps, displaying thickening-upward or thinning-upward features, characteristic of the Kesrouane Formation (Nahr Ibrahim Member). Fathi Higazi is located at the 230 m point of the log in Figure 7. Centimetre-thick intercalations of organic matter occur between the thinnest beds (arrowed). (d) Kleiat section. Characteristic appearance of the Kesrouane Formation (the truck is at about the 175 m point of the log in Fig. 5). (e) Nahr Ibrahim Valley. The outcrop exposes the top of the Kesrouane Formation (k; Nahr Ibrahim Member), the Bhannes Formation (b), here more than 100 m thick, the Bikfaya Formation (bk) and the Gre`s de Base Formation (g). The white line locates the section measured in the Bikfaya Formation (Fig. 4). Note the incomplete nature of the Bikfaya Formation (arrowed), here eroded by an incised valley filled by sandstones. (f) Bhannes Formation. Basalt flow near Bechare.

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Fig. 3. (g) Bhannes Formation. Accumulation of several tens of metres of volcano-sedimentary material beneath the village of Qartaba, Nahr Ibrahim Valley. (h) Panorama of the top of the Kesrouane Formation (k), overlain by the Bhannes Formation (volcano-sedimentary continental deposits, with two intercalations of marine limestone (foraminifer-bearing wackestone) are interbedded, indicating brief marine incursions (arrowed). (i) Incised valley (arrowed) at the top of the Bikfaya Formation near Tarchich (approximately 185 m point on log in Fig. 4). (j) Cross-bedded fluviatile sandstones near Tarchich. Gre`s de Base Formation (top of log in Fig. 5).

comprehensive information. Special attention was given to the sedimentary structures. All sections were regularly sampled, especially beds containing numerous microfossils so that they (microfossils) could be studied in thin-sections under the microscope. The facies were then interpreted in terms of their depositional processes and then in terms of their depositional environments. The stacking pattern of beds was also used to understand changes in relative sea-level. The biostratigraphy of the study sections was based on a biostratigraphical framework of benthic foraminifera and calcareous algae (Fig. 5).

Results Stratigraphy In the following, the results of the micropalaeontological study of several samples from the investigated area are outlined. The most fossiliferous samples belong to the carbonate rocks of the

Kesrouane and Bikfaya Formations, which crop out widely in the studied area. The analysis focused on the benthic foraminifera – with many species exhibiting complex endoskeletal structures – and the calcareous algae. These organisms are typical of carbonate platform environments. Identification of certain stratigraphically significant species allowed the formations to be dated (e.g. the Tarchich section, Fig. 5) and the sections to be correlated (Fig. 4). The Kesrouane and Bikfaya Formations are described below in terms of their micropalaeontological content. Kesrouane Formation. The oldest part of the formation occurs in the Kleiat section and in the lower part of the Bechare section. The most significant taxa are shown in Table 1a. This micropalaeontological assemblage suggests a Late Bathonian –Callovian age. The younger part of the formation occurs in the lower part of the Faitroun section, in the upper

0

BECHARE SECTION

m 30 20

?

330

10 0

320

SAFSAFE SECTION

250

30

240

KLEIAT SECTION

?

m 290

210 200 190 180 170 160 150

84 140

74

130 250

120

50 110 100

upper Bathonian to Callovian

Sequence boundary

upper Bathonian to Callovian

200 m 0

Upper Bathonian to Callovian

? m 0

220

m 0

90

DOGGER

N

350 m 340

Oxf.- l. Kim.

10

LACLOUCQ SECTION up. Kim. Tith.

upper Kim.

20

NAHR IBRAHIM SECTION

230

120 104

0

m 30

Oxfordian to Kimmeridgian

196

?

l. Oxf. - Kim.

Oxfordian to Kimmeridgian

Kimmeridgian 250

50

m 40

up. Kim.-Tith.

FAITROUN SECTION BESKINTA SECTION

286

l. Kim.

TARCHICH SECTION

S

Upper Oxfordian to Kimmeridgian

Oxfordian to (Lower ?) Kimmeridgian

MALM JURASSIC

upper Kimmeridgian Cretato Tithonian ceous

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Cretaceous Cretaceous

232

?

80 70 60 50 40 30 20

150

10 0

Relative sea level fall 100

Maximum flooding surface Relative sea level rise Transgressive surface

Grès de Base Fm. (continental) 50

Unconformity

Bikfaya Fm. (marine) Bhannes Fm. (continental)

Position of the samples

Kesrouane Fm. (Nahr Ibrahim Member; marine) 0

Fig. 4. Correlation of the study sections based on biostratigraphy of benthic foraminifera and calcareous algae, and chronostratigraphy of the various formations. The hemi-cycles recording short-period relative sea-level changes are also shown for each section. The two periods of major regression correspond to the exposure of the Levantine margin, located in the Bhannes Formation and the Gre`s de Base Formation.

part of the Bechare section, in the Tarchich, Quartaba and Beskinta sections and, probably, in the Safsafe section. The most significant species identified are summarized in Table 1b and allowed us to date this part of the Kesrouane Formation as Oxfordian–Kimmeridgian. Bikfaya Formation. This formation crops out in the Lacloucq section, in the middle–upper part of the Qartaba section, at the top of the Tarchich section and, in the Bechare section. The micropalaeontological assemblage (Table 1c) suggests a Late Kimmeridgian –Tithonian age. At the top of the Bechare section, oolitic limestones contain several species of the genus Trocholina, including Neotrocholina infragranulata (Noth). Ultimately, for the first time in Lebanon, a series of sections running from north to south is presented (Fig. 4) for which fauna, relevant for stratigraphic interpretation, are set against carefully surveyed logs. Accordingly, a precise stratigraphic

framework is proposed for the Jurassic formations under study. The Kesrouane Formation (here the Nahr Ibrahim Member) covers a time span from the Late Bathonian to parts of the (Early?) Kimmeridgian and the Bikfaya Formation extend from the (Late?) Kimmeridgian to lower part of the Tithonian. The volcanic episode (Bhannes Formation) is therefore clearly Kimmeridgian p.p. in age. It is to be noted that in the Anti-Lebanon (Yanta, Bakaa region), a yellow marl unit, recognized only in southern Lebanon, has been dated as midOxfordian by ammonite (Perisphinctes; determined by P. Courville, Univ. Rennes 1, France).

Sedimentology and depositional environments of the Jurassic formations of Lebanon Detailed sedimentological analysis of each of the sections has led to the characterization of eight facies identified in the Jurassic formations of Lebanon. Based on these, a summary sedimentological model

Cretaceous

Kimmeridgian Bhannes Fm.

Grès de

Bikfaya Fm. base Fm.

Fm

Kesrouane Fm. (Nahr Ibrahim Member)

286

250

196

120

104

84

74

0m

Praekurnubia crusei Nautiloculina circularis Redmondoides primitivus Permocalculus sp. Trocholina sp. Kurnubia cf. morrisi Siphovalvulina variabilis Redmondoides lugeoni Redmondoides gr. medius Kurnubia gr. palastiniensis Salpingoporella annulata Thaumatoporella parvovesiculfera Nautiloculina sp.

Trocholina elongata ? Heteroporella lemmensis Permocalculus inopinatus Kurnubia sp. ? Sanderella sp. Redmondoides cf. primitivus Cladocoropsis mirabilis

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

TARCHICH SECTION Levantinella egyptiensis Kurnubia cf. variabilis ? Glomospira sp. Clypeina jurassica cf. Rectocyclammina sp. Everticyclammina virguliana Thyrsoporella pseudoperplexa

233

Facies 1: bioturbated nodular limestone Description. These highly bioturbated wackestones look nodular; in outcrops they form indented steps. The microfauna consists mostly of benthic

Fig. 5. Distribution of benthic foraminifera and calcareous algae in the Tarchich section dating the top of the Kesrouane Formation and of the Bikfaya Formation. The enclosed Bhannes Formation most likely belongs to the Kimmeridgian.

has been constructed (Fig. 6). This has made it possible to identify fine variations in the palaeoenvironment within these thick and supposedly monotonous lithological units.

Jurassic Malm Upper Oxfordian - Kimmeridgian

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Table 1. List of the benthic foraminifera and calcareous algae of the Kesrouane and Bikfaya Formations (a) Kesrouane Formation (base) Callorbis minor Metzger Siphovalvulina variabilis Septfontaine Nautiloculina circularis (Said & Barakat) Pfenderalla arabica Redmond Redmondoides medius (Redmond) Redmondoides primitivus (Redmond) Redmondoides lugeoni (Septfontaine) Protopeneroplis striata Weynschenk, Satorina apuliensis Fourcade & Chorowicz Palaeopfenderina salernitana (Sartoni & Crescenti) Palaeopfenderina trochoidea (Smout & Sugden) Conicopfenderina mesojurassica (Maync) Praekurnubia crusei Redmond Kilianina blanketiformis Tasli Permocalculus inopinatus Elliott Kurnubia cf. variabilis Redmond Kurnubia palastiniensis Henson Trochamijiella gollsstanehi Athersuc, Banner & Simmons Alzonella cuvillieri Bernier & Neumann Nautiloculina oolithica Mohler *Salpingoporella annulata Carozzi. (b) Kesrouane Formation (top) Permocalculus inopinatus Elliott Cladocoropsis mirabilis Felix Nautiloculina circularis (Said & Barakat) Praekurnubia crusei Redmond Kurnubia cf. morrisi Redmond Kurnubia palastiniensis Henson Nautiloculina oolithica Mohler Trocholina elongata (Leupold) Trocholina cf. palastiniensis Henson Pseudolithocodium carpaticum Misik Redmondoides lugeoni (Septfontaine) Verneuilina cf. pharaonica Said Levantinella egyptiensis (Fourcade, Arafa & Sigal) Parurgonina caelinensis Cuvillier, Foury & Pignatti-Morano *Salpingoporella annulata Carozzi *?Heteroporella lemmensis (Bernier) *Thyrsoporella pseudoperplexa Garnier & Braik *Clypeina jurassica Favre (c) Bikfaya Formation Nautiloculina circularis (Said & Barakat) *Salpingoporella annulata Carozzi *?Heteroporella lemmensis (Bernier) Permocalculus inopinatus Elliott Everticyclammina virguliana (Koechlin) Alveosepta jaccardi (Schrodt) Redmondellina powersi (Redmond) Cladocoropsis mirabilis Felix Kastamonina abanica Sirel Campbelliella striata (Carozzi) Pseudospirocyclina maynci Hottinger Valvulina alpina Dragastan. *Calcareous algae

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

235

Wind Plus hautes mers

(6)

Plus basses mers

(5) (4) (6) (3)

Facies 1: Facies 2: Facies 3: Facies 4: Facies 5: Facies 6: Facies 7:

(2)

(1)

Bioturbated nodular limestone Limestone with benthic foraminifera and calcareous algae Limestone with green algae (dasycladales), carophytes and benthic foraminifera Coral limestone Fine-grained limestone with streaks of bioclastic limestone Oolitic limestone with megaripples Fine-grained limestone with birds-eyes

Fig. 6. Sedimentological model and distribution of various carbonate platform facies defined for the Kesrouane (Nahr Ibrahim Member) and Bikfaya Formation of the Jurassic of Lebanon.

foraminifera and sponge spicules (porifera). Echinoderm bioclasts are common. Interpretation. The wackestone texture reflects a comparatively calm depositional environment. Intense bioturbation is indicative of an environment with low sedimentation rates. The general absence of dasycladales (in comparison with the following facies), and the abundant benthic foraminifera indicate a relatively shallow, but poorly illuminated, shelf environment. A deep lagoon-type would provide such an environment. Facies 2: limestone with benthic foraminifera and calcareous algae Description. The white, grey or beige wackestones or packstones are very massive. They contain a microfossil assemblage composed mostly of benthic foraminifera and calcareous green algae such as dasycladales. Sponge spicules (porifera) are common, ostracods are present. Pellets are common to abundant. The macrofauna consists mostly of echinoderms bioclasts, bivalves and gastropods. Late dolomitization (isolated euhedral dolomite crystals) is observed locally. Interpretation. The wackestone or more rarely packstone texture is indicative of a comparatively calm depositional environment conducive to the accumulation of micrite particles. The microfauna assemblage with benthic foraminifera, calcareous

algae, and rare dasycladales suggests a warm depositional environment, somewhat shallower than that of facies 1, akin to a lagoon-type environment. Facies 3: limestone with green algae (dasycladales) charophytes and benthic foraminifera Description. The white, grey, or beige packstones to wackestones have a massive appearance. Dasycladales are very abundant and charophytes are frequent. Benthic foraminifera are always present, as pellets. The remainder of the fauna is composed of ostracods, gastropods, and rare bivalves. Commonly centimetres thick beds with very high organic content (here plant-debris accumulation) are observed. Interpretation. The packstones represent a relative high-energy environment where the micrite particles are partly winnowed. The dasycladales are very abundant reflecting a shallow, warm and well-illuminated environment. The charophytes and layers with plant-debris accumulation may indicate the existence of exposed areas (islands) near to these generally calm sedimentation zones. This facies is indicative of a shallow lagoon-type depositional environment. Facies 4: coral limestones Description. The white, beige or grey limestones contain branched or massive coral heads in life position or tilted (boundstones). Coarse reworked

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bioclastic facies were not observed. The corals are distributed in small patches or occur as biostromes several decimetres to one metre thick. The matrix is composed of micrite containing abundant benthic foraminifera, calcareous algae, mostly dasycladales, and pellets. Interpretation. The occurrence of coral patches or biostromes is indicative of a shallow platform environment, as is the presence of calcareous algae, mostly dasycladales. The abundant micrite matrix between patches indicates a comparatively low-energy environment. The depositional environment is similar to that of facies 4, the only difference is the local development of small coral build-ups. Facies 5: fine-grained limestones with streaks of bioclastic debris Description. The wackestones contain benthic foraminifera and calcareous algae, mostly dasycladales. They exhibit interspersed streaks of coarse bioclastic grainstone or packstone with an erosional base. Erosion surfaces are sometimes observed at the top of the beds. Interpretation. The streaks of coarse bioclastic grainstone or packstone are interpreted as spillover-wash. The wackestone texture reflects a generally calm depositional environment, influenced by high-energy currents (spill-over-wash and erosion surfaces). This, together with the microfauna assemblage, indicates a very shallow, lagoon-type depositional environment. Facies 6: oolithic limestones with megaripples Description. The grainstones are made up mostly of ooids. Sedimentary structures are common: trough crossbedding, planar crossbedding, wave ripples and high-energy horizontal laminations. Interpretation. The grainstone texture indicates shallow, high-energy depositional environments influenced by relatively high-energy, uni-directional currents in the fair-weather wave zone. These features indicate a more open shelf setting than for facies 1–5. Facies 7: fine-grained limestone with birds-eyes Description. The white, beige, or grey wackestones are characterized by the presence of birds eyes. The microfauna is represented mostly by benthic foraminifera, calcareous green algae, and sponge spicules (porifera); the remainder of the fauna, which is less abundant, is made up of echinoderms bioclasts, bivalves and gastropods. Pellets are common. Interpretation. The wackestone texture and the microfauna indicate a calm, warm environment. The birds-eyes are characteristic for temporary

subaerial exposure. These features point to a very shallow lagoon-type depositional environment in the intertidal zone. Facies 8: basalt flows and pyroclastic material Description. This facies is associated with basaltic volcanic activity. It is composed mostly of eruptive material represented either by basalt flows or by build-ups of pyroclastic material. The basalt flows often exhibit columnar structures; weathering may lead to the formation of crumbly balls. However, no pillow lavas have been recognized (Abdel-Rahman 2002). The pyroclastic material is generally very fine but coarse streaks are observed locally. Within these volcanic strata palaeosols occasionally occur containing traces of roots, pedogenic nodules, and amber. In addition, fine-grained limestone intercalations containing benthic foraminifera occur locally (e.g. in the Bakaa section, Fig. 3h). Conglomerates and sands containing volcanic material or carbonate are sometimes interspersed in basalt flows. Interpretation. The absence of pillow lavas and the presence of palaeosols indicate that the volcanic material was deposited in a continental environment as envisaged by other workers (e.g. Abdel-Rahman 2002). The intercalated limestone with foraminifera are probably indicative of brief marine incursion. Thus, based on a sedimentological analysis, the carbonate facies in the Middle and Upper Jurassic formations of Lebanon indicate relatively calm, warm-water, shallow-platform environments (facies 1–5 and 7). Judging from the rock texture and from the flora and fauna, the environments were lagoonlike (Fig. 6). More specifically, the extension of these shallow (0–15 m) environments over several tens or even hundreds of kilometres is characteristic of an epicontinental shelf. Facies 6, which is far less common and existed only for limited periods of time, denotes areas that were slightly more open and so subject to waves and currents (Fig. 6).

Interpretations and discussion Evolution of depositional environments, relative sea-level changes, and palaeogeography of the Levantine margin In each section, detailed identification of the various facies and environments made it possible to trace their vertical and lateral evolution. Taken in conjunction with a stratonomic analysis, it also allowed us to reconstruct relative sea-level changes. These results are summarized in Figure 4. Evolution of depositional environments, stacking patterns and relative sea-level changes. Relative

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

sea-level changes of longer periods and greater amplitudes (sequences of several tens of metres thick) are clearly visible in the evolution of the facies, interpreted in terms of evolution of the depositional environments (e.g. Bechare section, Fig. 7). The top of the Kesrouane Formation clearly exhibits shallowing upward facies (uppermost Oxfordian –lowermost Kimmeridgian) reflecting a global relative sea-level fall (Fig. 5), observed in all the study sections. The regression corresponds to the time when the continental volcanic strata were emplaced across Lebanon (Fig. 4), and was followed by a marine transgression corresponding to the deposition of the Bikfaya Formation (uppermost Kimmeridgian –Tithonian?). Above, the Upper Jurassic erosional surface observed all along Lebanon corresponds to a major regional regression. This major unconformity at the Jurassic –Cretaceous boundary is consistent with that observed all along the Levantine margin (e.g. Mouty 2000; Hirsch 2005). In an epicontinental shelf context with facies that are vertically and laterally highly monotonous, only a very detailed study of the facies and of the stacking pattern (bed thickness, frequency of erosional structures, etc.) can indicate shorter termed and smaller amplitude changes. For example, when the individual beds are grouped into thinningupward bundles, this stacking pattern can be used for the interpretation of thinning-upward depositional sequences and of relative sea-level changes (Strasser et al. 1999). Thus, in the Kesrouane Formation (Lower Bathonian p.p. – Kimmeridgian p.p.), the few metres to a few tens of metres thick sedimentary sequences (Fig. 4) often exhibit an upward thinning of beds, without changes in facies, but with erosion structures at their tops. If relative sea level dropped below the depositional surface and unconsolidated sediment was eroded, the erosional surface is considered as the sequence boundary (Strasser et al. 1999). Sometimes, the sequence boundaries are defined by subtidal, intertidal (birds-eyes), or supratidal surfaces that exhibit the shallowest facies of the sedimentary sequence. Here, ore sporadic relative sea-level rises are recorded, though, within sedimentary units of less than 1 m thick; these are generally marked by highly bioturbated facies (facies 1). In shallow-water carbonate environments, a thin transgressive lag at the base of the sequence is common, resulting from the time that is needed to re-establish the carbonate production, after emersion for example. In our study, very often, the transgressive deposits are absent, and only the transgressive surface occurs, directly overlain by the regressive deposits. In such environments, thin or absent deepening-upward and thick shallowingupward depositional sequences may be found

237

(Jones & Desrochers 1992; Pratt et al. 1992). The shallowing-upward trend is due to the high carbonate production and accumulation, which, as long as environmental conditions are favourable, easily outpaces the relative sea-level rise (Schlager 1981). In the Kesrouane Formation, the deepeningupward sequences are absent and thick shallowingupward sequences are observed. A thinning-upward trend of shallowing-upward sequences (or elementary and/or small scale sequences sensu Strasser et al. 1999) may be interpreted as a general loss in accommodation. The top of the Kesrouane Formation (Oxfordian – Kimmeridgian p.p.) is marked by more amalgamated shallowing-upward sequences than those identified for the Callovian (e.g. Tarchich, Kleiat, and Bechare sections; Fig. 4). This probably reflects a slow-down in the creation of accommodation in the course of the Oxfordian and Kimmeridgian compared with the Callovian, before the exposure stage of the Levantine margin during the Kimmeridgian. Later, the deposition of the Bikfaya Formation corresponds to a brief increasing of the creation of accommodation before a loss of accommodation (regression and exposure) at the Jurassic– Cretaceous boundary. The same succession is recognized in northern Israel (Hirsch 2005). In Syria, the Lower Jurassic is very thin and shallow marine, or absent. Here, only the Lacloucq section, which is the only one where the Kimmeridgian and Tithonian could be separated, displays a more complex pattern, suggesting that cycles of the same order or of a higher order are poorly recorded in the other sections where the Kimmeridgian and Tithonian strata could not be separated. Palaeogeography of Lebanon and the Levantine margin. From the end of the Bathonian to the onset of the Kimmeridgian (Kesrouane Formation), epicontinental shelf environments persisted over much of Lebanon. More local observations, for example, further north in Val D’Enfer and in the Sir Ed Danie´ region, confirm that the same facies occur here as in central Lebanon. The same is true for the Anti-Lebanon, further south (Yanta, Bakaa region), at least until the lowermost Oxfordian, whereas during mid-Oxfordian a yellow marl unit with ammonites characteristic of a deep outer-shelf environment is recognized. This is overlain by a coral limestone formation deposited in a shallowshelf environment, probably from the end of the Middle to Late Oxfordian to Early Kimmeridgian. This is representative of part of the deeperenvironment sedimentary series recognized on the eastern slopes of Mount Hermon (Vautrin 1934; Dubertret 1955), indicative of deeper palaeogeographic areas in this sector. However, the presence of the more recent, large transform Yammouneh

Grés de base fm.

mG mP G P WM

BECHARE SECTION m 350

Facies number

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E1

E2

E3

E4

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E6

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ooid

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green algae pedogenesis echinoderm bivalve sponge peloid

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organic matter birds-eyes

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corals pillow lava

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spill over wash ripple mark

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current ripple 110

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Dogger

ENVIRONMENTS

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26 27 28

50 40

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E1 - Continent E2 - Very shallow lagoon (intertidal) E3 - Very shallow lagoon (subtidal) E4 - Shallow lagoon (subtidal) E5 - Lagoon (subtidal) E6 - Open shallow shelf (subtidal) CARBONATE TEXTURES G - Grainstone mG - Micrograinstone

20

P - Packstone mP - Micropackstone

10 0

W - Wackestone

34 35

M - Mudstone

Fig. 7. Bechare section. Log, stratigraphy, sedimentology, and evolution of the depositional environments. The stacking pattern of the different facies defined in this study (Fig. 6) allows the longer-period relative sea-level changes to be defined.

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

Fault confuses the exact palaeogeographic situation of this sector of Lebanon in the Jurassic. By comparison, in Israel, a 1000 m thick sedimentary unit stretches from the Coastal Chain to Galilee and has been ascribed to the Upper Liassic to Oxfordian. This Haifa Formation is composed of limestones, dolomitic limestones, and dolostones (Hirsch & Picard 1988; Hirsch 2005; Krascheninnikov 2005) deposited in shallow marine shelf environments. The uppermost part of this marine formation is the equivalent of the Kesrouane Formation in Lebanon. In the Coastal Ranges of Syria, the Middle Jurassic sedimentary succession is thick and formed by a monotonous stack of massive, fine-grained marine limestones with abundant benthic foraminifera, interstratified with dolomites, and dolomitic limestones (Mouty 2000; Alme´ras & Mouty 2001). All of these marine sedimentary formations of the Middle Jurassic are evidence of a vast carbonate shelf that extended at least from Syria (Coastal Ranges) to northern Israel (Coastal Chains, Galilee). In the Sinai and Negev, deeper marine environments occurred (Hirsch 2005). During the Kimmeridgian, in Lebanon the carbonate shelf regime ended suddenly with a volcanic episode that was contemporaneous with regional exposure, associated with the emplacement of more or less extensive and persistent continental environments. Whereas palaeosols occurred locally, elsewhere there were brief marine incursions, denoting the locally temporary character of the exposure phase. Similarly, the top of the Kesrouane Formation exhibits local signs of karstification (Nader & Swennen 2004). Further south, in northern Israel (Galilee), similar to Lebanon, the Upper Jurassic exhibits a break in the marine sedimentation regime with the emplacement of volcanic units associated with a period of marked regression and total or partial exposure of the platform (Hirsch & Picard 1988; Hirsch 2005; Krascheninnikov 2005). Further south, in southern Israel, marine sediments persisted during this period. At the end of the Kimmeridgian and for part of the Tithonian, a new marine transgression occurred across Lebanon (Bikfaya Formation). A more or less internal, shallow carbonate platform developed with components (oolites, calcareous algae, mostly dasycladales) that attest to a warm climate. In Syria, the Upper Jurassic is found in the Coastal Chains alone and exhibits the same characteristics as in Lebanon; the Tithonian is absent (Mouty 2000). In Galilee, in the course of the marine transgression, a shallow platform developed. These features indicate the re-establishment of a relatively shallow, large carbonate shelf extending from Syria to northern Israel. Further south, deeper marine environments are reported (Hirsch 2005).

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Over the entire Levantine margin, the end of the Jurassic is marked by a marine regression of large amplitude, leading to the subaerial exposure of the entire region. This phase is associated locally with intense erosion of Jurassic sediments, marked occasionally by incised valleys. This exposure persisted until the onset of the Cretaceous with the deposition of the Gre`s de Base Formation.

Development of Lower Jurassic volcanism and geodynamic implications In the Phanerozoic, the fracturing and thinning of the continental crust in the Levant region reportedly results from several phases of extension. Various ages have been proposed for the opening of the NeoTethyan basin, ranging from the Late Permian, the Triassic, or the Early Jurassic (Freund et al. 1975; Garfunkel 1998; Stampfli & Borel 2002), the Jurassic (Ginzburg & Gvirtzman 1979) through to the Cretaceous (Dercourt et al. 1986). In this context, the first major tectonic event occurred in the Triassic or Liassic, depending on the region (Garfunkel & Derin 1984; Brew et al. 2001). This first major phase extended all along the entire Levantine margin in the course of the Liassic. The second episode of major extension occurred in the Early Cretaceous (e.g. Sawaf et al. 2001). However, these two rifting episodes, that were responsible for significant thinning of the continental crust, do not seem to have led to the emplacement of any oceanic crust. Between these two events, the time span from the Late Liassic to Late Jurassic is often considered as a period of great tectonic stability (Cohen et al. 1990). Observations made in Lebanon and in the adjacent countries point, in part, in the direction exposed above. The persistence of a large epicontinental carbonate shelf (Kesrouane Formation) during the Middle and Late Jurassic (Bathonian– Kimmeridgian p.p.) reflects constant tectonic and geodynamic stability of the Levantine margin in the course of the period. However, for some workers (e.g. Walley 1998; Nader et al. 2007), this stability was interrupted by the development of volcanism during the Kimmeridgian, which was supposedly associated with a tectonic phase. AbdelRahman (2002) alone argues that this volcanism is to be related to an extensional phase and the subsequent rise of magmatic material through the continental crust. To answer the problem, it should be recalled that: (1) the Kesrouane Formation is of a relatively constant thickness and the top of the formation, wherever it has been dated, is of the same age; (2) the depositional environments of the Kesrouane Formation reflect platform

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morphology of little contrast. The top of the platform seems to be a near-planar shelf stretching across northern Israel, Lebanon, and much of Syria; the volcanic-sedimentary deposits of the Bhannes Formation display great thickness variations ranging from 0 m to more than 100 m; and the Bikfaya Formation, which is generally marked by a limestone cliff, also displays great thickness variations ranging from 0 m to 30 m.

As regards the thickness variations of the latter two formations, the maximum extent of the Bhannes Formation generally coincides with that of the Bikfaya Formation. Conversely, where the Bhannes Formation is very thin or absent, the same is true of the Bikfaya Formation. This is particularly apparent in the Nahr Ibrahim valley where, at the NE end of the valley, the Bhannes Formation, made up of pyroclastic material and volcano-clastic sediments, is about 150 m thick, as is the Bikfaya Formation. At the SW end of the valley, the two formations are absent and the

Qartaba

1

2 1 km

N

J 6 - Kesrouane Fm. (Upper Bathonian - Kimmeridgian p.p.) Qartaba B J6 - Bhannes Fm. (Kimmeridgian p.p.) J 6a - Bikfaya Fm. (Kimmeridgian p.p. - Tithonian p.p.) C 1 - Grès de Base Fm. (basal Cretaceous) C2 à C4 - Cretaceous Formations Quaternary superficial deposits Major fault Uncertain fault

Fig. 8. Excerpt from the simplified Qartaba geological map (1:50 000). Formation notations are from the geological map. Formation names have been adapted to the terms used in this paper. (1) Probable incised valley eroding Jurassic sediments and filled with the Gre`s de Base. (2) The unconformity observed in the field displays a progressive thinning of the Bhannes Formation and Bikfya Formation to the SE, then their complete disappearance, with the contact of the Gre`s de Base Formation directly overlying the Kesrouane Formation. A few kilometres to the SE the two formations are observed again and are several tens of metres thick; they thin out again progressively to the SE and totally disappear.

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION

Cretaceous Gre`s de Base Formation overlies directly the top of the Kesrouane Formation (Fig. 8). Moreover, it appears that the transition between these two extremes includes lateral variations in thickness and lateral facies transitions in

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the Bikfaya Formation. The Gre`s de Base displays strong thickness variations (Dubertret 1955), with the thickest development directly above the thickest Upper Jurassic formations and vice versa. Some kilometres further to the SW of the valley, the (D)

N

S Earliest Kretaceous

Regression

(C) N

Bikfaya Fm.

S Kimmeridgian pars to Tithonian (?)

Transgression

(B) N

Bhannes Fm.

S Kimmeridgian pars

Regression

(A) N

S

Kesrouane Fm.

Late Bathonian to Kimmeridgian pars

Basalts and volcano-sedimentary deposits (Bhannes Fm.)

Grès de Base Fm.

Limestone of epicontinental shelf (Kersouane Fm.)

Limestone of agitated to protected shelf (Bikfaya Fm.)

Fig. 9. Diagram summarizing the main sedimentary and tectonic events in the Lebanon from the Late Bathonian to the onset of the Cretaceous. After a long period of geodynamic stability (Bathonian to Kimmeridgian), a regression is associated with the formation of titled blocks corresponding to the initiation of a rifting phase, in conjunction with intra-plate volcanism during the Kimmeridgian. Subsequently, the deposition of the Bikfaya Formation (Kimmeridgian– Tithonian) reflects a brief return to marine conditions before the major regression at the end of the Jurassic. During the earliest Cretaceous, fluviatile sedimentation prevailed and in some cases largely eroded Jurassic formations. The rifting phase initiated before seems to have continued with low intensity.

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Bikfaya and Bhannes Formations are observed again at their maximum thicknesses. This arrangement reflects a tilted block morphology, with the Gre`s de Base capping the underlying topographies (Fig. 9). This is consistent with the observations of Nader et al. (2007), who report several preCretaceous faults in the Kesrouane Formation. This observation holds for many valleys of Mount Lebanon and in the Bakaa region, on the western slopes of the Anti-Lebanon. The rifting phase reported here for the Late Jurassic is consistent with the interpretation of Abdel-Rahman (2002) and Laws & Wilson (1997), who conclude that the Lebanese Mesozoic alkali basalts developed in an intra-plate setting, related to a rifting stage, possibly related to a mantle plume, and occurring in association with the formation of the Levant margin. According to Hirsch (2005), in Israel, prior to the Early Cretaceous, the Jurassic was tilted and followed by a phase of erosion. These observations as a whole show that the rifting period, initiated at the end of the Triassic, came to a halt in the Middle and Late Jurassic, before resuming in the Kimmeridgian and Tithonian (Fig. 9). This phase represents a fairly small amplitude resumption of rifting before the major Cretaceous episode.

Conclusions Sedimentological and stratigraphical studies of the Jurassic sedimentary succession of Lebanon (Levantine margin) have led to the development of a depositional model for the Kesrouane (Nahr Ibrahim Member), Bhannes and Bikfaya Formations, and to their positioning in a stratigraphic framework. Based on the study of assemblages of benthic foraminifera and calcareous algae, it has been defined that: (1)

the Kesrouane Formation extends from the Late Bathonian probably to the Early Kimmeridgian; (2) the Bhannes Formation corresponds to parts of the Kimmeridgian (Early/Late Kimmeridgian boundary?); and (3) the Bikfaya Formation is of (Late?) Kimmeridgian to Tithonian age. The carbonate deposits of the marine Kesrouane and Bikfaya Formations are characteristic of shallow environments with water depths, not more than a few metres, that were sheltered and low energy, warm and well-illuminated. Locally, more open and higher energy environments existed. The Bhannes Formation, interbedded between the previous two, is made up of volcanic rocks (basalt lava flows and pyroclastic material) and developed in a continental setting.

Two regressive phases are recognized: one in the course of the Kimmeridgian and a major one at the end of the Jurassic, continuing into the base of the Cretaceous with the deposition of the Gre`s de Base Formation. No major transgressive phase was recognized within the Kesrouane Formation. By contrast, the deposition of the Bikfaya Formation is associated with a transgression clearly recorded over all this part of the Levantine margin in the (Late?) Kimmeridgian–Tithonian. The north–south extension of the Kesrouane Formation over several hundred kilometres, from Syria to northern Israel, suggests there was a large epicontinental shelf that suddenly terminated with the volcanic event recognized across Lebanon. The persistence of this epicontinental platform reflects the great geodynamic stability of the region in the Bathonian, Oxfordian and parts of the (Early?) Kimmeridgian. The Kimmeridgian volcanic event is associated with a major geodynamic event during which the Levantine margin was dissected into tilted blocks. This phase represents a fairly rifting resumption phase before a much larger episode in the Cretaceous (Homberg et al. 2010). This work was supported by the Middle East Basin Evolution programme. We are grateful to A. Lethiers, F. Delbe`s and C. Abrial for technical support. The authors thank anonymous Lebanese people for their help in the field. We are grateful to F. Fu¨rsich and one anonymous reviewer, and to M. Bachmann and C. Homberg (invited Editors) for comments on a previous version of this paper.

References Abdel-Rahman, A. M. 2002. Mesozoic volcanism in the Middle East: geochemical, isotopic and petrogenetic evolution of extension-related alkali basalts from central Lebanon. Geological Magazine, 139, 621–640. Almeras, Y. & Mouty, M. 2001. Les brachiopodes du Jurassique de Syrie. Revue de Pale´obiologie, Gene`ve, 20, 9 –17. Beydoun, Z. R. 1988. The Middle-East: Regional Geology and Petroleum Resources. Scientific Press, London, 292. Brew, G., Barazangi, M., Al-Maleh, A. K. & Sawaf, T. 2001. Tectonic and geologic evolution of Syria. Geoarabia, 6, 573– 615. Butler, R. W. H. & Spencer, S. 1999. Landscape evolution and the preservation of the tectonic landforms along the northern Yammouneh Fault, Lebanon. Journal of the Geological Society, London, Special Publications, 162, 1 –14. Cohen, Z., Kaptsan, V. & Flexer, A. 1990. The tectonic mosaic of the southern Levant: implications for hydrocarbon prospects. Journal of Petroleum Geology, 13, 437–462.

JURASSIC STRATIGRAPHY AND SEDIMENTARY EVOLUTION Dercourt, J., Zonenshai, L. P. et al. 1986. Geological evolution of the Tethys belt from the Atlantic to the Pamirs since the Lias. Tectonophysics, 123, 241– 315. Dubertret, L. 1955. Carte Ge´ologique du Liban au 1/ 200000 avec notice explicative. Re´publique Libanaise, Ministe`re des Travaux Publics, Beiruth. Dubertret, L. 1975. Introduction a` la Carte Ge´ologique au 1/50 000 du Liban. Notes et Me´moires sur le Moyen-Orient, 13, 345– 403. Freund, R., Goldberg, M., Weissbrod, T., Druckman, Y. & Derin, B. 1975. The Triassic-Jurassic structure of Israel and its relation to the origin of the eastern Mediterranean. Bulletin – Geological Survey of Israel, 65, 26. Garfunkel, Z. 1998. Constraints on the origin and history of the eastern Mediterranean basin; collision-related processes in the Mediterranean region. Tectonophysics, 298, 5–35. Garfunkel, Z. & Derin, B. 1984. Permian-early Mesozoic tectonism and continental margin formation in Israel and its implications for the history of the eastern Mediterranean. Geological Society Special Publications, 17, 187– 201. Ginzburg, A. & Gvirtzman, G. 1979. Changes in the crust and in the sedimentary cover across the transition from the Arabian platform to the Mediterranean basin: evidence from seismic refraction and sedimentary studies in Israel and Sinai. Sedimentary Geology, 23, 19– 36. Gomez, F., Meghraoui, M. et al. 2001. Coseismic displacements along the Serghaya Fault: an active branch of the Dead Sea Fault System in Syria and Lebanon. Journal of the Geological Society, London, 158, 1 –4. Hirsch, F. 2005. The Jurassic of Israel. In: Hall, J., Krasheninnikov, V. A., Hirsch, F., Benjamini, C. & Flexer, A. (eds) Geological Framework of the Levant. Volume II: The Levantine Basin and Israel, Jerusalem, 361– 391. Hirsch, F. & Picard, L. 1988. The Jurassic facies in the Levant. Journal of Petroleum Geology, 11, 277–308. Homberg, C., Barrier, E., Mroueh, M., Muller, C., Hamdan, W. & Higazi, F. 2010. Tectonic evolution of the central Levant domain (Lebanon) since Mesozoic time. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 245– 268. Jones, B. & Desrochers, A. 1992. Shallow platform carbonates. In: Walker, R. G. & James, N. P. (eds) Facies Models, Response to Sea-Level Change. Geological Association of Canada, 277 –301. Kuss, J. 1990. Middle Jurassic calcareous algae from the circum-Arabian area. Facies, 22, 59–86. Krasheninnikov, V. A. 2005. Jurassic: stratigraphy and lithology of the sedimentary cover. In: Hall, J., Krasheninnikov, V. A., Hirsch, F., Benjamini, C. & Flexer, A. (eds) Geological Framework of the Levant. Volume I: Cyprus and Syria, Jerusalem, 221–235. Laws, E. D. & Wilson, M. 1997. Tectonics and magmatism associated with Mesozoic passive continental margin development in the Middle East. Journal of the Geological Society, London, 154, 459–464.

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Mouty, M. 2000. The Jurassic in Syria: an overview. Lithostratigraphic and biostratigraphic correlations with adjacent areas. In: Crasquin-Soleau, S. & Barrier, E. (eds) Peri-Tethys Memoir 5: New Data on Peri-Tethyan Sedimentary Basins. Memoire du Museum National d’Histoire Naturelle, Paris, 182, 159– 168. Mouty, M. & Zaninetti, L. 1998. Le Jurassique du mont Hermon (Anti-Liban). De´couverte de Trias et de Lias. Archives Scientifiques, Gene`ve, 51, 295– 304. Nader, F. H. & Swennen, R. 2004. Petroleum propects of Lebanon: some remarks from sedimentological and diagenetic studies of Jurassic carbonates. Marine and Petroleum Geology, 21, 427– 441. Nader, F. H., Swennen, R. & Ottenburgs, R. 2003. Karst-meteoric dedolomitisation in Jurassic carbonates, Lebanon. Geologica Belgica, 6, 3– 23. Nader, F. H., Swennen, R. & Ellam, R. 2004. Stratabound dolomite versus volcanism-associated dolomite: an example from Jurassic platform carbonates in Lebanon. Sedimentology, 51, 339– 360. Nader, F. H., Swennen, R. & Ellam, R. 2007. Field geometry, petrography and geochemistry of a dolomitization front (Late Jurassic, central Lebanon). Sedimentology, 54, 1093– 1119. Noujaim Clark, G. & Boudagher-Fadel, M. K. 2001. The larger benthic foraminifera and stratigraphy of the Upper Jurassic/Lower Cretaceous of Central Lebanon. Revue de Micropale´ontologie, Gene`ve, 44, 215– 232. Noujaim Clark, G. & Boudagher-Fadel, M. K. 2004. Larger benthic foraminifera and calcareous algae of the upper Kesrouane Limestone Formation (Middle/Upper Jurassic) in Central Lebanon. Revue de Pale´obiologie, 23, 477–504. Pratt, B. R., James, N. P. & Cowan, C. A. 1992. Peritidal carbonate. In: Walker, R. G. & James, N. P. (eds) Facies Models, Response to Sea-Level Change. Geological Association of Canada, 203– 322. Renouard, G. 1951. Sur la de´couverte du Jurassique infe´rieur (?) et du Jurassique moyen au Liban. Compte-Rendu de l’Acade´mie des Sciences, Paris, 232, 992–994. Saint-Marc, P. 1980. Le passage Jurassique-Cre´tace´ et le Cre´tace´ infe´rieur de la re´gion de Ghazir (Liban Central). Ge´ologie Me´diterrane´enne, 7, 237–245. Sawaf, T., Brew, G. E., Litak, R. K. & Barazangi, M. 2001. Geologic evolution of the intraplate Palmyride Basin and Euphrates fault system, Syria. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & Crasquin-Soleau, S. (eds) Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basin and Passive Margins, Memoire du Museum National d’Histoire Naturelle, Paris, 186, 441 –467. Schlager, W. 1981. The paradox of drowned reefs and carbonate platforms. Geological Society American Bulletin, 92, 197 –211. Stampfli, G. M. & Borel, G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth and Planetary Science Letters, 196, 17– 33. Strasser, A., Pittet, B., Hillga¨rtner, H. & Pasquier, J. B. 1999. Depositional sequences in shallow

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tectonic evolution of the Levantine region. Tectonophysics, 298, 37–62. Walley, C. D. 2001. The Lebanon passive margin and the evolution of the Levantine Neotethys. In: Ziegler, P. A., Cavazza, W., Robertson, A. H. F. & Crasquin-Soleau, S. (eds) Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basin and Passive Margins, Memoire du Museum National d’Histoire Naturelle, Paris, 186, 407– 439.

Tectonic evolution of the central Levant domain (Lebanon) since Mesozoic time C. HOMBERG1*, E. BARRIER1, M. MROUEH2, C. MULLER3, W. HAMDAN2 & F. HIGAZI2 1

Universite´ Pierre et Marie Curie, ISTeP, CNRS, Case 129, 4 place jussieu, 75252 Paris Cedex 05, France 2

Universite´ libanaise, Faculte´ d’Agronomie, B.P. 13-5368 Chourane, Beyrouth 1102-2040 Lebanon 3

Independent researcher

*Corresponding author (e-mail: [email protected]) Abstract: The tectonic history of the central part of the Levant domain (Lebanon) is re-evaluated. Examination of the tectonic structures and mechanical analysis of the meso-scale brittle deformation indicate that Lebanon has experienced four major tectonic events since Late Mesozoic time. The first was an Early Cretaceous extensional phase orientated north– south to NNE– SSW. It produced WSW–ENE to WNW– ESE normal faults with offsets up to several hundreds of meters and led to the development of an approximately WNW–ESE-trending basin. A second extension, with similar driving stresses, occurred during Eocene time and persisted, perhaps until Oligocene times. The Early Neogene period marked a dramatic change in the structural evolution of Lebanon after which strike– slip and reverse faulting and folding dominated. During Early Miocene times, an east– west compression produced moderate folding and faulting. A second, but much more severe, folding event occurred during Late Miocene time owing to a NNW– SSE compression. This new tectonic history allows the discussion of several aspects of the Eastern Mediterranean basin development and the later deformation of its continental margin and surroundings, in particular: (1) the driving mechanisms of the Levant basin opening; (2) the inversion of its adjacent margin; and (3) the age, origin, and evolution of the restraining bend of the Dead Sea Transform in Lebanon.

The Levant domain encompasses the eastern-most part of the Eastern Mediterranean in a region where several crustal plates interact, the major ones being the Arabian, African and Eurasian plates (Fig. 1). Deformation in this complex area started during Late Palaeozoic –Early Mesozoic time with the development of the Eastern Mediterranean basin and its associated passive margins (Freund et al. 1975; Ben-Avraham 1989; Cohen et al. 1990; Hirsch et al. 1995; Garfunkel 1998; Ben-Avraham et al. 2002; Gardosh & Druckman 2006). It was later faulted and folded during the closure of the western Neo-Tethys and as a result of the initiation of the Dead Sea left-lateral Transform boundary (DST) (Freund et al. 1970; Garfunkel 1981; Atallah 1992; Moustafa & Khalil 1994; Ben-Gai & Ben-Avraham 1995; Vidal et al. 2000; Gardosh & Druckman 2006; Schattner et al. 2006a). A crucial part of the Levant domain is the Lebanese sector, where Lower Jurassic to Quaternary deposits are well exposed offering key data to reconstruct the tectonic history of the Levant domain since Mesozoic time. However, only sporadic studies have

been undertaken during the period of political instability in the second half of the 20th century. As a result, the understanding of the geological history of Lebanon is rather incomplete. A few recent tectonic studies have served to quantify recent vertical and horizontal movements and they leave no doubt that Lebanon and its western offshore extension are active parts of the DST (Gomez et al. 2003; Daeron et al. 2004; Morhange et al. 2006; Nemer and Meghraoui 2006; Elias et al. 2007; Gomez et al. 2007). Many aspects of the tectonic evolution of Lebanon remain uncertain. The geology of Lebanon was mapped at the 1:50 000 scale by Dubertret and collaborators (1955, 1975) during the 1950s and 1970s. Interpretation of the geological structures of Lebanon within a coherent structural-tectonic scheme was offered by Quennel (1984), Garfunkel (1981), Butler et al. (1998) and Walley (1988, 1998, 2001). These models agree on the general tectonic evolution of Lebanon but often disagree on specific details. For example, the timing and significance of the Lebanese folds is highly controversial and the

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 245–268. DOI: 10.1144/SP341.12 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Geologic and tectonic structures in Lebanon. (a) Kinematic framework. Af, Ar, Eu: African, Arabian, Eurasian plates. DSF: Dead Sea fault, main fault of the Dead Sea transform plate boundary. RS, CA and CZ: Red Sea rift, Cyprian Arc and Arabia-Eurasia collision zone. LB: Levant basin. The Lebanese restraining bend is indicated with a dashed square. (b) Simplified geologic and tectonic map of Lebanon after Dubertret and collaborators (1955). Faults: Yammouneh fault (YF); Roum fault (RoF); Rachaya fault (RaF); Sergaya fault (SF); Mount Lebanon thrust (MLT, after Elias et al. 2007) and Akkar thrust (AT). Yp: Yammouneh pull-part. Folds: Mount Lebanon anticline (MLA); Bekaa syncline (BS); Mount Anti-Lebanon anticline (MALA); Cities: Beirut (Be); Balbeck (Ba); Tripoli (Tr); Zahle (Za). White stars indicate from north to south Meten, Chouf, Jezine areas. All bed dip data are from this study. (c) Cross section through the Mount Lebanon anticline (see Fig. 1b for location) modified from Walley (1998). Legend: 1, Jurassic; 2, Neocomian –Albian; 3, Cenomanian–Turonian; 4, Senonian; 5, Eocene; 6, Miocene (m).

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age of the major folding episode(s) in Lebanon are thought to be either Late Cretaceous and Eocene (Walley 1998), pre-Messinian (Butler 1998), or Pliocene (Quennel 1984). Lack of constraint is particularly pronounced for the pre-Neogene period when the history is based primarily on interpretation of large-scale stratigraphic relationships in the absence of associated structural features. In order to better constrain the tectonic history of Lebanon, field data have been acquired within the framework of the Middle East Basins Evolution (MEBE) Programme and a cooperative programme between the University Pierre and Marie Curie and the Lebanese University. The goal was to reexamine the tectonic structures with modern geological tools and to establish the age and nature of the main tectonic events that have deformed Lebanon since Mesozoic time. In this paper, we present the data acquired during this four year field campaign and discuss their implications for the tectonic evolution of the Levant domain. These new data led us to significantly refine the tectonic history for Lebanon and to address fundamental issues concerning the development of the Eastern Mediterranean basin and later deformation of its eastern margin.

Geological overview of Lebanon Faults and folds The Lebanese sector of the Levant margin exhibits folds at various scales (Fig. 1). The three largest are, from west to east, the NNE– SSW Mount Lebanon anticline (MLA), the NE– SW Bekaa syncline (BS) and the NE –SW Mount Anti-Lebanon anticline (MALA). The MLA can be followed along c. 150 km, from Tripoli in the North to the Israeli border in to the South. Its axis rotates progressively southward from a NNE–SSW to a north–south direction and it shows Lower Jurassic rocks in its core (Fig. 1). In its northern and central parts, it has a steeply dipping western limb (up to 908) that is referred to as the Lebanon flexure. The MALA is a roughly symmetric fold with Middle Jurassic rocks in its core. The BS is filled by poorly dated Neogene continental sediments and its two limbs dip at about 308. Smallerscale folds have been documented locally, with similar trends as the main ones and with WSW – ESE axes in a few places, like along the limbs of the BS. Near Tripoli, these WSW– ESE folds have shorter wavelengths and involve the Neogene sequence. Regional fold-related shortening has been estimated to be about 10 km (Butler et al. 1998). The main fault in Lebanon is the c. 150 km long, NNE–SSW-trending Yammouneh fault (YF) that cuts the eastern flank of the MLA (Fig. 1). Pull-apart basins, such as the Yammouneh pull-apart, link left

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stepping segments of the YF and attests to its left lateral motion. However, only 7– 11 km of left slip have been documented along the YF (Dubertret 1970; Hancock & Atyia 1979; Quennel 1984), a value much smaller than the 107 km suggested for offset along the southern DST (Quennel 1958; Freund et al. 1970). However, the YF is the only through-going structure and links the Jordan Valley fault and Ghab fault, the two major faults of the southern and northern DST. Therefore, it is generally considered as the main fault of the DST in Lebanon. Its Late Pleistocene–Holocene slip rate is estimated at 3.8–6.4 mm a21 (Daeron et al. 2004). The three other important left-lateral faults from west to east are, the north –south Roum fault, the NE–SW Rachaya fault and the Sergaya fault. According to Butler et al. (1998), the Rachaya and Sergaya faults have maximum offsets of a few kilometres (see Walley 1998 for a larger estimate) and are presently inactive. However, Gomez et al. (2003) documented a left-lateral slip rate on the Sergaya fault during the Holocene of 1.4 mm a21. The Roum fault accumulated 8–30 km of leftlateral slip (Walley et al. 1988; Butler et al. 1998) and its mean recent slip rate was 0.9 mm a21 (Nemer & Meghraoui 2006). Another major fault system, referred to as the Mount Lebanon Thrust (MLT), is the east-dipping offshore thrust system first mapped by Elias et al. (2007). It is situated about 25 km west of the Lebanese coast and runs NNE –SSW along the base of the continental slope. Its northern part is interrupted by a lateral ramp the inland continuation of which near Triopli probably connects with the onshore WSW– ESE Akar thrust (Elias et al. 2007). Numerous shorter faults, a few tens of kilometres in length, also exist. Some are confined to the Jurassic to Early Cretaceous cores of the anticlines, whereas others, especially in southern Lebanon, cut the Palaeogene sedimentary sequence. According to Hancock & Atiya (1979), these are minor dextral and sinistral strike –slip faults related to Neogene transform tectonics, although offsets mapped by Dubertret and collaborators (1955) suggest that significant vertical movements are present also.

Age and origin of the tectonic structures in Lebanon According to Walley (1998), the tectonic structures in Lebanon resulted from three main tectonic events: (1) a Late Triassic to Early Jurassic extensional phase associated with the opening of the Levant basin (easternmost Mediterranean) followed by a possible late Early Cretaceous rejuvenation; (2) a multi-stage inversion phase lasting from Late Cretaceous to Late Oligocene time, during which most folds developed driven by the progressive

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closure of the Neo-Tethys ocean to the north; and (3) a Late Cenozoic dominantly strike–slip phase associated with the development of the DST. Other models of the late history have been proposed, however owing to the fact that extensive data could not be collected, they remain poorly constrained. In detail, the early tectonic history of Lebanon is also quite poorly constrained. An Early Mesozoic NNE–SSW structural fabric is generally assumed given the westward thickening of the Albian – Turonian sequences and facies evolution from shallow marine to deep pelagic sediments (Saint-Marc 1974). Owing to the fact that this trend resembles the well documented NNE –SSE to NE–SW faults that developed onshore and offshore Israel during Early Mesozoic time (e.g. Cohen et al. 1990; Gardosh et al. 2006), Walley (1998) concluded that the present-day Mount Lebanon western limb marks the margin of the Levant basin in this area. Although this interpretation is quite coherent with general knowledge of the Levant basin development, it should be noted that normal faults associated with this Early Mesozoic extensional tectonic have not been observed in Lebanon to date. Furthermore, other regional geological features are difficult to interpret within this scheme, including the southward decrease of the Late Jurassic to Early Cretaceous volcanism and the abrupt thickening of the Early Cretaceous sandstones (Dubertret 1975). There is a consensus that the Late Cretaceous to Cenozoic tectonic history of Lebanon began with Syrian Arc tectonism and continued through the Late Cenozoic with the initiation of the DST. Syrian Arc tectonism refers to a major phase of inversion in the southern Levant domain and northern Africa (Israel, Egypt and offshore) where broad folds and high-angle reverse faults are associated with breaks in deposition and angular unconformities (e.g. Bartov et al. 1980; Mimran 1984; Lewy 1991; Moustafa & Khalil 1994; Lu¨ning et al. 1998; Gardosh & Druckman 2006). These structures developed in response to regional WNW –ESE compression (Eyal 1996) and although the age of development is still debated (see Shalar 1994; Walley 1998 for a review), there is a general consensus that the Syrian Arc tectonism includes two main episodes, one during Senonian time and one during Late Eocene to Late Oligocene time (possibly continuing up into Early Miocene time). The curvature of the Syrian Arc structures from an east –west to WSW–ESE belt from Sinai to Negev and their interpretation as reactivated Early Mesozoic structures led workers to extend this belt into Lebanon. However, to date no Syrian Arc structures have been documented there. Using a number of logical arguments, Walley (1998) claimed that the folds in Lebanon could not solely be the result of Cenozoic transpressive tectonics, and must have largely

developed earlier and thus during Syrian Arc compression. Most authors (Freund et al. 1970; Hancook & Atiya 1979; Garfunkel 1981; Walley 1988, 1998; Butler 1998) agree that the Late Cenozoic deformation of Lebanon is related to the development of the DST, which today has a restraining bend rotated approximately 308 clockwise relative to the other north –south DST segments. Although the amount of slip is uncertain, it is widely accepted that the Cenozoic movements include a combination of left-lateral strike –slip faulting and thrusting along north–south to NE faults, as well as folding along these same trends. These faults and folds are now the most prominent tectonic structures in Lebanon (Fig. 1 and see two previous sections for detailed description of tectonic structures). A number of significant unresolved differences exist regarding the evolution and nature of deformation associated with the DST in Lebanon. The first point concerns the exact timing of DST development in Lebanon, which varies from Early Miocene (Quennel 1958; Walley 1988) to Late Miocene (Butler 1998; Walley 1998). The second concerns the importance and age of transpressive deformation, which according to Butler (1998), was already ongoing by Messinian time and transpression thus existed early in the DST history. Alternatively, following the kinematic approach of Garfunkel (1981), Walley (1998) advocated a Late Miocene pure transform motion along the central DST. In this model, strike –slip movement occurred first and folding did not start until Pliocene time. In contrast, Hancock and Atiya (1979) suggested that the YF is younger than the folding. Walley (1998) stressed the importance of Syrian Arc inversion in shaping the Lebanese folds relative to minor Neogene transpression. The final issue is the present-day NNE –SSW trend of the YF, which strikes oblique to the southern and northern segments of the DST. According to Quennel (1984), the DST initiated at the Red Sea Rift and propagated northward along a small circle, with an almost north– south direction. As such, the modern NNE – SSW trend of the YF would result from a later regional clockwise rotation in Lebanon during Pliocene or later time caused by WSW–ENE dextral simple shear across Lebanon and into the Palmyrides. Others claimed that the bend of the YF in Lebanon was a primary feature of the plate boundary architecture (Butler 1998).

Data and method In order to identify the main tectonic events and tectonic structures that have deformed Lebanon since Mesozoic, we combined a field study of the large-scale structural features with a mechanical

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analysis of meso-scale faults. This approach aimed to both constrain the age of the tectonic structures in Lebanon and establish how the style of deformation has evolved since Mesozoic time.

Large scale tectonic structures and dating arguments Available 1:50 000 surface geological maps of Lebanaon show faults with lengths from a few hundred metres to hundreds of kilometres (Fig. 1). In the best cases, correlation between mapped stratigraphic units allow us to determine the apparent fault offsets, however, the true nature of the faults is not fully defined. Therefore, we examined in the field the relationship between faults and offset beds and measured kinematic data on well-preserved fault planes (Fig. 2). The timing of fault movement was evaluated using sediment thickness variations across the faults and the periods of activity were defined using appropriate cross-cutting relationships (Fig. 3). We also examined the continuity of the sedimentologic units, with particular attention paid to the presence of angular unconformities. Across the

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unconformities, bedding attitudes were systematically measured in order to quantify the amount of tectonic tilting. Samples were also collected in the sequences bounding the angular unconformities and their nannofossil contents were examined in order to refine the ages of these sequences. This allowed us to better subdivide the previously undifferentiated Senonian sequence and to clarify relationships within the monotonous chalky Senonian to Palaeogene units. In addition, we also identified previously unknown Oligocene deposits within the study area (Mu¨ller et al. 2010).

Meso-scale faults and inferred stress field In order to determine how the stress field evolved in Lebanon, fault–slip data were collected on mesoscale faults cutting the Mesozoic to Cenozoic sequences with offsets between a few millimetres to a few metres. Approximately 2700 faults were measured across 103 sites, and fault –slip data were collected in small volumes where the stresses could be considered to be homogeneous. The local stress states were inferred from the inversion of faults using the method of Angelier et al. (1990).

Fig. 2. Early Cretaceous normal fault plane cutting the Jezine formation. Site of Kafer Hachno (see fig. 8 for location). Modified from Homberg et al. (2009). The striae or slip vector (thin arrows) is aligned with the maximum dip and indicates a normal movement (thick shear arrow). View is to the N8308E. Meso-scale faults collected in this site and corresponding stress state are shown (lower hemisphere, equal-area projection). Traces: fault planes. Slickenside lineations in dots with double arrows for strike–slip motion and with an outward-directed single arrow for normal motion. White stars with 5, 4, 3 arms: s1, s2 and s3, respectively. Divergent large black arrows show directions of s3.

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Fig. 3. Aptian–Albian growth fault. Modified from Homberg et al. (2009). Inset shows how the growth fault splits upward into two branches, the southern segment making a counterclockwise angle with the single deep fault. Jf, Jezine Formation (latest Early Aptian). See Figure 8 for location.

This method is based upon the Wallace –Bott hypothesis that faults slip in the direction of the resolved shear traction. It allows for determining the orientation of the three principal stresses, s1, s2 and s3, (s1 . s2 . s3, compression positive), as well as a shape ratio between the principal stress differences. The correct agreement of each fault with the calculated stress states is expressed by two misfit estimators, a and RUP. These are the mean angle and vector between the computed and measured striae, respectively, and vary from 08 to 1808 and 0% to 200% when the computed shear stress rotates from a vector with the same to the opposite sense as the measured striae. In general, the uncertainty on the orientations of the stress axes depends on the three-dimensional (3D) distribution of the fault–slip data, so we paid particular attention to measuring faults of variable orientation and systematically searched for conjugate faults in order to tightly constrain the stress orientation. In such cases, the accuracy of stress plunge and direction is about 108. In most of the sites, the fault population included strike–slip, normal, and sometimes reverse faults, indicating that several stress regimes occurred through time. For such heterogeneous fault populations, the inversion process failed to determine a unique stress state with acceptable misfits, thereby confirming the likely superimposition of different

stress states (Fig. 4). Sometimes, mechanical incompatibilities also existed between the same fault type, such as for the strike–slip faults shown in Figure 4. The stress states were computed by performing an inversion on each of the homogeneous subsets. Classification of the fault population into homogeneous subsets is guided by field observation, such as different slips on a same fault strike, overlapping compressive and extensive dihedra defined by each fault, relative chronology between faults, as well as high misfits between actual fault–slips and slips predicted by the computed tensor. The total dataset resulted in the determination of 209 stress tensors: 127 strike –slip regimes; 48 normal regimes; and 34 reverse regimes. The succession of the local superimposed stress states rested on relative chronological data between brittle structures (e.g. cross-cutting relationships between faults or between superposed striae on fault surfaces) or criteria establishing the age of stress state relative to folding. The age of the stress state with respect to folding was obtained from the examination of relations between the stress inclination and the local bedding orientation. Measurement of present-day stresses (e.g. Cornet & Burlet 1992; Brudy et al. 1997) and palaeostress reconstruction (Homberg et al. 2002) support the hypothesis that stresses follow the Anderson model, in which two of the principal stresses are approximately

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(a)′

(a)

Pre-folding Back-tilting

: 094 11 : 248 78 3: 003 05

: 271 18 : 165 40 3: 019 44

1

1

2

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(c)

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Post-folding

: 344 00 : 252 75 3: 074 15

: 178 03 : 269 17 3: 077 73

1

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Fig. 4. Example of meso-scale fault analysis. Site of Aitoun, located in Cenomanian limestones of northern Lebanon. The fault population (Lower hemisphere, equal-area projection) includes strike–slip faults (a) and (b) and reverse faults (c) which were analysed separately. Mechanical incompatibilities between the strike–slip faults (overlapping compressive and extensive dihedra) led to distinguish two fault subsets (a) and (b). (a)0 corresponds to faults of group (a) that have been back-tilted. Stress states responsible for each subset are shown. Direction and inclination of the principal stresses are also indicated. Same legend as for Figure 3, plus inward-directed single arrow for reverse motion. Dashed lines: local bedding planes. Convergent and divergent large black arrows show directions of s1 and s3, respectively. Lower values of misfit were obtained with back-tilted faults [a ¼ 298, RUP ¼ 57% in (a) and a ¼ 188, RUP ¼ 49% in (a)0 ]. D is the mean deflection of the principal stresses relative to the vertical and horizontal. Likewise, lower D angle was obtained in back-tilted state [358 in (a) and 98 in (a)0 ]. All this suggests that faults of group A pre-date folding (see text and Fig. 5 for further explanation). The strike– slip and reverse regimes in (b) and (c) show s1 trending respectively 1648E and 1788E. They both correspond to a NNW– SSE compression that postdates folding.

horizontal, the third one is approximately vertical. If this criterion is fulfilled for principal stresses calculated with rotated faults (rotation around the local bed strike by the amount of tilting), the stress state occurred before folding (Fig. 5). If it is fulfilled with the present attitude of faults, the stress state post-dated folding. Considering the uncertainties on the calculated stress orientations, a reliable dating of faulting relative to folding requires at least a 208 tilt of the beds. According to this logic, we were able to distinguish pre-folding and postfolding stress states in 42 of the 103 studied sites.

The pre-Neogene extensional tectonic events Evidences for early normal faulting Study of the large-scale and meso-scale faults revealed that the Jurassic to Eocene sequences are

cut by numerous normal faults of various lengths. Faults with lengths of several kilometres to several tens of kilometres are particularly numerous in the Meten, Chouf and Jezine areas where they trend WSW –ENE to WNW –ESE (Fig. 1). In these regions, the strata are sub-horizontal and show significant vertical offsets along these faults (Fig. 6). Observation of fault planes, when visible (c. 50% of the faults), revealed that the striae inclination is systematically close to 908, indicating that these faults are dip–slip normal faults (Fig. 2). As far as the small scale brittle structures are concerned, normal faults represent 25% of the measured population and indicate that part of the brittle deformation in Lebanon resulted from an extensional stress regime. Several lines suggest that most of these normal faults are early features. First, these faults were found in rocks of Middle Jurassic to Middle Eocene age and they are largely absent within the Neogene units. Second, out of the 43

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(a)

(b)

Fig. 5. Folding–faulting chronology. Example of site of Qsaibe located in Eocene chalks of southern Lebanon. Normal faults measured in this site (and related stress state) are shown in their present-day position in (a) and after rotation around the local bed strike of the amount of tilting (back-tilting) in (b). Same legend as in Figure 4. Direction and inclination of the principal stresses are indicated. a and RUP are the two misfit estimators of the stress inversion (good agreement of the faults with the calculated stress state for low values) and D is the mean deflection of the principal stresses relative to the vertical and horizontal (good agreement with the Anderson model for low value). Lower a, RUP and D values are obtained with inversion performed on back-tilted faults. The inferred normal stress regime therefore predates folding. See also Figure 4 for an example of strike– slip regime.

sites affected by the meso-scale normal faults, the faulting–folding chronology could be established in 10 localities and (with one exception) normal faults always pre-date folding (see Figs 5, 8–10). In order to define the number and duration of the tectonic events that produced these faults, the whole Mesozoic –Cenozoic sedimentary succession was examined in detail and two extensional events were identified.

Early Cretaceous extension Examination of the Mesozoic units exposed in the cores of the Mount Lebanon and Anti-Lebanon

anticlines led us to identify a major extensional event during Early Cretaceous time. In the central part of Mount Lebanon and in the southern AntiLebanon, good outcrops allowed us to document normal growth faults within the Early Cretaceous sequence. The oldest series where such faults were found are the Chouf sandstones (Fig. 7). In this sequence, we observed growth faults with offsets of several metres. According to Dubertret (1975), these fluvial deposits are of Neocomian– Barremian age, but recent investigations suggested that the first strata of this sequence post-dates Early Valanginian time (Ferry, personal communication). No growth faults were recognized in the Middle to Late Jurassic levels. However, such structures may be difficult to document in this poorly bedded sequence. While we do not exclude normal faulting during Late Jurassic time, we believe that significant vertical movements did not start before Early Cretaceous time. The widespread Kimmeridgian basalts and tuffs (see Collin et al. 2010) attest to a regional volcanic event that immediately predates the intense Early Cretaceous faulting. Faultcontrolled thickness variations are also common in the Aptian to Albian sequences of Mount Lebanon, and in particular in the southern Chouf (south of Der al Qamar) and Jezine areas (Figs 3 & 6). Here, the fault offsets range from several tens of metres to hundreds of metres. Together these observations indicate that normal faults developed during over a long period from Valanginian to Albian time. When outcrop conditions allowed observation of the entire Early to mid-Cretaceous sequence, the Cenomanian beds sealed most of the normal faults, as illustrated in Figures 3 and 6. A few faults cut basal Cenomanian strata that generally forms the top of the cross-sections. According to the geological maps of Dubertret, these faults do not extend for more than a few hundred metres from the Albian –Cenomanian boundary. Saint-Marc (1970) showed that the first Cenomanian levels are in fact Late Albian in age. Therefore, the major episode of Early Cretaceous normal faulting ended just prior to Cenomanian time. Our observations in the younger series confirm tectonic quiescence through the remainder of Late Cretaceous time. The growth faults described above have a similar orientation whatever the age of the series they cut. They typically strike between WNW– ESE to WSW–ENE and have a mean dip of 608, either to the North or to the South (Figs 2, 3, 6 & 7). This suggests that the tectonic process that produced these faults did not vary with time and reflects a long-lived extensional phase that started during Valanginian time and ceased at the beginning of Cenomanian time. The stress calculation obtained on fault–slip data measured in Jurassic and Lower Cretaceous rocks indicates that this tectonic event

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Fig. 6. Normal faults cutting the Jurassic to Early Cretaceous sequences. Jezine region. The latest Albian beds seal the normal faults. Jf: Jezine formation. See Figure 8 for location.

was characterized by a north–south to NNE–SSW extension, which varied slightly in direction from one locality to another (Fig. 8). The stress field shown in Figure 8 was obtained from various faults, with or without chronological data. In four

Fig. 7. Growth faults in the Chouf sandstones. Outcrop along the western limb of the southern Anti-Lebanon anticline. See Figure 8 for location.

sites, the fault data were measured on growth faults. Because faulting generally implies a variety of fracture scales, neighbouring meso-scale faults were also included in the inversion as they likely reflect the same mechanism as the large-scale faulting. In theses cases, there was no ambiguity on the age of the fault –slip and related stress tensors. Fault inversion in these four sites yielded normal stress tensors in which s2 and s3 were subhorizontal and s1 was sub-vertical. The direction of extension trends between N1698E and N0648E (azimuthal range is described in a clockwise sense). While this range is rather large, it is consistent with a mean NNE–SSW extension direction. A similar direction was recognized in 13 other sites situated in the Jurassic to Lower Cretaceous sequences. Here, chronological arguments are absent, precluding the establishment of the absolute age of the normal faulting. Nevertheless, the consistent direction of s3 (direction of extension) led us to attribute the calculated stress states in the remaining 13 sites to the Early Cretaceous extension, although we cannot firmly exclude that some of them may reflect a later tectonic event. Notably, in two sites where the beds have a significant inclination, the stress states pre-date folding. Considering the 17 stress states together, 71% of them have a direction trending between N1608E and N0408E (Fig. 8) and are compatible with the strike of the

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Fig. 8. Stress field during the Early Cretaceous extension in Lebanon. Bars indicate the direction of extension (s3) obtained from fault–slip data inversion. Examples of fault– slip data and stress calculation are shown. When faulting– folding chronology could be established, the stress state systematically predates folding (one exception). The rose diagram of s3 directions is shown. N is number of stress states. The mean direction of the Early Cretaceous extension trends north– south to NNE– SSW, with local stress deflections. F2, F3, F6 and F7 indicate location of Figures 2, 3, 6 and 7. Same legend as in Figures 1 and 2.

large-scale WSW– ENE to WNW–ESE growth faults. Therefore, we conclude that the Early Cretaceous normal faults formed owing to a NNE –SSW extension, with minor local deflections in the stress direction.

The Early Cretaceous extension is of significance in Lebanon as it generated a structural grain and is responsible for large-scale thickness variations within the Lower Cretaceous sedimentary sequence. In the southern Chouf and Jezine areas,

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the Early Cretaceous normal slip along the numerous and several kilometres long WSW– ENE to WNW– ESE faults is unambiguous thanks to observations of both fault striae and fault-related thickness variations within the sequences. In other areas, like in northern Mount Lebanon and Mount Anti-Lebanon, these faults show an apparent normal movement. Their age could not be accurately established but several lines suggest that most of the WSW–ENE to WNW –ESE faults in Lebanon (with exception of those of southern Lebanon that cut the Palaeogene sequences) slipped during the Early Cretaceous extensional phase. First, these faults form a consistent fault system over the entire country with a combination of WSW– ENE and WNW– ESE trends. Second, most of them are restricted to the Jurassic –Lower Cretaceous sequences and do not cut younger formations. Third, small-scale normal growth faults (with offsets of a few tens of metres) have similar trends as the larger ones in many places in Lebanon. These factors suggest that most of the WSW – ENE to WNW –ESE faults (faults of southern Lebanon excluded) are Early Cretaceous in age. At the regional scale, the Lower Cretaceous sequence exhibits a progressive northward thinning. It is particularly spectacular for the Chouf sandstones, thickness of which reaches 300 m in central Lebanon (Chouf area), and is reduced to a few tens of metres in northern Lebanon. This regional trend underlines the development of a roughly WNW –ESE striking basin during Early Cretaceous time in Lebanon, a direction consistent with the mean east –west trend of normal faults and with the north –south to NNE–SSW direction of extension inferred from our fault –slip data inversion. We estimate that the northern margin of the basin was situated c. 50 km to the south of Tripoli, where fluvial and shallow marine sedimentation was strongly reduced and sometimes replaced by lavas flows. The southern margin is not visible owing to later burial beneath Palaeogene deposits in southern Lebanon.

Eocene extension Our observations of the brittle deformation in the Mesozoic–Cenozoic sedimentary sequence revealed that the Late Cretaceous to Eocene strata are cut by numerous meso-scale normal faults with consistent east– west trends. These faults comprise 17% of the total data. Among the 24 sites where such faults were documented, nine cut Cenomanian – Turonian limestones, two offset Senonian marls and 13 affect Eocene chalks. Normal faults are almost absent in Neogene rocks suggesting that these post-Cenomanian faults pre-date Miocene time. Indeed, in the eight sites where the significant

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dip of the bedding allows definition of a faulting – folding chronology, the fault inversion indicates that these brittle features systematically pre-date folding (Fig. 5). This is true for all the localities where these faults were found, that is, all over the country (Fig. 9). Because we believe that folding in Lebanon started in Early Miocene time (see following section), the normal faults discussed above are pre-Neogene features. We examined the entire Late Cretaceous to Palaeogene sedimentary succession for normal growth faults in order to define the duration of extension. Such faults were documented in the Eocene units that are well exposed in the southern part of the country where uplift and later erosion was moderate (Fig. 1). Fresh outcrops are, however, rare and are generally situated in quarries or along road cuts. The very best cross-sections are found along the coastal highway where we documented normal growth faults with offset of several metres (Fig. 10). All growth faults recognized in the Middle Eocene sequence showed a consistent east – west mean strike. No growth faults were found in the Senonian or Palaeocene formations, suggesting no tectonic activity during this period. However, it may be difficult to recognize syn-tectonic movements within this monotonous chalky and marly interval. Nevertheless, we suggest that, in Lebanon, normal faulting after the important Early Cretaceous extension did not re-commence before Eocene time. Revision of the ages of the Palaeogene series in Lebanon (Mu¨ller et al. 2010) revealed previously unknown Late Oligocene deposits, however, crosssections in this sequence are not big enough to observe growth faults so we could not establish directly if the normal faulting that started during Eocene time continued into Oligocene time. At three sites, Eocene growth faults bear wellpreserved striae and extensive fault –slip data could be collected. Inversion of these data allowed the determination of the stresses that drove the Eocene faulting, which correspond to a normal stress regime in which the direction of extension (s3) trends NNE–SSW (Fig. 10). Meso-scale normal faults cutting Cenomanian to Eocene rocks in 21 other sites indicate the same s3 orientation (Fig. 9). Their systematic pre-folding age and the likely absence of fault development during the Senonian led us to attribute them to the Eocene extensional phase. In detail, the Eocene stress states show s3 directions varying between N1418E and N0488E (azimuthal range described in a clockwise sense) and for a large majority of them (63%), s3 strikes between N1708E and N0308E. The major peak trends between N0108E and N0208E (26% of the stress states). It should be noted here that this direction is close to that of the Early Cretaceous extension. This similarity makes it difficult to determine

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Fig. 9. Stress field during the Eocene extension in Lebanon. Bars indicate the direction of extension (s3) obtained from fault– slip data inversion. Examples of fault –slip data and stress calculation are shown. When faulting– folding chronology could be established, the stress state systematically predates folding. The rose diagram of the s3 directions is shown. N is number of stress states. The mean direction of the Eocene extension trends NNE–SSW. F 10 indicates location of Figure 10. Same legend as in Figures 1 and 8.

which of these two tectonic events produced the normal faults collected in pre-Late Cretaceous rocks where syn-tectonic arguments are absent. In other words, some of the stress states shown in Figure 8 could reflect either the Early Cretaceous or the Eocene stress fields. Stress states shown in

Figure 9 are, however, clearly Eocene in age because they were established on the basis of fault–slip data collected in Late Cretaceous to Eocene rocks. Although small-scale brittle deformation developed all over Lebanon during the Eocene extension, large-scale normal faulting is not obvious.

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The Cenozoic deformation of the Levant Margin Evidence for poly-phase Neogene tectonism

Fig. 10. Growth fault in the Middle Eocene sequence. Outcrop in southern Lebanon (see Fig. 9 for location). Note that unit A is thicker in the hanging wall block. Fault– slip data collected in this site indicate a normal stress regime with a NNE– SSW direction of extension. Same legend as in Figure 2.

Documenting such large-scale faults is difficult because the Palaeogene and Neogene series have been largely eroded in Lebanon so that crosscutting geometries are not preserved. A set of NE–SW to WNW– ESE fault with length of several tens of kilometres exist in southern Lebanon, which juxtapose different Late Cretaceous rocks and sometimes Late Cretaceous and Eocene rocks. The sub-horizontal attitude of these sequences indicates a significant vertical component of the fault movement, which was confirmed for two faults by observation of well-preserved sub-vertical and normal striae. Late activity appears very likely along the few faults that have a modern topographic expression, however we do not exclude an earlier Eocene movement. At the regional scale, the Eocene sequence evolves from chalky marine deposits in the coastal zone to shallow marine carbonate facies inland and no abrupt north –south thickness variation have been identified to date in these series. Bedding is often chaotic, suggesting that slumping occurred even in the eastern-most part of the country. Such apparent instability could be controlled by fault movement, but may also reflect the regional deepening of the platform towards the Levant basin. Whatever the importance of the Eocene vertical movements, faulting during Eocene time is no doubt related to a extensional event across all of Lebanon. This same event is also observed in Syria (Al Abdalla et al. 2010) and its regional significance will be discussed later.

After the dominant extensional tectonism lasting from Early Cretaceous to Eocene time (and maybe even until Oligocene time), the Neogene period marks a dramatic change in the tectonic evolution of Lebanon, and in general cross the entire Levant domain. This latter period was dominated by strike –slip and compressional deformation related to the interaction of the Africian, Arabian and Eurasian plates. Meso-scale brittle structures observed in Lebanon are dominantly strike –slip and reverse faults (86 of the 103 visited localities), and represent the vast majority (75%) of collected data. These faults were found in rocks of Jurassic to Middle Miocene age, but no fault–slip data could be measured extensively in the relatively incompetent and incohesive Pliocene and younger units. Inversion of these fault data yielded the determination of 161 stress tensors, with strike–slip regimes dominant (127 of the 161 stress states) over reverse regimes (the remaining 34 stress states). A few normal faults that post-dated folding were also found but they represent a very minor part of the brittle deformation in Lebanon. As they were recognized at only few places, they likely reflect relative minor local deformation and will not be discussed further here. Analysis of the fault– slip data indicates that the strike –slip and reverse meso-scale faults resulted from several stress fields. Indeed, in 50 of the 86 sites showing strike –slip and reverse faults, the inversion process failed to explain the fault population with a single stress tensor (Fig. 4). Several deformation periods must, therefore, be proposed for these slips. This poly-phase origin is obvious considering the azimuthal distribution of the strike –slip faults observed (Fig. 11a). Of the leftlateral strike –slip faults 86% trend between N1108E and N0408E and 85% of the right-lateral strike –slip faults have trends between N0608E and N1408E (azimuthal range described in a clockwise sense). Opposite slips on N1108E to N1408E faults were thus frequently observed suggesting a large variation in the direction of s1. The distribution of s1 directions show two main groups: the first (34% of the data) with s1 trending between N0708E and N1108E and a second one (42% of the data) with s1 striking between N1408E and N1708E (strike– slip and reverse regimes taken together) (Fig. 11b). Therefore, brittle faulting possibly resulted from two compressions, which strike NNW –SSE and east –west, respectively. Several lines of evidence suggest that the east – west compression pre-dates the NNW– SSE one.

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N

First, WNW–ESE to NW–SE fault planes often show sinistral striae overprinted by dextral striae. Second, stress states with s1 trending between N0708E and N1208E (implying a mean east –west compression) form the majority of the pre-folding stress states (56%), but they are poorly represented in the post-folding data (17%) (Fig. 11c). Finally, in the 13 sites where the two compressions are superposed and where the faulting –folding chronology could be established, the east –west compression systematically pre-dates folding. Conversely, the NNW– SSE compression shows an inconsistent temporal relationship with respect to folding. Together, these observations indicate that the principal stress directions rotated clockwise during Late Cenozoic time, from an east –west to a NNW–SSE compression, and this rotation occurred before the major folding episode. It follows that the mechanism of deformation evolved from Neogene to Present time. This period included several tectonic events defined on the basis of their durations and different palaeo-stress orientations.

N

N

Early Miocene compression

N

N

Fig. 11. Strike– slip and reverse faults with associated stress states. A: rose diagram of fault strike for left-lateral (A1) and right lateral (A2) faults. Note the presence of both left- and right-lateral faults in the ESE directions. B: rose diagram of s1 direction for strike–slip and reverse regimes. Note two separate peaks. C: rose diagram of s1 direction for pre-folding (C1) and post-folding (C2) strike– slip and reverse regimes. Folding occurred after shift from east–west to NNW– SSE compression (see text). N indicates number of data.

Study of the angular unconformities within the Neogene sequence and dating of the sedimentary sequences indicate that an initial folding episode occurred during Early Miocene time. According to the geological maps, the Neogene units in Lebanon rest on formations ranging in age from Senonian (undifferentiated) to Eocene (Fig. 1). The Neogene sequence is preserved in a few areas along the coastal zone and a Vindobonian age (Langhian to Tortonian) has been attributed to these shallow water marine sediments (Dubertret 1975). It also outcrops along the Bekaa syncline and consists of poorly dated continental marls and conglomerates. In general, good outcrop conditions in Lebanon allow observation of the internal bedding relationships, but the contact between the Neogene and older units is rarely seen. In four outcrops (UA1, UA2, UA3 and UA4, see Fig. 13 for location), we observed the Miocene package resting on older formations above a minor angular unconformity (Fig. 12). The angle between the two sequences is typically less than 108. We refer to this unconformity herein as Unconformity A (UA). Three of the outcrops described above are situated in the Neogene marine units along the Lebanese coast near to Tripoli, Ras Checka and Saida, while the fourth is located on the eastern flank of the Bekaa syncline, near to Zahle. In the three first outcrops, we could define the age of the units on both sides of UA, whereas in the last one we could only date the underlying sequence. The basal levels above UA are of Langhian age whereas the units below UA range in age from Late Campanian to

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Fig. 12. First Neogene unconformity (UA). (a) outcrop along the Lebanese coast at Raas Cheeka. Miocene shallow water carbonates overlay a chalky Lower Eocene sequence. Note the moderate angle between the two units. UA is marked by thick line. (b) and (c) cross-sections (modified from Dubertret and collaborators 1955) showing the unconformity of the continental Neogene formations along the western and eastern flanks of the Bekaa syncline, respectively. 1, Jurassic; 2, Neocomian; 3, Aptian; 4, Albian; 5, Cenomanian; 6, Turonian; 7, Senonian; 8, Eocene; 9, Continental Neogene; 10, Quaternary. Cenomanian and Turonian are not distinguished in Figure 12b. See Figure 13 for location and in text for discussion.

Middle Eocene. Therefore, the tectonic event responsible for UA occurred between Oligocene and Early Miocene time. Previously unknown Late Oligocene deposits were identified in western Lebanon (Mu¨ller et al. 2010) but their relationship to the regional Neogene unconformity is unclear. In one area in southern Lebanon (UA5), we could clearly observe the sub-horiztonal, erosive contact between Late Oligocene chalky deposits and Middle Miocene units (age from Dubertret). Since no angular unconformity was recognized within the Eocene –Late Oligocene chalks, this erosional surface likely marks the end of the tectonic event that produced UA. Therefore, the tectonic event responsible for the regional Miocene unconformity UA likely occurred during Early Miocene time. It should, however, be noted that the age of UA was define mainly in western Lebanon and in the Bekaa syncline, a Middle Eocene age was obtained for the series below UA, but our investigation could not accurately define the age of the undifferentiated Neogene continental units above UA. It is possible, therefore, that the UA unconformity in eastern Lebanon could be of a slightly different age, but without strong evidence for diachronism we infer the same Early Miocene age prevailed all across Lebanon. At the regional scale, the Middle Miocene and later sequences are either sub-horizontal or dip in

the same direction as the older sequences (Fig. 1). In consequence, the post-Early Miocene tectonism did not reverse the dip direction of the units and the present-day attitude of the pre-Miocene beds correctly reflects the regional dip direction at Early Miocene time. From west to east, the preMiocene sequence dips westward along the Lebanese coast (Fig. 12a), eastward along the western limb of the BS (Fig. 12b) and westward along its eastern limb (Fig. 12c) defining a succession of large folds. Given that chronological criteria suggests that the east– west compression is the oldest event, it likely characterizes the first Neogene episode of deformation. The Early Miocene stress field (Fig. 13) was constructed by selecting stress states in the dataset showing s1 close to east –west and using the local stress succession established previously. Post-folding east –west stress states determined in significantly inclined layers were excluded because folding during the Early Miocene was moderate. The east–west compression was recognized in 50 of the 103 studied sites widely distributed over the entire country (Fig. 13). It is expressed by 47 strike –slip and 13 reverse regimes where the s2 and s3 are respectively horizontal and vertical. In 9 sites, the strike– slip and reverse regimes were recorded together and the compared s1 directions make an angle of 208, a value slightly above the

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Fig. 13. Stress field during the Early Miocene tectonics in Lebanon. Bars indicate the direction of compression (s1) obtained from fault–slip data inversion. Examples of fault– slip data and stress calculation are shown. Most of the calculated stress states predate folding. The mean direction of the Early Miocene compression trends east– west. White squares indicate sampling location for dating of Unconformity A (UA). See in text for further explanation. Same legend as in Figures 1 and 4.

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uncertainty on the stress orientation. The similar orientation of s1 suggests that the strike–slip and reverse regimes are linked through a switch in the s2 and s3 principal stresses, and we interpret these two regimes within a single tectonic phase characterized by east– west compression. In detail, the Early Miocene compression varies between N0518E and N1268E (defined in a clockwise sense) and the vast majority of these stress states (85% of the data) have s1 trends between N0808E and N1108E. Because the east –west compression was found in rocks as young as Langhian, this stress field prevailed at least until the beginning of Middle Miocene time. We stress here that the east –west compression could not drive sinistral slip on the Yammouneh fault and suggest that the Yammouneh fault did not exist during Langhian time.

Late Miocene compression The local dip of the Middle Miocene units (Fig. 1) indicates that folding occurred after their deposition. In order to date this second folding, we examined the geometrical relations within the Neogene and later sequences and collected samples for biostratigraphical dating. No Late Miocene rocks have been found in Lebanon suggesting the nondeposition or subsequent erosion of these units. Pliocene sediments are preserved in only a few areas. The largest outcrop is located in the northern part of the country, near to Tripoli (Fig. 1) and ages obtained on these units indicated an Early Pliocene age. According to the geological map of Dubertret, these rest on undifferentiated Miocene beds, however unfortunately, the geometrical relationships between these two sequences could not be observed in this lowland area. Previously un-defined Early Pliocene deposits were documented north of Beirut in the Keb valley, on the western limb of the Mount Lebanon anticline (Fig. 14). Here, the flat lying Early Pliocene beds overlay the 508 westdipping Cretaceous units defining a spectacular unconformity, referred to herein as Unconformity B (UB). To the west, the internal bedding of the Palaeocene –Early Eocene sequences (ages from this study) is not well preserved, but this sequence seems to flatten westward and is overlain through UA by Vindobonian reef formations that dip gently to the west (age from Dubertret). Because, the youngest age obtained on the Vindobonian rocks of Lebanon is Langhian, the UB unconformity post-dates early Middle Miocene time but must predates Early Pliocene time. UB defines a folding episode that occurred during late Middle to Late Miocene time and is referred to below as the Late Miocene tectonic event. In the Keb valley, sealing of the western limb of the MLA by Early Pliocene rocks (Fig. 14) leaves no

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doubt that this anticline is a pre-Pliocene structure. Because folding in Lebanon started in Early Miocene time but was moderate during this first tectonic event (see previous subsection), we suggest that the MLA was mostly formed during the Late Miocene tectonic episode. Because Late Cenozoic units in eastern Lebanon are continental and poorly dated, the timing of folding in this sector could not have been tightly constrained. However, similar geometries in western and eastern Lebanon suggest that fold development was driven by similar factors across the entire country. Indeed, a second major unconformity (also referred to as UB) was recognized along the eastern flank of the Bekaa syncline. It separates flat-lying, un-dated and poorly rounded conglomerates from underlying Neogene conglomerates dipping to the west at 308. This observation indicates that the Bekaa syncline and MALAs are not recent structures. Furthermore, the angle between the sequences below and above the unconformities is much larger for UB than UA suggesting the Late Miocene tectonic episode is also a more important event here. The second generation of strike– slip and reverse meso-scale faults indicates a NNW –SSE compression (Fig. 4), which makes a c. 708 azimuth difference with the Early Miocene compression (compare Figs 13 & 15). The NNW –SSE compression (Fig. 15) was recognized in 72 of the 86 sites cut by strike –slip and reverse faults and represents 66% of such faults and 50% of the meso-scale structures in Lebanon. Because the NNW –SSE compression post-dates the east–west one, we associated it with the major Late Miocene tectonism. This correlation is supported by faulting –folding chronologies in highly dipping strata where the NNW –SSE compression both predates and post-dates folding while the east –west compression systematically pre-dates folding. Where bed rotation clearly occurred during the Late Miocene tectonism, the compression that drove folding trended NNW–SSE. Just as for the Early Miocene event, the Late Miocene stress field includes both reverse and strike–slip regimes, with strike –slip regimes dominating (80 strike–slip versus 21 reverse). These regimes were recognized together in 17 of the 72 sites where the NNW– SSE compression was documented and they show a similar orientation in s1. They are, therefore, likely linked through a switch between the s2 and s3 principal stresses within the same Late Miocene event. In detail, the Late Miocene stress states have s1 trends between N1188E and N0218E (defined in a clockwise sense), but for the very large majority, fluctuation in the stress direction does not exceed 508 (85% of the stress state have s1 trends between N1208E and N1708E) and the most common trends are between N1408E

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Fig. 14. Second Neogene unconformity (UB). (a) Detailed view of Keb valley outcrop. The flat-lying Lower Pliocene sequence rests unconformably on the 508 west dipping Late Cretaceous sequence. This unconformity (UB) is marked by a thick line. The Mount Lebanon anticline is thus a pre-Pliocene structure. (b) Larger cross-section through the Keb valley (modified from Dubertret and collaborators 1955). In the original section, the Pliocene formations were thought to be Vindobonian in age. Basis of the Pliocene sediments according to our observations and dating. 1, Lower Pliocene (this study); 2, Vindobonian (Langhian to Tortonian); 3, Eocene to Senonian; 4, Turonian; 5, Cenomanian; 6, Albian; 7, Aptian. See Figure 15 for location and in text for discussion.

and N1708E (66% of the data). Moving eastward, the stress field shows a regional and moderate counterclockwise rotation of s1 from a mean NNW–SSE direction to a mean NW–SE direction.

This is in good agreement with the observed change in the direction of outcropping folds axes. Although brittle deformation was driven by a combination of strike–slip and reverse regimes

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Fig. 15. Stress field during the Late Miocene tectonics in Lebanon. Bars indicate the direction of compression (s1) obtained from fault–slip data inversion. Examples of fault –slip data and stress calculation are shown. The mean direction of the Late Miocene compression trends NNW–SSE. White squares indicate sampling location for dating of Unconformity B (UB). See in text for further explanation. Same legend as in Figures 1 and 4.

during the Neogene, the stress rotation from a east – west to NNW–SSE direction indicates a major change during Late Miocene time, which we interpret as a re-organization of the plate interaction in the Near-East. Considering that the NNW–SSE compression is compatible with YF kinematics, this re-organization likely corresponds to the

initiation of the DST development in Lebanon. If this is correct, the large-scale folding in Late Miocene time suggests an early transpressional character to the plate boundary and implies that the bend of the YF in Lebanon was primary and not the result of later rotation. This conclusion is also supported by palaeomagnetic data (Henry et al. 2010).

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Pliocene and later deformation Documenting Pliocene deformation in Lebanon faces the difficulty that Late Cenozoic rocks are rarely preserved here. This period was not investigated in this study, but comparison between Neogene deformation and recently published data on the Pleistocene-Recent tectonics allow to us to investigate the late structural evolution of the DST. Two aspects of sub-recent deformation are relevant in this discussion. First, sinistral strike– slip motion occurred during Pleistocene and Holocene times on several north–south to NE–SW faults, such as the Sergaya and Roum faults (Gomez et al. 2003; Nemer & Meghraoui 2006), but the 6.4 mm a21 slip rate along the YF (Daeron et al. 2004) suggests that this fault is the major plate structure of the modern DST in Lebanon. Second, folding and faulting occurred offshore along a major NNE –SSW thrust system, the Mount Lebanon thrust (Fig. 1). Together these indicate that the transpressive character of the DST in Lebanon persists until the present day. However, our data provided compelling evidence that significant regional folding did not occurred onshore after Early Pliocene time. This suggests that the faulted and folded belt in Lebanon evolved such that shortening migrated out of onshore Lebanon to the west. Variations in shoreline elevation suggest that the MLT played an important role in Late Holocene uplift of the coastal area (Mohrange et al. 2006). Similarly, a genetic relation between the post-Early Pliocene folding near the Tripoli area, and the eastward continuation of the MLT seems likely. The late stress field in Lebanon is difficult to constrain because meso-scale faults are poorly preserved in the non-cohesive Pliocene sediments and in the basalts. However, post-folding faults indicate that the NNW–SSE compression persisted after Late Miocene time. A significant stress re-orientation following initiation of the DST appears unlikely because it would have produced faults inconsistent with the regional NNW –SSE compression. No such faults were recognized although we paid attention to measure all brittle data whatever their orientation. Therefore, the Pliocene (and maybe later) stress field should be characterized by a general NNW–SSE compression.

Discussion and conclusion Study of tectonic structures in Lebanon and mechanical analysis of the meso-scale brittle deformation allows us to document the major tectonic events that deformed Lebanon since Late Mesozoic time and to determine how the stress field evolved during this period. The first major event is an extensional

tectonic phase, which started in Valanginian (or slightly later) time and ceased at the beginning of Cenomanian time. This NNE– SSW Early Cretaceous extension produced WSW–ENE to WNW – ESE normal faults with offsets of several tens to several hundreds of metres. A second extension with similar driving stresses occurred during Eocene and persisted maybe until Oligocene time. The Early Neogene period marks the first significant change in the structural evolution of Lebanon after which strike –slip and reverse faulting and folding dominated. Moderate folding and faulting occurred during Early Miocene time in response to an east –west compression. A second re-organization occurred during Late Miocene time (or maybe as soon as late Middle Miocene time), since which deformation in Lebanon was driven by NNW– SSE compression. The Lebanese folds developed during this last period and were in large-part formed before the Pliocene. Several lines of evidences suggest that the Early Cretaceous tectonic event that we have documented in Lebanon has a regional significance throughout the Middle East. First, major tectonic elements developed in the Levant domain during Early Cretaceous time, and in Lebanon, these elements are a WNW –ESE-trending basin cut by WSW–ENE to WNW –ESE normal faults with length of several tens of kilometres. The depocentre of this basin was situated in the Chouf area where the Lower Cretaceous sedimentary sequence reached it maximum thickness, and then thins northwards where fluvial and shallow marine sedimentation dimishes and is replaced by lavas flows. While the southern margin of this basin is not visible owing to burial beneath later Palaeogene deposits, the much thinner Lower Cretaceous sequence in Israel suggests the southern margin of the Lebanese basin should be located in southern Lebanon. The Lower Cretaceous sequence also thickens to the west suggesting that the WNW– ESE Lebanese basin extends into the offshore Levant basin (LB), although its structural continuity may have been slightly disrupted by offshore Cenozoic faulting. The acme of faulting in Lebanon was during Early Cretaceous time, but moderate movements during Late Kimmeridgian and Tithonian times are also suspected (Collin et al. 2010). Similar extensional structures have also been recognized in offshore Sinai, they are ENE –WSW half grabens with expanded thicknesses of Late Jurassic to Lower Cretaceous strata (Yousef et al. 2010). Basins of the same age also exist onshore in Egypt, both in Northern Sinai and the Western Desert (Moustafa 2010). The cessation of fault activity is marked by a strong unconformity (Camera et al. 2010). In addition to the areas mentioned above, such faults have been recognized in Israel (Lang & Mimran

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1985), in the Palmyrides (Chaimov et al. 1992) and along the Syrian coastal range (Al Abdalla et al. 2010). In these areas, fault movements were moderate, but consistent with a regional extensional event that might also have be responsible for the Early Cretaceous volcanic activity recognized all around the Levant Margin (with precursor events also in the Late Jurassic; Bartov et al. 1980; Lang & Mimran 1985). Following the widely accepted view that the Levant Basin formed as the result of Tethyan rifting during Mesozoic time (Freund et al. 1975; Ben-Avraham 1989; Garfunkel 1998), the Early Cretaceous event likely corresponds to one of these rifting episodes. Well documented NE –SW faulting and grabens in onshore and offshore Israel attest of a previous rifting event that lasted from mid-Triassic to Liassic time (Cohen et al 1990; Gardosh & Druckman 2006). While the earliest history of the basin is poorly constrained, Permian extension is known in the Palmyrides (Brew et al. 2001). The mid-Triassic –Liassic faulting extended along a NE–SW-trend from the SE corner to the Levant Basin in Israel to the Palmyrides rift (the latter being now situated farther north as the result of Cenozoic sinistral movement along the DST (Garfunkel 1998). Unfortunately, we could not find direct evidences of mid-Triassic– Liassic (or older) tectonics structures in Lebanon because sequences of this age rarely outcrop here. Superimposed trends, however, exist in Lebanon where Early Cretaceous faulting occurred along both WSW–ENE and WNW –ESE faults. The WSW– ENE faults likely reflect an earlier structural fabric that is well-developed to the east in Syria within the Permian to Early Mesozoic Palmyride Trough (Brew et al. 2001) and thus associated with the first rifting event. In Lebanon, the variability of the local extension direction during Early Cretaceous time is probably related to this complexity in fault orientations, which induced local deflections of the regional NNE–SSW extension. The relative importance of each of these rifting episodes cannot be constrained with the available data, but the mid-Triassic –Liassic and Early Cretaceous rifting events were both important phases in the growth history of the Levant Basin. The directions of extension were almost perpendicular, trending NW–SE during the first episode (Garfunkel 2004; Gardosh & Druckman 2006) and NNE–SSW during the second (this study; Yousef et al. 2010). Importantly, the WNW –ESE Early Cretaceous belt was composed of at least two segments, a southern one along the Egyptian coast and a northern one within Lebanon. The second tectonic event seen in Lebanon is a NNE–SSW extension active at least during Middle Eocene time. Enigmatically, this period is

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generally regarded as time of regional compression driving the later phase of Syrian arc inversion (Eyal & Reches 1983; Chaimov et al. 1992; Ponikarov et al. 1996; Walley 1998). Our data and the data of others (Al Abdalla et al. 2010) do not support this hypothesis. Increasing evidence of Eocene normal faulting in the Levant area, however, suggests a phase of regional extension during Eocene and maybe even Oligocene time. Similar east –west-trending Eocene normal faults have now been recognized in the Eocene units of NW Syria (Al Abdalla et al. 2010), in Israel (Hardy et al. accepted) and in northern Egypt (Mostafa 1999). The youngest levels cut by faults associated with this NNE–SSW extension are Oligocene in age located in Syria (Al Abdalla et al. 2010) suggesting that, at least, extension continued into Oligocene time. These observations suggest that following the Late Cretaceous inversion, compressive deformation did not return before Early Miocene time. In this revised context, we re-interpret the high and low areas inferred from thickness and facies variation of Eo-Oligocene sequences as not being evidence of Eocene folding, but suggest instead that they relate to extensional basins, such as those seen in offshore Israel along the Carmel fault (Schattner et al. 2006b). Here, a NW–SE half graben (the Haifa basin) developed during Late Oligocene time (and continued later) with additional minor faulting along east –west trends. While the origin of this Eo-Oligocene tectonism is not well understood, we propose that it could be related to increasing slab pull of the subducting African –Arabian slab prior to collision. Considering the possible block rotation associated with the later Cenozoic tectonism, it is important to stress that the Mesozoic to Palaeogene structures discussed above were described using their present-day trends and orientations. In addition, the stress maps and structures herein were reported in their present-day plate configuration and not in their palaeogeographic positions with rotation during plate motion not taken into account. In Lebanon, recent palaeomagnetic data indicate that about 288 of counterclockwise rotation took place after Albian time (probably occurring during Neogene –Recent time; Henry et al. 2010). This rotation, therefore, affected the preNeogene structures but it did not modify the angular relationships between the early tectonic elements because Lebanon behaved as a single rigid block during rotation. Subtracting this counterclockwise rotation suggests that the Early Cretaceous and Eocene extension actually trended NE –SW. Neogene time marks a significant change in the tectonic evolution of Lebanon to a period that was dominated by compressive and strike–slip

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deformation. We stress here that such movements did not start before Early Miocene in Lebanon and that this area, unlike Israel and Egypt, did not experience Late Cretaceous Syrian Arc inversion (Gardosh et al. 2010; Moustafa 2010; Yousef et al. 2010). The same applies for NW Syria where the Mesozoic– Palaeogene extension was only interrupted by a discrete compressive period during which ophiolites of the Neo-tethys ocean were obducted onto the Arabian platform (Al Abdalla et al. 2010). In view of these new observations, we feel it is relevant to revisit the meaning of the Syrian Arc deformation. Syrian Arc tectonism is referred to herein as the inversion of the Levant basin and its margins during Late Cretaceous time. We suggest that Palaeogene – Early Neogene structures should not be included because they probably resulted from a different driving mechanism. This definition corresponds to the original Syrian Arc Phase I of Walley (1998), the Syrian Arc is an east –west to WSW– ESE belt located in the southwestern corner of the LB. The northern Levant domain thus escaped the Syrian Arc inversion. Moderate Early Miocene folding in Lebanon occurred synchronous with the emplacement of a major NE –SW fold-and-thrust system in NW Syrian in response to NW–SE compression (Al Abdalla et al. 2010). Notably, after removing the later counter-clockwise rotation, the Early Miocene compression trended NW –SE in Lebanon and had a similar direction to that in Syria. This direction is almost perpendicular to the Eurasian– African collisional plate boundary in NW Syria and we therefore believe that Early Miocene deformation in the northern Levant domain was related to ongoing collisional processes along the plate boundary to the north. Given the flat-lying Eocene sequences in northern Sinai and the Negev Desert, this convergent deformation appears to decrease southward along the Levant margin. We interpret the major folding episode occurring in Lebanon during Late Miocene time to be driven by the initiation of transform tectonics along the DST plate boundary. Given that this age is younger than that of the southern segment of the DST, we conclude that the DST initiated in the south and propagated northwards into Lebanon. The maximum age of the YF, the main fault of the plate boundary in Lebanon, must be Late Langhian because the Early Miocene stress state that was active at least until Early Langhian time would imply it must have experienced dextral strike–slip motion. The regional 288 counterclockwise rotation observed in Lebanon (Henry et al. 2010) rules out the possibility that the YF originally trended north– south, similar to the other main fault branches of the DST. In its present-day

configuration, it trends at a 308 angle relative to the Jordan Valley and Gab fault. Accepting the fact that 178 of counter-clockwise rotation occurred after Miocene time (Henry et al. 2010), the relative deflection of the YF in Lebanon was at least of 478 and may have been larger if all 288 of the rotation occurred late in the DST history. We feel that the origin of this deflection is to be found in crustal rheological constrats within the Levant domain, probably formed during the Early Mesozoic extension. In Lebanon, the transform tectonics included sinistral strike– slip on north–south to NE –SW faults, as well as folding and thrusting along similar orientations in order to accommodate the obliquity of motion relative to the through-going transform boundary. Because the major folds in Lebanon are largely pre-Pliocene structures and because transpression persisted until the present day, a re-organization in deformation drove a transfer of shortening into the offshore portion of the Levant margin. A model for the structural evolution of central segment of the DST is proposed in a companion paper (Henry et al. 2010). In this model, the transform motion was first accommodated during Late Miocene time by folding in Lebanon and the Palmyrides, which was then later transferred onto the Yammouneh and associated faults. This model implies that 118 of the counterclockwise rotation is associated with the first phase, and 178 occurred during the second phase. Our data support a certain degree of stability in the stress field since the transform tectonics began with stress data indicating that a NNW–SSE compression prevailed along the central DST. Depending of the exact time during which the regional rotation occurred, the compression was oriented in a more or less north– south direction. We are grateful for the thoughtful reviews by P. Bentham and L. Csontos. This work was supported by the MEBE (Middle East Basins Evolution) Program, the French and Lebanese CNRS, and the Lebanese University.

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fans. Earth and Planetary Science Letters, 227, 105– 119. Dubertret, L. 1970. Review of the structural geology of the Dead Sea and surrounding areas. Philosophical Transactions of the Toyal Society, London, A, 267, 9– 22. Dubertret, L. 1975. Introduction a` la carte ge´ologique a 1/50 000e du Liban. Notes et Me´moires sur le Moyent-Orient, 13, 345– 403. Dubertret, L. And COLLABORATORS. 1955. Carte ge´ologique du Liban au 1/200 000. Re´publique libanaise, Ministe`re des Travaux publiques, Institut ge´ographique national. Elias, A., Tapponnier, P. et al. 2007. Active thrusting offshore Mount Lebanon: source of the tsunamigenic A.D. 551 Beirut-Tripoli earthquake. Geology, 35, 755– 758. Eyal, Y. 1996. Stress field fluctuations along the Dead Sea rift region since the Late Cretaceous based on mesostructures. Tectonics, 2, 167–185. Eyal, Y. & Reches, Z. 1983. Tectonic analysis of the Dead Sea Rift region since the Late Cretaceous LC based on mesostructures. Tectonics, 2, 167– 185. Freund, R., Garfunkel, Z., Zak, I., Goldberg, M., Weissbrod, T. & Derin, B. 1970. The shear along the Dead Sea Rift. Philosophical Transactions of the Toyal Society, London, 267, 107–130. Freund, R., Goldberg, M., Weissbrod, T., Druckman, Y. & Derin, B. 1975. The Triassic–Jurassic structure of Israel and its relation to the origin of the Eastern Mediterranean. Bulletin Geological Survey of Israel, 65, 26. Gardosh, M. A. & Druckman, Y. 2006. Seismic stratigraphy, structure and tectonic evolution of the Levantine Basin, offshore Israel. In: Robertson, A. H. F. & Mountrakis, D. (eds) Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, 260, 201– 227. Gardosh, M. A., Garfunkel, Z., Druckman, Y. & Buchbinder, B. 2010. Tethyan rifting in the Levant Region and its role in Early Mesozoic crustal evolution. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 9 –36. Garfunkel, Z. 1981. Internal structure of the Dead Sea leaky transform (rift) in relation to plate kinematics. Tectonophysics, 80, 81–108. Garfunkel, Z. 1998. Constrains on the origin and history of the Eastern Mediterranean basin; collision-related processes in the Mediterranean region. Tectonophysics, 298, 5 –35. Garfunkel, Z. 2004. Origin of the Eastern Mediterranean basin: a reevaluation. Tectonophysics, 391, 11–34. Gomez, F., Meghraoui, M. et al. 2003. Holocene faulting and earthquake recurrence along the Serghaya branch of the Dead Sea fault system in Syria and Lebanon. Geophysical Journal International, 153, 658– 674. Gomez, F., Karam, G. et al. 2007. Global Positioning System measurements of strain accumulation and slip transfer through the restraining bend along the Dead Sea fault system in Lebanon. Geophysical Journal International, 168, 1021– 1028.

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Hancock, P. L. & Atiya, M. S. 1979. Tectonic significance of mesofracture systems associated with the Lebanese segment of the Dead Sea transform fault. Journal of Structural Geology, 1, 143–153. Hardy, C., Homberg, C., Eyal, Y., Barrier, E. & Mu¨ller, C. Stresses driving faulting in the southern Levant since Mesozoic. Tectonophysics, accepted. Henry, B., Homberg, C., Mroueh, M., Hamdan, W. & Higazi, F. 2010. Rotations in Lebanon inferred from new palaeomagnetic data and implications for the evolution of the Dead Sea Transform system. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 269–285. Hirsch, F., Flexer, A., Rosenfeld, A. & Yellin-Dror, A. 1995. Palinspatic and crustal studies of the eastern Mediterranean. Journal of Petroleum Geology, 18, 149– 170. Homberg, C., Bergerat, F., Philippe, Y., Lacombe, O. & Angelier, J. 2002. Structural inheritance and Cenozoic stress fields in the Jura fold-and-thrust belt (France). Tectonophysics, 327, 137–158. Homberg, C., Barrier, E., Mroueh, M., Hamdan, W. & Higazi, F. 2009. Basin tectonics during Early Cretaceous in the Levant margin, Lebanon. Journal of Geodynamics, 47, 218–223. Lang, B. & Mimran, Y. 1985. An Early Cretaceous volcanic sequence in central Israel and its significance to the absolute date of the base of the cretaceous. Journal of Geology, 93, 179– 184. Lewy, Z. 1991. Periodicity of Cretaceous epeirogenic pulses and the disappearance of the carbonate platform facies in the Late Cretaceous times (Isr). Israel Journal of Earth Science, 40, 51– 58. Lu¨ning, S., Marzouk, A. M., Morsi, M. & Kuss, J. 1998. Sequence stratigraphy of the Upper Cretaceous of central-east Siani, Egypt. Cretaceous Research, 19, 153– 196. Mimran, Y. 1984. Unconformities on the eastern Flank of the Fari’a anticline, and their implications on the structural evolution of Samaria (Central Israel). Israel Journal of Earth Science, 33, 1– 11. Mohrange, C., Pirrazoli, P. A., Marriner, N., Montaggioni, L. F. & Nammour, T. 2006. Late Holocene relative sea-level in Lebanon Eastern Mediterranean. Marine Geology, 230, 99– 114. Mostafa, M. S. 1999. Evolution tectonique de la plateforme. Africaine en Egypte depuis le Me´sozoı¨que a` partir de l’analyse des de´formations cassantes. PhD, Univ. Paris VI. 326. Moustafa, A. R. 2010. Structural setting and tectonic evolution of North Sinai folds, Egypt. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 37– 63. Moustafa, A. R. & Khalil, M. H. 1994. Structural characteristics and tectonic evolution of north Sinai fold belts. In: Said, R. (ed.) The Geology of Egypt. A.A. Balkema, Rotterdam, 381–389. Mu¨ller, C., Higazi, F., Hamdan, W. & Mroueh, M. 2010. Revised Stratigraphy of the Upper Cretaceous and Cenozoic Series of Lebanon based on

nannofossils. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 287– 303. Nemer, T. & Meghraoui, M. 2006. Evidence of coseismic rupture along the Roum fault (Lebanon): a possible source for the AD 1837 earthquake. Journal of Structural Geology, 28, 1483– 1495. Ponikarov, V. P., Kazmin, V. G. et al. 1966. The Geological Map of Syria. Scale: 1.000 000. Ministry of Industry, Syrian Arab Republic. Quennel, A. M. 1958. The structure and the evolution of the Dead Sea Rift. The Quarterly Journal of the Geological Society, London, 64, 1 –24. Quennel, A. M. 1984. The western Arabia rift system. In: Dixon, J. E. & Robertson, A. H. F. (eds) The Geological Evolution of the Eastern Mediterranean. Blackwell, Oxford, 775–788. Saint-Marc, P. 1970. Le cre´tace´ infe´rieur et moyen du bord occidental du Jabl Sannine (Liban). Notes et Me´m. Moyen Orient, 12, 217– 226. Saint-Marc, P. 1974. Etude stratigraphique et micropale´ontologique de l’Albien, du Ce´nomanien et du Turonien du Liban. Notes Me´moires Moyen-Orient, 8, 8– 342. Schattner, U. Z., Ben-Avraham, M. & Reshef, M. 2006a. Tectonic isolation of the Levant basin offshore Galilee-Lebanon-effects of the Dead Sea fault plate boundary on the Levant continental margin, eastern Mediterranean. Journal of Structural Geology, 28, 2049– 2066. Schattner, U. Z., Ben-Avraham, M. & Reshef, M. 2006b. Development of parallel rifts across the northeastern African plate during the Oligocene-Mid Miocene times. Tectonophysics, 419, 1 –12. Shalar, J. 1994. The Syrian arc system: an overview. Palaeogeography, Palaeoclimatology, Palaeoecology, 112, 125 –142. Vidal, N., Alvarez-Marron, J. & Klaeschen, D. 2000. Internal configuration of the Levantine Basin from seismic reflection data (Eastern Mediterranean). Earth and Planetary Science Letters, 180, 77– 89. Walley, C. D. 1998. Some outstanding issues in the geology of Lebanon and their importance in the tectonic evolution of the Levantine region. Tectonophysics, 298, 1, 37–62. Walley, C. D. 1988. A braided strike-slip model for the northern continuation of the Dead Sea Fault and its implications for Levantine tectonics. Tectonophysics, 145, 63– 72 Walley, C. D. 2001. The Lebanon passive margin and the evolution of the Levantine Neo-Tethys. In: Ziegler, P. A., Cavazza, P. A. W., Robertson, A. H. & Crasquin-Soleau, S. D. (eds) Peri-Tethyan Rift–Wrench Basins and Passive Margins IGCP 369 Results. Me´moires du Muse´um National d’Histoire Naturelle, Paris. Peri-Tethys Me´moire 6. Yousef, M., Moustafa, A. R. & Shann, M. 2010. Structural Setting and Tectonic Evolution of Offshore North Sinai, Egypt. In: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 65–84.

Rotations in Lebanon inferred from new palaeomagnetic data and implications for the evolution of the Dead Sea Transform system B. HENRY1*, C. HOMBERG2, M. MROUEH3, W. HAMDAN3 & F. HIGAZI3 1

Pale´omagne´tisme, IPGP and CNRS, 4 avenue de Neptune, 94107 Saint-Maur cedex, France 2

Universite´ Pierre et Marie Curie, ISTeP, CNRS, Case 129, 4 place Jussieu, 75252 Paris cedex 05, France 3

Lebanese University, Faculte´ d’Agronomie, B.P. 13-5368 Chourane, Beyrouth 1102-2040 Lebanon *Corresponding author (e-mail: [email protected])

Abstract: A study carried out on widespread sites of the Aptian– Albian formations in Lebanon led to two palaeomagnetic directions corresponding to the primary magnetization (N ¼ 37 sites, D ¼ 307.18, I ¼ 23.78, k ¼ 18 and a95 ¼ 5.58 after tilt correction and to a post-folding remagnetization (N ¼ 18 sites, D ¼ 346.38, I ¼ 49.28, k ¼ 108 and a95 ¼ 3.28 before tilt correction). Comparison of these data with previous palaeomagnetic results for the Jurassic age in Lebanon and expected directions from African apparent polar wander path yields evidence of three different counter-clockwise regional rotations, of the order of 338 before Aptian deposition, of 118 during Late Miocene times, and of 188 since Miocene period. The two last rotations are related to the relative displacement of the African and Arabian plates. A model is proposed for the evolution of this particular Middle East area, in which the Dead Sea Transform shows a strong deviation relative to its main north–south orientation.

The Arabian plate is bounded to the west by a c. 1100 km transform fault system, the Dead Sea Transform (DST). The DST connects the Red Sea Rift in the south to the Arabia– Africa–Eurasia triple junction to the north. This left-lateral transform plate boundary developed during the Cenozoic since the Arabian plate moved to the north faster than the African plate. The DST consists of three segments, a north–south southern segment that separates Jordan and Israel, a central NNE–SSW segment across Lebanon, and a north–south northern segment across Syria. The southern and northern segments correspond to a narrow deformation zone with one major fault, the Dead Sea Fault (DSF). Instantaneous plate models (Jestin et al. 1994; Chu & Gordon 1998; Westaway 2004) predict that the Arabia –Africa relative plate motion is almost parallel to the DSF – or with a slight tranpressive component – in these southern and northern segments. In Lebanon, the main fault, the Yammouneh fault (YF), strikes N0308E, and thus corresponds to a restrained bend of the DST. Depending on the Eulerian poles, the YF makes a 308 to 48 angle with the plate motion. This angle would even reach 708 when using the Eulerian pole of DeMets et al. (1994), which should imply transpression all along the DST. The transpressive character of the central DST is thus widely accepted but the question concerning

the origin of the 308 curvature of the DST in Lebanon, that is, primary or resulting from a late tectonic rotation, remains open. Quennell (1984) claimed that the DSF initiated at the Red Sea Rift and then propagated northward along the small circle. According to him, the plate boundary originally trended north–south and a regional clockwise rotation later brought Lebanon into transpression. Walley (1988) also suggested that the proto-YF was at a smaller angle to the plate motion. Others like Butler et al. (1998) and Homberg et al. (2007) claimed that transpression occurred early in the plate boundary history and that the bend of the DSF in Lebanon is primary. Available palaeomagnetic data in Lebanon (Van Dongen et al. 1967; Gregor et al. 1974) do not allow testing of these contradictory hypotheses. They all indicate that Lebanon underwent a regional counter-clockwise rotation of 55 to 708 since Jurassic time. Assuming global rotation of the Lebanon and existence of a north –south DST before the rotation, the YF should have been deviated toward a NW –SE to WNW –ESE trend, totally different from its present NNE– SSW orientation. The chronology of the different events and mode of rotations has, therefore, to be reconsidered. A mechanism of local block rotations has been already suggested (Ron 1987). On the another hand, previous palaeomagnetic data have been obtained in restricted

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 269–285. DOI: 10.1144/SP341.13 0305-8719/10/$15.00 # The Geological Society of London 2010.

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areas in Lebanon and are so far to have been obtained with the present methods of measurement and interpretation. In order to quantify tectonic rotations in Lebanon and to understand their relation with the structural evolution, a new palaeomagnetic study has been carried out in numerous and widespread sites. After describing the main Lebanese geological structures, the sampling and the analysis procedure, the results of our palaeomagnetic study are presented and their structural implications discussed.

Geological setting The Lebanese sector of the DST plate boundary encompasses a broad 100 km wide deformed area. Whereas the plate boundary includes a single main fault in Israel and Syria, the DST in Lebanon splits into a number of splayed faults. This fault system includes three main faults that are from west to east the north –south Roum fault, the

NNE– SSW YF, and the NNE–SSW RachayaSergaya fault (Fig. 1). Although only 7 –11 km of left slip has been documented along it (Dubertret 1970; Hancock & Atiya 1979) – a much lower value than the 70 –75 and 105 km of total slip evidenced along the northern and southern DST respectively (e.g. Dubertret 1955; Freund et al. 1970; Garfunkel 1981; Quennell 1984) – the YF is generally regarded as the main strike –slip branch of the DST in Lebanon. It is a 150 km long fault trending NNE–SSW that includes several segments separated by small pull-apart, like the Yammouneh pull-apart. In detail, markers charting the strike– slip displacement have not yet been identified so that the fault–slip distribution is poorly constrained in Lebanon. Walley (1998) estimated that the YF slipped left-laterally by 47 km and that the Sergaya and Roum faults respectively accumulated 20 km and 8 km of slip. Others quoted different estimates. According to Butler et al. (1998), the Rachaya –Sergaya faults did not slip more than a few kilometres whereas the Roum

36° BA

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Fig. 1. Location of the sampling sites in Lebanon: On the left, sites in their geological context (dark grey, Jurassic and Lower Cretaceous formations; white, Upper Cretaceous formations; light grey, Cenozoic formations; YF, Yammouneh fault; RoF, Roum fault; RaF, Rachaya fault; SF, Sergaya fault). Geological contours and faults from Dubertret & collaborators (1955). Aptian (squares), Albian (diamonds) and Miocene (star) sites. On the right, sites (open squares) according to palaeomagnetic results: sites allowing determination of the magnetization component B (small full circles) and of the magnetization C (crosses).

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

fault accommodated at least 15 to 30 km left-lateral displacement. Among the three Lebanese faults, the YF is the only structure that has the characteristics to account for the plate motion. First, it is a well expressed geomorphological feature that can be followed through the whole country. Second, the YF links the other segments of the DST: the Jordan Valley fault to the south to the Missyaf segment of the Gab fault to the north and is thus the only through-going structure that makes the continuity of the plate boundary across Lebanon (Fig. 1). Recent palaeoseismological studies evidenced a 3.8–6.4 mm a21 late Pleistocene to Holocene slip rate (Daeron et al. 2004) and thus confirm that the YF is the major plate structure. The Sergaya fault shows a 1.4 mm a21 left lateral slip rate during the Holocene (Gometz et al. 2003). Tapponier et al. (2001) claim that the Roum fault is also active today and evolves offshore into a curved thrust that enters again inland in front of the east– west Tripoli anticline. WSW–ENE to WNW – ESE faults also exist on both sides of the YF. These faults are generally assumed to accommodate right-lateral slip and are thus considered as part of the DST. New field investigations allowed recognition of an Early Cretaceous and Eocene normal slips along these faults (Homberg et al. 2009) and large horizontal offsets have not been yet identified. Whatever the exact slip amount along each fault in Lebanon, the large-scale folds and associated topographic elevations also argue for the transpressive character of the central DST. The three main fold structures are from west to east the Mount Lebanon anticline, the Bekaa syncline, and the AntiLebanon anticline (Fig. 1). The Mount Lebanon anticline exhibits a north–south axis in its southern part, which progressively rotates northward to a NNE–SSW direction. The Bekaa syncline and the Anti-Lebanon anticline axes trends NNE– SSW. The two anticlines culminate at 3088 m and 2814 m and consist of Middle Jurassic sequences in their cores. The Bekaa syncline is filled by continental Neogene sediments. To the east, the Lebanese folds pass into the Palmyrides fold-and-thrust-belt that strikes NE –SW. Folding and faulting of the cover continues offshore, as evidenced by recent seismic profiles (Elias 2006). Additional smaller scale folds exist in the Bekaa syncline and near Tripoli. The age of structures of the DST in Lebanon has been a matter of debate (e.g. Quennell 1984; Butler et al. 1998; Walley 1998). The lack of markers precludes dating initiation of strike –slip faulting. Recent revision of the age of the angular unconformities within the Cenozoic sequence in Lebanon provides constraints on the fold development. Homberg et al. (2007, 2010) showed that the Lebanese folds initiated during Early Miocene and

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were mostly shaped during Late Miocene. Because the moderate Early Miocene folding is associated with an east –west compression that cannot drive left-lateral slip on the YF, they claim that the fault activity did not start before Late Miocene. This age is slightly younger than the 15–19 Ma age estimated for the initiation of the southern DST segment, based on offset of 19 Ma basaltic dykes (Eyal et al. 1981).

Sampling and analysis procedures The Meso-Cenozoic geological series in Lebanon includes very few facies favourable for palaeomagnetic studies. Jurassic and Cretaceous volcanic levels are often altered. Their structural position cannot always be precisely determined and they are present only in a reduced part of the Lebanon. Moreover, to be usable for structural purposes, palaeomagnetic studies in volcanic rocks have to concern several flows of slightly different ages to eliminate the effect of the secular variation of the Earth’s magnetic field. On the another hand, most of the sedimentary deposits are limestone and sandstone, which are very poor in magnetic minerals, as we determined in a preliminary study on unoriented blocks. Only the Aptian –Albian formation gave reliable stable magnetizations. This could be owing to the presence in the underlying levels of volcanic rocks, which were probably partially eroded and contributed locally derived magnetic minerals. The Aptian– Albian formation moreover is present in different areas in Mount-Lebanon and Anti-Lebanon regions. We then decided to focus our sampling on this formation, corresponding to a large window of time (20–27.5 Ma, Harland et al. 1989; Odin 1994; Gradstein et al. 2004). In the field, however, it appeared that only a few facies in this formation are actually suitable for palaeomagnetism. With these constraints, 38 widely distributed sites (Fig. 1) have been found for sampling. 325 cores (a mean of eight per site) were obtained with a gasoline-powered portable drill and oriented with a compass. At least one core per site has been oriented also with a Suncompass to look for possible local deviation of the Earth’s magnetic field. 20 sites are within the Aptian part of the formation, and the other 18 sites are of Albian age. As a test, one additional sampling site (site K, eight samples) has been chosen in the Miocene formation. One to three specimens of standard size (cylinders of 11 cm3) were cut from each core, allowing both thermal and alternating field (AF) demagnetization and rock-magnetism studies to be performed. Prior to any demagnetization analysis, the specimens were stored in a zero magnetic field for

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at least one month in order to reduce a possible viscous magnetization. The remanant magnetization was measured using a JR4 spinner magnetometer (AGICO, Brno, Czech Republic). Laboratory-made equipments were used for AF-demagnetization (Le Goff 1985) and thermal treatment. During thermal demagnetization, variations of the magnetic susceptibility at room temperature were monitored using a Kappabridge KLY3 (AGICO, Brno, Czech Republic). In order to correctly isolate and identify the magnetization components, numerous demagnetization steps were performed on pilot-specimens with increments ranging from 50 8C in the lowest temperatures to 20 8C in the highest ones, whereas increments range from 2.5 mT to 5 mT during AFdemagnetization (maximum field intensity: 90 mT). The results of the demagnetization process were studied on orthogonal vector plots (As & Zijderveld 1958; Zijderveld 1967). The direction of magnetization components was determined by principal component analysis (Kirschvink 1980). Mean site characteristic directions were calculated using Fisher (1953) and bivariate (Le Goff 1990; Le Goff et al. 1992) statistics.

Rock magnetism analysis The intensity of the natural remanent magnetization (NRM) is mostly low (in average of the order of 1023 A/m, the lowest value being as weak as 1025 A/m, and the highest one up to 4  1022 A/ m). This intensity is sometimes relatively variable within the same site. We will see that, according to the samples, two kinds of behaviour of remanent magnetization during demagnetization have been found, whatever the NRM intensity. This argues for the presence of the same ferrimagnetic mineralogy in all these rocks. Hysteresis loops were measured on small samples (about 3 cm3) using a translation inductometer within an electromagnet. Figure 2 shows the two extreme kinds of obtained hysteresis loops. Sample F4 has high coercivity (coercive force Hc ¼ 0.077 T; remanent coercive force Hcr ¼ 0.45 T). In contrast sample B9 has a much lower coercivity (Hc ¼ 0.01 T; Hcr ¼ 0.04 T). Both curves show that magnetization saturation is not reached at 0.8 T and they are slightly waspwaisted. That indicates in both cases the presence of at least two magnetic phases, one of them having high coercivity. The same conclusion is obtained with AF-demagnetization. Some samples, like B9, can be easily fully demagnetized, while only 5–10% of the NRM can be eliminated for samples like F4. In a few sites, very different coercivity can be observed in different samples.

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Fig. 2. Hysteresis loops after correction for paramagnetism for samples B9 (a) and F4 (b) – M in 1023 Am2/kg and H in T.

Thermomagnetic curves (low-field susceptibility as a function of temperature) were determined using CS3 equipment – Kappabridge KLY3 (AGICO, Brno, Czech Republic) on crushed samples. Only one of the studied samples (F4) presents weak alteration during heating, occurring only at a temperature higher than 580 8C (Fig. 3a). For this sample, the thermomagnetic curve in air suggests the presence of mainly magnetite. For sample B9, mineralogical alteration during heating in air occurred since about 300 8C (Fig. 3b). This alteration gives a decrease of susceptibility, and therefore does not correspond to formation of magnetite. The Curie temperature of 580 8C on the heating curve indicates the presence of magnetite, therefore preexisting in this sample. Measured curves for other samples indicate stronger mineralogical alteration during heating. Example of sample Z7 is presented in Figure 3c (experiment in air) and 3d (experiment in argon atmosphere). In both curves, alteration started at about 400 8C, but the newly formed minerals are not quite identical. Monitoring of the mean susceptibility at room temperature during progressive thermal demagnetization gives additional information. Weak variations of susceptibility after 100 8C heating probably reflects stress release or grain breakage of magnetic minerals owing to fast heating rather than true mineralogical changes. Important susceptibility increase after heating at 200 8C is clearly owing to the beginning of mineralogical alterations. Up to 400–4308 the susceptibility decreases in most

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

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samples, but at higher temperatures, strongly increases for almost all of them. Analyses of the difference in hysteresis loops measured after stepwise heating (Henry et al. 2005) made on sample Z7 show that the susceptibility decrease during thermal demagnetization from 200 to 400 8C corresponds to the disappearance of a high-coercivity component. At higher temperature, a moderate coercive magnetic phase is formed (probably a Ti-poor titanomagnetite). These rock-magnetism studies show that the initial magnetic mineralogy corresponds to at least two minerals, one of moderate coercivity corresponding to magnetite, and the other with much higher coercivity. They also indicate strong and complex changes occurring for the magnetic mineralogy during thermal demagnetization at a relatively low temperature.

Palaeomagnetic results Magnetization analysis NRM directions are mostly scattered, in a northward to north-westward plunging plane, between a direction close to the present day magnetic field (D ¼ 3.88; I ¼ 49.88) and a NNW to WNW moderately dipping direction. Detailed analysis of the NRM components was made first on pilot-specimens (two analyses per site, one using thermal treatment, the other one with AF-demagnetization).

During thermal demagnetization, a component can be determined from 150 8C to 340– 430 8C steps (Fig. 4a), but lines corresponding to this component on an orthogonal vector plot are often not very well defined. This is probably owing to acquisition of parasitic magnetization at low temperatures or to demagnetization of several components at the same time. On orthogonal vector plots, the lines do not converge toward the origin, and it is clear that at least another magnetization component resists demagnetization at the 340 to 430 8C steps. However, at a higher temperature, the direction of remaining magnetization becomes erratic and does not permit the definition of a stable vector component. Parasitic magnetizations acquired by newly formed minerals during thermal treatment are likely more important than the initial high temperature component(s). Sometimes, the orthogonal vector plot shows, for high temperature steps, small straight lines for three sequential steps, but the corresponding directions are very scattered and without any coherence within a site. They are attributed to parasitic magnetizations acquired during the previous heating steps. AF-demagnetization gives much better results. A small magnetization component is generally eliminated after 5 to 10 mT (component A). It probably corresponds to a viscous magnetization. A second magnetization component is then clearly defined on orthogonal vector plots between 10 and 40 –55 mT (component B). As indicated in the

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(b)

(c)

West - Up

West - Up

2.5 mT

(a)

West - Up 0.5 10–9 A m2

150 °C 340 °C

North

A3 20 mT

590 °C

490 °C

X1

North 55 mT North

–8

10 A m

2.5 mT 55 mT

2

BC2 –9

2 10 A m

Down

2

Down

20 mT

Down

(d)

West - Up

(e)

2.5 mT

West - Up North

25 mT BB1

55 mT 55 mT

45 mT North

–9

10 A m

20 mT

–10

5 10 A m2

10 mT

2.5 mT

2

Down

BI1 Down

Fig. 4. Orthogonal vector plots (full squares, horizontal plane; open squares, vertical plane) in stratigraphical coordinates for (a, sample A3) thermal (temperature in 8C) and (b, c, d, e) AF-demagnetization (field in mT) for samples dominated by moderate (b, sample BC2) and high (c, sample X1) coercive components and for intermediate cases (d, sample BB1; e, sample BI1).

rock-magnetism section, this component represents a very different percentage of the NRM in different samples (Fig. 4b–e). The lines on the orthogonal vector plot do not converge toward the origin in specimens where a significant part of the NRM still remains after 55 mT AF-treatment. For higher peak-AF amplitudes, magnetization intensity becomes relatively stable, with only weak parasitic anhysteretic magnetizations starting to appear. Application of thermal demagnetization after 55 mT AF-treatment does not yield any change of the magnetization up to about 400 8C treatment and only parasitic magnetizations are acquired for the higher temperature steps. Components isolated at a moderate temperature (150 to 340 –430 8C) and at a moderate AF (10 to 40–55 mT), therefore, represent the same part of the NRM. This magnetization component B is carried by the magnetite, the high coercive mineral(s) being the carrier(s) of the component(s) C that cannot be analysed because of the development of parasitic magnetization during demagnetization at high temperature. AF-demagnetization was then chosen to analyse the remanent magnetization of all the other

specimens. In some specimens, component B was too weak to be reliably determined. For two sites (BP and the Miocene site K), NRM had too low intensity to be analysed. In all the other sites (Fig. 1), the magnetization component B carried by magnetite has been isolated at least in part of the samples (Table 1). The problem when magnetization component B is isolated while another magnetization component C is not, is to determine if magnetization C started to be demagnetized at low treatment levels, that is, whether magnetization component B results from a partial superimposition with magnetization C (Bouabdallah et al. 2003). Here, in sites where the proportion of high and moderate coercive minerals is variable, the consistency of component B directions indicates that there is no, or only limited, superimposition. An interesting observation concerns the direction of the remaining magnetization C after AF-demagnetization. For the samples where this magnetization represents an important part of the NRM, this direction has a coherent orientation before tilt correction with a mean inclination (49.28)

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

275

Table 1. Palaeomagnetic data for the magnetization component B Site

Age

N

Dg (8)

Ig (8)

Ds (8)

Is (8)

k

a95 (8)

A B C D E F G H I L M N O P R S T U V W* X Y Z BA BB BC BD BF BG BH BI BJ BK BL BM BN BO

Aptian Albian Aptian Albian Albian Albian Albian Albian Albian Albian Albian Aptian Aptian Aptian Aptian Aptian Albian Aptian Aptian Aptian Aptian Albian Aptian Aptian Albian Albian Aptian Aptian Albian Aptian Aptian Aptian Albian Aptian Aptian Albian Albian

9 9 4 7 8 7 7 8 9 7 9 8 6 7 8 7 7 9 7 10 7 8 7 7 7 8 8 4 5 7 5 6 8 6 7 8 7

307.0 316.4 297.7 344.2 318.6 292.1 310.7 314.1 319.1 330.0 251.6 303.4 304.7 298.7 310.1 312.0 310.7 287.8 297.4 312.4 305.0 313.0 298.9 293.4 325.1 299.0 277.4 333.1 340.0 317.4 358.1 284.1 284.9 297.6 296.9 299.1 334.4

32.0 35.8 12.7 42.4 24.6 56.0 18.6 16.9 55.2 47.5 45.1 32.8 24.7 31.5 237.7 28.4 2.9 30.3 19.0 2 6.3 31.8 40.8 33.7 23.4 48.5 57.1 31.4 37.5 35.8 38.8 40.6 59.0 21.9 23.1 25.1 25.6 39.2

307.0 309.7 297.7 325.7 319.0 286.0 296.3 317.0 319.9 318.1 246.0 304.6 308.7 296.0 301.2 311.2 319.0 289.5 298.8 309.2 311.6 312.8 300.4 294.7 321.6 296.2 279.4 327.8 332.7 311.8 347.3 288.4 287.5 294.4 296.9 299.1 307.4

32.0 12.7 12.7 25.8 34.6 19.7 28.2 18.5 12.2 37.8 33.9 11.5 17.1 20.8 6.2 27.0 36.7 1.9 13.9 22.7 22.1 25.3 22.0 18.4 20.9 23.2 7.0 35.8 41.7 21.2 30.6 37.3 15.5 17.8 26.0 25.6 20.8

82 101 50 77 130 106 651 86 171 165 67 68 75 21 77 211 87 202 94 50 42 33 60 174 153 154 119 166 566 16 225 64 69 41 71 179 32

5.1 4.6 9.9 6.0 4.3 5.1 2.1 5.3 3.6 4.1 5.7 6.0 6.6 11.5 5.6 3.6 5.7 3.2 5.5 6.3 8.2 8.6 6.8 4.0 4.3 4.0 4.5 5.4 2.6 13.2 4.2 7.1 6.0 8.9 6.3 3.7 9.4

Note: Site, number of retained samples N, Declination Dg and inclination Ig in geographical coordinates, Declination Ds and inclination Is in stratigraphical coordinates, Fisher (1953) parameters k and a95. *All reversed directions changed to normal polarity.

close to that of the present day Earth magnetic field (49.88) and is always of normal polarity. This magnetization C has been obtained in widely distributed sites in Lebanon (Fig. 1 right).

(a)

(b)

Palaeomagnetic tests The directions of magnetization C are much more clustered before (N ¼ 18 sites, D ¼ 346.38, I ¼ 49.28, k ¼ 108 and a95 ¼ 3.28) than after (D ¼ 337.58, I ¼ 39.68, k ¼ 30 and a95 ¼ 6.18) tilt correction (Fig. 5). During progressive untilting, the maximum k value (113), obtained for 10% untilting, is very close to the k value for 0% untilting (108),

Fig. 5. Mean palaeomagnetic directions for the magnetization C before (a) and after (b) tilt correction (stereographic projection – open symbols for the upper hemisphere and full symbols for the lower hemisphere).

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B. HENRY ET AL.

but much higher than the k value for 100% untilting (30). Because it was impossible to demagnetize this high coercivity magnetization C, the latter could be composite. However, its high clustering, consistent normal polarity, and the positive fold test strongly argue for a single component acquired after folding. The mean direction of magnetization component B is slightly different before (N ¼ 37 sites, D ¼ 308.98, I ¼ 33.38, k ¼ 11 and a95 ¼ 7.08) and after (D ¼ 307.18, I ¼ 23.78, k ¼ 18 and a95 ¼ 5.58) tilt correction, with a better clustering after the correction (Fig. 6). The mean directions for Aptian (N ¼ 20 sites, D ¼ 306.68, I ¼ 26.68, k ¼ 10, a95 ¼ 10.68 and D ¼ 305.88, I ¼ 20.58, k ¼ 17, a95 ¼ 7.78 before and after tilt correction, respectively) and Albian (N ¼ 17 sites, D ¼ 312.68, I ¼ 38.08, k ¼ 13, a95 ¼ 9.68 and D ¼ 308.88, I ¼ 28.08, k ¼ 19, a95 ¼ 7.88 before and after tilt correction, respectively) sites both show a better clustering after tilt correction. Progressive tilt correction for all the 37 sites gives the best clustering (maximum k value ¼ 17.8) for 85% of untilting, but without significant difference with 100% untilting (k ¼ 17.6). A MacFadden (1990) fold test yields also optimal untilting at 86%. Bivariate statistics shows that the component B directions after tilt correction have an elongated distribution, roughly along a NE–SW direction. This corresponds to a scattering mainly in declination, indicating presence of differential rotations with vertical axis according to the sites. Classical fold tests are then partly biased. Inclination and scattering in declination being moderate, bias for the mean inclination (overestimated value) remains weak. In fact, after tilt correction, the mean value of inclination (22.68) is close to mean inclination (23.78) from Fisher (1953) statistics. A fold test based on inclinations only (Enkin & Watson 1996) mainly considers, for moderate inclination, the variation of the standard deviation of the inclination distribution

(a)

(b)

σ

19 18 17 16 15 14 13 12 11 10

P

9 0

10

20

30

40

50

60

70

80

90

100

Fig. 7. Standard deviation s of the distribution in inclination (in degrees) of the magnetization component B (mean site directions) as a function of the untilting percentage P.

during progressive untilting (Fig. 7). Minimum standard deviation (s ¼ 9.6) corresponds to about 85% untilting, but its difference with 100% untilting (s ¼ 10.0) is not significant. That is not the case for 0% untilting: the standard deviation for 85% untilting is almost half of its value for 0% (s ¼ 18.1). Aptian –Albian times are characterized by a normal polarity of the Earth magnetic field, except at least for one relatively short period of the Early Aptian times (Reversed Chron M0). For the magnetization component B, normal polarity was obtained in almost all our sites. Only one of the Lower Aptian sites gave both normal (seven samples) and reversed (three samples) polarities. A positive reversal test (McFadden & McElhinny 1990) has been obtained for this site W: the angle g between the mean normal and reversed directions is only 7.58 while the critical value gc is 10.98. This is not an unquestionable argument of presence of a primary magnetization (Henry et al. 2004a), but a positive reversal test for a secondary magnetization remains an exceptional case.

Interpretation of the results and discussion Apparent polar wander path (APWP).

Fig. 6. Mean palaeomagnetic directions for the magnetization component B before (a) and after (b) tilt correction (stereographic projection – open symbols for the upper hemisphere and full symbols for the lower hemisphere).

Because the large majority of the plate movement occurred on the YF, the area situated on the western side of this fault can be considered in first approximation as being the northern continuation of the African plate. We therefore used, as a reference, the African APWP of Besse & Courtillot (2002) who chose the Harland et al. (1989) time scale. Figure 8 presents the expected declination and inclination of the Earth’s magnetic field at Beirut since 200 Ma calculated from this APWP. Because some of our sites are east of the YF, we also determined declination and inclination from

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST 200

b

100

0

Ma 4

330 a b

2 300

3

1

b

D

a

100

b

200

Ma

0 a

a b

I 1 30

3 2 b 4

60

Fig. 8. Declination (D) and inclination (I) in degrees, expected at Beirut during geological times (Age in Ma – Harland et al. 1989), assuming no relative movements between Lebanon and Africa, calculated from African APWP (Besse & Courtillot 2002), of magnetization component B (after tilt correction) for Aptian sites (1), Albian sites (2) and all the sites (3), and of magnetization C (4 – before tilt correction) from this study. Data ‘a’ from Van Dongen et al. (1967) and ‘b’ from Gregor et al. (1974).

the Arabian APWP (determined from the African APWP after correction for the Red Sea opening). The difference of declination and inclination values from African and Arabian APWPs is weak and only values from the African curve are indicated in Figure 8.

Age of the magnetizations Highly coercive minerals, carriers of the magnetization C, either are not stable at a temperature higher than 300 –430 8C or their magnetization is hidden by parasitic magnetization appearing from heating at these temperatures. Sample F4 (Figs 2 & 3) does not show evidence of transformation at these temperatures but presents strong parasitic magnetization at temperatures higher than 400 8C. The other samples show strong alteration during heating, but we do not know which minerals were altered. Curie temperatures of the highly coercive minerals are higher than 430 8C. These minerals cannot be Fe-hydroxides or sulphides. They are then probably Ti-iron oxides. Such minerals are not frequent in sedimentary formations. Here, a probable origin is titanomaghemite or titanomagnetite from eroded volcanic rocks. As shown by the fold test, these Ti-iron oxides now carry a secondary

277

magnetization acquired after the main folding, that is, after the Late Miocene. The inclination of magnetization C before tilt correction (direction 4 in Fig. 8) corresponds to the expected inclination in Lebanon in the Upper Cenozoic from the African APWP. This confirms the age of the magnetization C. Owing to the reversal frequency during the Upper Cenozoic, the normal polarity observed in all the sites argues for acquisition of the remagnetization C in a relatively short period, for example, during a fluid migration (Henry et al. 2004a). The magnetite carrying component B probably also came mainly from the volcanic rocks. Contrary to the Ti-iron oxides, magnetite was not altered. The fold test shows that this magnetization could be prefolding or have been acquired just at the beginning of folding. The inclination of component B is not compatible with the expected inclination of the magnetic field in Lebanon during the Upper Cenozoic (Fig. 8), which is the period of folding in this area (folding began during Early Miocene; Homberg et al. 2010). Component B therefore cannot be synfolding. Comparison with data from the African APWP shows that the inclination of component B (direction 3 in Fig. 8) is compatible with the expected inclination of the Cretaceous magnetic field in Lebanon. The directions for Aptian and Albian sites both yield the same observation. We also note that inclination increases from Aptian data (direction 1 in Fig. 8) to Albian data (direction 2 in Fig. 8), like the expected inclination from African APWP. The relative evolution from Aptian declination (direction 1 in Fig. 8) to Albian declination (direction 2 in Fig. 8) in our data is also equivalent to that for the reference curve from African APWP. Finally, the positive reversal test in site W strongly argues for a primary magnetization. All these arguments confirm the primary character of component B. Magnetite in these sediments is then the carrier of a primary detrital remanent magnetization.

Previous studies Several palaeomagnetic studies (Table 2) in the Levant area offer the possibility to discuss rotations at the plate scale. Unfortunately, no information is available on the number of different flows studied, except for the data of Roperch & Bonhommet (1986) – 32 flows – and of Nur & Helsley (1971) – 18 flows for the Pliocene lavas. Palaeomagnetic analytical precision improved considerably with the use of principal component analysis (Kirschvink 1980). This is particularly important in the cases of complex remanence structure, as with our sedimentary Aptian –Albian formation. The determination of the palaeomagnetic direction by principal component analysis has been used by Ron (1987),

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Table 2. Previous palaeomagnetic results in the Levant area Stratigraphical age Triassic Lower Jurassic Middle Jurassic Upper Jurassic Upper Jurassic Lower Cretaceous Lower Cretaceous Lower Cretaceous Lower Cretaceous Lower Cretaceous Lower Cretaceous Upper Cretaceous Upper Cretaceous Upper Cretaceous Miocene Miocene Miocene Miocene Pliocene Pliocene Pliocene Pliocene Pliocene Quaternary

Rock type Sediments Sediments Sediments Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Sediments Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks Volcanic rocks

– Limestones – Sediments

– Sediments – Sediments

D (8)

I (8)

a95 (8)

Location

Ref.

252 352 349 93 95 314 122 332 173 91 340 325 343 354 290 312 179 5 170 2 344 175 4 356

– 20 27 41 10 21 9 2 8 – 43 9 12 –6 6 19 32 47 – 34 33 – 35 46 41.8 – 49 41 56

– – – 3 11 6 9 9 – 13 7 5 – 7 (9) 11 6 8 10 8 22 5 13 12

Israel Israel Israel Lebanon Lebanon Lebanon Lebanon Israel Israel Israel Israel Israel Israel Israel Israel Israel Syria Syria Syria Lebanon Israel Israel Syria Syria

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

Note: Stratigraphical age, rocks type, mean declination D and inclination I, radius of confidence angle at 95% a95, location and reference (Ref.): 1, Van Dongen et al. (1967); 2, Gregor et al. (1974); 3, Helsley & Nur (1970); 4, Nur & Helsley (1971); 5, Freund & Tarling (1979); 6, Ron et al. (1984); 7, Roperch & Bonhommet (1986); 8, Ron (1987); 9, Ron & Baer (1988); 10, Abou Deeb et al. (1999). Ron et al. (1984) gave also palaeomagnetic data from Cretaceous rocks, but directions are very different according to the studied areas.

Ron & Baer (1988), Abou-Deeb et al. (1999), and probably (although not documented in their text) by Ron et al. (1984) and Roperch & Bonhommet (1986). All other studies have chosen the direction remaining after demagnetization (i.e. corresponding to the recent remagnetization C in our sedimentary samples) as the palaeomagnetic result. Van Dongen et al. (1967) and Gregor et al. (1974) gave for volcanic rocks orthogonal vector plots showing convergence to the origin, and these volcanic rocks likely contained no, or only weak, secondary high coercivity magnetization components. That could also be the case for the studies of Helsley & Nur (1970) and Nur & Helsley (1971) based also on volcanic rocks, but no indication was given. The results of Freund & Tarling (1979) were obtained on few samples (mainly sediments), which, moreover, gave within each site large variations of palaeomagnetic direction. The latter then probably results partly from secondary magnetization. On the another hand, the result of Van Dongen et al. (1967) is based on few samples per site, and no indication about tilt correction is given in Gregor et al. (1974). Ron et al. (1984) have mainly chosen sites close to fault zones, with the aim of structural

applications, but they obtained coherent results in several neighbouring sites. As a synthesis, the results of Van Dongen et al. (1967) and Gregor et al. (1974) probably represent a correct approximation of the true palaeomagnetic directions. Those of Roperch & Bonhommet (1986), Ron (1987), Ron & Baer (1988) and Abou-Deeb et al. (1999) are quite reliable. Though relatively scattered and possibly related to local rotations, data of Ron et al. (1984) yields important information. A serious doubt remains concerning the data of Helsley & Nur (1970) and Nur & Helsley (1971), particularly for the Miocene data from the southern Mount Hermon area (possibly from a single flow). According to the most reliable studies, counterclockwise rotation of Lebanon of the order of 408 occurred between the emplacement of the Jurassic and Cretaceous basalts. Since the Cretaceous time, counterclockwise rotation of the order of 308 affected the studied sites in Lebanon and part of the north of Israel, but not in the south of Israel. Since the Miocene period, no rotation affected the Syrian sites, but the questionable result of Nur & Helsley (1971) suggests that the southern Mount Hermon area could have been

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

affected by the strong counter-clockwise rotation of the order of 708. No rotation affected the Pliocene sites, except locally in the region around the Galilee Sea. Ron et al. (1984) highlighted in fact the complicated behaviour of the faults in this region, giving various rotations in different areas. For Cretaceous sites, they obtained a rotation similar to that determined in Lebanon for only three sites of the central part of northern Galilee. Other Cretaceous sites indicate no rotation (2 sites in the western part of northern Galilee and 4 sites in northern Carmel) or clockwise rotation (14 other sites in Galilee and Carmel). In the Tiberias province, the 13 Miocene sites indicate strong counterclockwise rotation, and the 17 Pliocene sites, a less important but significant counterclockwise rotation. In the Golan and east of the Galilee Sea, the 12 Pliocene sites on the contrary do not indicate any rotation.

Rotations in Lebanon As demonstrated above, the magnetization measured in the Aptian and Albian rocks of Lebanon includes two components, a component B acquired during deposition and a late component C acquired after folding of the cover (i.e. during the Pliocene or later). From comparison with the reference curve of the declination of the magnetization component B (Fig. 8), a mean counterclockwise rotation R2 of 28.0 + 6.48 (Demarest 1983) affected of all our sites since the Early Cretaceous period. Because the Aptian and Albian mean palaeomagnetic declinations differ only by 38, we did not distinguish Aptian and Albian sites in Figure 9. The magnetization directions are homogeneous in almost all the studied sites (Fig. 9), indicating that this R2 rotation is a regional feature. Our component B data are in good agreement with the Cretaceous results (Fig. 8) of Van Dongen et al. (1967) and Gregor et al. (1974). According to Gregor et al. (1974), no significant rotation occurred in Lebanon from Pliocene lavas data. However, comparison with the reference curve of the declination of the magnetization C indicates that at least part (17.7 + 3.88 – Demarest 1983) of the rotation R2 occurred after the Miocene folding. In this case, the chronology of the different events should be folding, acquisition of the magnetization C, rotation, and ending lava emplacement. In detail, the cumulative R2 rotation is smaller in southern Lebanon than in the northern part of the country. In three sites, the rotation exceeds the regional 288 value. These sites are located close to the Sergaya fault, to the Rachaya fault, and at the intersection of the Roum and Yammouneh faults, and the rotation reaches respectively 518, 608 and 938.

279

The comparison with the reference curve of the Cretaceous declination and of the Jurassic declination from Van Dongen et al. (1967) and Gregor et al. (1974) shows that a strong counterclockwise rotation R1 (of the order of 338) affected the Lebanon between Upper Jurassic and Middle Cretaceous. Similar deviation from the reference curve for declination of our Aptian and Albian data (Fig. 8) indicates that no rotation occurred during this period. The rotation R1 therefore predates the Aptian period.

Structural implications Concerning this rotation R1, few related geological facts are available. † Important and widespread volcanic activity characterizes the Late Jurassic and the Early Cretaceous. † Presently trending east –west faults were active, at least during the Early Cretaceous. † No compressive events are known during Late Jurassic and Early Cretaceous. We propose that the presently trending east –west faults already existed during the Early Jurassic, but with a NNW– SSE orientation. Similarly to the model of Ron (1987), dextral strike –slip between different small blocks owing to east–west to NW –SE extension could have produced the counterclockwise rotation. This assumption has to be, however, considered with care since no indisputable geological arguments, in favour or against it, actually exist. Palaeomagnetic data from Jurassic rocks are moreover insufficient to determine if rotation was different in the diverse parts of Lebanon. Concerning the rotation R2, it largely exceeds the regional value in three sites located close to faults. Here, local rotations clearly occurred, probably as a result of a simple shear along faults as it is generally the case along major faults (Henry et al. 2004b). Our palaeomagnetic analysis from the other sites indicates that Lebanon underwent a mean cumulative 28.08 (+6.48) counterclockwise rotation R2 since the Late Cretaceous. Although the amount of rotation slightly varies through Lebanon (the local values of rotation in Lebanon are smaller in the southern part of the country and larger in the northern part), the low dispersion of our data suggests that this rotation reflects a large scale process. A 17.78 (+3.88) of rotation occurred after folding, that is during the Pliocene or later. This Late Cenozoic rotation is thus coeval to the transform tectonics. To the east, no rotation has been found for this period in the Damas area in Syria. That is not the case to the south for Galilee (Ron et al. 1984), where there are complex

280

B. HENRY ET AL. v v v

CZ

v

v v

v v

v

v

v

v

v v

v

v v

v

v

v v

v v

v

Tripoli

Ar

YF

DSF Nu

v

36°

Eu

RS

34°

Baalbek

SF Ra F

RoF

Beirut

Fig. 9. Declination (single arrows) obtained in the different sites for magnetization component B. Open arrow on the bottom right indicates the expected mean declination from the African APWP for Aptian–Albian times (see Fig. 1). Inset presents regional context (DSF, Dead Sea Fault; Nu, African; Ar, Arabian; Eu, Eurasian plates; CZ, collision zone; RS, Red Sea).

rotations. Galilee probably represents a transition zone between the rotated Lebanon and the nonrotated southern part of Israel. We propose a model of the accommodation of the differential movement of the African and Arabian plates in Lebanon based on the fact that the deformation mechanism fundamentally evolved

since the very beginning of the plate boundary, from a widespread mainly folding mode during the Late Miocene to a mixed shortening – strike–slip mode from the Pliocene. In this model, during the Late Miocene, the locking of the propagation of the southern segment of the DST to the north led to the accommodation

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

of the plate relative movement by the rotation or shortening of deformable blocks (Fig. 10). Block boundaries are controlled by contrasts of crustal thickness and by faults inherited from Mesozoic tectonics. The Palmyride block is bounded to the north by the Triassic –Jurassic Palmyride trough. The boundary between the Palmyride and Lebanon blocks corresponds to the present-day Sergaya – Rachaya fault, probably reflecting deep Mesozoic faults. The western limit of the Lebanon block is probably associated with the boundary between the lithosphere of the Lebanese platform and the thinned lithosphere of the Levant basin. The Lebanon block was affected by a counter-clockwise rotation of about 118, while the Palmyride block underwent a northward translation. These movements were accommodated by a shortening in the two blocks by folding. To model this evolution,

281

we artificially applied to the Palmyride block two successive opposite 118 rotations (see Fig. 10): the first is the same as the Lebanon block (counterclockwise around the northern end of the Jordan Valley fault) and the second is clockwise around the western end of the Palmyrides block and above the Palmyride trough. This mechanism led to absorption of 40 km of the plate motion by folding in Lebanon and in the Palmyride (Fig. 10). Note that the remaining relative displacement, about 67 km, is in line with that (70–75 km) proposed along the northern prolongation of the DST in Syria (Dubertret 1955; Freund et al. 1970). In Lebanon, in agreement with the elevation of the topography, shortening increases northward to a maximal value of 25 km. The second stage that occurred since the Pliocene corresponds in this model to the propagation

Fig. 10. Miocene mechanism of deformation in the central Dead Sea plate boundary. The displacement d1 between Nubia (Nu) and Arabia (Ar) is accommodated by folding in Lebanon and Palmyrides without strike– slip motion on the Yammouneh or other faults. Modelling was made assuming artificially two successive opposite rotations for the Palmyride block (schema on the right-hand bottom corner). In Lebanon, 118 of rotation occurs in a counterclockwise sense, which corresponds to the first stage of the rotation R2: Hatched areas indicate the amount of shortening. The Lebanese blocks (light grey), the Palmyride block (dark grey), and the main folds (dark lines with double arrows) are shown in their palaeogeographic positions before the displacement on the Yammouneh and Roum fault (see Fig. 11). Their present-day positions are in dashed lines. Line l and line p illustrate the displacements and rotations in the Lebanese and Palmyrides blocks. B, D, Tr, Ta: cities of Beirut, Damas, Tripoli, Tartous. DSF, Dead Sea fault.; JF, Jhar fault. See text for further explanations.

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B. HENRY ET AL.

Fig. 11. Post Miocene mechanism of deformation in the central Dead Sea plate boundary. (a) initial stage; (b) sinistral slip (dr) on the Roum fault (RF); (c) counterclockwise rotation in the Lebanese block (second stage of rotation R2); (d) sinistral slip motion (dy) on the Yammouneh fault (YF). Nu, Ar, Nubian and Arabian plates. Steps b, c and d may be synchronous. 70 km of plate motion is accommodated by displacement on the Roum and Yammouneh faults. Slip on the Sergaya– Rachaya fault (SF) is neglected. Be, Ba, T: cities of Beirut, Balbeeck, Tripoli. See text for further explanations.

ROTATIONS IN LEBANON AND IMPLICATIONS ON THE DST

of the DST in Lebanon. Plate movement is accommodated by strike –slip motions along the Roum and Yammouneh faults, as well as local folding and thrusting (Fig. 11). Various amounts of total slip have been proposed for the Lebanese faults (Butler et al. 1998; Walley 1998). In our model the Roum and Yammouneh faults respectively absorbed 20 km and 50 km of the plate motion, movement onto the Sergaya –Rachaya fault being neglected. Slip on the Roum fault gave the thrusting and folding observed in the Tripoli area. Folding also occurred under the sea as a result of the transpressive character of the plate boundary owing to the obliquity of the YF. Owing to the angle between the DST in Lebanon and in Syria, the left-lateral movement implied rotation of the western block and could explain the 188 counterclockwise rotation determined in Lebanon. Important volcanism north of Lebanon and normal faulting in Galilee (Freund 1970) could be related to local extension related to this rotation. Whatever the mechanism of the regional rotation in Lebanon, our data exclude the possibility that a proto-YF was first aligned with the Arabia –Africa slip vector and that its present-day NNE– SW strike results from a later rotation. Such phenomenon would require a regional clockwise rotation of about 308, which is in contradiction with our palaeomagnetic data. The bend to the right of the YF relative to the Jordan and Gab segment is thus original. It highlights that transform faults may propagate with a trend oblique to the plate motion. The origin of the bend of the YF may result from several mechanisms. The crust may contain NE–SW low shear strength discontinuities, like faults, and strike– slip movement will thus preferentially occurred on these discontinuities. Such NE– SW faults are well-known in Israel and in the Sinai and are Late Triassic to Liassic in age. The scarcity of the Lower Jurassic outcrops and of deep drilling precludes evidence of the development of such faults in Lebanon. Alternatively, the bend to the right of the YF may reflect mechanical contrasts along the western Arabia crust.

Conclusion Palaeomagnetic data obtained in the Aptian –Albian formation clearly documents large counterclockwise vertical-axis rotations that affected the whole Lebanon during the Mesozoic times and after the Early Miocene period. The mechanism for these two events is probably very different. The second regional rotation is clearly connected with the relative displacement of African and Arabian plates. Its two different stages could reflect the effect of the locking, and later the unlocking, of the propagation of the DST to the north. Such a locking in the

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first stage highlights the key importance of inherited structures within a transform fault area. This work was supported by the programme ‘Middle East Basins Evolution’ Programme and the Lebanese University. We are very grateful to R. Hamzeh for help in the field and to G. Borradaile and A. Morris for detailed and useful review.

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tectonic evolution of the Levantine region. Tectonophysics, 298, 37–62. Westaway, R. 2004. Kinematic consistency between the Dead Sea fault zone and the Neogene and Quaternary left-lateral faulting in SE Turkey. Tectonophysics, 391, 203–237. Zijderveld, J. D. A. 1967. AC demagnetisation of rocks: analysis of results. In: Collinson, D., Collinson, W., Creer, K. M. & Runcorn, S. K. (eds) Method in Paleomagnetism. Elsevier, Amsterdam, 254– 286.

Revised stratigraphy of the Upper Cretaceous and Cenozoic series of Lebanon based on nannofossils ¨ LLER1*, FATHI HIGAZI2, WALID HAMDAN2 & MUSTAFPHA MROUEH2 CARLA MU 1

6 bis Rue Haute 92500 Rueil-Malmaison, France

2

Universite´ libanaise, Faculte´ d’Agronomie, B. P. 13-5368 Choorane, Beyrouth 1102-2040 Lebanon *Corresponding author (e-mail: [email protected])

Abstract: The aim of these investigations was to revise the geological map of Lebanon. Calcareous nannofossils were studied from a predominantly marly Senonian– Maastrichtian and Cenozoic series. About 900 spot samples were collected from exposed strata. This lithologically homogenous stratigraphic interval is poorly subdivided on the existing geological maps. The present study allows the precise determination of hiatuses and tectonic events. Palaeocene, Upper Eocene, Upper Oligocene, and Lower Miocene units were identified for the first time in Lebanon.

In the framework of a revision of the geological map of Lebanon, biostratigraphic studies of calcareous nannofossils of the Upper Cretaceous and Cenozoic series were carried out. About 900 spot samples, taken from exposed strata throughout Lebanon, were studied (Fig. 1). The sample localities are given in Table 1. The aim of this investigation was to date the series and to identify hiatuses. Several sections were sampled on both flanks of the Bekaa valley (Fig. 1). Age determinations and subdivision of the Upper Cretaceous are based on the nannoplankton range charts published by Thierstein (1976) and Perch-Nielsen (1985) as well as our own observations (Fig. 2). The zonations established by Martini (1971) and Martini & Mu¨ller (1986) were used to subdivide of the Cenozoic. The results are summarized in Figures 3 and 4. Areas of most intensive sampling are indicated in Figure 1. On the 1:50 000 maps (Dubertret 1945– 1953) and the 1:200 000 map (Dubertret 1955) the Upper Cretaceous is not subdivided. It was mapped as ‘Senonian’, often including Paleocene and Eocene deposits. This is due to the fact that the Senonian and the Palaeogene both consist of a predominantly marly series, with occasional intercalations of chert layers or nodules that make subdivision of the series practically impossible without dating. Ages and lithologies are summarized in Table 2. At the time, when Dubertret mapped Lebanon, microand specially nannofossil biostratigraphy were just beginning to be developed. However, Dubertret (1945–1953) subdivided the Eocene on the 1:50 000 maps in those areas where the Middle Eocene is partly characterized by limestones. This

subdivision was based on nummulites. He determined the Lower Eocene (e1 ¼ Ypresian) and the Middle Eocene (e2 ¼ Lutetian), which in certain areas can be separated into Lower Lutetian (e2a) marl and Upper Lutetian (e2b) limestone.

General geological setting A simplified geological map showing the main stratigraphic units is given in Figure 1 (after Dubertret 1955 and simplified by Walley 1998). Lebanon is constituted of two mountain chains – the Mount Lebanon in the west and the Anti-Lebanon in the east – both with a general NNE –SSW trend (Fig. 1). They are separated by the plain of the Bekaa syncline, which is mostly filled by Miocene to Quaternary continental sediments. Jurassic limestones and dolomites are exposed in the core of the mountain belts, overlain by Lower Cretaceous clastics (Barremian to Aptian) and mid-Cretaceous (Albian to Turonian), predominantly limestone (Dubertret 1975). The Upper Cretaceous (Coniacian to Maastrichtian), object of this study, is represented by marl and marly limestone deposited in an open-marine environment with varying amounts of chert layers or nodules. Regarding the distribution of Senonian– Cenozoic deposits in Mount Lebanon, it is possible to distinguish a northern and southern area. The separation lies south of Beirut and corresponds to a change in structural style of Mount Lebanon (Walley 1998). Based on the new datings of the present study, we can propose the following chronology of

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 287–303. DOI: 10.1144/SP341.14 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Generalized geological map of Lebanon (Dubertret 1955) simplified (Walley 1998) with indication of areas of main sampling (B.Q.F., Beit ed Dine–Qabb Elias Fault).

Table 1. Localities of the studied samples Longitude

338280 338560 338570 338180 338180 338190 338180 338280 338560 338190 338190 348250 348250 348250 338570 338560 338570 338560 338280 338280 348370 348370 348370 348300 348240 338520 348200 348200 348240 348250 338340 338340 338340 338340 338340 338350

358200 358370 358370 358170 358150 358150 358170 358200 358370 358160 358150 358550 358550 358560 358360 358360 358360 358370 358200 358200 348370 348370 348370 358570 358540 358530 358530 358530 358520 358520 358490 358490 358490 358490 358490 358430

25.4 92.3 56.2 31.7 59.5 03.7 09.4 21.8 57.5 37.1 01.7 23.0 46.4 10.0 01.0 59.3 00.6 58.8 25.4 25.4 04.7 04.7 04.7 30.0 52.0 09.9 10.2 29.3 59.3 03.6 30.3 58.1 58.8 59.2 59.3 16.4

08.0 32.5 35.0 00.1 39.2 42.1 48.0 03.9 08.3 07.9 38.9 51.0 48.7 24.0 50.1 54.8 59.4 02.7 08.8 08.8 04.7 04.7 04.7 42.9 31.0 56.0 15.7 37.6 08.5 05.9 59.1 09.1 09.0 08.9 08.3 23.7

Age E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc. E. Plioc.

Latitude 338270 338200 348260 348270 348270 338300 338270 338270 338270 338270 338260 338260 338260 338260 338270 338270 338160 348280 348280 338170 338280 338560 338300 338280 338280 338280 348260 348260 348260 348270 348270 348220 348220 348220 338560 348210

23.5 40.2 33.4 27.8 09.3 53.2 57.3 39.7 26.0 13.9 00.6 02.5 06.1 04.1 11.2 12.8 51.1 06.4 06.4 30.5 25.4 90.1 49.2 43.3 44.6 41.9 32.7 32.7 41.7 33.1 07.0 40.0 38.4 37.4 57.8 50.6

Longitude 358200 358180 358520 358530 358540 358220 358190 358200 358200 358200 358160 358160 358160 358160 358200 358200 358140 358190 358190 358140 358200 358350 358220 358210 358210 358210 358520 358520 358540 358530 358530 358460 358460 358460 358360 358540

11.3 07.0 54.3 39.4 01.9 28.8 52.6 02.5 07.7 36.8 40.9 41.5 42.9 46.9 56.5 36.4 47.9 33.8 33.8 53.7 08.8 65.5 28.5 00.9 01.3 00.5 54.6 54.6 58.8 53.7 55.5 14.6 15.5 13.0 58.2 36.5

Age M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc.

Latitude 338220 338200 338200 338570 338570 338570 338280 338280 338280 338280 338170 338190 338190 338190 338280 338280 338280 338570 338190 338280 338280 338280 338260 338280 338280 338260 338280 338280 338280 338280 338280 338270 338270 338270 338280 338170

21.4 19.3 23.6 17.7 07.9 42.2 43.6 42.2 42.2 43.0 42.5 44.1 45.1 43.5 43.2 40.9 43.2 17.7 31.6 26.1 28.3 30.6 13.6 43.4 43.3 08.8 41.5 44.1 42.8 40.2 15.2 59.0 55.7 29.2 27.9 44.1

Longitude 358160 358160 358160 358360 358360 358370 358200 358210 358210 358210 358150 358150 358150 358150 358200 358210 358210 358360 358160 358210 358210 358210 358210 358210 358210 358170 358210 358210 358210 358210 358220 358210 358210 358220 358210 358150

19.4 30.5 50.3 54.7 52.1 09.2 59.8 00.4 00.4 16.3 42.7 56.2 53.9 54.0 57.8 08.7 11.8 54.7 11.0 17.6 19.6 21.6 25.2 14.3 13.9 51.9 24.1 27.0 32.1 40.4 13.7 39.0 43.1 33.8 58.0 45.7

Age L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc.

289

(Continued)

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

Latitude

290

Table 1. Continued Longitude

338290 338350 348250 348250 348250 348250 338270 338280 338300 338270 348270 348270 338310 338310 338310 338310 338310 338310 338270 338280 338210 338210 338200 338250 338170 338280 338270 338280 338280 338280 338280 338280 338280 338280 338280 338280 338280 338240 338100

358200 358450 358500 358500 348500 358500 358200 358190 358220 358200 358550 358530 358220 358220 358220 358220 358230 358230 358200 358210 358160 358170 358180 358160 358160 358210 358210 358210 358210 358210 358210 358210 358210 358200 358220 358220 358220 358280 358130

09.9 30.4 29.0 29.0 29.0 29.0 26.3 14.5 46.0 05.0 14.8 07.0 48.8 45.9 36.1 38.0 21.8 04.4 13.4 44.2 34.7 08.5 51.6 59.7 52.6 04.3 58.6 11.6 11.3 26.2 24.0 15.4 24.0 18.4 15.1 11.0 02.7 04.2 40.1

27.6 13.3 52.0 52.0 52.0 52.0 12.6 71.4 30.6 48.0 25.4 55.5 27.6 30.8 28.6 38.8 03.7 17.1 09.3 40.8 29.7 39.0 01.5 49.4 03.4 36.9 43.0 44.3 48.2 51.3 56.1 54.7 55.6 04.5 14.3 16.1 54.4 43.0 40.9

Age E. Plioc. E. Plioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. M. Mioc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L, Eoc.

Latitude 338180 338230 338190 338190 338190 338200 338310 338170 338200 338570 338280 338280 338280 338280 338280 338300 338300 338300 338300 338240 338240 338230 338230 338230 338320 338220 338220 338310 338310 338300 338300 338300 338290 338300 348230 338210 338190 338120 338140

31.4 57.0 37.1 51.1 45.4 19.3 07.7 29.2 18.6 70.9 45.7 43.0 37.0 47.6 08.5 41.2 39.8 40.9 44.1 04.5 06.6 44.0 32.6 28.5 08.0 25.0 23.4 58.1 49.5 51.1 32.5 15.4 34.3 25.1 00.5 41.6 26.3 02.8 29.4

Longitude 358170 358380 358160 358150 358150 358170 358230 358150 358160 358370 358210 358210 358210 358210 358210 358220 358220 358220 358250 358170 358170 358160 358170 358160 358170 358170 358170 358220 358220 358230 358240 358240 358220 358210 358460 358160 358250 358250 358220

00.4 35.8 07.9 13.8 32.4 00.2 21.5 10.2 56.2 13.5 10.4 11.4 02.0 18.3 24.9 22.4 23.3 31.8 21.2 03.1 08.3 58.9 03.2 59.7 37.4 55.4 49.7 35.9 27.6 19.6 10.2 11.5 54.9 54.9 59.4 23.4 27.6 08.7 25.6

Age M. Mioc. M. Mioc. E. Mioc. E. Mioc. E. Mioc. E. Mioc. E. Mioc. E.Mioc. E, Mioc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. L. Oligoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc.

Latitude 338170 338170 338170 338160 338160 338300 338300 338200 338200 338210 338280 338280 338170 338170 338170 338160 338100 338270 338280 338270 338270 338200 338200 338200 338210 338210 338210 338510 338160 338280 338280 338570 338570 338570 338570 338570 338570 338500 338500

52.8 31.9 25.5 45.1 59.6 45.3 41.5 19.7 27.7 07.9 40.3 41.5 42.5 42.5 47.6 23.7 42.7 22.7 40.9 27.7 16.3 30.6 19.4 47.4 15.4 17.5 19.2 29.8 18.7 39.7 35.6 34.6 34.2 33.4 32.6 52.1 40.2 05.4 08.2

Longitude 358160 358180 358170 358170 358170 358220 358220 358160 358170 358170 358210 358210 358150 358150 358150 358170 358130 358220 358210 358220 358220 358250 358250 358260 358280 358280 358280 358530 358240 358210 358210 368010 368010 368010 368010 368010 368010 368020 368020

07.0 28.5 31.7 22.0 18.3 12.4 10.8 01.5 02.2 38.1 18.9 32.5 42.7 42.7 55.1 15.7 33.1 31.3 18.9 31.7 35.0 00.8 18.7 35.3 50.0 49.2 54.3 37.1 45.8 51.2 50.4 34.8 33.6 32.3 30.9 16.1 42.9 34.2 31.1

Age L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. L. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc.

¨ LLER ET AL. C. MU

Latitude

69.2 54.8 34.3 11.8 04.9 02.3 03.8 08.0 08.0 20.6 28.0 47.7 32.5 28.7 04.2 00.5 55.2 10.6 23.6 21.2 08.8 06.7 20.1 48.3 17.8 11.6 02.2 46.7 41.5 43.3 43.9 46.5 13.1 37.4 09.8 69.2 03.7 95.0 39.5 03.1 15.6

358370 358160 358150 358460 358460 358460 358450 358450 358220 358220 358220 358160 358180 358170 358160 358170 358170 358180 358170 358150 358180 358180 358180 358160 358170 358530 358530 358130 358130 358130 358130 358130 358140 358130 358140 358370 358160 358460 358230 358200 358200

28.0 52.3 85.9 70.6 43.6 23.9 82.3 58.1 47.4 05.7 03.9 26.4 48.2 13.1 50.6 52.2 57.5 45.5 58.9 52.1 55.5 56.1 30.7 24.7 15.5 16.3 28.8 29.8 30.1 38.4 37.6 50.8 06.4 58.9 31.7 28.0 43.5 91.7 21.2 47.9 56.2

M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc.

338140 338140 338120 338120 338120 338120 348230 338320 338270 338260 338250 348050 348010 338260 338190 348210 348210 348210 348210 348210 348210 338580 338580 338580 338580 338580 338580 348580 338260 338230 338200 338200 338270 338200 338270 338330 338330 338330 338330 338200 338120

20.4 43.1 09.7 07.2 03.1 21.4 16.5 42.1 46.4 21.2 17.9 13.4 49.5 04.2 33.0 07.0 08.0 29.0 38.0 38.0 38.0 15.6 15.6 15.6 15.6 15.6 45.4 45.7 55.2 50.7 53.6 49.6 01.6 40.0 25.1 06.2 08.2 14.8 21.8 19.5 23.3

358200 358180 358260 358260 358290 358290 358470 358270 358380 358180 358170 368150 368110 358170 358160 358460 358470 358470 358460 358460 358460 368120 368120 368120 368120 368120 368110 348110 358230 358280 358270 358270 338200 358250 358220 358420 358420 358420 358420 358230 358250

20.4 11.4 25.5 31.4 00.2 03.0 37.6 06.3 13.6 32.9 58.7 09.2 33.2 38.8 43.9 41.8 00.0 50.0 55.0 55.0 55.0 03.5 03.5 03.5 03.5 03.5 45.3 45.2 25.6 03.8 06.4 01.6 36.1 34.4 36.6 54.5 57.4 57.9 58.6 02.6 45.0

M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc.

338500 338210 338210 338210 338210 338210 338210 338190 338190 338180 338180 338150 338130 338130 338120 338300 338270 338260 338260 338270 338140 338140 338170 338280 338270 338270 338270 338280 338280 338270 338270 338270 338400 338400 338360 348180 348170 348160 348160 338270 338270

15.4 41.6 41.6 41.6 41.6 39.4 38.9 29.1 29.8 35.3 07.2 30.1 50.7 26.6 18.2 10.7 31.2 56.8 28.1 10.0 08.0 01.2 12.2 32.5 33.3 26.8 55.9 17.4 01.5 51.7 24.5 30.5 65.0 62.9 95.0 08.8 11.5 65.8 05.2 54.1 48.6

368020 358160 358160 358160 358160 358160 358160 358250 358250 358250 358250 358250 358240 358240 358240 358270 358180 358180 358180 358180 358150 358150 358190 358210 358230 358230 358220 358220 358230 358200 358220 358230 358520 358520 358500 358420 358420 358410 358410 358230 358230

15.2 22.8 22.8 22.8 22.8 22.6 23.2 04.6 11.9 39.9 38.5 09.2 58.9 25.1 37.8 43.8 57.3 55.1 39.7 56.1 09.5 44.0 56.0 17.5 18.9 30.2 06.9 05.1 04.4 59.0 10.6 29.5 60.8 56.4 36.0 36.2 01.4 53.3 69.5 10.9 16.1

M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. M. Eoc. E. Eoc. E. Eoc. E. Eoc. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal.

291

(Continued)

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

338570 338230 338200 348210 348210 348210 348210 348210 338280 338280 338280 338170 338240 338240 338240 338220 338210 338210 338220 338200 338270 338270 338260 338170 338160 348200 338420 338100 338100 338100 338100 338100 338110 338110 338120 338570 338230 348220 338270 338300 338290

292

Table 1. Continued Longitude

338340 348220 348230 348200 348200 348200 348200 338220 338230 338220 338220 338220 338360 338360 338370 338380 338380 338380 338380 338240 338240 338560 338560 338570 338150 348200 348160 348160 348160 348160 348180 348170 338420 338420 338420 338420 338420 338200 338230

358230 358460 358470 358470 358470 358470 358460 358360 358360 358350 358360 358350 358460 368460 358500 358510 358510 358510 358510 358170 358160 358360 358360 358370 358170 358530 358400 358410 358410 358410 358460 358480 358530 358530 358530 358530 358530 358150 358170

53.0 54.4 11.1 18.4 11.0 47.8 47.6 56.1 03.5 15.6 15.4 43.2 13.9 16.8 36.3 31.6 30.9 25.2 26.2 58.7 05.8 51.8 57.8 13.2 23.1 13.8 26.4 29.7 30.6 33.6 36.2 55.7 18.9 11.8 08.1 06.8 09.9 30.8 26.3

34.7 49.3 29.0 05.2 22.4 08.8 57.1 25.0 13.8 56.4 05.7 58.9 15.9 16.8 20.4 52.7 49.7 38.6 34.2 28.2 30.4 53.3 50.3 02.2 42.8 14.7 58.3 00.3 01.2 04.2 20.3 41.8 32.7 39.1 51.1 52.3 52.5 24.0 03.5

Age E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc.

Latitude 338330 338330 338330 338450 338480 348040 338560 348200 338240 338350 338340 338340 338580 338580 338580 338220 338280 338350 338350 338350 338350 338350 338350 338350 338350 338350 338350 338350 338350 338350 338510 338330 348170 348170 348210 338280 338300 348010 348010

39.9 03.0 13.4 41.0 56.9 54.6 57.2 04.8 05.6 06.9 53.6 51.1 15.6 15.6 15.6 10.7 25.0 31.0 26.1 26.4 26.1 25.6 27.3 27.4 27.4 28.9 29.1 30.0 29.9 30.5 32.3 29.0 12.9 12.9 07.0 50.0 05.6 50.0 50.0

Longitude 358250 358250 358260 358570 368020 368160 358360 358530 358160 358230 358230 358230 368120 368120 368120 358280 358200 358480 358450 358450 358450 358450 358450 358450 358450 358450 358450 358450 358450 358450 358530 358200 358410 358410 358460 358200 358280 368110 368110

23.3 48.1 05.0 11.4 05.0 11.6 49.5 04.0 30.2 45.8 34.8 34.1 03.5 03.5 03.5 41.4 48.8 20.6 55.5 54.4 50.6 47.4 47.8 36.0 33.7 28.4 26.5 22.7 22.9 21.2 28.0 38.5 48.8 48.8 41.8 48.8 08.0 46.4 30.4

Age E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc.

Latitude 338270 338270 338290 348200 348200 358200 348200 348170 338250 338250 338240 338240 338230 338220 338360 338370 338380 338250 338560 338560 338120 348270 348200 348200 348200 348170 348170 348160 348160 348180 338420 338420 338330 338330 338330 338340 338340 338220 338310

39.5 39.5 04.2 42.4 39.7 32.2 38.4 00.3 20.8 42.8 45.0 22.3 47.3 59.8 06.0 01.6 38.3 10.8 56.3 32.8 05.1 12.8 07.4 08.9 45.9 35.2 33.1 58.5 58.9 53.7 07.5 01.8 38.6 36.7 34.2 01.6 02.8 59.1 28.0

Longitude 358230 358230 358220 358460 358460 358450 358450 358410 358400 358400 358390 358390 358390 358360 358460 358500 358520 358190 358360 358360 358180 358550 358530 358530 358540 358420 358420 358420 358420 358450 358540 358540 358480 358480 358480 358480 358480 358160 358230

21.2 21.2 29.7 44.4 42.6 46.2 08.5 53.7 26.9 34.3 49.8 34.3 32.9 24.3 08.2 25.2 05.0 52.5 56.1 38.3 18.7 38.3 16.3 15.2 20.0 12.1 09.7 11.6 12.6 50.7 02.2 12.1 37.9 38.6 40.8 48.3 51.0 23.6 25.0

Age L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal.

¨ LLER ET AL. C. MU

Latitude

358170 358460 358470 368010 368010 368010 368020 358420 358420 358420 358250 358220 358220 358230 358230 358230 358230 368010 358570 358420 358430 358420 358470 358470 358470 358230 358250 358250 358240 358230 358260 358260 358310 358300 358250 358250 358230 358540 358360 358230 358230

03.5 54.4 26.7 28.9 27.7 36.9 37.4 58.1 58.1 55.7 17.0 14.2 53.2 04.1 04.3 23.1 14.0 21.7 46.2 59.1 03.5 28.3 15.6 16.8 17.7 55.1 39.7 08.4 22.4 20.1 09.7 09.7 01.5 52.1 35.0 45.0 29.7 14.6 40.0 20.1 43.3

E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. E. Eoc. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal.

338280 338400 338380 348200 348200 348210 348210 348170 348170 348170 348370 338300 338530 348200 338290 338330 338340 338310 348170 338240 338240 338250 338300 348330 348200 348170 348160 338330 338190 338350 338290 338280 338300 338570 338570 338570 338570 338550 338330 338190 338100

30.9 51.0 51.0 95.9 90.5 18.1 04.1 90.1 90.5 98.3 04.7 51.8 58.0 46.2 45.1 35.3 22.5 52.6 69.7 26.4 24.7 58.6 51.8 04.0 44.1 04.1 57.7 59.2 35.4 11.5 53.4 56.7 08.3 28.6 30.3 30.2 26.1 39.1 03.6 44.8 13.1

358080 358520 358510 358450 358450 358450 358440 358420 358420 358420 348370 358460 358570 358540 358210 358240 358230 358250 358420 358390 358390 358220 358460 368460 358540 358420 358420 358480 358150 358250 358240 358230 358220 368010 368010 368010 368010 368010 358420 358230 358310

30.0 68.9 81.1 64.2 54.1 04.2 88.4 28.2 25.7 35.6 04.7 47.3 46.2 19.8 59.6 02.6 20.4 30.5 39.9 21.6 21.6 02.0 47.3 55.2 15.6 13.0 15.4 57.7 43.5 24.3 19.4 53.7 13.1 18.3 15.5 12.6 12.2 28.6 28.3 54.2 58.4

E. Eoc. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. E. Pal. E. Pal. E. Pal. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr.

338330 338330 348140 348160 338300 338300 338300 338300 338290 338300 338100 338230 338270 338330 338120 338120 338120 338580 338580 338580 348270 348200 338200 348200 348270 348190 338280 338110 338460 338290 348200 348160 338280 338230 338350 338520 338520 338570 348170 348170 348170

47.7 44.2 50.9 50.0 56.6 56.6 57.8 18.8 44.5 00.2 56.9 11.4 51.8 55.9 34.0 13.6 13.6 15.6 15.6 15.6 15.0 15.3 45.2 05.3 17.3 52.7 28.6 46.5 57.1 48.5 15.9 63.4 30.3 43.8 38.7 55.7 08.5 62.7 75.7 03.6 09.7

358230 358240 368240 358410 358230 358230 358230 358240 358220 358250 358140 358170 358230 358260 358170 358180 358180 368120 368120 368120 358550 358450 358250 358530 358550 358530 358220 358240 358590 358210 358450 358410 358220 358380 358450 358540 358530 538370 358420 358420 358420

53.6 48.9 19.5 41.2 20.0 21.2 02.4 31.8 46.7 17.0 16.1 15.4 35.7 20.1 45.9 01.5 01.5 03.5 03.5 03.5 13.7 26.8 01.8 17.8 06.5 01.4 19.2 54.4 27.0 55.5 34.6 73.2 21.0 40.3 16.0 42.7 02.9 37.8 35.2 59.0 48.7

L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. E. Maastr. E. Maastr. E. Maastr. E. Maastr. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. L. C. / E. M. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. Maastr. L. Camp. L. Camp. L. Camp. L. Camp.

293

(Continued)

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

338230 19.3 348220 55.8 348230 09.6 338570 32.7 338570 32.5 338570 43.5 338500 00.4 338330 04.2 338330 04.2 338330 04.5 338300 00.2 338300 03.6 338270 58.7 338270 49.4 338270 53.0 338270 52.2 338270 54.3 338570 30.5 338530 57.5 338330 22.8 338330 26.7 338330 03.6 338310 39.1 338310 38.2 338310 37.7 338200 10.0 338190 25.6 338120 12.3 338130 40.6 338140 17.9 338120 10.4 338120 10.4 338100 29.7 338100 35.5 338120 17.8 338120 23.3 338100 36.1 348270 04.7 338560 43.2 338330 40.6 338330 46.2

294

Table 1. Continued Longitude

338220 338330 338460 338460 338460 338460 338480 338570 338320 338330 348370 348370 348370 348370 348160 348160 348160 348170 348170 348170 348170 348170 348170 348370 348160 348160 348160 348250 348250 348170 348190 348190 348160 338220 348150 348150 348150 348150 348140

358270 358230 358580 358580 358590 358590 368100 368100 358250 358230 348370 348370 348370 348370 358420 358420 358420 358420 358420 358410 358410 358420 358420 348370 358480 358480 358480 358580 358580 358420 358450 358450 358480 358160 368240 368240 368240 368240 368240

21.8 40.6 44.0 42.4 53.9 49.5 02.8 53.8 07.3 36.4 04.7 04.7 04.7 04.7 55.7 55.7 55.7 00.3 00.3 09.6 09.6 04.5 11.4 04.7 48.5 49.9 51.7 27.6 31.1 41.7 36.1 17.6 05.3 42.3 10.3 13.3 13.3 08.9 51.2

24.4 20.1 43.8 45.0 08.1 09.0 39.8 10.4 13.9 27.6 04.7 04.7 04.7 04.7 11.2 11.2 11.2 11.2 11.2 59.3 59.3 10.8 13.1 04.7 42.0 39.0 37.9 12.0 09.5 20.5 09.2 37.1 03.6 18.1 53.6 53.6 53.6 55.5 25.7

Age L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Pal. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp.

Latitude 338100 338190 338310 338140 338130 338130 338130 338130 338130 338130 348160 348160 348160 348160 348170 348170 348170 348170 348170 348170 338250 348270 348170 348170 338210 338210 348200 348200 348200 348200 348200 348200 348200 348200 348200 338350 338350 348180 338220

36.1 03.0 11.4 13.6 31.8 31.8 22.1 20.1 16.1 14.2 51.1 51.1 57.7 57.7 04.2 17.0 23.0 44.3 50.9 58.2 51.5 27.2 56.8 51.9 21.0 30.7 0.05 12.8 13.6 13.5 14.2 15.5 16.7 20.2 22.7 39.3 39.4 43.4 36.5

Longitude

Age

358230 29.7 358170 57.9 358270 29.5 358140 48.6 358140 54.6 358140 54.6 358150 19.5 358150 17.6 358150 20.0 358150 24.6 358420 13.6 358420 13.6 358420 11.2 358420 11.2 358420 12.8 358420 16.1 358420 19.1 358420 27.0 358420 29.3 358420 31.3 358200 48.3 358540 26.8 358420 24.0 358420 22.9 358270 23.0 358260 17.3 358530 16.9 358530 38.8 358530 37.5 358530 36.1 358530 35.5 358530 34.6 358530 36.2 358530 36.5 358530 42.0 35845’ 14.1 358450 15.7 358510 45.2 358280 23.3

L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. L. Maastr. E. Maastr. E. Maastr. E. Maastr. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp.

Latitude 348170 348170 348170 348140 348150 348200 348200 348200 348230 338270 348180 338380 338270 338270 338160 338160 338570 348200 348200 348200 348170 348170 348160 348160 348170 348160 348160 348190 338330 338330 338190 348230 348160 338440 338440 338430 338300 338280 338250

13.0 20.9 23.5 84.8 60.2 37.7 37.7 37.7 13.2 27.6 43.4 47.4 01.7 00.2 51.6 37.6 38.0 05.7 43.3 41.8 41.4 06.2 31.6 44.6 17.0 44.0 53.1 31.1 33.9 46.0 57.4 00.6 01.4 48.2 02.9 08.3 11.4 34.9 58.5

Longitude 358420 358420 358420 358430 358400 358440 358440 358440 358470 358230 358510 358520 358200 358200 358210 358220 358370 358530 358540 358540 358490 358480 358490 358480 358500 358480 358480 358440 358480 358480 358160 368250 358400 358270 358270 358270 358220 358220 358220

46.2 39.4 33.3 00.7 07.5 70.7 70.7 70.7 40.0 38.9 45.2 17.9 52.2 54.6 08.8 19.8 23.5 16.8 23.3 23.4 21.4 50.7 11.6 47.1 03.9 48.5 35.0 27.5 54.4 58.1 09.8 09.2 36.5 44.7 18.3 13.2 02.4 45.7 15.7

Age L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp.

¨ LLER ET AL. C. MU

Latitude

51.6 51.6 01.6 14.1 10.9 09.5 40.3 41.5 39.1 40.9 15.6 52.2 24.4 36.4 51.6 18.4 25.3 25.3 52.5 39.4 06.2 08.9 46.5 00.1 51.6 58.8 20.0 26.9 45.4 47.8 35.1 35.1 03.0 11.5 57.8 19.4 48.1 49.5 10.4 15.6 03.3

368240 368240 358400 358420 358420 358420 358270 358270 358270 358270 358550 358240 358210 358210 338450 358200 358250 358250 358170 358230 358150 358150 358250 368160 358260 358260 358250 358250 358250 358240 358240 358240 358230 358290 358280 358120 358420 358420 358420 368120 358220

27.7 27.7 24.7 17.2 18.7 20.5 18.5 32.3 45.1 42.0 15.8 30.8 53.7 48.1 08.9 10.7 39.0 39.0 59.7 19.9 53.5 44.8 25.6 48.4 29.0 11.3 36.2 30.9 08.2 25.8 18.8 18.8 53.4 54.2 53.5 10.8 29.1 20.8 36.5 03.5 39.7

L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Camp. L. Sant.

338230 338250 338260 348150 348150 348150 348150 338570 338570 348170 348160 348160 348160 348160 348160 348160 348140 348140 348140 348160 348160 338280 338290 238280 338170 338190 338180 338180 338350 338240 338250 338260 338270 338270 338250 338110 338430 338170 348170 348170 338140

31.3 08.8 41.3 53.7 55.0 56.5 59.5 76.5 60.7 37.3 64.9 82.6 77.3 86.6 07.6 97.7 72.4 73.4 72.9 65.1 75.6 40.3 17.5 31.7 27.3 12.9 58.7 32.0 03.7 58.1 05.3 55.0 01.2 02.3 35.8 50.7 08.2 27.8 19.8 17.3 31.2

358250 358230 358220 358400 358400 358400 358400 358370 358370 358490 358420 358420 358420 358420 358420 358420 358420 358420 358430 358420 358420 358220 358230 358230 358140 358160 358160 358170 358240 358400 358400 358210 358210 358210 358210 358180 358270 358140 358490 358500 358200

21.9 05.1 07.2 22.8 23.8 24.4 25.5 91.1 50.6 99.8 61.4 82.5 69.4 67.2 69.7 67.3 91.4 86.1 10.0 05.9 02.2 35.9 31.7 39.6 44.2 36.2 09.3 50.8 02.9 22.5 17.8 20.5 26.6 43.6 54.2 16.8 13.3 46.3 58.0 01.0 11.3

L. Camp. L. Camp. L. Camp. E. Camp. E, Camp. E, Camp. E, Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. L. Sant.

338570 338570 338540 348160 338190 338180 338160 338330 338310 348370 348160 348160 348160 338230 338210 338200 338200 338270 338360 338300 338270 338110 348250 348220 338210 348250 338270 338230 338250 338260 338270 338290 338300 338130 338270 338270 338210 338570 348160 348160 338170

29.6 32.2 00.4 20.1 32.1 39.4 33.2 52.0 35.8 04.7 40.5 39.7 40.0 27.9 03.6 57.2 40.4 05.8 19.2 00.8 08.3 51.5 33.0 58.8 10.0 20.6 03.0 01.2 15.0 08.8 17.2 32.9 50.4 14.0 18.8 16.7 20.9 72.6 64.0 64.0 36.2

368010 368010 358570 358490 358250 358180 358150 358260 358250 348370 358420 358420 358420 358280 358260 358250 358290 358220 358450 358240 358210 358190 358580 368250 358300 358570 358210 358210 358220 358240 358230 358290 358260 358150 358220 358220 358250 358370 358420 358420 358160

09.1 05.2 48.0 11.2 41.9 15.1 45.4 27.0 31.8 04.7 37.6 31.1 30.1 00.4 06.8 43.6 05.5 33.4 52.8 45.2 21.1 17.7 30.0 18.7 08.7 43.6 45.6 04.7 54.8 08.8 49.1 42.4 25.9 22.3 55.3 55.7 57.2 91.7 64.4 64.4 56.6

E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E. Camp. E.Camp. Camp. Camp. Camp. Camp. Camp. Camp. Camp. Camp. Camp. Camp. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant.

295

(Continued)

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

348140 348140 348150 348170 348170 348170 338440 338440 338340 338340 348270 338290 338300 338300 338330 338200 338190 338190 338140 338100 338100 338100 338330 348050 338330 338330 338340 338340 338340 338340 338340 338340 338350 338290 338290 338130 348160 348160 348180 338580 338270

296

Table 1. Continued Longitude

338270 338270 338170 338170 338170 338220 338360 338360 338360 338360 338350 338380 338100 338100 338300 338300 338280 338280 338270 338270 338250 338250 338240 338240 338240 338240 338250 338220 338260 348160 348160 348160 338330 338310 338200 338190 338180

358230 358230 358140 358170 358170 358370 358450 358450 358450 358450 358450 358520 358140 358140 358360 358460 358250 358240 358230 358230 358270 358280 358290 358280 358280 358280 358230 358290 358210 358490 358490 358490 358430 358470 358190 358250 358250

07.6 05.5 30.3 31.9 33.2 12.8 52.2 46.4 44.5 32.9 46.5 17.4 54.3 43.8 43.1 49.1 25.5 21.5 50.8 14.4 48.1 04.8 51.6 42.5 32.1 27.2 03.3 48.5 55.7 05.1 05.6 05.6 42.2 36.3 38.4 14.6 38.1

11.8 16.6 48.4 26.2 09.6 57.3 54.2 59.4 54.3 56.9 27.8 21.7 36.7 32.8 43.9 46.9 13.2 50.6 54.8 36.2 26.2 12.8 08.5 41.5 39.6 41.9 02.1 27.2 21.7 06.2 05.5 05.5 04.4 25.5 46.9 49.6 59.5

Age L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant.

Latitude 338190 338580 338270 338260 338260 338270 338260 338260 338250 338250 338230 338230 338240 338260 338220 338210 338370 338370 348140 348140 348140 348140 338420 338200 338200 338200 348160 338230 338230 348250 348250 348260 348280 338510

29.7 15.6 12.2 56.2 85.6 25.6 38.0 13.4 11.9 08.4 57.8 45.0 24.3 32.7 03.4 55.4 00.4 14.7 29.9 08.6 74.1 95.5 29.1 08.6 10.5 21.2 62.5 46.1 27.1 53.2 56.3 19.6 31.0 51.6

Longitude 358160 368120 358230 358230 358230 358220 358230 358240 358240 358240 358260 358270 358270 358230 358260 358250 358460 358460 358430 358430 358430 358440 358270 358180 358180 358190 358420 358250 358250 358590 358590 368000 368020 358530

56.0 03.5 40.8 24.8 04.8 56.2 25.9 11.7 17.8 36.2 01.4 06.6 23.6 08.5 16.6 36.7 03.4 23.1 34.3 27.5 61.8 33.1 26.2 19.7 41.7 03.9 08.1 12.9 44.9 27.3 38.1 42.3 09.0 21.9

Age L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. Cenom. Cenom. Cenom. Cenom. Cenom.

Latitude 338170 338170 348270 348150 338410 338410 338260 348260 348190 338390 338390 338350 348270 338290 338280 338280

36.2 08.3 14.0 25.3 23.9 02.9 49.6 10.2 14.4 15.1 01.8 02.2 13.9 32.4 42.4 36.2

Longitude 358160 358170 358540 358470 358540 358540 358220 358590 358440 358260 358260 358270 358540 358240 358250 358250

56.6 19.9 44.6 59.8 15.6 01.2 37.3 26.4 11.4 25.4 46.0 17.6 44.2 35.1 38.1 45.7

Age L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant. L. Sant.

¨ LLER ET AL. C. MU

Latitude

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

297

Table 2. Age and lithology of the Upper Cretaceous (Senonian –Maastrichtian) and Cenozoic series Lower Pliocene

Lower Eocene

NN 12–NN 15 Sandy mudstones with ophiomorphs Yellowish marls with thin limestone layers Argillaceous marls

NP 10 – NP 13 Alternation of marls and calcareous marls White marls with chert layers and nodules Calcareous marls thinly bedded Marls and marly limestones with or without chert

Upper Miocene Continental-lacustrine light reddish to yellowish marls with pisolites and locally conglomerates (no nannofossil datings) Lower-Middle Miocene NN 5 Alternation of calcareous silt, marls and indurated marly limestones Limestones with some rare thin marl layers and a reefal nodular interval rich in Lithothamian and macrofossils White chalky marls with locally conglomerates at the base NN 4 Argillaceous marls with thin limestone layers

Upper Paleocene NP 5 – NP 9 Argillaceous marls thinly bedded bioturbated compact marls Argillaceous marls with thin chert layers Alternation of marls and marly limestones with chert Lower Paleocene NP 3 Grey argillaceous marls Cretaceous Maastrichtian

NN 3 Marls with very thin layers of limestone breccia

White chalky marls with more or less chert Very thinly bedded silificified limestones with minor marl layers Alternation of marls and marly limestones

Upper Oligocene

Upper Campanian

NP 24–NP 25 Bedded marls Compact marls

Alternation of marls very bioturbated with indurated calcareous marls Compact marls Alternation of marls and marly limestones with chert nodules

Upper Eocene NP 18–NP 20 Compact marls Alternation of marls and indurated calcareous marls Middle Eocene NP 17 Thick beds of indurated marly limestones with thin marl layers Hard limestones sometimes silicified with nummulites NP 16–NP 15 Marls with chert layers Alternation of marls and marly limestones Limestones with nummulites and few marl layers Compact marls rich in radiolarians NP 14 Compact marls Marls with chert layers or nodules Marls with limestone beds and chert

Lower Campanian White compact calcareous marls Marls with dark brown argillaceous marl layers Alternation of marls and calcareous marls Santonian Compact calcareous marls Argillaceous marls Alternation of marls and indurated calcareous marls

298

¨ LLER ET AL. C. MU

Fig. 2. Nannofossil datings obtained from the Upper Cretaceous of Lebanon.

deposition for the studied stratigraphic interval. The depocentre of the Senonian –Maastrichtian sediments is located NE of Ras Chekka in northern Mount Lebanon, where Maastrichtian or older strata are overlain by uppermost Palaeocene (zone NP 9) up to lowermost Middle Eocene (zone NP 14b) series. These are the youngest Palaeogene sediments identified in northern Mount Lebanon. Above a hiatus, representing the interval of late Middle Eocene to earliest Miocene time, Lower Miocene (Burdigalian) shallow-water limestone was deposited probably in a transgressive setting. These observations indicate a first inversion and elevation of northern Mount Lebanon at latest Maastrichtian to Early Paleocene time, and a second elevation pre-Early Miocene. Upper Oligocene deposits related to the important transgression

within nannoplankton zone NP 24 (Hardenbol et al. 1998) were found only locally in the area of the Kelb Valley north of Beirut (Fig. 1). This allows a more precise dating of the second elevation of northern Mount Lebanon around the Palaeogene – Neogene boundary. The elevated position of northern Mount Lebanon persists during Neogene time as indicated by deposition of about 300 m of shallow-water reefal limestone in the area of Tripoli and southward of the Kelb Valley, whereas around Beirut and further to the south the Lower –Middle Miocene sediments were deposited in a somewhat deeper environment. Also a more complete Palaeogene series, including the Middle–Upper Eocene and Upper Oligocene, have a wide distribution in southern Mount Lebanon. The important

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

transgression at the base of the Pliocene (Hardenbol et al. 1998) reached the coastal area of both regions.

Nannofossil studies

299

trifidum and Quadrum gothicum. Arkhangelskiella cymbiformis is common and of large size. The Upper Maastrichtian was determined by the occurrence of Lithraphidites quadratus and Micula murus within the uppermost part of the series.

Senonian – Maastrichtian The nannoplankton markers used for dating and subdivision of the Senonian– Maastrichtian series are given in Fig. 2. A distinct facies change in Lebanon took place between the Cenomanian–Turonian limestones, and the overlying predominantly chalky and marly series of the Senonian –Maastrichtian, separated by a hiatus including the time interval of at least Coniacian and Early Santonian. The Upper Santonian marls are the oldest biostratigraphic level dated by nannofossils overlying the ‘Middle Cretaceous’ limestone. This hiatus was described earlier by Dubertret (1955) based on macrofossils only. Our age determination is based on the presence of Lithastrinus grillii, Marthasterites furcatus, Reinhardtites anthophorus, Lucianorhabdus cayeuxii and Broinsonia parca expansa typical for the Upper Santonian. The sediments are rich in calcareous nannofossils that are often recrystalized and broken due to diagenesis. The series mainly consist of argillaceous to calcareous marls. The contact between the Cenomanian– Turonian and the overlying Senonian was sampled at different localities throughout Lebanon. The Santonian–Campanian boundary was determined by the first occurrence of Broinsonia parca. The Upper Campanian is characterized by the presence of Quadrum trifidum, Quadrum gothicum, Eiffellithus eximius and Reinhardtites anthophorus. In general, an increase through time in carbonate content can be observed resulting in a series of alternating light gray, strongly bioturbated marl, and marly limestone including chert layers or nodules. In northern Lebanon, a 1 to 2 m thick dark chert layer within the Upper Campanian has been observed in several places near Tripoli. It can be followed to the south into the area east of Ras Chekka, but it is missing in southern Mount Lebanon. The origin of the chert throughout the studied interval is considered to be biogenetic owing to early diagenesis of siliceous microfossils. The Campanian–Maastrichtian boundary is determined by the extinction of Eiffellithus eximius, Reinhardtites anthophorus and Broinsonia parca. Sometimes it is difficult to recognize the exact position of this boundary, because these species become very rare within the uppermost part of the Campanian. Maastrichtian sediments generally consist of white chalk. Nannofossils are broken and strongly destroyed by diagenesis. The Lower Maastrichtian is characterized by the presence of Quadrum

Cenozoic– Palaeogene Palaeocene. The Cretaceous –Cenozoic boundary was not recognized in any of the studied localities. A regional hiatus comprises at least the lowermost Palaeocene (NP 1–NP 2) and possibly the uppermost Maastrichtian (Fig. 3). The Palaeocene was not separately mapped by Dubertret (1955). The lower part of the succession consists of thin-bedded argillaceous marl overlain by compact marl or an alternation of marl and indurated marly limestone with chert layers or nodules. In the area of Sidon, the Upper Danian (NP 3), corresponding to a sea-level high stand according to Hardenbol et al. (1998), was dated by the presence of Cruciplacolithus tenuis and Chiasmolithus danicus, and the lower Upper Palaeocene (NP 5) by Fasciculithus tympaniformis and Chiasmolithus danicus. Zone NP 6 (Upper Palaeocene) was determined by the occurrence of Heliolithus kleinpellii. It is exposed on the eastern flank of the Bekaa valley. Uppermost Paleocene (NP 9) sediments are widely distributed in the northern and southern Mount Lebanon (Fig. 3) corresponding to an important global sea-level highstand (Hardenbol et al. 1998); nannofossils are common. The assemblages are characterized by Discoaster multiradiatus, Neochiastozygus junctus, Sphenolithus anarrhopus and Fasciculithus tympaniformis. At several localities in the field sediments belonging to zone NP 9 lie directly on the Upper Maastrichtian or older strata without any visual indication of a hiatus, which nevertheless represents a time interval of at least 8 Ma years. Eocene. The different Eocene lithologies are summarized in Table 2. Sediments of Early Eocene age (zones NP 10 –NP 14a) are widely distributed in Lebanon. The lowermost Eocene (zones NP 10 –NP 11) was only determined in samples from southern Mount Lebanon and in the Bekaa plain, where the Paleocene/Eocene boundary is marked by a limestone layer up to 2 m thick at the base of the Eocene, which can be easily followed in the field. This biostratigraphic interval is characterized by the occurrence of Marthasterites contortus, M. tribrachiatus, Discoaster diastypus and D. binodosus. In the same sediments reworked Upper Palaeocene species were observed. The Early Eocene zones NP 12 and NP 13 have a regional extension, they have been identified in northern and southern Mount Lebanon (Fig. 3).

300

¨ LLER ET AL. C. MU

Fig. 3. Palaeogene nannofossil zones and hiatuses determined in Lebanon.

Zone NP 12 is determined by the presence of Discoaster lodoensis and Marthasterites tribrachiatus. The Middle Eocene (Lutetian and Bartonian) is represented by the zones NP 14b–NP 17. In Lebanon, zone NP 14b is widely distributed. It is characterized by the occurrence of Rhabdosphaera

inflata together with Discoaster sublodoensis and corresponds to a sea-level highstand according to Hardenbol et al. (1998) (Fig. 3). The Middle Eocene marls (zones NP 15 –NP 16) are sometimes rich in siliceous microfossils (radiolarians). Zone NP 15 was identified by the occurrence of

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

301

Fig. 4. Neogene nannofossil zones and hiatuses determined in Lebanon.

Chiasmolithus gigas. An Upper Lutetian limestone unit was previously identified by Dubertret (1975) using nummulites. Because this limestone is often silicified, we could not recognize the nannofossils in this unit. However, our datings show that it overlies strata of different ages (Late Palaeocene NP 9 to Middle Eocene NP 16), thus indicating that it might be diachronous. This nummulitic limestone is missing in northern Mount Lebanon, where the youngest Palaeogene sediments were dated as NP 14. It also developed only locally in the central and southern part, with a greater thickness

(several tens of metres) in the southern Bekaa valley. The Upper Eocene as well as the Upper Oligocene, so far unknown in Lebanon, were dated at several localities in the south. Dubertret (1975) described a hiatus for this stratigraphic interval, which, however, seems to be restricted to the Lower Oligocene. Upper Eocene deposits (zone NP 18 –NP 20) cover wide areas in southern Mount Lebanon. The nannofossil assemblages are characterized by Chiasmolithus oamaruensis, Isthmolithus recurvus

302

¨ LLER ET AL. C. MU

and Helicosphaera reticulata. The sediments consist predominantly of white marl. Oligocene. Lower Oligocene sediments (zones NP 21 –NP 23) were not identified. The Lower Oligocene seems to be represented by a regional hiatus (Fig. 3). Upper Oligocene sediments (zone NP 24–NP 25) were sampled from several localities in southern Mount Lebanon (Table 2) and dated by the occurrence of Helicosphaera recta, Cyclicargolithus abisectus, Sphenolithus distentus and Sphenolithus ciperoensis. The northernmost Upper Oligocene exposure lies about 20 km north of Beirut in the area of the Kelb Valley. This stratigraphic interval corresponds to the important transgression within zone NP 24 according to Hardenbol et al. (1998).

Cenozoic – Neogene Miocene. In the southern and central part of Lebanon, Neogene series are more or less restricted to the coastal areas (Beirut, south of Sidon), whereas they have a wider distribution in the northern part. The oldest marine Miocene sediments in southern Lebanon were dated Early Miocene –Burdigalian (zone NN 3–NN 4) in age as shown on Fig. 4. Aquitanian series seem to be missing. Deposits of zone NN 4 have a very local distribution. The deposits are marl and alternation of marl and marly limestone within the upper part, which already belongs to the Langhian (zone NN5) as indicated by the presence of Sphenolithus heteromorphus and the absence of Helicosphaera ampliaperta. They have been deposited in a deeper environment (middle– outer shelf) with a shallowing upward trend during Late Langhian time, whereas in northern Mount Lebanon the Burdigalian is represented by up to 300 m of shallow-water massive limestone with intercalations of algal reefal limestone. They were not dated by nannofossils. However, within the upper part of these limestone a few marly layers occur, and were dated Langhian in age (zone NN 5). We interpret these differences in facies between the northern and southern parts of Mount Lebanon as related to a more elevated position of northern Mount Lebanon that has persisted at least since Late Palaeogene time. In southern Mount Lebanon the Langhian (zone NN 5) lie transgressive upon series of different ages (Campanian, Late Palaeocene, Early–Late Eocene, Late Oligocene). This unconformity and the occurrence of abundant reworked (older) species within Langhian sediments, which were not observed in the Burdigalian deposits, might indicate a tectonic event within the uppermost Burdigalian. Late Middle Miocene series have not been identified in the studied area, however, undated

continental sediments attributed to the Middle – Upper Miocene by Dubertret (1975) are distributed east of Tripoli and in the northern Bekaa valley, where they reach a thickness of several hundreds of meter. They consist of marl with algal pisolites, with intercalated conglomerates and layers of fresh-water limestone. The increasing thickness of the conglomerates (more than 200 m) in northern Lebanon might indicate an important uplift during this time in connection with tectonic activity along the Aqaba–Dead Sea rift system. Pliocene. The Pliocene transgression at the base of zone NN 12 (Hardenbol et al. 1998) is well represented south of Sidon, in the Kelb Valley (north of Beirut), as well as in the area of Tripoli. The silts and sandy silts are rich in nannofossils with discoasters and Amaurolithus delicatus. The Upper Pliocene was not identified.

Summary Dating and subdivision of this predominantly marly series without distinct lithological differences, of Late Cretaceous and Cenozoic age, is based on calcareous nannofossils. The Paleocene, Upper Eocene, Upper Oligocene, and Lower Miocene so far unknown in Lebanon, were identified. All Upper Cretaceous stages were identified, with exception of a regional hiatus comprising the Coniacian and Lower Santonian, which was recognized throughout the study area. Regional hiatuses were also identified at the top of the Maastrichtian to lowermost Paleocene, the Lower Oligocene, and the lower part of the Lower Miocene (Aquitanian–lowermost Burdigalian). All known Eocene nannofossil zones were identified in southern Mount Lebanon. Local gaps are related to palaeo-reliefs. Main regional transgressive cycles can be correlated with well-known global sea-level highstands (Hardenbol et al. 1998). Marine Miocene sediments were deposited from Late Burdigalian to Langhian times. The Lower Miocene (Burdigalian, NN 3), dated with confidence in the south, is the first transgression followed by the more extended Langhian transgression (NN 5) overlying strata of different ages. Above this, continental deposits are probably of late Middle to Late Miocene age. An important transgression marks the basal Pliocene, when the sea invaded the coastal areas of both regions of Lebanon. This study allows a more precise dating of the structural evolution of northern and southern Mount Lebanon. For support of this study we thank the University of Lebanon.

NANNOFOSSIL STRATIGRAPHY OF THE LEBANON

References Dubertret, L. 1945–1953. Carte ge´ologique au 50 000 de la Syrie et du Liban. 21 sheets and notes. Damas and Beirut, Ministe`res de Traveaux Publiques. Dubertret, L. 1955. Carte ge´ologique du Liban 1:200 000. Beirut (1955) Dubertret, L. 1975. Introduction a` la carte ge´ologique a` 1:200 000 du Liban.-Notes. Me´moire Moyen-Orient, 13, 345–403. Hardenbol, J., Thierry, M. B., Farley, T., Jacquin, P., De Graciansky, P. C. & Vail, P. 1998. Mesozoic and Cenozoic sequence chronostratigraphic framework of European Basins. In: de Graciansky, P. C., Hardenbol, J., Jacquin, T. & Vail, P. (eds) Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. SEPM Special Publication, 60, 3 –13.

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Martini, E. 1971. Standard tertairy and quaternary calcareous nannoplankton zonation. Proceedings of the 2nd Plankton Conference, Roma 1970, 2, 739 –785. Martini, E. & Mu¨ller, C. 1986. Current tertiary and quaternary calcareous nannoplankton stratigraphy and correlations. Newsletters on Stratigraphy, 16, 99–112. Perch-Nielsen, K. 1985. Mesozoic calcareous nannofossils. In: Bolli, H. M., Saunders, J. B. & PerchNielsen, K. (eds) Plankton Stratigraphy. Cambridge, 329– 426. Thierstein, H. R. 1976. Mesozoic calcareous nannoplankton biostratigraphy of marine sediments. Marine Micropaleontology, 1, 325– 362. Walley, C. D. 1998. Some outstanding issues in the geology of Lebanon and their importance in the tectonic evolution of the Levantine region. Tectonophysics, 298, 37–62.

Late Cretaceous to Cenozoic tectonic evolution of the NW Arabian platform in NW Syria ABDELKARIM AL ABDALLA1, ERIC BARRIER2*, ANIS MATAR3 & CARLA MULLER4 1

Department of Geology, Techrine University, Lattakia, Syria

2

ISTEP, Universite´ P & M Curie, UMR 7193 CNRS, 4 place Jussieu, 75252 Paris cedex 05, France 3

Department of Geology, Aleppo University, Aleppo, Syria 4

6 bis Rue Haute, 92500 Reuil-Malmaison, France

*Corresponding author (e-mail: [email protected]) Abstract: We associate a brittle tectonic analysis and a stratigraphic study of the NW Arabian platform in Syria in the northern Coastal Range, Lattakia basin, and Baer-Bassit areas. These complementary approaches allowed characterizing the tectonic and palaeostress evolutions of this region since the Late Cretaceous. In Mesozoic and Palaeogene, before the Arabia– Eurasia collision, essentially developed extensional tectonics. A major extensional phase, characterized by a NE–SW directed extension, was recognized during the Senonian. It is associated with the activity of the Euphrates graben. The Eocene– Oligocene period is marked by a north–south directed extension associated with minor east– west trending normal faults, associated with syn-depositional structures. The compressional deformation initiated at the end of Oligocene north of Baer-Bassit. The major phase of shortening is Early Miocene in age in NW Syria. The related brittle structures and folding were resulted from a 1108 –1358 oriented compression. During this event, the Baer-Bassit is thrusted over the Coastal Range platform along the SE vergent Lattakia thrust. This major thrusting induced the flexure of the Arabian platform and the formation of the Middle to Late Miocene Lattakia basin. From the end of the Miocene, and until Present, the region experienced a NNW–SSE directed regional compression. We also evidenced an east–west trending compression near the Dead Sea Fault (DSF) area, coeval with the NNW– SSE one, associated with the north–south folding of the Coastal Range. It likely corresponds to a stress-field deflection in relationship with the DSF activity. From latermost Miocene to Present, the left-lateral displacement along the northern segment of the DSF can be estimated to 30– 40 km, from the offset of the Early Miocene deformation front.

Northwestern Syria is located near the active northern and western boundaries of Arabian plate (Fig. 1a), that is, the Dead Sea Fault (DSF) to the west, and the Anatolia –Arabia collision zone to the north. This collision zone approximately extends over 3000 km between the studied area at the Syria –Turkey border in the west, to the Zagros belt to the east in Iran. The northern Arabian margin experienced the obduction of the Tethyan ophiolites onto the Arabian platform during the Late Cretaceous (Ricou 1971; Parrot 1977; Thuizat et al. 1981). The DSF is the transform boundary between Arabian and African plates (Garfunkel 1981; Hempton 1987; Walley 1988; Matar & Mascle 1993). This north–south 1200 km long strike– slip fault accommodated about 100 km of left-lateral displacement since Miocene time (Quennel 1958; Freund et al. 1970; Garfunkel 1981). It connects the Gulf of Aqaba to the south with the Karasu Rift in Turkey along the

east-Anatolian fault to the north (Sengor et al. 1985; Bozkurt 2001; Rojay et al. 2001). It is associated with several Pliocene– Quaternary pull-apart basins such as the Dead Sea or the Al-Ghab basin in Syria (Brew et al. 2001a). The inception age of the DSF is still under debate. For example, according to Freund et al. (1970), the fault initiated in Early Miocene time, while Steckler & Ten Brink (1986) argued for a Early–Middle Miocene age and proposed that the displacement significantly increased during the Pliocene–Quaternary period. According to Hempton (1987), the Arabian plate started to separate from the African plate as early as Oligocene time with the initiation of rifting in the Gulf of Aden and the Red Sea, while its northern margin was already colliding with the Eurasian margin in Zagros since the Late Eocene. Although the tectono-stratigraphic evolution of the Arabian platform in NE Syria has been studied using different approaches such as structural

From: Homberg, C. & Bachmann, M. (eds) Evolution of the Levant Margin and Western Arabia Platform since the Mesozoic. Geological Society, London, Special Publications, 341, 305– 327. DOI: 10.1144/SP341.15 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Geological context of NW Syria. (a) Index map of the studied area within the NW Arabian plate and situation of the main structural domains. Topography from GTOPO30; (b) Simplified geological map of northwestern Syria redrawn from the Geological Map of Syria (Ponikarov 1966). L2 (Lattakia-2), and F1 (Fdio-1) are location with depth of wells (see logs in Fig. 3). DSFS and EAF: Dead Sea and East Anatolian faults. Location of cross-section A–B (Fig. 2) is indicated.

geology and stratigraphy, there have been few attempts to analyse the stress regime succession (Matar 1990; Zanchi et al. 2002; Al-Abdalla 2008). In this paper, we aim to characterize the polyphase tectonic evolution of the NW Arabian plate in Syria since Late Cretaceous time, and more particularly during Neogene time. We examine the brittle structures that cut the Upper Cretaceous to Cenozoic sequences of the northwestern Arabian platform. Finally, we reconstruct the regional palaeostress and tectonic evolutions and integrate them into the geodynamical context of the Middle East.

Geological setting The investigated domain of NW Syria comprises three main sectors (Figs 1 & 2): (1) the Mesozoic – Cenozoic Coastal Range platform, gently dipping toward the NW, (2) the slightly deformed Middle–Late Miocene Lattakia basin and (3) the

Baer-Bassit region, characterized by the Late Senonian obduction of the Tethyan ophiolites. The northernmost segment of the DSF separates this domain from the stable Aleppo plateau to the east (Fig. 1).

Stratigraphy In the Coastal Range, the oldest sediments are the Norian–Rhetian platform carbonate (Fig. 3a) of the Jweikhate Formation (Mouty 1997) that crop out in the southern part of the chain. The oldest rocks exposed in the northern Coastal Range are Middle–Upper Jurassic limestone. They are uncomfortably overlain by the lagoonal to shallow marine greenish marls of the Aptian Bab Janneh Formation (Fig. 3a). This latter formation marks the beginning of the Cretaceous marine transgression that occurred in Albian time (Mouty 1967). The Cretaceous carbonate sequence of the Coastal Range platform consists of a 500 m thick alternation of marls,

Fig. 2. NW– SE cross-section through the NW Arabian platform in Northern Syria. The Baer-Bassit overthrusts the Coastal Range platform. The Middle–Late Miocene Lattakia basin seals the major Lattakia thrust. Location in Figure 1.

LATE CRETACEOUS TO CENOZOIC TECTONIC EVOLUTION OF THE NW ARABIAN PLATFORM 307

limestones and marly limestones deposited during the Albian to Coniacian –Santonian times (Ponikarov 1966; Mouty 1967; Filak 2002). Sedimentation during Senonian time includes widespread pelagic facies such as the Campanian marly deposits of the Shiranish Formation (Fig. 3a). An intra-Late Maastrichtian angular unconformity has been evidenced in ophiolite bearing sediments cropping out near Hafa town in the northern Coastal Range (Al-Abdalla 2008). This intra-formational unconformity is probably related to the flexure of the northern Arabian platform in response to the obduction of the Tethyan Ophiolites onto the northern Arabian margin during the Maastrichtian. The Palaeocene –Lower Eocene chalky deposits, less than 150 m thick, are locally preserved in the Coastal Range, while in the Baer-Bassit they are more than 600 m thick (Figs 3b & 4). This difference in thickness indicates that in Early Cenozoic time the Baer-Bassit was a basin with respect to the Costal Range platform. The Middle Eocene deposits of the Coastal Range display nummulitic shallow-marine facies, while in Baer-Bassit cherty marly limestones and chalks dominate. The Middle Eocene sequence is characterized by channels, intra-formational angular unconformities, and slump structures, which evidence syndepositional instabilities (Fig. 5). Upper Eocene deposits are less abundant owing to the later Early Neogene erosion (Fig. 4). A stratigraphic revision, based on nannofossil study, enabled us to evidence Oligocene sediments (mainly of Late Oligocene age, NP 24 –25). They consist in white marls with interbedded grained clastics derived from the erosion of the ophiolites and radiolarites of the Baer-Bassit allochtonous units. This sequence is particularly well preserved near the Archouk village, 17 km NE of Lattakia (Al-Abdalla 2008). The Oligocene deposits conformably overlay the Eocene chalky sequence. In the Neogene Lattakia basin, well data, as well as field evidences (Fig. 3b), show that beneath the Middle Miocene basal deposits, part of the Upper Cretaceous to Palaeogene sequence is missing, probably as a consequence of the Early Miocene erosion (Leonov 1985). Early Miocene sediments are rare in the studied region. Our new biostratigraphic data indicate that the Aquitanian deposits described by Ponikarov (1966) are Late Oligocene (NN 24 –25) in age (Al-Abdalla 2008). Only few Aquitanian (NN 1 and 2) and Burdigalian (NN 3 and 4) deposits have been evidenced and consist of marls intercalated in medium-bedded bioclastic limestones. The basal part of the Lattakia basin sequence is Middle Miocene in age (Langhian–Serravalian, NN 5– 6). These deposits rest unconformably on the Upper Cretaceous to Oligocene sequence. They consist in conglomerates, sandstones, marls

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Fig. 3. Stratigraphy of the Coastal Range and Lattakia basin. (a) Generalized stratigraphic column for the Coastal Range; (b) stratigraphic log of the Lattakia basin from well and field data (modified from Leonov 1985). Location of wells in Figure 1b. Two kilometres of Middle–Upper Miocene marine deposits have been penetrated during drilling in the Lattakia basin while these deposits are thinnned or missing in the Coastal Range and Baer-Bassit areas. The transgressive Pliocene sediments unconformably overlay all the older deposits.

Fig. 4. Deformations and erosion beneath the Lattakia basin. (a) Middle Eocene vertical strata in the Early Miocene shear zone near the Lattakia thrust (coastal cliffs, south of Lattakia); (b) Middle Miocene carbonates of the base of the Lattakia basin sequence uncomfortably overlying the Turonian carbonates in the northern Coastal Range (Kfaryah Quarry, NE of Lattakia). The Senonian–Palaeogene sequence has been eroded during Early Miocene.

LATE CRETACEOUS TO CENOZOIC TECTONIC EVOLUTION OF THE NW ARABIAN PLATFORM 309

Fig. 5. Syn-depositional normal faults in the Middle Eocene chalky deposits (Railway cut, south of Lattakia). This pre-tilting conjugate system of normal faults is sealed by younger sedimentary layers (dashed line) of the same formation. Note the thickness change in syn-tectonic layers.

and bioclastic limestones (Safkoun formation) exposed near the Safkoun village, 20 km NE of Lattakia (Al-Abdalla 2008). The Lower Miocene sequence shows slumps, intra-formational unconformities, and chaotic layers. It contains boulders and fragments of older formations, the more recent being dated as Late Oligocene (NN 24– 25). The lowermost part of the Upper Miocene sequence (Tortonian) presents the same characters as the underlying Middle Miocene one. The alternation of Tortonian limestone and marly limestone beds (NN 9) observed beneath the Messinian gypsiferous deposits constitutes the top of the sedimentary pile of the Lattakia basin (Ponikarov 1966; Al-Abdalla 2008). Our datings show that the Lattakia basin was filled up during Middle– Late Miocene time. The transgressive Lower Pliocene (NN 12 –14) deposits unconformably overlay the Upper Miocene sequence and the older formations of the northern Coastal Range, Baer-Bassit, and Lattakia basin regions (Figs 1b & 3b). The shallow marine Pliocene deposits consist of clays, conglomerates, and alternation of limestones, marly sandstones and marly limestones. Their thicknesses vary from 100 m to 635 m (see Lattakia-2 well in Fig. 3b).

Structures The Coastal Range is a broad north–south trending asymmetrical anticline, delimited in the east by the DSF (Fig. 1). Its maximum altitude is of 1500 m. This 100 km long chain is constituted by a slightly deformed Mesozoic to Cenozoic sedimentary sequence. In the studied area, the northern Coastal Range

forms a monocline dipping toward the WNW of about 108 beneath the Neogene Lattakia basin. The NE –SW oriented Lattakia basin separates the Coastal Range to the east and the Baer-Bassit to the west. It is 15 –20 km wide and narrows toward the NE to 6–10 km. It further extends towards the SW into the Mediterranean Sea. The Early –Middle Miocene infill of the Lattakia basin rests unconformably on the Cretaceous –Palaeogene sequence of the northern Coastal Range in the south (Fig. 3b), and on the ophiolites and their Late Maastrichtian to Late Oligocene cover on the Baer-Bassit in the north. The stratigraphic and structural data, as well as the observations at cartographic scale (Ponikarov 1966), show that the whole Middle–Upper Miocene sequence of the Lattakia Basin is gently tilted towards the west, like the Mesozoic –Cenozoic platform constituting the north-western flank of the northern Coastal Range. The Lower Pliocene subhorizontal deposits (Fig. 1) unconformably cover this slightly deformed Middle –Upper Miocene sequence. The low angle angular unconformity that separates these two sequences implies that the north –south folding of the Coastal Range has already been achieved at least in Early Pliocene time, and possibly before the deposition of the Messinian evaporites. However, it has not been clearly established yet if the Messinian gypsum of the western Lattakia basin constitutes the top of the Middle–Late Miocene sequence filling up the Lattakia basin or the base of the overlying Pliocene sequence. These structural and stratigraphic data indicate that the north–south striking Coastal Range anticline is younger than the Middle Miocene Lattakia basin and formed before Early Pliocene time, and possibly before latest Miocene time. The Baer-Bassit platform is dominated by the Senonian ophiolite obduction process (Parrot 1977; Al-Riyami & Robertson 2002). The obducted ophiolites are remnants of Upper Triassic to Lower –Middle Jurassic oceanic crust of the NeoTethyan oceanic domain (Ponikarov 1966; Parrot 1977; Al-Riyami & Robertson 2002). The Maastrichtian ophiolitic me´lange, thrusted onto the Maastrichtian Baer-Bassit platform, is unconformably overlain by a Late Maastrichtian to Palaeogene marine sequence. The whole Baer-Bassit, including both the ophiolitic melange and its overlying sedimentary cover, were folded during the Neogene into NE– SW trending large open anticlines and synclines. Blankenhorn (1891) first proposed that the SE limit of the Baer-Bassit region is the Alpine deformation front. This hypothesis was later confirmed by recognition of a major SE-vergent thrust fault (Figs 2 & 4), the Lattakia thrust underlying the Lattakia basin. This latter cuts all the sedimentary sequence up to the Early Miocene (Al-Abdalla

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2008), and extends toward the NE in the Afrin region and offshore toward the SW. The Baer-Bassit is thrusted over the Coastal Range platform along this major thrust (Fig. 2).

sedimentary layers in all sites in order to reconstruct the palaeo-horizontal and thus the pre-folding/ tilting attitudes of faults. The majority of faults (80%) belong to Anderson-type conjugate patterns, facilitating field interpretation and data separation. One major problem in palaeostress reconstruction is the heterogeneity of the data, which commonly results from two or more successive tectonic events, with sometimes local stress permutations during the same tectonic phase. In these cases, we separated the data into mechanically homogeneous subsets of faults that fit with different palaeostress attitudes related to distinct tectonic events. Field observations related to crosscutting relationships between faults and folded strata are the main criteria used for ordering and relative dating of events, which are subsequently used for subset determinations. Crosscutting relationships between different brittle structures and successive striae observed on fault planes indicate fault reactivation under a new stress state.

Brittle tectonics methodology Analysis of brittle deformation is based on the study of fault –slip data and other related tectonic features, which includes millimetric to kilometric scale tectonic features, essentially faults, and, in some cases folds and joints. The common practice in these analyses involves data collection in the field, data separation, which means separation of data belonging to different deformation events, dating of events, reconstruction of stress configurations, and characterization of tectonic events and deformation phases (Figs 6 & 7).

Fault – slip measurements The majority of the brittle structures that we collected in the field comprises fault–slip data. These data include the orientations of fault planes and slickenside lineations, and sense of motion. We systematically measured the bedding attitudes of

Stress tensor determination Our palaeostress determinations systematically use the Direct Inversion method (INVD) of Angelier

N

(a)

N

(b)

64

32

(c)

N 33

Late Cretaceous

(d)

N 24

Palaeogene

(e)

N 7

Neogene

Fig. 6. Stress fields in NW Syria: statistical presentation. Top: (a) Rose distribution of the s1 stress axis trends for strike– slip and compressional stress regimes determined throughout the studied area; (b) Rose distribution of the s3 stress axis trends in extensional stress regimes throughout the studied area. Bottom: (c), (d) and (e) Rose distribution of the s3 stress axis trends for extensional stress regime determined in Upper Cretaceous, Palaeogene, and Neogene deposits, respectively. Number of determined palaeostress tensors is shown in each diagram.

LATE CRETACEOUS TO CENOZOIC TECTONIC EVOLUTION OF THE NW ARABIAN PLATFORM 311

Fig. 7. Distribution of the s3 stress axes related to the Late Cretaceous and Eocene– Oligocene extensional stress regimes. Pairs of hatched and black arrows indicate the s3 reconstructed stress axes for Late Cretaceous and Eocene– Oligocene extensional regimes in the studied area; site numbers refer to Table 1. Example of fault data and calculated stress tensor are shown in stereoplots (lower hemisphere Schmidt’s projections). For pre-tilting stress tensors, faults are shown both in their present-day and back-tilted position (the two stereoplots are linked by a curved arrow). Solid lines are fault planes; solid dots with arrows are striae; dashed lines are bedding planes; open dots are poles to bedding; stars with five, four and three branches are the principal palaeostress axes s1, s2 and s3, respectively; pairs of divergent large black arrows indicate the directions of extension (s3 directions). Detail of stress tensor parameters in Table 1. Rose diagram: distribution of the s3 stress axis trends determined in all the formations of the studied area; hatched and black arrows indicate the mean Late Cretaceous (N0308– N0458) and Eocene–Oligocene (N0008– N0158) directions of extension, respectively. The geology is redrawn and simplified based on the geological map of Syria of Ponikarov (1966).

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(1984, 1990). The principle of this method is based on the Wallace-Bott hypothesis (Wallace 1951; Bott 1959), which states that slip occurs along the maximum resolved shear stress on the fault plane. This implies that the observed slip lineations are parallel to the maximum resolved shear stress on the fault plane. The INVD method is based on the minimization of the angle (a) between the real striae (s) and the calculated relative shear stress (t). The method seeks to find the best fit between observed directions and senses of slip on faults and theoretical shear stress from which a reduced stress tensor is determined. The method allows reconstructing the stress configuration (or palaeostress) that prevailed at the time of faulting. An important parameter in fault –slip data inversion is the minimum misfit level required for defining acceptable agreement between the data and the calculated reduced stress tensor. In the INVD method, this parameter is the criterion RUP (Angelier 1990), which ranges from 0% (perfect fit) to 200% (total misfit). The highest bound involves maximum shear stress acting in the direction opposite to slip. In this study, we accepted numerical solutions with average value of the parameter RUP beneath 65%. A summary of palaeostress reconstruction is provided in Table 1, in which calculated reduced stress tensors are defined by direction and plunge of the principal stress axes (s1, s2 and s3), number of data (N ), ratio of stress magnitude differences (F), reduced misfit angle (a), quality criterion (RUP), and relative chronology between the reconstructed palaeostress tensor and the folding (i.e. pre-, post-tilting, or unknown).

Chronology and dating of the tectonic events In deformation zones, the sedimentary layers may have been subjected to faulting during and after the sedimentation, and/or before and after folding (Fig. 8). The stratigraphic ages of affected rocks and the regional and local unconformities are important criteria for dating the tectonic events. However, tectonic events cannot accurately be dated only by means of stratigraphic observations, especially for the polyphase deformation. It is thus indispensable to pay attention to geometrical and relative relationships between folds and brittle structures as well as to syn-depositional evidences. The relative chronology between faulting and folding leads us to distinguish pre- and post-folding fault populations, and consequently to determine a relative chronology. When the age of folding is well constrained, this relative chronology allows to clarify the age of tectonic events. Where the strata have remained horizontal, no relative chronology of faulting can be established with respect to folding. The differences between pre- and post-fold

faulting are evident where the strata is significantly tilted. For pre-tilting fault population, beddingparallel, or bedding-perpendicular striae can be observed on bedding-perpendicular inclined faults for strike –slip regimes (e.g. site 12 in Fig. 9), and normal and reverse regimes respectively. For posttilting fault population, horizontal striae can be observed on sub-vertical fault planes for strike – slip regime (e.g. site 32 in Fig. 9), and sub-vertical or oblique ones on various oriented fault planes for normal and reverse regimes, regardless to the dip of tilted strata. Sites 45, 52 (Fig. 7), and 27 (Fig. 10) present good examples of a conjugate normal fault systems in the present-tilt and backtilting positions. Where fault populations are clearly identified as pre-tilting, we systemically back-tilted the strata to reconstruct the pre-folding orientation of the stress axes in order to separate brittle structures that formed before and after folding. Several other criteria allow to date the tectonic episodes or to establish a relative chronology between the successive palaeostress fields and tectonic episodes. Angular unconformity allows dating tectonic phases at regional scale, more particularly the compressive ones. When a formation sealed a deformed older formation, it provides a time interval for the tectonic episode. This is the case for the folded Cretaceous to Lower Miocene deposits of SE Baer-Bassit, unconformably overlain by the sub-horizontal Middle–Upper Miocene sequence of the Lattakia basin (Fig. 2), providing a late Early Miocene age for this folding. The syn-depositional faults constitute one of the most efficient mean to date the extensional tectonic episodes, when accurate stratigraphic information is available. We used syn-depositional evidence within formation where normal faulting has been developed. South of Lattakia, abundant syn-depositional normal faults were measured in the Middle Eocene chalky deposits and are covered and sealed by layers of same age. These syn-depositional faults evidence a north–south extension of this age (Figs 5, 8 & 11). For the outcrops where the age of pre-tilting normal faults could not be constrained by syntectonic sedimentation, the stratigraphic age of affected rocks indicates a maximum age for corresponding extension. In the case of NE Syria, the tilted normal faults in the main folds have been developed prior to the major Neogene folding.

Palaeostress reconstruction in NW Syria Between 2004 and 2006, 1563 fault–slip data of 64% of normal, 30% of strike–slip, and 6% of reverse slip character were measured in 57 sites. Of these, 26 sites are located in the northern Coastal Range, 20 in Lattakia basin, and 12 in

Table 1. Palaeostress orientation from fault–slip data inversion Site name

1

Abo Said1 Abo Said2 Wadi Knd1 wadi Knd2 Machkas Machkita1 Machkita2 Blouran Om Al Tior1 Om Al Tior2 Brj Islam Shabtlia Shab2-1 Shab2-2 Brneh Sud Blor Cornish1 Cornish2 Ayob1 Ayob2 Ayob3 Archock Al Rabiah1 Al Rabiah2 Tarkia Laban1 Laban2 Laban3 E Ayob Zytoneh 4Archok 3Archok 1Jrymkia Jndairia1 Jndairia2 Jndairia3

2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24

Age Middle Eocene Middle Eocene Palaeocene Palaeocene Eocene Eocene Eocene Palaeocene Maastrichtian Maastrichtian Middle Miocene Middle Eocene Middle Eocene Middle Eocene Maastrichtian Maastrichtian Middle Eocene Middle Eocene Oligocene Early Miocene Early Miocene Oligocene Early Miocene Early Miocene Early Miocene Oligocene Oligocene Oligocene Early Miocene Middle Miocene Late Eocene Oligocene Middle Miocene Oligocene Oligocene Oligocene

N

s1 strike

s1 dip

s2 strike

s2 dip

s3 strike

s3 dip

V

a

RUP

30 13 11 7 24 27 18 24 12 5 24 42 10 8 13 18 12 8 17 9 11 18 23 11 19 37 12 17 10 9 18 27 17 8 8 5

168 65 151 135 313 4 275 33 10 137 13 108 92 229 185 344 310 300 300 119 296 91 129 186 9 24 359 305 338 251 54 171 28 4 39 159

1 82 74 10 86 87 76 68 87 11 78 78 68 17 84 8 13 2 74 7 1 17 5 74 78 80 80 52 87 7 81 87 74 72 83 4

78 284 269 328 210 200 24 230 168 244 192 257 258 331 17 232 211 31 201 211 26 357 4 10 206 222 197 112 207 160 201 20 223 120 274 47

13 6 8 79 1 3 4 21 3 57 12 10 21 34 6 71 37 12 3 16 2 11 81 16 12 10 10 38 2 12 7 3 16 8 4 78

261 193 1 225 120 110 115 138 258 41 282 348 350 118 287 77 56 199 110 5 171 236 219 279 115 132 106 207 117 10 291 289 132 212 183 250

77 5 14 2 4 1 13 6 1 31 0 6 5 51 1 18 50 78 15 73 88 69 8 1 4 3 3 6 2 76 5 1 4 16 6 11

0.39 0.38 0.45 0.3 0.3 0.41 0.51 0.42 0.39 0.18 0.27 0.36 0.11 0.08 0.34 0.33 0.03 0.36 0.36 0.36 0.5 0.65 0 0.39 0.34 0.34 0.35 0.82 0.25 0.52 0.32 0.34 0.42 0.42 0.11 0.26

20 12 5 20 13 9 11 18 10 12 19 16 17 5 7 18 14 8 10 6 7 19 11 22 24 16 9 11 20 16 21 10 17 16 29 12

40 32 20 61 40 19 31 49 20 14 54 42 39 61 27 50 46 23 29 30 27 40 42 54 46 42 29 46 60 44 55 29 41 35 53 38

Age/tilting

Stress regime

postdates postdates unknown unknown postdates unknown unknown postdates unknown unknown unknown unknown unknown unknown unknown predates predates unknown postdates unknown unknown unknown unknown unknown unknown unknown unknown unknown unknown unknown postdates postdates unknown unknown unknown predates

compression extension extension strike – slip extension extension extension extension extension compression extension extension extension strike – slip extension compression strike – slip compression extension compression compression compression compression extension extension extension extension extension extension compression extension extension extension extension extension strike – slip (Continued)

LATE CRETACEOUS TO CENOZOIC TECTONIC EVOLUTION OF THE NW ARABIAN PLATFORM 313

S

314

Table 1. Continued Site name

Age

N

s1 strike

s1 dip

s2 strike

s2 dip

s3 strike

s3 dip

V

a

RUP

25 26 27 28

Al Bida 2Jrymkia 2Archok 1Archok1 1Archok2 1Archok3 Kastal 1 Kastal 2 Archock3 Kalaa1 Kalaa3 Al Ashra1 Al Ashra2 Ain Al Tin1 Ain Al Tin2 Ain Al Tin3 EAin Al Tin1 EAin Al Tin2 Al Najiah Slma Zaynia1 Zaynia2 Al Mahalba Saladain Samar1 Samar2 Shblow1 Shblow2 Znbora2 Znbora3 1Al Hafa

Middle Miocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Maastrichtian Maastrichtian Late Cenomanian Late Cenomanian Late Cenomanian Late Cenomanian Late Cenomanian Middle Eocene Senonian Oligocene Oligocene Maastrichtian Turonian Cenomanian Cenomanian Maastrichtian Maastrichtian Albian Albian Maastrichtian

11 13 31 9 9 10 8 15 15 11 10 6 17 9 29 9 16 5 22 10 8 25 9 21 10 5 6 7 9 8 16

94 103 208 173 173 272 309 92 45 155 22 105 115 290 358 265 223 316 84 145 122 328 257 72 139 314 309 35 274 51 269

75 17 61 11 73 12 16 73 64 9 81 70 2 74 84 78 71 69 88 80 77 9 75 80 80 24 24 52 81 87 83

192 248 46 279 342 6 71 322 226 59 245 332 217 69 135 169 70 95 316 239 277 91 58 338 234 129 161 185 35 306 120

2 69 27 55 17 17 62 11 26 35 6 14 80 12 4 1 17 16 1 1 12 74 14 1 1 66 62 34 5 1 6

282 10 312 76 73 147 212 229 136 258 155 238 25 162 226 79 338 189 226 329 8 236 149 248 324 223 45 285 126 216 30

15 11 7 33 3 69 22 12 0 53 6 14 10 10 4 12 8 13 1 10 5 14 4 10 10 2 31 15 8 3 4

0.41 0.38 0.4 0.57 0.29 0.42 0.11 0.48 0.49 0.29 0.34 0.27 0.09 0.46 0.4 0.62 0.33 0.55 0.28 0.37 0.26 0.49 0.31 0.26 0.54 0.49 0.36 0.22 0.29 0.5 0.29

18 14 13 7 43 9 1 9 15 28 16 9 14 16 10 7 12 10 10 14 18 22 5 8 11 8 12 12 12 4 17

40 44 45 25 55 25 30 24 50 54 40 35 58 34 27 23 32 33 41 41 37 47 15 29 28 45 42 47 45 26 48

29 30 31 32 33

34 35 36 37 38 39 40 41 42

Age/tilting

Stress regime

unknown predates predates postdates unknown postdates predates unknown predates postdates postdates predates postdates unknown unknown unknown unknown unknown unknown unknown unknown unknown predates unknown unknown unknown unknown unknown unknown unknown unknown

extension strike – slip extension strike – slip extension compression strike – slip extension extension compression extension extension strike – slip extension extension extension extension extension extension extension extension strike – slip extension extension extension strike – slip strike – slip extension extension extension extension

A. AL ABDALLA ET AL.

S

46 47 48 49 50 51 52 53 54

55 56 57

Al Jsr Sarnah Slnslma1 Slnslma2 Kassatel Knsaba1 Knsaba2 Knsaba3 Ghsania1 Ghsania2 Lylon Maroniat1 Maroniat2 Domzine1 Domzine2 Domzine3 Ain Lylon Gassania Tran et1 Tran et2 Tran et3 Tran et4 Viarge1 Viarge2 Hossamo 2 Al Hafa1 2 Al Hafa2 2 Al Hafa3

Middle Eocene Turonian Late Cenomanian Late Cenomanian Maastrichtian Maastrichtian Maastrichtian Maastrichtian Maastrichtian Maastrichtian Early Cenomanian Maastrichtian Maastrichtian Maastrichtian Maastrichtian Maastrichtian Senonian Maastrichtian Turonian? Turonian? Turonian? Turonian? Cenomanian Cenomanian Late Cenomanian Senonian Senonian Senonian

18 26 27 14 12 12 13 8 4 9 10 19 4 10 5 10 18 9 27 17 20 12 10 6 11 8 9 8

272 274 81 241 293 290 142 286 325 352 338 329 298 47 112 333 216 176 292 94 68 60 318 13 78 300 119 97

84 80 8 82 12 16 76 72 6 73 3 75 14 83 0 72 73 69 13 84 76 70 23 72 69 15 79 71

125 128 203 117 79 106 267 121 58 189 177 131 153 217 203 98 120 20 56 334 172 237 122 130 322 102 212 253

5 8 75 4 76 74 8 18 24 17 87 15 74 7 64 11 2 19 67 3 3 20 67 9 9 74 1 17

35 38 349 27 201 200 359 30 223 98 68 222 30 307 22 191 29 287 198 244 263 328 226 223 229 209 302 345

3 5 13 1 8 1 12 4 65 5 1 4 9 1 26 14 17 8 19 5 13 1 6 16 18 5 11 7

0.21 0.38 0.36 0.22 0.42 0.31 0.24 0.45 0.01 0.26 0.3 0.5 0.47 0.22 0.43 0.67 0.29 0.22 0.25 0.39 0.5 0.37 0.58 0.25 0.37 0.42 0.15 0.46

20 12 13 15 9 6 14 19 11 17 12 12 10 15 7 19 11 15 8 8 11 13 8 10 13 9 14 8

49 33 28 53 31 29 31 40 27 51 28 43 39 45 32 38 31 50 24 28 30 36 40 44 51 30 36 25

unknown unknown postdates predates unknown predates unknown unknown unknown unknown unknown unknown predates predates predates unknown predates unknown unknown unknown unknown unknown unknown unknown unknown unknown predates unknown

extension extension strike – slip extension strike – slip strike – slip extension extension strike-slip extension strike – slip extension strike – slip extension strike – slip extension extension extension strike – slip extension extension extension strike – slip extension extension strike – slip extension extension

Abbreviations: S, Site reference; N, Number of fault –slip data. Strike and dip, trends, and plunges of principal stress axes s1, s2 and s3 in degrees (s1 . s2 . s3 . 0); Ratio of stress magnitude differences, V ¼ (s2 – s3)/(s1 – s3). RUP (%) is the criterion of quality ranging from 0% (calculated shear stress maximum, parallel to actual striae, and in the same sense) to 200% (calculated shear stress maximum, parallel to actual striae but opposite sense), acceptable with RUP , 65%. a (8), average angle between observed slip and computed shear stress (acceptable with a ,258).

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43 44 45

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Fig. 8. Extensional syn-depositional structures in Middle Eocene chalky deposits south of Lattakia. (a) Slumps and debris flows. (b) Syn-depositional conjuguate normal faults and horst. The sedimentary layer outlined by a dashed line, that seals some of the normal faults, is in turn affected by a normal fault of same orientation.

Baer-Bassit. From a stratigraphic point of view, 29 of these sites are distributed within Senonian– Palaeocene sequences, 20 of them within Eocene – Oligocene sequences, and nine sites are located within the Neogene sequence. The location and the geological position of each site are indicated in Figures 7, 9, 10 and 12. The analyses performed in 57 sites allowed to determine 96 reduced stress tensors (Table 1). The reconstructed palaeostress tensors reveal several stress orientations of which characterize 3 different types of stress regimes having principal stress axes as follows: s1 vertical and s2 and s3 horizontal in extensional regimes; s3 vertical and s1 and s2 horizontal in compressional regimes; and s2 vertical and s1 and s3 horizontal in strike–slip regimes. The extensional regimes are the most pronounced one, constituting the 65% of the 96 reconstructed stress configurations. Almost all the data characterizing these regimes are obtained from conjugate-like normal faults. The stress configurations characterizing strike–slip regimes constitute the 29% of the constructed tensors and include both conjugate-like strike–slip faults, and other strike–slip faults developed owing to reactivation of pre-existing normal faults. The remaining 6% of the stress configurations belong to compressional stress regimes characterized by bedding-parallel reverse slips. Despite careful examination, no relative chronological evidence between the reverse and strike– slip faults was found even in the same site, thus precluding to establish the relative timing of these two stress regimes. However, the stress configurations with strike –slip and reverse characters have commonly similar s1 orientations and the field observations were not conclusive to separate strike–slip and compressive stress regimes as

distinct tectonic phases. Therefore, we concluded that these two stress regimes are better explained by local permutations between the s2 and s3 axes during a single compressional event.

Tectonic evaluation of palaeostress configurations The tectonic events that have affected the Arabian platform in NW Syria before and during the collision between the Arabian and Eurasian plates have produced a wide range of mesoscale and regional structures including faults and folds. Analyses of these structures allowed us to propose tectonic models for the kinematic evolution of the region since the Late Cretaceous.

Late Cretaceous extensions The Upper Cretaceous sequence of the northern Costal Range is cut by numerous NW –SE normal faults. A statistical analysis of s3 directions from 64 palaeostress tensors from the Upper Cretaceous sequences is given in the form of rose diagrams (Fig. 6c). The figure shows a dominant direction of extension, N0308 to N0458 oriented. These directions were not encountered in the younger formations as indicated on the rose diagrams (Fig. 6d, e). This suggests a pre-Cenozoic age for the NE–SW oriented extension. In addition, this extension phase is mainly represented by conjugate systems of normal faults (see sites 45, 47 and 52 in Fig. 7) and in the case of sites having tilted bedding planes, these conjugate normal faults were formed systematically before the tilting (e.g. sites 45 and 52, Fig. 7). This tilting event is believed to occur during the Neogene folding event. Therefore, the

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Fig. 9. Distribution of the s1 stress axes related to the Early Miocene compression. Pairs of convergent black arrows indicate the s1 reconstructed stress axes for strike– slip and compressional regimes obtained in Mesozoic to Upper Miocene rocks throughout the studied area; site numbers refer to Table 1. Example of fault data and calculated stress tensor are shown in stereoplots: same symbol as in Figure 7; pairs of convergent large black arrows are the directions of compression (s1 directions). Detail of stress tensor parameters in Table 1. Rose diagram: distribution of the s1 stress axis trends determined in all the formations of the studied area; black arrows indicate the mean direction of the Early Miocene (N0308– N0458) compression. The geology is redrawn and simplified based on the geological map of Syria of Ponikarov (1966).

NE–SW extension phase took place before the Neogene. In the absence of syn-depositional evidence, it is difficult to date the extensional events by direct

observation. However, in the case of the NE–SW extensional event, both the relative geometries with respect to the folded strata, and the stratigraphic age of the faulted rocks where this direction

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Fig. 10. Distribution of the s3 stress axes related to the Middle Miocene extensional stress regimes. Pairs of divergent black arrows indicate the s3 reconstructed stress axes for the Middle Miocene extensional regime obtained in Mesozoic to Middle Miocene rocks throughout the studied area; site numbers refer to Table 1. Example of fault data and calculated stress tensor are shown in stereoplots: same symbol as in Figure 7. Rose diagram: distribution of the s3 stress axis trends determined in all the formations of the studied area; black arrows indicate the mean direction of the Middle Miocene (N1008– N1308) extension. The geology is redrawn and simplified based on the geological map of Syria of Ponikarov (1966).

were obtained, suggest that this pre-folding event cannot be either older or younger than the Senonian. This Late Cretaceous NE–SW extensional event has already been reported in the Bishri region and

NE Palmyrides (Salel 1993; Caron et al. 2000; Brew et al. 2001b). In the Euphrates graben, based on sub-surface data, the NE –SW directed extension is associated with Campanian syn-depositional

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Fig. 11. Syn-depositional extensional structures in the Middle Miocene sediments of the Lattakia basin. (a) Syn-depositional normal fault associated with slumps. The slumps and debris flows in the hanging wall are related to a palaeo-submarine normal faulting. The related fault–slip pattern and palaeostress tensor are shown and indicate a N1108 direction of extension; Stereogram: same symbol as in Figure 7. (b) Intra-formational breccia and conglomerates in the Middle Miocene sequence in the lower part of the Lattakia basin sequence. The interbedded chalky pebbles (white dashed lines) are Oligocene to Early Miocene in age. (c) Intra-formational angular unconformity (black dashed lines) associated with slumps in the Serravalian deposits (Lattakia–Aleppo road).

faults (Litak et al. 1998; Brew et al. 1999; Caron et al. 2000). Therefore, we concluded that the N0308 to N0458 oriented extensions determined in the Upper Cretaceous rocks of NW Syria probably correspond to the regional Senonian rifting that also gave way to the development of the Euphrates and Azraq grabens in the NW Arabian plate. Apart from NE –SW extension, a less pronounced NNW –SSE directed extension was also observed in the Upper Cretaceous deposits (Fig. 6c). This extension direction is characterized by N1508–N1658 trends of s3 axes and predates the Early Miocene folding. Likewise, it is also

missing in the Palaeogene and Neogene sites (Fig. 6d, e). This stress state is strictly restricted to the northwestern Costal Range. We assume that this local extension could be related to the flexing of the northern Costal Range platform during the Late Maastrichtian in response to the obduction of the Tethyan ophiolites onto the Baer-Bassit platform. This deformation is also evidenced by intraMaastrichtian unconformities and by change in thickness of the Palaeogene sequence from more than 600 m in the north near Baer-Bassit region to only several metres in the south in the northern Coastal Range.

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Eocene –Oligocene extension Numerous outcrops throughout NW Syria exhibit syn-depositional Eocene –Oligocene normal faulting (Fig. 8), which is spectacularly developed in the Middle Eocene sequence cropping out in the coastal cliffs south of Lattakia (Figs 5 & 8) and in the Wadi Kandil area. The faults are mainly pure normal faults organized in conjugate pattern (Figs 5, 7 & 8; see sites 1, 2, 17 and 47, in Fig. 7). They are commonly associated with intraformational unconformities and slumps (Fig. 8a). As evidenced by post-faulting tilting, these normal faults predated the Early Miocene folding and were extensively observed in the Eocene to Oligocene formations (Fig. 6d) of the northern Coastal Range and the Baer-Bassit region. The youngest deposits where these faults have been determined are of Late Oligocene age (NN 25). Some of these east– west oriented normal faults reactivated later during the Neogene compressions as right lateral strike–slip faults. The distribution of the calculated stress axes for the corresponding fault–slip datasets shows a major N0008–N0158 trend for the s3 stress axes after back tilting (Figs 6d & 7). Most of these palaeostress reconstructions were based on Eocene –Oligocene syn-depositional structures that are normal faults striking east– west to WNW –ESE (see sites 1, 2, 17 and 47 in Fig. 7). Such an orientation of s3 is not observed in the overlying Neogene sequence (Fig. 6e). We also measured consistent brittle structures at some Eocene – Oligocene sites where no evidence of syn-tectonic deposition could be identified. Here the direction of calculated stress axes for the corresponding back-tilted fault –slip data shows the same N0008 –N0158 direction of s3 stress axes in a pre-fold attitude. Such a similarity in the results led us to infer that, even in the absence of syndepositional evidence, these pre-folding normal faults can be considered as belonging to the Eocene –Oligocene extensional deformation phase.

Early Miocene compression The folding of the Mesozoic –Lower Miocene sedimentary sequence marks the Arabian platform in NW Syria. This phase of deformation is the major tectonic event in the investigated area. We refer to this compression as the NW-compression event hereafter. In Baer-Bassit, the Upper Cretaceous to Upper Oligocene sequence forms a series of broad NE–SW oriented anticlines and synclines (Fig. 2). In the south, the whole Baer-Bassit sequence plunges towards the SE (Fig. 2). Outcrops in the southern bank of the Techrine Lake (20 km NE of Lattakia) display a NE– SW trending shear zone parallel to the main fold axes of Baer-Bassit. A

similar folded and sheared zone crops out in the coastal cliffs south of Lattakia (Fig. 3a). This shear zone outlines an intra-platform, SE-vergent major thrust along which the folded Baer-Bassit platform overthrusts the Coastal Range platform. In the Techrine Lake area, this shear zone is unconformably overlain by the sub-horizontal almost undeformed Middle Miocene deposits of Lattakia Basin. These sediments postdate the main phase of thrusting. Our new biostratigraphic data show that the youngest sheared sediments are of Early Burdigalian (NN 3), while the oldest undeformed deposits of the Lattakia Basin are of Middle Miocene (NN 5–6) age. Thus, the end of this deformation event can be clearly constrained as post-Early Burdigalian and pre-Middle Miocene, more specifically as intra-Burdigalian. In order to deduce the regional average of the stress orientations for this phase we provided a rose diagram for the distribution of the reconstructed s1 axes (Figs 6 & 9), including both pre- and post-folding palaeostress configurations. The distribution of all s1 axes, in both compressional and strike –slip regimes, shows a major NW–SE to WNW –ESE direction compression with a dominant trend of N1108 to N1358. Most of the data for these s1 NW–SE trends (90%) were collected from the Late Cretaceous to Palaeogene sediments. Only two palaeostress tensors have been reconstructed in Lower Burdigalian strata in sites 13 and 15 indicating that this stress state is not older than Early Miocene (Fig. 9). At many sites, there are rightlateral strike –slip reactivations along N0908 to N1108 striking pre-existing planes of mainly inherited Mesozoic to Palaeogene normal faults (e.g. sites 40 and 50 in Fig. 9). As shown in Table 1, most of the faulting has occurred under a strike –slip fault regime (77% of the stress states). Data separation based on the relative chronology/geometry between the brittle structures and folded strata allowed us to reconstruct the pre- and post-folding attitudes of the computed stress trends. However, because the strata are generally slightly tilted in the studied area, we were able to define this relative chronology in only 35% of the cases (see sites 12, 32 and 50, Fig. 9). According to our data, almost all of the pre-fold faulting has occurred under a strike– slip stress regime indicating an average NW–SE trend for the s1 stress axis (e.g. sites 5, 12 and 50, Fig. 9). A distinct post-fold NW– SE compression has been determined in only one site (site 32, Fig. 9). The pre- and post-folding stress trends are almost perpendicular to the NE – SW direction of the main fold axes in the Baer-Bassit and the Lattakia basin, suggesting that the stress and strain were coaxial during the folding process. Therefore, folding and related brittle structures resulted from a NW–SE compressive stress trend.

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In addition, there is no evidence of the NW –SE compression in the younger sedimentary rocks (i.e. Middle Miocene to Pliocene). In NW Syria, the Early Miocene NWcompression event is the major regional tectonic event associated with folding and thrusting. Here, the Baer-Bassit platform overthrusts onto the Coastal platform along the Lattakia thrust. This major intra-Arabian platform structure constitutes the Alpine deformation front during the first stage of the Arabia –Eurasia collision.

Middle Miocene extension in Lattakia basin In Lattakia basin and surroundings, we determined 22 palaeostress tensors corresponding to a well defined WNW–ESE extension. The rose diagram shown in Figure 10 indicates that the s3 directions related to these extensional stress regimes range between N1008 to N1208. In the Middle Miocene sequences of the Lattakia basin, syn-depositional structures such as slumps, intra-formational breccia, and unconformities were encountered in several sites (Fig. 11). These features are associated with north–south to NE–SW oriented syn-depositional normal faults. These normal faults do not exist in the overlying undeformed Late Miocene and Pliocene strata and are contemporaneous with the filling up of the Lattakia basin during the Middle Miocene. In the Upper Cretaceous to Palaeogene rocks of the eastern Baer-Bassit and northern Coastal Range, the conjugate normal fault populations related to this extension are generally post-folding, as indicated by the attitude of the reconstructed stress tensors with respect to the bedding orientation (sites 3, 7, 13 and 20 in Fig. 10). This extension has been determined in the Upper Cretaceous, Palaeogene, and Lower Miocene rocks as well (Figs 6d & 9). It postdates both the SE-vergent Lattakia thrust and the major folding. This WNW– ESE extension is perpendicular to the general trend of both the Lattakia thrust and the axis of the Lattakia basin. In addition, it is restricted to the Lattakia basin and surroundings. We attributed the N1008 –N1208 directed local extension as a consequence of the flexing of the Arabian platform during the Middle Miocene. The thrusting of the Baer-Bassit onto the Coastal Range platform along the major Lattakia thrust at the end of the Early Miocene induced flexing of the Arabian platform and a local extensional stress field in the basin (Fig. 2). In contrary to Leonov (1985) that argued that major normal faults bounded the Lattakia basin during the Late Oligocene –Early Miocene time, we proposed a flexural origin for this basin because no such faults have been observed in the basin and surroundings.

Late Neogene NNW – SSE and east – west compressions The rose diagrams shown in Figures 6 and 12 present the s1 directions determined from the data collected from the Upper Cretaceous to Neogene sequences in the study area. In addition to the main NW–SE trend discussed above (53% of the compressions), the rose diagrams indicate two minor peaks trending NNW–SSE and east –west and representing 28% and 19% of the s1 axes, respectively (Figs 6 & 12). The east –west compression, although not described before in this region, is clearly documented by both strike–slip and reverse faults at 6 sites (sites 13, 14, 19, 26, 28 and 45, in Fig. 12; see also Table 1). It is almost perpendicular to the north–south structural trend of the Coastal Range. The second minor compression direction trends N1608–N1708 and was determined from the fault–slip data at 9 sites (e.g. sites 1, 11, 23, 36, and 49 in Fig. 12, see also Table 1). The analysed fault –slip data show that for both east–west and NNW–SSE oriented s1 axes, faulting has occurred under both the strike –slip and compressional stress regimes, as indicated by the attitude of the s1 axes and fault populations at sites 23, 26, 36, 45, 49, 55 and 11, 14, 19, respectively. This suggests that strike–slip and compressive regimes are linked by permutations of s2 and s3 axes. The Middle Miocene age of the youngest rocks affected by the east –west and NNW–SSE compressions provides a maximum age for these two compressions. Defining a minimum age is more difficult, because the overlying Upper Miocene to Pliocene deposits are scarce and poorly dated. Both compressions were not encountered in these latter deposits. Therefore, we carefully consider hereafter the relationships between faulting and folding in order to establish a relative chronology between the different palaeostress states and the folding episode. In the sites where the east– west and NNW– SSE compressions were documented, strata are generally slightly tilted (,158), like in the Lattakia basin area (Fig. 12). In these cases, the angular differences in the stress orientations calculated with faults in their present-day attitude and with back-tilted faults lie in the range of measurement uncertainties and dispersion, thus making it difficult to establish faulting –folding chronology in most sites. Nevertheless, we were able to establish a relative chronology between the faulting and the north– south folding of the Coastal Range at eight sites out of 15 where these events have been determined. For the east –west compression, the analysis of the strike –slip fault populations in sites 26 and 45 show that faulting has occurred before and after

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Fig. 12. Distribution of the s1 stress axes related to the Late Miocene to Pliocene compressions. Pairs of convergent black arrows indicate the s1 reconstructed stress axes for strike–slip and compressional regimes obtained in Mesozoic to Pliocene rocks throughout the studied area; site numbers refer to Table 1. Example of fault data and calculated stress tensor are shown in stereoplots: same symbol as in Figure 7; pairs of convergent large black arrows are the directions of compression (s1 directions). Detail of stress tensor parameters in Table 1. Rose diagram: distribution of the s1 stress axis trends determined in all the formations of the studied area; black arrows indicate the mean directions of the Late Miocene to Pliocene compression (N0908 and N1608–1708). The geology is redrawn and simplified based on the geological map of Syria of Ponikarov (1966).

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tilting, respectively. Such opposite chronologies were also unambiguously recognized at three sites for faults associated with the NNW–SSE compression. For instance, at site 11, located in BaerBassit (Fig. 12), a N1648 trending s1 axis has been determined from a pre-tilting population of left lateral strike –slip faults and of conjugate reverse faults, whereas a similar compression was documented by post-folding faults cutting the Upper Oligocene deposits at site 36 (Fig. 12). At site 28 (Fig. 12), the east –west and NNW–SSE compressions were recognized together, but without relative chronology, and both postdated the tilting. The results of our fault analysis and structural study show that the east –west and NNW–SSE compressions both (1) pre- and post-date the north– south folding of the Coastal Range and (2) are not older than Middle Miocene. The east –west trending reconstructed compressions can be considered as coeval with the north –south Coastal Range folding because (1) the Coastal Range fold axis and the east –west compression are almost orthogonal and (2) the east –west compression is both pre- and post-tilting, suggesting that it is synchronous with the folding of the Coastal Range. Thus, we assume that this folding, accurately dated postEarly Tortonian (NN 9) and pre-Early Pliocene (NN 15) from stratigraphic and structural data, dates the east –west compression. The NNW –SSE event was active before the north– south folding and had initiated in the Middle–Late Miocene. Such a NNW–SSE compression direction was also determined in NW Syria by Zanchi et al. (2001) and interpreted as a Late Miocene regional compressional event. It was observed in the Afrin region (M. Khatib 2008, pers. comm.), in Aleppo plateau where it is the major post-Middle Miocene compression (AlAbdalla 2008), and in the northern Arabian plate in SE Turkey (Kaymakci et al. 2006, 2009). According to the relative chronology with respect to folding, the NNW–SSE compression also continued after the folding process. Insufficient direct binary relative chronology between the NNW–SSE and east –west events did not allow us to establish a clear chronology between these events. Although the number of sites is too small to establish a reliable relative chronology between the different post-Early Miocene compressions, we may propose a model of tectonic evolution since the Late Miocene that accounts for our results. We suggest that the NNW– SSE and east –west compressions were coeval. Accepting this hypothesis, and considering the regional significance of the NNW –SSE compression (see above), a 408 clockwise rotation of the s1 trajectories occurred in the Middle–Late Miocene in relationship with the inception of the DSF. This NNW–SSE oriented

compression is active until Present. Within this frame, the east –west compression recognized in our study likely corresponds to a stress-field deflection in relationship with the DSF activity. This stress orientation is coeval with the north–south folding of the Coastal Range. The absence of the east–west compression far from the DSF supports this conclusion (Al-Abdalla 2008).

Syntheses and conclusions The brittle tectonic analyses presented in this study bring new information for reconstruction of structural configuration of the Arabian platform prior to, and during, the Arabia– Eurasia collision. The present-day tectonics of the northern Arabian plate is a consequence of the Late Cenozoic Neo-Tethyan closure and continental collision under a NNW– SSE to north–south oriented plate convergence (Kaymakci et al. 2006). In NW Syria, the regional angular unconformities and the facies change within the Neogene sedimentary sequence show that the folding process initiated in the northern Arabian plate in Early Miocene time. The deformation front propagated from north to south from the Arabian margin to the platform before a major change in the kinematics of the collision occurred. In the following sections, we propose possible tectonic scenarios for the northwestern Arabian platform since Late Cretaceous.

The Late Cretaceous – Palaeogene pre-collision extensional events In Senonian (Fig. 13a), a NE– SW regional extension occurred in the Coastal Range of Syria as indicated by the abundant NW– SE trending normal fault populations observed in the study area. This Senonian extension was also determined in Aleppo Plateau (Al Abdalla 2008) and in the Palmyrides (Salel 1993) as well. This extension is associated with the opening of major NW– SE trending grabens such as the Azraq graben in Jordan, the Euphrates graben, and the Djebel Abdelaziz halfgraben in Syria (Kent Hickman 1997; Litak et al. 1998; Caron et al. 2000), and at the northernmost margin of the Arabian plate in southeastern Turkey (Kaymakci et al. 2009). In addition, this extension was also observed westwards in the northern African platform in northern Egypt (Mostafa 1999), and as far west as the Sirt Basin in Libya. The NE–SW extension is thus a major regional extensional event at the scale of the north-eastern African plate. In Maastrichtian, the northern Arabian plate was flexed owing to the overload of the Ophiolitic nappes obducted onto the Arabian margin and

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Fig. 13. Tectonic evolution and palaeostress history of the northwestern Arabian platform since the Late Cretaceous. The sketches illustrate six major steps of the regional tectonic evolution, before (a) and (b) and during (c) – (f) the Arabian–Eurasia collision. (a) the Campanian sketch shows the NE– SW Senonian extension that lasted from Coniacian to Early Maastrichtian; (b) the NNE–SSW extension prevailed from Early Eocene to the early Late Oligocene; (c) The NW– SE compression is the major deformation phase in NW Arabian plate, it is probably Burdigalian in age; (d) the Middle Miocene extension was restricted to the Lattakia basin area, while the general regional tectonic context remained compressional; (e) Late Miocene–Early Pliocene NNW–SSE compression, with east–west deflections near the DSF, and initiation of Levant Fault; ( f ) Pliocene to present NNW– SSE compression and the DSF progression. DSF and EAF: Dead Sea and East Anatolian faults. The 30 km offset of the Early Miocene deformation front is shown.

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platform. It caused the subsidence of the NW edge of the Arabian platform and a change in the style of sedimentation. In Baer-Bassit and Afrin regions, a thick Upper Maastrichtian– Palaeocene pelagic marly sequence unconformably overlays the Ophiolitic nappe. The obduction of the ophiolites onto the platform is marked in the Maastrichtian sequence by intra-formational unconformities and by ophiolite and radiolarite bearing deposits. The flexing of the platform was associated with a NNW–SSE Late Cretaceous extension restricted to the northern Coastal Range, which predates the Early Miocene folding. The Eocene –Oligocene period in NW Syria was characterized by a north–south to NNE–SSW regional extension (Fig. 13b). Syndepositional structures were commonly observed in the Eocene –Oligocene deposits of Baer-Basit and Coastal Range. This major regional extensional event has also been evidenced in the Palmyrides, Aleppo Plateau (Al Abdalla 2008), Lebanon (Homberg et al. 2007, 2010), and as far south as northern Egypt (Mostafa 1999). This sub-meridian extension was not associated with large-scale extensional structures such as large grabens, but it locally controlled the thickness and the pattern of sedimentation of the Eocene –Oligocene sequence. The Eocene –Oligocene regional extensional event strongly suggests that the Africa– Eurasia collision had not initiated before the end of Oligocene in the western part of the Northern Arabian promontory.

The Neogene compressional events related to the Arabia – Eurasia collision The brittle tectonic analysis reconstruction presented in the previous sections allowed us to identify and reconstruct the different fold and fault systems with the corresponding stress field configurations that developed during the Neogene deformation processes in NW Syria. The interaction between folding and faulting in the Neogene times gave way to structural complexity. The following results have been deduced from our Neogene palaeostress reconstruction in relation to the regional tectonic evolution of the NW Arabian platform in the frame of the Arabia –Eurasia collision. A major NW compressional event (1208– 1358 compression direction) was recognized in this study. It occurred and prevailed in NW Syria in the Early Miocene (Figs 9 & 13c). The Early Miocene period corresponded to the beginning of the Arabia–Eurasia collision in this region, which initiated northward in the Uppermost Oligocene. Our data document stress regimes with NW– SE s1 directions that were active both before and

after folding. The NW pre-fold strike –slip stress regime evolved into the compressional stress regime associated with the main stage of NE–SW folding probably at the end of the Early Miocene. The major NE– SW trending intra-platform Lattakia thrust associated with this regional NW event developed in the Arabian platform of NW Syria at this period (Fig. 13c), as well as the regional NE–SW folding in Baer-Bassit, Afrin, and probably Palmyrides regions. This tectonic phase is associated with a significant uplift and a strong erosion of the Mesozoic to Palaeogene sequence of the folded platform and inverted Palmyra Basin. In the Palmyrides, piggy-back basins developed in the main synclines, filled up with Miocene continental deposits (Ponikarov 1966; Gianne´rini et al. 1988; Salel 1993). Because of the flexing of the platform associated with the Early Miocene thrusting, a narrow NE –SW elongated through developed during the Middle –Late Miocene forming the Lattakia basin (Fig. 13d). Normal faults with small offsets developed in the basin and surroundings during the Middle Miocene in response to the flexure of the platform. In the Late Miocene (Fig. 13e) started a kinematic reorganization of the Arabia –Eurasia collision, (i.e. inception of the Dead Sea and East Anatolian faults). A NNW– SSE oriented compression initiated in the Middle–Late Miocene and prevailed in the NW Arabian platform (Gianne´rini et al. 1988; Zanchi et al. 2002; Kaymakci 2006; Kaymakci et al. 2006, 2010; Al Abdalla et al. 2007; Al-Abdalla 2008). This event post-dated the NW compression event described above and predated the north –south folding of the Coastal Range. This, about 408 clockwise rotation of the s1 stress axes (from N1258 to N1658), has been already described by Gianne´rini et al. (1988), in the light of oblique Miocene and Pliocene–Quaternary volcanic dykes and cone alignments in the western Arabian plate. We propose that this rotation is coeval with the inception of the DSF. Near the DSF, the stress field locally re-oriented and s1 trajectories become sub-perpendicular (east –west) to the DSF. This stress configuration was associated with the folding of the Coastal Range (north–south axis) during the Late Miocene. From latermost Miocene or Early Pliocene to Recent (Fig. 13f ), the NNW –SSE oriented compression has prevailed, as shows by the recent deformation patterns, except in the vicinity of the DSF. Sinistral offsets of the pre-uppermost Miocene structures, especially of the Early Miocene deformation front of the Baer-Bassit and Afrin area on either side of the northern segment of the DSF, is estimated to be around 30 –40 km since the latest Miocene (about 6 –8 Ma). Considering the 105 km displacement along the central segment of the

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DSF (Quennel 1958; Freund et al. 1970; Garfunkel 1981), the remaining 65 –75 km may be absorbed by the shortening in the Palmyrides and Lebanon. We are grateful to the Middle East Basins Evolution (MEBE) Programme for funding this study. We greatly appreciate the logistic support provided by Aleppo University (Syria) and the University Pierre & Marie Curie (Paris, France). The scholarship of A. Al Abdalla was sponsored by the Syrian–French studentship cooperation Program. We thank J. Hall and N. Kaymakci for reviewing the manuscript and providing valuable suggestions for its improvement.

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LATE CRETACEOUS TO CENOZOIC TECTONIC EVOLUTION OF THE NW ARABIAN PLATFORM 327 Matar, A. & Mascle, G. 1993. Cine´matique de la faille du Levant au Nord de la Syrie: analyse microtectonique du fosse´ d’Alghab. Geodinamica Acta, 6, 153–160. Mostafa, M. S. 1999. Evolution tectonique de la plateforme. Africaine en Egypte depuis le Me´sozoı¨que a` partir de l’analyse des de´formations cassantes. Me´moire de the`se de l’Universite´ de Paris 6, Paris, 326 (in French). Mouty, M. 1967. Results of the stratigraphical study of the Alaouite Mountains. Syrian Ministry of Petroleum, unpublished report (in Arabic). Mouty, M. 1997. Le Jurassique de la Chaıˆne Coˆtie`re (Jibal As-Sahilyeh) de Syrie: essai de biozonation par les grands foraminife`res. Comptes Rendu de l’Acade´mie des Sciences Paris, 325, 207–213. Parrot, J. F. 1977. Assemblage ophiolitique du BaerBassit et termes effusifs de volcano-se´dimentaire. Travaux et documents de l’O.R.S.T.O.M, 72 (in French). Ponikarov, V. P. 1966. The geology of Syria. Explanatory notes on the geological maps of Syria, scale 1:200 000. Ministry of Industry, Syrian Arab Republic. Quennel, A. M. 1958. The structural and geomorphic evolution of the Dead Sea Rift. Journal of the Geological Society, London, 114, 1 –24. Ricou, L. E. 1971. Le croissant ophiolitique pe´ri-arabe, une ceinture de nappes mise en place au Cre´tace´ supe´rieur. Revue Ge´ographique Physique et Ge´ologie Dynamique, 2, XIII, 4, 327– 350. Rojay, B., Heimann, A. & Toprak, A. 2001. Neotectonic and volcanic characteristics of the Karasu fault zone (Anatolia, Turkey): the transition zone between the

Dead Sea transform and the East Anatolian fault zone. Geodinamica Acta, 14, 197– 212. Salel, J. F. 1993. Tectonique de chevauchement et inversion dans la chaıˆne des Palmyrides et le Graben de l’Euphrate (Syrie); conse´quence sur l’e´volution de la plaque arabe. PhD thesis, Me´moire de the`se de l’Universite´ de Montpellier 2, France, 288 (in French). Sengor, A. M. C., Gorur, N. & Saroglu, F. 1985. Strike-slip faulting and related basin formation in zones of tectonic escape: turkey as a case study. In: Biddle, K. T. & Christie-Blick, N. (eds) Strike– slip deformation, basin formation, and sedimentation. Society of Economic Paleontologists and Mineralogists, Special Publication, 37, 227–265. Steckler, M. S. & Ten Brink, U. S. 1986. Lithospheric strength variations as a control on new plate boundaries: examples from the northern Red Sea region. Earth and Planetary Science Letters, 79, 120– 132. Thuizat, R., Whitechurch, H., Montigny, R. & Juteau, T. 1981. K–Ar dating of some infra-ophiolitic metamorphic soles from the Eastern Mediterranean: new evidence for oceanic thrustings before obduction. Earth and Planetary Science Letters, 52, 302– 310. Wallace, R. E. 1951. Geometry of shearing stress and relation to faulting. Journal of Geology, 59, 118–130. Walley, C. D. 1988. A braided strike-slip model for the northern continuation of the Dead Sea Fault and its implications for Levantine tectonics. Tectonophysics, 145, 63–72. Zanchi, A., Crosta, G. P. & Darkal, A. N. 2002. Paleostress analyses in NW Syria: constraints on the Cenozoic evolution of the Northwestern margin of the Arabian plate. Tectonophysics, 357, 255–278.

Index Page numbers in italic denote figures. Page numbers in bold denote tables. Acoustic sequences PSDM 90 AF-demagnetization 272, 273– 274, 274 Aitoun mesoscale fault analysis 251 Akar thrust 247 Albian 99, 250 Albian beds 253 al-Naqab Desert structural features 66 Anti-Lebanon 252 Jurassic sedimentary rocks 227 Apparent polar wander path (APWP) 276–277 Aptian 99, 250 Aptian–Albian formation 271 Aptian–Albian growth fault 250 Arabia–Eurasia collision 323 Arabian plate 269, 282 Levant margin global control 133– 165 mid-Cretaceous carbonate system 133–165 mid-Cretaceous sedimentary units distribution 134 Arabian platform 305 palaeostress orientation 313 –315 stress axes distribution 318 tectonic evolution and palaeostress history 324 Arkan Formation 136 Asher Basin 13 Asher volcanics 16 Baer–Bassit 320, 323, 325 Miocene deformation 326 overthrusts 307, 320 platform 309–310, 319, 320, 321 Bairdia 194– 195, 195 Barremian 99 Bairdiidae 195 Basal surface of forced regression (BSFR) 136 Bechare section 230 depositional environments 238 Bekaa syncline 258, 259 Betzet Valley basinal laminites facies type-1 144– 145 Bhannes Formation 230 benthic foraminifera 232, 233 calcareous algae biostratigraphy 232, 233 Qartaba geological map 240 volcano-sedimentary continental deposits 231 volcano-sedimentary materials 231 Bikfaya Formation 230, 231, 235 benthic foraminifera 233, 234 calcareous algae 233, 234 carbonate platform facies 235 deposition of 241 Qartaba geological map 240 sedimentological model 235 Bina Formation 216, 223, 225 folds location 218–291 landslides 224 sedimentary structures 213

Bouguer gravity anomaly 20, 26, 31 Brachcythere 200– 201, 201– 202 BSFR. See Basal surface of forced regression (BSFR) Bythocytheridae 196 Calcimetry 176– 177 carbonate values 178 power spectra 180 Campanian–Maastrichtian boundary 299 Candonidae 195–196 Carbonate cycles 176, 178 gypsum beds 182 Carboniferous erosion 12 Carbon isotope excursion (CIE) 176, 178, 181, 183 Carbon isotope record 182 Carbon isotope values Barremian–Albian 104, 105, 107 Carmel 135 Cenomanian SB 159 cross-section 154, 160 dolomites 136 facies-thickness changes 152 platform-basin configuration 152 sedimentary/structural configuration 156 CC. See Composite cycles (CC) CDP. See Common depth point (CDP) Cenomanian– Early Turonian ostracod zonal schemes 203 Cenomanian limestone 251 Cenomanian sequence-1 eustatic 161–162 Cenomanian succession cross-section 160 Cenomanian– Turonian boundary 180 lithostratigraphic scheme 135 Cenomanian– Turonian Eastern Levant platform biostratigraphy 172– 175 calcimetry and gypsum marker beds 176– 177 carbon isotope stratigraphy 171–184 geological setting 171–172 global carbon isotope events 181 high-resolution correlation with Pueblo 177–180 intra-platform basin correlation 171– 184 intra-platform basin sections lithological evolution 175–176 isotope and calcimetry correlation spectral analysis 177 isotope record 176 Jordan orbital time scale 171– 184 Levant carbonate platform demise timing 183–184 obliquity cycle 183 sea-level across C– T boundary 181–183 sea-level history 171– 184 stratigraphy isotope record implications 181– 183 third-order sea-level change 183

330 Cenomanian– Turonian succession of northern Israel facies types 138–140 position models 141 stratigraphic correlation of cyclic units 148 transition interval 5 UC units 145, 146 Cenozoic rocks 40, 41, 47 maximum thickness 41 Cenozoic series 287–303 age lithology and nannoplankton markers 297 nannofossils revised stratigraphy 287 –303 Central Jordan Middle Cenomanian –Early Turonian Levant Platform biostratigraphy 187– 207 isotope stratigraphy 190–191 location and stratigraphic setting 188 material and methods 188–189 ostracod systematic descriptions 191– 202 ostracod biostratigraphy 202– 204 ostracod biozone I 202 ostracod biozone II 202– 204 ostracod biozone III 204 ostracod biozone IV 204 ostracod biozone V 204 palaeobiogeography 206 palaeoecology 187– 207, 205– 206 palaeogeography based on ostracods 187–207 Central Levant domain Lebanon tectonic evolution geological overview 247– 248 faults and folds 247 Lebanon tectonic structures age and origin 247–248 large scale tectonic structures and dating arguments 249 Levant Margin Cenozoic deformation 257–264 Early Miocene compression 258– 261 Late Miocene compression 261–263 Pliocene and later deformation 264 poly-phase Neogene tectonism evidence 257– 258 mesoscale faults and inferred stress field 249–251 Mesozoic time and younger 245– 266 pre-Neogene extensional tectonic events 251– 257 Early Cretaceous extension 252–255 Early normal faulting evidences 251–252 Eocene extension 255– 257 Central Levant margin global control mid-Cretaceous carbonate system 133– 165 Chalky Lower Eocene sequence 259 Chouf sandstone growth faults 253 CIE. See Carbon isotope excursion (CIE) CIG. See Common image gathers (CIGs) Coastal Range Mesozoic to Cenozoic stratigraphy 308 Common depth point (CDP) 87 Common image gathers (CIGs) 87, 88 Composite cycles (CC) 137 Cretaceous 1 –8. See also Late Cretaceous Cythereis 198–199, 199 Cytherella 191– 195, 194– 195 Cytherellidae 191–195 Cytherideidae 196 Cytheruridae 197– 199

INDEX DDS. See Depth dependent stretching (DDS) Dead Sea Fault (DSF) 1, 3, 86, 269 affecting Levant area 12 declination indicated 280 digital topographical model 2 geological and tectonic structures 246 location description 305 Miocene mechanism of deformation 281 Dead Sea Transform (DST) 1, 11, 37, 60, 99, 248, 264, 266 and Miocene series 211 and Palmyra Trough 13 plate boundary 246 Miocene deformation mechanisms 281 Post Miocene deformation mechanisms 282 and Suez rift 53 Dead Sea Transform (DST) system 269– 283 age of magnetizations 277 APWP 276– 277 declination 280 declination and inclination in degrees 277 evolution implications 269– 283 geological setting 270 –271 hysteresis loops 272 Lebanese sector 270 Lebanon 269– 283 Lebanon rotations 279 magnetization component 276 orthogonal vector plots 274 palaeomagnetic data 275 palaeomagnetic directions 275, 276 palaeomagnetic results 273–276 magnetization analysis 273–275 palaeomagnetic tests 275–276 previous studies 277– 279 rock magnetism analysis 272–273 structural implications 279– 283 thermomagnetic curves 273 Deir Hanna Formation 143 Depth dependent stretching (DDS) 27, 28, 30, 31 Dishon sedimentary/structural configuration 156 Dolocytheridea 195, 196 DSF. See Dead Sea Fault (DSF) DST. See Dead Sea Transform (DST) Early Cenomanian maximum-flooding bed 142– 143 packstones facies type-7 142–143 peloidal grainstones facies type-7 142– 143 serpulids and bioclastic debris facies type-8 143 Early Jurassic rifting phase 24 tectonic pulse 30 Tethyan rifting 23–24 volcanic rocks 22 Early Mesozoic rifting activity 12 Early Miocene compression stress axes distribution 317 Early Turonian section 190 Eastern Mediterranean 95 deformation 86 tectonic elements 10

INDEX Egypt continental margin. See also specific areas of Egypt e.g. Sinai Desert acoustic basement 91–93 angular unconformity 60 data processing 87–88 geological cross section 94 Geological Survey 38 MEDISIS MCS survey 87 multichannel seismic data deep structures 85–95 sedimentary cover 90– 91 seismic stratigraphy 85–95 shaded-relief bathymetry 86 structural setting and tectonic evolution 37–61 systematic surveys 85 tectonic setting 85– 87 Ein El Assad Formation 114– 115 Eocene 292 Middle chalky deposits 309, 316 extensional syn-depositional structures 316 growth fault 257 syn-depositional normal faults 309 Eocene chalks Qsaibe 252 Eocene extension stress field 256 Eocene– Oligocene stress axes distribution 311 Eratosthenes 22 Facies classification and interpretation components composition and sedimentary structures, 108– 110 Fathi Higazi 230 FK. See Frequency-wavenumber (FK) Forced regressive systems tract (FRST) 136, 138, 149, 151, 154, 155, 163 Frequency-wavenumber (FK) 87 FRST. See Forced regressive systems tract (FRST) Gaash-Meged High 18 Galilee 135 biostratigraphical markers 107 correlation 111 cross-section 160 dolomites 124 facies-thickness changes 152 Late Cenomanian carbonate system 164 Levant Platforms depositional architecture 118 limestones 136 lithostratigraphic units 101 palaeoenvironmental evolution 164 platform-basin configuration 152 sequence-stratigraphic correlation 149 Gebel (mountains) 100 Gebel Amrar 105– 106 swell 123 Gebel-El Minsherah 51 fault 51, 53 Gebel Halal inverted structures 58 maximum thickness of Mesozoic and Cenozoic rocks 43

Gebel-Halal 60 anticline field-mapped faults 50 fault-slip data analysis 51 fold 50 Gebel-Maghara 60, 101 anticline 56, 59 Barremian-Upper Albian succession 101 Landsat TM image 46 outcrop thickness 43 seismic section 56 Gebel-Manzour 101 Aptian-Upper Albian succession 101 asymmetric syncline 47 Gebel-Meneidret El Etheili mesostructures 49 Gebel Raghawi 105– 106, 123 Gebel-Rishat El Themed folds field-measured faults 55 Themed fault 55 Gebel-Rishat Lehman fault 51 Gebel-Rishat Saada fault 51 vertical aerial photograph 52 Gebel-Yelleg 60 anticline 56, 59 structures 48 Geovector post-stack time migration 89 Gevim High 16, 18 interpreted seismic profile, 17 Palaeozoic vertical motions 22 Ghawr Al-Mazar section (GM3) 172, 187. See also Middle Cenomanian section biostratigraphy and carbon isotope 174 carbonate 176 carbonate cycles correlation 177 carbonate record 180 CIE 176 gypsum 176 lithological units 175 location 173 location map 189 ostracod biozones 202 Pueblo correlation 178 Global carbon isotope event 181 GM3. See Ghawr Al-Mazar section (GM3) Golan Heights biostratigraphical markers 107 Levant Platforms depositional architecture 118 lithostratigraphic units 101 succession correlation 111 Goliath structure 78 Gondwana Palaeozoic succession 11–12 Permian breakup and rifting 29 plate 12 Tethyan-rifted margins 31 Gre`s de Base Formation 230, 231 benthic foraminifera 232 calcareous algae biostratigraphy 232 Qartaba geological map 240 Hamayir-Amrar area inverted structures 58

331

332

INDEX

Hamra Valley Pelech sequence 158 Hidra Formation 114–115 Highstand systems tract (HST) 121, 136 Hummar Formation 204 Inversion-related folding 80 Isfiyye section 159 Israel 1, 3 –7, 193. See also Dead Sea Transform (DST); Northern Israel; Syrian Arc; specific named areas of Israel e.g. Galilee Asher Basin wells 24 belt of deep seated highs and lows 20 Cenomanian– Turonian 191 Coastal Chain of Galilee 239 coastal plain 12 crustal scale section 26 Early Jurassic volcanic rocks 22 Geological Survey 38, 41 grabens formed during rifting stage and block tilting 72 highest basement block 16 Lower Cretaceous sequence 264 Lower Jurassic section 13 margin onshore and offshore 10 multi-biostratigraphical framework 172, 188 Northern 99– 127, 133– 165 number of fault zones 27 offshore 26 onshore and offshore 29, 248, 265 ostracod biozonations 202– 204 Palmyrides intra-plate basin 229 prior to Early Cretaceous 242 sector of Levant margin 247 seismic depth maps 11 tectonic map 212 well data 31 Jezine formation 253 Early Cretaceous fault plane 249 faults 253 Jonah High marine seismic profile 21 Jordan 190 Karak– Silla intra-platform 172, 184 ostracod faunas 204 section GM3 and KB3 location 173 Judea Graben 13, 16, 18, 23 interpreted seismic profile, 17 Jurassic 14, 15, 178. See also Early, Late, Lower, or Upper Jurassic Anti-Lebanon 227 formation Lebanon 270 formation sedimentology 232– 235 Levantine margin 232– 235 Middle Levantine margin 227– 242 Shederot Formation 23 stratigraphy 227– 242 Tethyan rifting 23– 24 unconformity 17 Mount Hermon 227 Mount Lebanon 227 Neogene formation 259

reflector 57 rifting and structural evolution 24, 77 rocks 41, 42– 43, 57 Levant area 278 North Sinai folds 42 thickness 41, 43 sedimentary rocks 227 seismic section reflector 57 sequences and faults 253 strata 18 Kafer Hachno 249 KB3. See Kuthrubbah section (KB3) Keb valley 261 valley 262 Kesrouane Formation 231 benthic foraminifera 233, 234 calcareous algae 234 calcareous algae distribution 233 carbonate shelf 239 Cretaceous faults 242 karstification 239 marine formation 239 outcrops 230 Qartaba geological map 240 sedimentary sequences 237 sedimentological model and carbonate platform facies 235 Kimmeridgian carbonate shelf regime 239 Kishor section 155 clinoform unit 157 Kleiat section panorama 230 Krithidae 199 Kuthrubbah section (KB3) 175 biostratigraphy and carbon isotope 174 carbonate 176 cycles correlation 177 record 180 CIE 176 description 172 lithological units 176 location 173 Larger benthic foraminifer biozones (LFBs) 104 Late Barremian–Albian Levant Platform location 99 Late Cenomanian FRST 155 SB 163 Late Cretaceous 17, 58, 262 deformation in Levant margin 6 –7 extensions 316 –319 Northwest Arabian platform 305– 326, 324 Northwest Syria 305 –326 pre-collision extensional events 323–325 stress axes distribution 310, 311 tectonic evolution 305–326 Late Miocene compression 261–263 distribution and stress axes 322 Lattakia thrust 307 Levant Margin Cenozoic deformation 261 –263

INDEX Neogene compressional events 325– 326 tectonics and Arabian platform 324 tectonics and Lebanon 263 Lattakia basin 308, 309, 316, 319 deformation and erosion 308 folding 320 Lebanese blocks 281, 282 Lebanese coast at Raas Cheeka outcrop 259 Lebanese folds 264 Lebanon 251, 252, 256 Cenozoic series 287–303 DST 283 Early Cretaceous extension 254 Early Miocene tectonics 260 faulting 264 folding 266 geological map 228, 288 geological setting 287– 299 geologic and tectonic structures 246 kinematic framework 246 Levantine margin 227– 242, 241 geological and regional setting 229 localities of nannofossils samples 289 –296 Mesozoic carbonate platforms 229 nannofossils stratigraphy 287– 303 Cenozoic–Neogene 302 Cenozoic–Palaeogene 299–302 Eocene 299–302 geological settings 287–299 Oligocene 302 Palaeocene 299 Pliocene 302 Senonian–Maastrichtian 299 zones and hiatuses 300, 301 new marine transgression 239 palaeogene nannofossil zones and hiatuses 300, 301 rotations inferred from new palaeomagnetic data 269–283 age of magnetizations 277 APWP 276– 277 DST system evolution implications 269– 283 Lebanon rotations 279 palaeomagnetic 273 –276 magnetization analysis 273–275 palaeomagnetic tests 275 –276 previous studies 277– 279 rock magnetism analysis 272–273 sampling and analysis procedures 271–272 structural implications 279– 283 sedimentary and tectonic events 241 stress field 254, 260, 263 tectonic event 265 Upper Cretaceous 287 –303 nannofossils revised stratigraphy 287–303 Levant area Bouguer gravity map 11 crust thickness 27 digital topography model 2 Early Jurassic tectonic pulse 30 palaeomagnetic 278 rifting 30 seismic depth map 11

333

Tethyan extensional structures 13 Tethyan rift system 11 Levant Basin 19, 21, 265 crustal scale section 26 crustal thinning 28 faults 15 geological cross-section 15 physical characteristics 29 rifting-drifting model 30 southern border 85 stratigraphic successions 15 structural blocks 15 Levant domain strike-slip and reverse faults 258 Levant hinge-belt 134 Levant margin 1– 7, 6 basin and adjacent areas post-rift evolution 5 –6 Levant platform 5 crustal scale section 26 general evolution 2– 3 geological cross-section 15 global control 133– 165 Jurassic 227– 242 benthic foraminifera 232 Bikfaya Formation 232 calcareous algae biostratigraphy 232 depositional environments 232– 235 depositional environments evolution 236–239 stacking patterns 236–237 facies 1 bioturbated nodular limestone 233–235 facies 2 limestone with benthic foraminifera and calcareous algae 235 facies 3 limestone with green algae, charophytes, and benthic foraminifera 235 facies 4 coral limestones 235 –236 facies 5 fine-grained limestones with bioclastic debris streaks 236 facies 6 oolithic limestones with megaripples 236 facies 7 fine-grained limestone with birds-eyes 236 facies 8 basalt flows and pyroclastic material 236 formation sedimentology 232–235 geodynamic implications 227– 242, 239 –242 geological and regional setting 229 Kesrouane Formation 231– 232 Lower Jurassic volcanism development 239– 242 Middle and Upper Jurassic stratigraphy 227 –242 palaeogeography 236– 239 relative sea-level changes 236– 239 sedimentary evolution palaeoenvironmental 227– 242 stratigraphy 231– 232 Late Cretaceous to Cenozoic deformation in Levant 6– 7 magma intrusions 28 Middle East Basin Evolution Levant Group 1 –2 rifting and tectono-stratigraphic sequences 4 –5 stratigraphic successions, faults, and structural blocks 15 Levant Margin Cenozoic deformation 257–264 Early Miocene compression 258–261 Late Miocene compression 261 –263 Pliocene and later deformation 264 poly-phase Neogene tectonism evidence 257– 258

334

INDEX

Levant Platform 5 coastal onlap 3 development 101 global sequences 182 shallow marine facies belts 3 Levant Region 25 Early Mesozoic crustal evolution 9 –31 Tethyan rifting 9–31 activity timing 22– 24 Early to Middle Jurassic 23–24 Late Palaeozoic 22–23 Middle to Late Triassic 23 main tectonic events shaping Levant region 11– 13 shaping crust 25– 29 basinal domain 28 coastal area and nearby slope 28 continental margin deep structure 25– 26 margin extension discrepancy 27– 28 passive continental margin formation 27 Red Sea analogue 28– 29 subsurface rifting structures 13–22 Eratosthenes High 20–22 inner basins 13– 17 Mediterranean coastline highs 16–17 offshore highs and lows 17–20 tectonic reconstruction 24–25 well data 14 LFB. See Larger benthic foraminifer biozones (LFBs) Lower Jurassic section Israel 13 Lower Jurassic strata 18 Lower Jurassic volcanism development 239– 242 Lower Miocene sedimentary sequence folding 320 Lowstand systems tract (LST) 136 LST. See Lowstand systems tract (LST) Maastrichtian 297, 299, 313–315 chalk 59 tectonosequence 2 67–69 Mahgara area. See also Gebel-Maghara inverted structures 58 Manara section Galilee cross-section 154 Mansour– Raghawi fault 123 Marine platform 5 Maximum-flooding bed bivalve shell-bed facies type-8 143 Maximum-flooding interval (MFI) 137 Maximum-flooding surface (MFS) 121 MCS. See MEDISIS MCS survey MEBE. See Middle East Basin Evolution (MEBE) Programme Mediterranean map 228 MEDISIS MCS survey 85, 87, 95 reflection data 95 Menuha Formation 216 Late Turonian age palaeokarst 223 Mesozoic passive continental margin 85 Mesozoic rocks 41 maximum thickness 41 Mesozoic sedimentary sequence folding 320

Metacytheropteron 197–199, 198–199 MFI. See Maximum-flooding interval (MFI) MFS. See Maximum-flooding surface (MFS) Mid-Cretaceous carbonate system Northern Israel 133–165 Arabian Plate 133– 165 central Levant margin global control 133– 165 facies evolution 133– 165 margin global control 133–165 tectonosedimentary configuration 133– 165 Middle Cenomanian section ammonite zones and stable carbon isotope record, 190 calcareous nannoplankton zones 190 formations 190 FRST 151 global isotope event 191 Middle East Basin Evolution (MEBE) Programme 1, 247 Mid-ocean range (MOR) 30 Milankovitch cycles 182 Miocene 79, 293 Middle extensional stress regimes 318 sediments 319 syn-depositional extensional structures 319 Mitla Pass en echelon doubly plunging anticlines 54 en echelon folds 58 Mizeraa reverse fault 46 Mohilla Formation 23 Monoceratina 195, 196 MOR. See Mid-ocean range (MOR) Mount Gamal clinoform unit palaeo-calcrete fabric 157 Mount Hermon 237 Jurassic sedimentary rocks 227 Mount Lebanon 252, 271 anticline 262 elevated position 298 Jurassic sedimentary rocks 227 Senonian-Cenozoic deposits 287 thrust 247 Mount Sannine cross-section 160 Muhraqa Formation limestones 157 turbidite 157 Mujib Dam 188 Nabi Said Formation 114– 115 Nahr Ibrahim Valley 230 Nannofossil stratigraphy Lebanon 287– 303, 297, 298 Cenozoic-Neogene 302 Cenozoic-Palaeogene 299– 302 Eocene 299–302 geological settings 287–299 Oligocene 302 Palaeocene 299 Pliocene 302 Senonian-Maastrichtian 299 studies 299–302 zones and hiatuses 300, 301

INDEX Naqb Desert 37 geological map 38 structural form-line map 44 Natural remanent magnetization (NRM) 272 NDSF. See Nile deep-sea fan (NDSF) Negev Desert 23, 266 Neocomian 78 Neocyprideis 196–197 Neogene 3, 7, 302 compressions 321–323, 325–326 folds location 218–291 formation 259 marine units 258 stress fields 310 tectonism 257–258 unconformity 262 Neogene Lattakia basin 307 Neotethys ocean 25 Nile deep-sea fan (NDSF) 87 NMO. See Normal move out (NMO) Normal move out (NMO) 87 Northern Israel 100, 134. See also Galilee facies-thickness changes 151 low-order CC-1-8 147 mid-Cretaceous carbonate system 133–165 Cenomanian –Turonian succession cycles and cyclic hierarchy 137–141 type-1 low-order CC-1 137 type-2 low-order CC-2 137 type-3 low-order CC-3 137 type-4 low-order CC-4 137–141 type-5 low-order CC-5 141 type-6 low-order CC-6 141 type-7 low-order CC-7 141 type-8 low-order CC-8 141 central Levant margin global control 133– 165 controlling mechanisms of sedimentation 161–165 Cenomanian sequence-1 eustatic 161–162 latter sequence-2 eustatic 163 lower Cenomanian sequence-1 eustatic 161 oceanic anoxic event-2 163–165 palaeoenvironmental controls 161–162 sequence-2 beginning eustatic 162– 163 correlation across Arabian Plate 161–165 facies evolution 133–165 facies types and cyclic patterns 137– 141 geological setting 135– 136 Cenomanian –Turonian lithostratigraphy 135– 136 Cretaceous Levant margin tectonic setting 135 Levant margin tectonosedimentary framework 160 –161 sequence stratigraphic subdivision 141–160 sequence 1 Early –Middle Cenomanian 141–153 Early Cenomanian maximum-flooding interval 143– 149 Early Cenomanian transgressive system tract 143 Early-Middle Cenomanian regressive system tract 149–151 mid-Cenomanian sequence boundary 151–153 sequence 2 Middle–Late Cenomanian 153–155 Late Cenomanian regressive system tract 154–155

335

Late Cenomanian transgressive system tract 153– 154 Middle Cenomanian lowstand system tract 153 sequence 3 Late Cenomanian SB 155– 158 sequence 3 Pelech sequence 155 –158 sequence 4 Early and Middle Turonian 158–160 tectonosedimentary configuration 133– 165 Upper Barremian–Albian Levant Platform 99–127 North Sinai anticlimal Syrian Arc structures 101 Aptian–Albian palaeoenvironmental map 100 Barremian/Aptian boundary 107 Barremian–Lower Aptian shallow shelf 105–106 basin 16 subsurface structural elements 75 tectonostratigraphy 72 time-thickness map 73, 74 carstification horizons 125 characterizations 99, 103, 104, 123 comparison with Tethyan ammonite zonation 103 compressional deformation 60 continental crust thinning 60 correlations of North Sinai and Galilee-Golan Heights succession 111, 115 deformation phases 61 Gebel Raghawi ammonites 104 growth extensional faults 77 inversion structures 81 Late Barremian flooding 120 lateral facies 112 Levant Platforms depositional architecture 118 limestone deposits 117 lithostratigraphic units 68, 101 low-energy environments 123 low-energy sub-basin development 122 Mango structure 71 offshore basin development and inversion 78– 80 basin inversion evidence 70–78 structural evolution 72–78 deformation timing 76–77 Goliath-inverted structure 73– 76 Jurassic-Early Cretaceous rifting 77 Neogene extension 78 Post-Santonian–Middle Miocene basin inversion 77– 78 tectonic structures 70–72 Goliath structure 71 Mango structure 70–71 North Sinai 21-2 structure 71–72 Tineh structure 72 Ziv structure 72 fold development and basin inversion mechanism 80 regional setting 65–67 structural deformation 81 structural setting and tectonic evolution 65– 81 tectonosequences 67– 70 Tectonosequence 1 (Upper Jurassic-Santonian) 67 Tectonosequence 2 (Campanian-Maastrichtian) 67–69 Tectonosequence 3 (Palaeocene-Upper Oligocene) 69

336

INDEX

North Sinai (Continued) Tectonosequence 4 (Miocene) 69–70 Tectonosequence 5 (Pliocene-Pleistocene) 70 palaeoenvironmental maps 116– 117 platform stages 112– 117 record 119–120 rose diagrams strikes 45 SB dating 121 sea-level change 116, 122, 125– 126, 135 second-order sequences Mid-Cretaceous MCL 120– 122 sedimentation rates and rising sea level 124 seismic section 73 seismic sections and boreholes 66 shallow marine and continental sedimentary rocks 65 shallow shelf schematic transect 116– 117 stratigraphical interpretation 107 stratigraphy 41–44 structural control on Late Barremian–Albian 119 structural features 66 structural form-line map 44 structural setting 44– 54 succession 107 tectonic evolution 54–58 tectonic map 100 tectonostratigraphic units features 70, 81 TS dating 122 Upper Barremian– Albian Levant Platform 99–127 Upper Barremian– Albian succession 102 lithostratigraphical subdivision 101–103 North Sinai folds 39 Cenozoic rocks 43– 44 compressional deformation 58–60 Cretaceous rocks 43 Gebel–Halal 47–48 Gebel–Yelleg folds 47 geometry 38 Jurassic rocks 42 thickness changes 42– 43 orientation 37– 38 Sinai hinge belt 48–53 themed fault 53–54 Triassic rocks 42 Northwest Arabian platform in Northwest Syria brittle tectonics methodology 310– 312 fault-slip measurements 310 stress tensor determination 310–312 tectonic event chronology and dating 312 cross-section 307 geological setting 306– 310 stratigraphy 306– 309 structures 309– 310 index map 306 Late Cretaceous-Palaeogene pre-collision extensional events 323– 325 Eocene– Oligocene period 325 Maastrichtian 323–325 Senonian 323 Late Cretaceous to Cenozoic tectonic evolution 305–326 Neogene compressional events 325– 326 Late Miocene 325 Late Miocene to Recent 325–326 NW compressional event 325

palaeostress configurations tectonic evaluation 316– 323 Early Miocene compression 320 –321 Eocene–Oligocene extension 320 late Cretaceous extensions 316 –319 Middle Miocene extension in Lattakia basin 321 Syria palaeostress reconstruction 312– 316 Northwest Syria. See also Northwest Arabian platform in Northwest Syria geological context 306 Late Cretaceous to Cenozoic tectonic evolution 305– 326 stress fields 310 NRM. See Natural remanent magnetization (NRM) Nubian plates 282 Oceanic anoxic event (OAE) 5, 107, 187 Oceanic anoxic event 1a (OAE 1a) 107 Oceanic anoxic event 2 (OAE2) 171, 172 carbonate 181 CIE 183 northern Israel mid-Cretaceous carbonate system 163– 165 ostracod species 187 Oligocene 41, 79, 297, 302, 313– 315 distribution of stress axes 311 extension 320 NW Syria characterization 325 tectonosequence 3 69 Oren Valley facies-thickness changes 156 Ostracods biostratigraphy 202–204 biozones GM3 202 faunas in Jordan 204 geographical distribution 193 palaeogeography 187–207 stratigraphical distribution 192 zonal schemes 203 Palaeozoic basin interpreted seismic profile, 18 Palaeozoic epeirogenic movements 12 Palmyra Trough 13 Palmyride block 281 Palmyrides 22 Paracypris 195– 196, 195 Parakrithe 198 –199 Pararithe 199 Peloriops 199, 200– 201 Perissocytheridea 197, 198– 199 Permian Triassic and Lower Jurassic strata, 18 Permian Arkov Formation 18 Plenus Marl 181 Pleshet Basin 18, 19, 22, 24 marine seismic profile 20 Pleshet depocentre 19 Pliocene 70, 278, 297 compressions 322 deformations in Lebanon 264 distribution and stress axes 322 formations 40, 262 rocks 220

INDEX tectonosequence 5 70 transgression at base of zone NN 12 302 Podocopa 191–202 Positive inversion 80 Pre-Neogene extensional tectonic events 251–257 Early Cretaceous extension 252– 255 Early normal faulting evidences 251– 252 Eocene extension 255 –257 Pre-stack depth migration method (PSDM) 87, 89, 91 acoustic sequences 90 processing steps 88 Ray and Born migration technique 88, 89 Profile MD06 extracts 92 PSDM 91 Profile MD08 post-stack time migration 93 PSDM 93 PSDM. See Pre-stack depth migration method (PSDM) Qartaba 231 Qsaibe folding-faulting chronology 252 Rajabiah Formation 56 Rakefet section 159 Rama Formation 114–115 Ray and Born pre-stack depth migration technique 89 processing steps 88 Reactivation 80 Red Sea Mesozoic Levant Basin 31 rifting 60 seismic refraction data 28–29 Red Sea Rift 269 Reticulocosta 199, 200– 201 Rifting phase 242 Rifting structures 4 Rizan Aneiza 105– 106, 123, 124 Roum fault 281, 282, 283 SAFB. See Syrian Arc fold belt (SAFB) Sakhnin Formation debrite breccias facies type-3 143 distal debrites facies type-3 143 dolomitized calcisilts facies type-2 145 vadose pisoids facies type-20 145 Santonian 78, 297 boundary 299 SB. See Sequence boundary (SB) Schmidt’s projections 311 Schuleridea 198– 199 Second Neogene unconformity 262 Senonian Menuha Formation folds location 218–291 Senonian period sedimentation 3 Sequence(s) Kesrouane Formation sedimentary 237 Late Cenomanian SB 155– 158 Middle-Late Cenomanian 153 –155 Pelech sequence 155 –158 second-order 120–122 sedimentary 232

stratigraphic subdivision 141–160 tectono-stratigraphic 4 –5 Sequence boundary (SB) 121, 136 Sergaya-Rachaya fault 282 Sinai 239. See also North Sinai benthic organisms 104 block 87 coast 37 transform fault 24 Desert 38 geological map 38 Lower Eocene rocks 44 offshore 6, 265 southwest segment 211 stratigraphical subdivisions 99, 100 underlying sediments 121 Wadi Feiran 163 Sinai hinge belt 39, 40, 58 deformation 58 dextral transpressive deformation 61 fault-slip data analysis 52 oriented en echelon faults 51 southern boundary 45 stratigraphic sections 40 Sinai Peninsula 37, 53 geological cross section 39 Sirius package 89 Southern Carmel facies-thickness changes 152 platform-basin configuration 152 Southern Lebanon fault-slip data 257 Syrian Arc 211–226 deformation 266 faults 213, 215, 216, 221 discordant to monocline trend 211–226 fold axes 220 fold orientation 215 folds 213– 215 hinge zone evolution 222 karst development related to fold hinge zone 215–216 monocline flank Mesoscale folds 211–226, 213 open folds 214 stereographic projection 221 tectonism 248, 266 Syrian Arc fold belt (SAFB) 7, 65, 67, 211 Syrian Arc System 80 Talme Yafe Group 134 Tarchich section benthic foraminifera 233 calcareous algae distribution 233 cross-bedded fluviatile sandstones 231 Techrine Lake 320 Tethyan composite marine seismic profile 19, 21 extensional structures 19, 21 fault-controlled rift basin 21 ophioliths 3 rifting 9, 13, 25 belt structure distribution 29 structures activation 23

337

338 Tethyan (Continued) structures timing 24 tectonic activity 9 tectonic reconstruction 25 Tineh area Goliath structure 76 TOC. See Total organic carbon (TOC) Total organic carbon (TOC) 187 Trachyleberididae 199–202 Transgressive surface (TS) 121 Transgressive systems tract (TST) 120, 121, 122, 136 Triassic rifting 30 TS. See Transgressive surface (TS) TST. See Transgressive systems tract (TST) Turonian Bina Formation disruption 224 Turonian transgressive system tract 158 Unconformity A (UA) 258– 259 Undividable cycles (UC) 137 Upper Aptian-Albian succession 107 Upper Barremian-Albian Levant Platform biostratigraphy 103 controls and facies patterns evolution 99–127 external controlling factors 125– 126 climate 125 second-order sea-level change 125– 126 tectonics 126 facies analysis 103 geological framework 101 investigated sections 101 –103 lithostratigraphical concepts 101– 103 palaeoenvironment reconstruction 103 platform development 111–120 deposits structural control 119–120 facies zones platform environments 111– 112 Galilee/Golan Heights platform stages 117–119 PS I 117 PS II 118 PS III 119 PS IV 119 PS V 119 North Sinai platform stages 112– 117 PS I 112 PS II 112–113 PS III 113 PS IV 113–117 PS V 117 PS I Late Barremian-earliest Early Aptian 122 –123 PS II Early Aptian-earliest Late Albian 123 PS III Earliest Late Aptian 123

INDEX PS IV Late Aptian-Middle Albian 123– 124 PS V Late Albian 125 second-order sequences 120–122 MCL-1 120–121 MCL-2 121–122 MCL-3 122 sequence stratigraphical interpretation 103 stable isotopes 103 stratigraphy 103–111 ammonites 104 benthic organisms 104 stable isotopes 104– 107 stratigraphical interpretation and section correlation 107 –111 Upper Barremian-Albian strata facies and bedding-patterns 114–115 Upper Barremian–Upper Albian biostratigraphic subdivision 102 succession 101 Upper Cretaceous 287–308 nannofossil datings 298 nannofossils revised stratigraphy 287–308 Cenozoic–Neogene 302 Cenozoic–Palaeogene 299–302 Eocene 299–302 nannofossil studies 299– 301 Oligocene 302 Pliocene 302 Senonian-Maastrichtian 229 nannoplankton markers 297 Upper Jurassic stratigraphy 70, 227–242 tectonosequence 67 Upper Maastrichtian chalk seismic section 59 Veeniacythereis 200– 201, 201 Volcanic igneous rocks 44 Wadi Al Karak northern slope 173 Western Arabia platform evolution since Mesozoic 1– 7 Yammouneh faults 281, 282, 283 Yanuch Formation karst surface 157 Yelleg area inverted structures 58 Ziv seismic section 73

This volume combines original data in various fields from the offshore Levant Basin and adjacent continental slopes and platforms. The first group of papers document the tectonic structures and sedimentological patterns associated with the development of the Levant Basin. They identify the successive rifting events from the Late Palaeozoic to the Early Cretaceous, followed by a moderate tectonic activity. The contribution of external factors like global sea-level and climate changes to the sedimentation processes during the Mid-Cretaceous is discussed in the second set of papers. The final group presents new kinematics and age constraints on the Late Cretaceous to Neogene tectonic phases and discusses the relationship of the structures with the closure of the Neo-Tethys and separation of the Arabia plate. This collection of research papers demonstrates new concepts on the opening and crustal thinning of the Levant Basin and gives updated interpretations of the latter tectonic structures of the Levant.