European Lithosphere Dynamics (Geological Society Memoirs No. 32)

European Lithosphere Dynamics

The Geological Society of London Books Editorial Committee Chief Editor B. PANKHURST ( U K )

Society Books Editors J. GREGORY (UK) J. GRIFFITHS ( U K ) J. HOWE ( U K ) P. LEAT ( U K ) N. ROBINS ( U K ) J. TURNER ( U K )

Society Books Advisors M . BROWN ( U S A ) E. BUFFETAUT ( F r a n c e ) R. GIER13 (Germany) J. GLUYAS ( U K ) D. STEAD (Canada) R. STEPHENSON (Netherlands)

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

It is recommended that reference to all or part of this book should be made in one of the following ways: GEE, D. G. & STEPHENSON, R. •. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32.

ARTEMIEVA,I.M. THYBO, H. & KABAN,M. K. Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga. In: GEE, D. G. & STEPHENSON, R. A. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 11-41.

GEOLOGICAL SOCIETY MEMOIR NO. 32

European Lithosphere Dynamics EDITED

BY

D. G. GEE University of Uppsala, Sweden and R. A. STEPHENSON Vrije Universiteit, Amsterdam, Netherlands

2006 Published by The Geological Society London

THE G E O L O G I C A L SOCIETY

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Contents Preface

vii

Introduction GEE, D. G. & STEPHENSON,R. A. The European lithosphere: an introduction ARTEMIEVA, I. M., THYBO, H. 8z KABAN, M. K. Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga

ll

ZIEGLER, P. A. & DEZES, P. Crustal evolution of Western and Central Europe

43

STAMPFLI, G. M. 8z KOZUR, H. W. Europe from the Variscan to the Alpine cycles

57

COCKS, L. R. M. & TORSVIK, T. H. European geography in a global context from the Vendian to the end of the Palaeozoic

83

Europe: Alpine to Present ZIEGLER, P. A., SCHUMACHER, M. E., DEZES, P., VAN WEES, J.-D. & CLOETINGH, S. Post-Variscan evolution of the lithosphere in the area of the European Cenozoic Rift System

97

CLOETINGH, S., ZIEGLER, P. A., BEEKMAN, F., ANDRIESSEN,P. A. M., HARDEBOL, N., VAN WIJK, J. & DI~ZES, P. Thermo-mechanical controls on Alpine deformation of NW Europe

113

KISSLING, E., SCHMID, S. M., LIPPITSCH, R., ANSORGE, J. • F~3GENSCHUH,B. Lithosphere structure and tectonic evolution of the Alpine arc: new evidence from high-resolution teleseismic tomography

129

WILSON, M. & DOWNES, H. Tertiary-Quaternary intra-plate magmatism in Europe and its relationship to mantle dynamics

147

HARANGI, S. DOWNES, H. & SEGHEDI, I. Tertiary-Quaternary subduction processes and related magmatism in the Alpine-Mediterranean region

167

HORVATH, F., BADA, G. SZAFIAN,P., TARI, G., /~D~,M,A. & CLOETINGH, S. Formation and deformation of the Pannonian Basin: constraints from observational data

191

CLOETINGH, S., BADA, G., MATENCO, L., LANKREIJER, A., HORV,A,TH, F. & DINU, C. Modes of basin (de)formation, lithospheric strength and vertical motions in the Pannonian-Carpathian system: inferences from thermo-mechanical modelling

207

VERGES, J. & FERNANDEZ, M. Ranges and basins in the Iberian Peninsula: their contribution to the present topography

223

ROBERTSON, A. H. F. Contrasting modes of ophiolite emplacement in the Eastern Mediterranean region

235

BEN-AVRAHAM, Z., WOODSIDE, J., LODOLO, E., GARDOSH, M., GRASSO, M., CAMERLENGHI,A.r VAI, G. B. Eastern Mediterranean basin systems

263

SAINTOT, A., BRUNET, M.-F., YAKOVLEV, F., St~BRIER, M., STEPHENSON, R., ERSHOV, A., CHALOT-PRAT, F. & MCCANN, T. The Mesozoic-Cenozoic tectonic evolution of the Greater Caucasus

277

Mesozoic-Palaeozoic Europe PHARAOH, T. C., WINCHESTER, J. A., VERNIERS, J., LASSEN, A. & SEGHEDI, A. The Western Accretionary Margin of the East European Craton: an overview

291

GREGERSEN, S., Voss, P., SHOMALI,Z. H., GRAD, M., ROBERTS, R. G. & TOR WORKING GROUP. Physical differences in the deep lithosphere of Northern and Central Europe

313

WINCHESTER, J. A., PHARAOH, T. C., VERNIERS, J., IOANE, D. & SEGHEDI, A. Palaeozoic accretion of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories

323

FRANKE, W. The Variscan orogen in Central Europe: construction and collapse

333

vi

CONTENTS

SIMANCAS,J. F., CARBONELL,R., GONZALEZ LODEIRO, F., PgREZ ESTAON, A., JUHLIN, C., AYARZA, P., KASHUBIN,A., AZOR, A., MART[NEZ POYATOS,D., SAEZ, R., ALMODOVAR,G. R., PASCUAL, E., FLECHA,I. & MART1,D. Transpressional collision tectonics and mantle plume dynamics: the Variscides of southwestern Iberia

345

MCCANN, T., PASCAL, C., TIMMERMAN,M. J., KRZYWIEC, P., L6PEZ-GOMEZ, J., WETZEL, A., KRAWCZYK, C. M., RIEKE, H. & LAMARCHE, J. Post-Variscan (end Carboniferous-Early Permian) basin evolution in Western and Central Europe

355

OKAY, A. I., SATIR, M. & SIEBEL, W. Pre-Alpide Palaeozoic and Mesozoic orogenic events in the Eastern Mediterranean region

389

BROWN, D., PUCHKOV, V., ALVAREZ-MARRON,J., BEA, F. & PEREZ-ESTAI)N,A. Tectonic processes in the Southern and Middle Urals: an overview

407

MATTE, P. The Southern Urals: deep subduction, soft collision and weak erosion

421

KASHUBIN, S., JUHLIN, C., FRIBERG, M., RYBALKA, A., PETROV, G., KASHUBIN, A., BLIZNETSOV, M. & STEER, D. Crustal structure of the Middle Urals based on seismic reflection data

427

BOSCH, D., BRUGUIER,O., EFIMOV,A. A. & KRASNOBAYEV,A. A. U - P b Silurian age for a gabbro of the Platinum-bearing Belt of the Middle Urals (Russia): evidence for beginning of closure of the Uralian Ocean

443

SLIAUPA, S., FOKIN, P., LAZAUSKIENE,J. & STEPHENSON,R. A. The Vendian-Early Palaeozoic sedimentary basins of the East European Craton

449

STEPHENSON, R. A., YEGOROVA,T., BRUNET, M.-F., STOVBA, S., WILSON, M., STAROSTENKO,V., SAINTOT, A. & KUSZNIR, N. Late Palaeozoic intra- and pericratonic basins on the East European Craton and its margins

463

SAINTOT, A., STEPI-IENSON,R. A., STOVBA,S., BRUNET, M.-F., YEGOROVA,T. & STAROSTENKO,V. The evolution of the southern margin of Eastern Europe (Eastern European and Scythian platforms) from the latest Precambrian-Early Palaeozoic to the Early Cretaceous

481

GEE, D. G., BOGOLEPOVA,O. K. & LORENZ, H. The Timanide, Caledonide and Uralide orogens in the Eurasian high Arctic, and relationships to the palaeo-continents Laurentia, Baltica and Siberia

507

Precambrian Europe

KOSTYUCHENKO,S., SAPOZHNIKOV,R., EGORKIN,A., GEE, D. G., BERZIN, R. & SOLODILOV,L. Crustal structure and tectonic model of northeastern Baltica, based on deep seismic and potential field data

521

HJELT, S.-E., KORJA, T., KOZLOVSKAYA,E., LAHTI, I., YLINIEMI,J. & BEAR AND SVEKALAPKO SEISMIC TOMOGRAPHYWORKING GROUPS. Electrical conductivity and seismic velocity structures of the lithosphere beneath the Fennoscandian Shield

541

KORJA, A., LAHTINEN,R. & NIRONEN, M. The Svecofennian orogen: a collage of microcontinents and island arcs

561

DALY, J. S., BALAGANSKY,V. V., TIMMERMAN,M. J. & WHITEHOUSE, M. J. The Lapland-Kola orogen: Palaeoproterozoic collision and accretion of the northern Fennoscandian lithosphere

579

BOGDANOVA, S., GORBATSCHEV,R., GRAD, M., JANIK, T., GUTERCH, A., KOZLOVSKAYA,E., MOTUZA, G., SKRIDLAITE, G., STAROSTENKO, V., TARAN, L. & EUROBRIDGE AND POLONAISE WORKING GROUPS. EUROBRIDGE: new insight into the geodynamic evolution of the East European Craton

599

SLABUNOV,A. I., LOBACH-ZHUCHENKO,S. B., BIBIKOVA, E. V., SORJONEN-WARD, P., BALAGANSKY,V. V., VOLODICHEV, O. I., SttCHIPANSKY,A. A., SVETOV, S. A., CHEKULAEV,g. P., ARESTOVA, N. A. & STEPANOV, V. S. The Archaean nucleus of the Fennoscandian (Baltic) Shield

627

CLAESSON, S., BIBIKOVA, E., BOGDANOVA,S. & SKOBELEV, g. Archaean terranes, Palaeoproterozoic reworking and accretion in the Ukrainian Shield, East European Craton

645

Index

655

Preface This Memoir, 'European Lithosphere Dynamics', has a history that goes back more than 20 years. At the International Geological Congress (IGC) in 1984 in Moscow, leading Earth Scientists from Western Europe and the Soviet Union agreed to start a new pan-European project, 'EUROPROBE', focused on the European Lithosphere and modelled on the European Geotraverse (EGT, Blundell, Freeman and Muller; see Introduction for references). The latter had started a few years before and was dedicated to integrated geological and geophysical studies of a north-south transect across Europe from the Barents Sea coast of northwestern Norway to Italy and the central Mediterranean. At the 1984 IGC, EUROPROBE was conceived as a comparable, east-west profile, from the border zone between Asia and Europe in the middle Urals to the Iberian Peninsula and the Atlantic margin. EUROPROBE took a few years to climb out of the cradle, but as 'glasnost' and 'perestroika' took over the latter part of the 1980s, the opportunity for wider East-West collaboration was recognized by the President of the International Lithosphere Programme (ILP), Karl Fuchs. At that time we were involved in ILP's Global Geoscience Transects programme and the EUROPROBE initiative was seen in this context. The European Science Foundation (ESF) funded early preparatory meetings in Russia, Poland, Denmark and Germany, and advice was provided from many sources, not least EGT. It was clear by 1990 that support for EUROPROBE was to be found in many European countries and a suitable programme, directly benefiting the many nations, was designed. The simple long-range transect model was abandoned and EUROPROBE emerged 'dedicated to carrying out a new generation of major projects that will improve our understanding of the tectonic evolution of the Earth's crust and mantle, and the dynamic processes which controlled this evolution through time' (Gee & Zeyen 1996). Nine target areas (see Fig. 1) were selected for the main research activities, each run by a research team with a high level of autonomy, and all dedicated to applying integrated geological, geochemical and geophysical methods to understand surfacedepth relationships and to interpret the processes leading to the formation of major features of the European lithosphere. Most of the latter were orogens, ranging in age from Archean to the Present; intra-cratonic rifting was also prominent. For ten years, from 1992 to 2001, EUROPROBE received support from the European Science Foundation at a level that financed

Fig. 1. EUROPROBEprojects.

meetings for Science and Management committees and allowed each of the ten projects to run a workshop. These annual activities were essential for the health of the individual projects and the cohesion of the general programme. There were no central research funds other than a small budget devoted to programme travel and exchange of scientists. The individual projects defined their own goals and leadership and obtained funding for research. Those involving EastWest cooperation were often successful in obtaining support from INTAS (The International Association for the Promotion of Co-operation with Scientists from the New Independent States, NIS, of the former Soviet Union). By 2002, the time had come to review a decade of ESF sponsored research (Gee & Artemieva EUROPROBE 1992-2001). At a meeting held in conjunction with the award ceremony for the Crafoord Prize (Prize-winner Dan McKenzie) in Stockholm, a large group of EUROPROBE geoscientists agreed on the ambitions of a final Memoir. The book would not be confined to the achievements of the ESF programme, but target European Lithosphere Dynamics in general, probing the main features of the European subcontinent to give the reader an overview of the whole development through time. About 30 countries and many hundreds of geoscientists were involved in ESF's decade of EUROPROBE research. About 80 workshops were held all over Europe, from Ekaterinburg to Lisbon and Ankara to Lammi; these were great years of multinational collaboration and the publications that resulted were innumerable. Support for EUROPROBE in eastern Europe was widespread and we particularly remember the encouragement and guidance of the Academy Vice-President Alexander Yanshin in Russia, Academician Vitaly Starostenko in Ukraine, and Academician Radim Garetsky in Belarus. Dr Andrey Morozov at the Russian Ministry of Natural Resources and now with Rosnedra has been the 'foundation' for much of EUROPROBE's research in Russia; he cannot be thanked enough. And for all the workshops and research logistics in the former USSR, Elena Gornaya and, later, Nadezhda Timofeeva were vital communicators and advisers, organizers and entertainers. EUROPROBE operated from a secretariat in Uppsala where Chris Juhlin, Herman Zeyen, Monica Beckholmen and Irina Artemieva kept the programme rolling through the 1990s and Olga Bogolepova since then. We are hugely indebted to our EUROPROBE colleagues, not only for the success of the programme, but also that a foundation for international collaboration in lithosphere science has been established that recognizes no end. Many of the major projects that were planned towards the end of the EUROPROBE's ESF Programme have been realized during the last five years, e.g. IBERSEIS, CAUCASUS, POLAR URALS, KOLBAKAR, FIRE, to name but a few. This Memoir has been monumental job for the Geological Society Publishing House in Bath, UK. We thank Angharad Hills and her colleagues warmly for taking it on and seeing it through, despite many hic-ups and a hip-out en route. Our 200 authors have built a comprehensive overview of the European Lithosphere and helped us with some of the reviewing, for most of which we thank the following: U. Achauer, A. Adam, J. Ansorge, I. Artemieva, G. Bertotti, S. Bogdanova, O. Bogolepova, T. Brewer, F. Beunk, C. Biermann, D. Brown, M.F. Brunet, J. Carney, F. Chalot-Prat, R. Cocks, M. Comas, S. Daly, J. Davidson, C. Doglioni, M. Friberg, K. Fuchs, M. Gaetani, J. Golonka, R. Gorbatschev, A. Gubanov, D. Harper, A. Hegedus, R. Huismans, L. Jolivet, A. Jones, J. Knapp, Y. Mart, P. Matte, A. Mauffret, J. Mosar, F. Neubauer, S. Nielsen, A. Okay, T. Pharaoh, A. Robertson, R. Rutland, A. Saintot, J. Tait, T. Torsvik, D. White, J. Winchester, G. W6rner, T. Yegorova and P. Ziegler. D. G. GEE and R. A. STEPHENSON

The European lithosphere: an introduction DAVID G. GEE 1 & RANDELL A. STEPHENSON 2

1Department of Earth Sciences, Uppsala University, Villavagen 16, SE-75236 Uppsala, Sweden (e-mail: david.gee @geo. uu.se) 2Netherlands Centre for Integrated Solid Earth Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands (e-mail: randell.stephenson @falw. vu.nl)

Europe provides on outstanding field laboratory for studying lithospheric processes through time: for tracing the results of plate movements from the present back into the early Precambrian. This book has been designed to focus on tectonic processes in the European lithosphere through these three billion years and how they may have changed during this time. Two things are particularly striking: the importance of plate tectonics far back through the Proterozoic into the Archaean, and the significance of tectonic inheritance, older structures and rheologies guiding, even defining, the younger evolution. Basement structure has a profound influence on subsequent basin evolution and the distribution of geo-resources. The economic importance of understanding these processes cannot be overestimated. Understanding the dynamics responsible for the construction of continental lithosphere requires integrated interpretation of geological, geophysical and geochemical observations. Hypotheses often benefit from testing by numerical and analogue modelling. In practice, one technology--multi-channel, near-vertical reflection profiling--has played a leading role in connecting surface observations to the deep crust and mantle structure. Combined with other geophysical methods, deep reflection profiling has guided the interpretation of the processes that created the architecture of the lithosphere. The European part of the Eurasian continent, reaching from the Ural Mountains in the east to the Iberian Peninsula in the west and from the Mediterranean into the high Arctic, has a lithosphere that

can readily be treated in two parts, east and west (Fig. 1). Most of eastern Europe is dominated by the old, cold East European Craton (EEC), partly covered by little deformed Phanerozoic and MesoNeoproterozoic rift and platform successions. Flanking the EEC to the east are the late Neoproterozoic Timanides and both this orogen and the craton are abruptly truncated by late Palaeozoic Uralian sutures, marking the border to Asia. Northernmost Europe is dominated by the Caledonides and Timanides. The southeastern edge of the EEC, from the northern parts of the Black and Caspian seas to the southern Urals, is less easily defined, an older history of Neoproterozoic accretion and Palaeozoic tectonics being overprinted by Alpine deformation and uplift, the latter being displayed most prominently in the Caucasus, a mountain belt crowned by Europe's highest peak, Mount Elbrus (5 642 m). Western Europe, with minor exceptions, is composed of thinner, warmer, dominantly Phanerozoic lithosphere, accreted to the EEC during Palaeozoic and younger orogenesis. A broad zone of suturing, reaching from the North Sea to the Black Sea and Anatolia, separates the Craton from Phanerozoic accreted terranes, Caledonide in the north, Variscide in central regions, and Alpine in the south. The term 'Trans-European Suture Zone' (TESZ) was coined by EUROPROBE (see Preface) for this wide zone of deformation, with faulting and tectonic reactivation, many strands of which involve large displacements of the craton margin, where the latter tapers westwards beneath a Palaeozoic platform cover.

Fig. 1. Tectonic map of Europe, showing the distributionof the East European Craton and main orogens.

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 1-9. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

2

D.G. GEE & R. A. STEPHENSON

Fig. 2. Magnetic map of Europe (by courtesy of S. Wybraniecand H. Thybo).

Defining the Palaeozoic sutures, as they occur beneath the Mesozoic and younger cover of the TESZ, and tracing them through the crust into the upper mantle is more difficult along this western edge of the EEC than in the Urals, where post-Palaeozoic deformation does not obscure first-order Uralian structures, including sutures. The contrast in structure and composition of the crusts of eastern and western Europe is clearly seen in the magnetic (Fig. 2) and gravity (Fig. 3) maps of the region. In the Early Palaeozoic, the EEC formed the core of a continent, Baltica, largely covered by a wide shelf sea. Faunal evidence, supported by paleomagnetic data, has provided the foundation for Baltica's inferred Palaeozoic independence, apparently surrounded by oceanic crust, at least during the Cambrian and Ordovician. The shape of Baltica is usually considered to have been roughly circular, comparable to Australia in size. However, the shelf-edge and sutures surrounding Baltica are not all well defined and uncertainties remain, particularly in the far north and in the south, the latter affected by the opening and

closing of Palaeo-Tethys and subsequent convergence and collisional events that have left the Palaeozoic history of this margin obscure. The oldest rocks yet found in Europe occur in Ukraine and are mid-Archean in age (c. 3700 Ma). They comprise part of the Sarmatian segment of the EEC. Two other major segments of the EEC have been recognized, Fennoscandia and Volga-Uralia, and these three segments were assembled in the late Palaeoproterozoic, prior to uplift and erosion and a long subsequent period of mild intraplate deformation. The configuration of this proto-EEC in the Proterozoic and its relation to other continent-bearing plates and to larger assemblies of continents (such as Rodinia), remain to be clearly defined. Fragments of the proto-EEC were certainly rifted off the craton during the Palaeozoic and Mesozoic and possibly also during the younger Proterozoic. Grenville-age orogenesis, so prominent in other parts of the Precambrian world, is clearly defined in Europe only along the western (Sveconorwegian) margin of the EEC in southern

INTRODUCTION

Fig. 3. Gravity map of Europe (from Wybranic et

al.

3

1998), Bouguer anomaliesover land areas and Free Air over the seas, shaded relief image NE illuminated.

Norway and south-western Sweden. However, c. 1000 Ma signatures have been detected on several parts of the craton margin, from Baltica-derived allochthons of the Scandinavian Caledonides to the Timanides of the high Arctic and the Urals. Either Grenville-age terranes were accreted to the EEC during c. 1000 Ma orogeny (e.g. in the Scandes) or, subsequently, during Neoproterozoic orogeny. While there is evidence of intracratonic rifting within the EEC, at least locally during the Neoproterozoic, the eastern flank of what was then proto-Baltica, from Novaya Zemlya in the north to westernmost Kazakhstan in the south, was influenced by Timanian orogenesis. From the southern Urals westward to the Scythian Platform of southern Russia and Ukraine, there is also indirect evidence indicating that Timanian-age orogenesis may have preceded the development of an Early Palaeozoic platform. And further west within the Trans-European Suture Zone, there is local evidence of Neoproterozoic accretion of outboard terrains to the EEC prior to deposition of Cambrian strata. Only along the Baltoscandian margin of the craton is there a well-defined depositional

history that apparently denies the influence of proximal orogeny in the Neoproterozoic. Outboard of Baltica, in the Palaeozoic and younger terranes of western Europe, there is abundant evidence of Neoproterozoic, tectono-thermal activity, referred to as the Cadomian Orogeny. From the Avalonian terranes of the British Isles and northern France to the internal 'basement' fragments of the Variscide and Alpine foldbelts, there is evidence of Cadomian convergent-margin tectonics; this tectono-thermal activity apparently occurred along the subducting margin of Gondwana prior to being transferred across to the Baltica margin during closure of intervening oceans in the Early-Mid-Palaeozoic. These Cadomian terranes carry with them evidence of earlier Proterozoic lithosphere, with age signatures that are not characteristic of the western (today's co-ordinates) EEC. Closure of Tethyan ocean systems and collision of Africa with Eurasia resulted in the complexities of southern Europe's Mediterranean world, with development of Alpine fold belts, from the Pyrenees to the Caucasus. Subduction is continuing today, with

4

D.G. GEE & R. A. STEPHENSON

volcanic arcs and back-arc basins, major thrusting and transcurrent and normal faulting. The signatures of these processes in the deeper crust and mantle are well seen in seismic tomography, potential field and thermal anomalies. All this convergence was going on while the North Atlantic was opening, while the passive margins, volcanic and non-volcanic, of Europe were developing, and the Iceland plume migrating to beneath the midAtlantic ridge. 'European Lithosphere Dynamics' is not a comprehensive account of the European lithosphere, but provides an overview of many of the more important aspects of the European crust and mantle. It is arranged to lead the reader, via introductory chapters treating large parts of the subcontinent, into the Alpine world and then backwards in time through the Variscides, Uralides, Caledonides, and Timanides into the early Proterozoic and Archaean of the East European Craton. The Alps--and associated orogenic belts such as the Pyrenees, Carpathians, and Caucasus--represent the results of the most recent phase of mountain-building in Europe, the one related to the collision of the African and Arabian plates with the European (Eurasian) plate in Cretaceous and Cenozoic times. Associated with Alpine collision is the European Cenozoic Rift System, one of the main components of which is the Rhine Graben. Ziegler et al. utilizing in part inferences about the crustal structure of western and central Europe derived from the Moho compilation presented in the Introductory section of this Memoir (Ziegler & D~zes), argue that the manifestation of Alpine tectonics is to some extent an 'accident' of what was left of Variscan crustal/ lithospheric structure. This was, in turn, inherited from Variscan tectonic processes and how they had interacted with older (Cadomian and Caledonian) structures. Thus, what remains of the results of Early Palaeozoic and Neoproterozoic collisional processes in Europe is the consequence of their convolution with Variscan processes and what remains of the results of Variscan processes is, in turn, related to the extent of their interaction with Alpine processes. The record of this multiphase evolution of the lithosphere, north of the Alps, is discussed further by Cloetingh et al. [a] who show how the thermo-mechanical structure of the lithosphere, in part defined by its tectonic memory, may also control Late Neogene and neotectonic anomalies in crustal subsidence and uplift, linking these with surface processes and topography evolution. While these insights are derived primarily from modelling studies, Kissling et al. have come to similar conclusions by looking at the present-day lithospheric structure of the Alps from tomographic data constrained with high resolution seismic models of crustal structure. They show that substantial differences in the structure of the deep crust appear between the western, central, and eastern Alps. Again, it seems likely that this is a manifestation of the inheritance of particular lithospheric architectures from earlier accretionary (Variscan and older) events. Tertiary-Recent anorogenic intraplate magmatism was widespread in Europe and is spatially and temporally associated with Alpine-Pyrenean collisional tectonics, the development of the Cenozoic rift system in the northern foreland of the Alps (e.g. Ziegler et al.), and, locally, with uplift of Variscan basement of the Massif Central, Rhenish Massif and Bohemian Massif. These are much the same regional structures identified by Cloetingh et al. [a] that are related to broad thermo-mechanical lithospheric heterogeneities, at least in part due to inheritance. Wilson & Downes conclude that the partial melting of the mantle leading to volcanism was induced by adiabatic decompression of the asthenosphere, locally in small-scale, plume-like diapirs that welled up from c. 400 km depth. Tertiary-Recent volcanism in Europe may therefore be the surface expression of a 'warm' European upper mantle interacting with a compositionally heterogeneous overlying lithosphere, the latter 'filtering' the former into its diffuse pattern.

The Alpine-Mediterranean area is characterized by a system of arcuate Cenozoic orogenic belts and extensional basins, both of which can be explained by the roll-back of subducted slabs and retreating subduction zones, in the convergence zone between Eurasia and Africa. Harangi et al. summarize the main characteristics of Tertiary-Quaternary 'subduction-related' magmatism in the Mediterranean region and argue that its compositional variability can be explained by having a strongly inhomogeneous mantle--related to 'accidents' in the nearby upper mantle such as the central Atlantic plume--accompanied by crustal contamination. Ben-Avraham et al. describe the extensional basin systems developed in the central and eastern Mediterranean Sea. They show that there is a fundamental change of style at the end of the Miocene, when major adjustments in the Africa-Europe convergent plate boundary occurred, related to the collision of the Arabian plate with Eurasia and the development of the Anatolian and Aegean terranes as independent microplates. The remnants of Neotethys-related (mainly Late Cretaceous and Tertiary) ocean basins found in the deformed zones of the eastern Mediterranean are addressed by Robertson who concludes that not all ophiolites were emplaced as a result of large-scale horizontal tectonic transport, but that a strike-slip/transpressional tectonic setting may have dominated in some cases. The evidence suggests that the mode of ophiolite emplacement was strongly influenced by first-order inheritance such as the relative orientations of ophiolite emplacement vis-~-vis the adjacent continental margin. Horvath et al. discuss continental collision and back-arc basin evolution as one single, complex dynamic process, with the minimization of potential and deformational energy as the driving principle, as exemplified by the Pannonian Basin, which formed from a collapsing, over-thickened Alpine lithosphere in the Neogene. A key requirement was the presence of a 'free boundary' offered by the rollback of the subducting Carpathian slab, thus allowing orogen parallel crustal extrusion towards the east. Modelling results suggest that, as a whole, the Pannonian Basin area has displayed pronounced lithospheric weakness since Cretaceous times (Cloetingh et al. [b]) and, therefore, has been prone to repeated tectonic reactivation. Pronounced lateral variations in lithospheric strength, at least in part related to geological inheritance, have strongly influenced the thrust load kinematics and post-collisional tectonic history along the adjacent Carpathians Mountains and their foreland. Mountain-building at the European extremes of the Alpine belt--the Pyrenees (and associated topographic features of the Iberian Peninsula) and the Greater Caucasus--are discussed by Verg6s & Fern~ndez and Saintot et al. [a], respectively. Once again tectonic inheritance plays a key role. The distribution of modern topography on the Iberian Peninsula (basins and mountain ranges) is argued to be the consequence of crustal and lithospheric thickening during Tertiary compression and upper mantle thinning during the Neogene-Quaternary, superimposed upon variations in crustal (and possibly mantle) densities. These are understood to be a legacy of Late Palaeozoic orogenesis and lithospheric accretion. Similarly, the Greater Caucasus (GC) fold-and-thrust belt, developed in response to Tertiary ArabiaEurasia collision, represents the structural inversion of a deep marine Mesozoic basin. Changes in tectonic style along the GC--in which varying degrees of thick- and thin-skinned deformation are displayed--may be related to heterogeneity developed in lithosphere that was accreted or modified much earlier, in Late Palaeozoic and Mesozoic times. A key process controlling how Alpine tectonics became convolved with the older Variscan framework of Europe was the closure of the Palaeotethys and opening of Neotethys ocean systems and the development of an array of south Eurasian back-arc basins, followed or accompanied by the break-up of Pangaea and the early development of the Central Atlantic. This stage of European tectonic history can be broadly referred to as 'Cimmerian'. While Stampfli & Kozur review the plate kinematic

INTRODUCTION

record during this time, focusing on accretionary events along the south Eurasian margin, McCann et al. consider the instability and re-equilibration of western and central European intraplate lithosphere that occurred after the Variscan Orogeny. An extensive phase of Permo-Carboniferous magmatism was accompanied by transtensional activity that led to the formation of more than seventy rift basins across the region (see Ziegler et al.). These basins can be characterized according to their position relative to the Variscan Orogen and its structural trends. The geological record of Variscan and Cimmerian orogenesis in the Eastern Mediterranean-Balkan region, with a focus on Anatolia, is described by Okay et al. The pre-Alpide evolution of this region is one of episodic growth of Europe by the accretion of oceanic terranes and Gondwana-derived micro-continents, as outlined by Stampfli & Kozur. While the Palaeozoic history of the Balkans and Pontides resembles that of Central Europe, its Mesozoic (Cimmerian) evolution--because of the consolidation of Pangaea to the west--diverges. Saintot et al. [b] investigate the history of the zone between Anatolia and the EEC--from the TESZ and Carpathians in the west to the southern Urals in the east--perhaps the least known part of the European lithosphere. These authors conclude that, although this area was dominantly an oblique convergent plate margin from the Late Palaeozoic through the Mesozoic, there is no compelling evidence for accretionary orogenesis of Variscan (Carboniferous-Permian) age. Rather, the available data are more consistent with an interpretation in which the crust of the Scythian Platform, from the Pre-Dobrogean Depression across the Crimean Peninsula to the North Caucasus area, represents the thinned margin of the Precambrian continent, reworked by mainly extensional Late Palaeozoic-Early Mesozoic tectonic events (in the hanging wall of a transform to obliquely convergent plate margin; cf. Fig. 4). They suggest that the Precambrian crust may have been accreted to the EEC during the Neoproterozoic, roughly contemporaneously with the Timanian Orogeny along its northeastern and eastern margin, (cf. Gee et al.). Thus, Baltica in the Early Palaeozoic was probably not as circular as previously thought, but had prolongations to the north- and southeast comprising Neoproterozoic accreted terranes. Palaeozoic orogenesis dominated the tectonic evolution both of western Europe and the eastern margin of Baltica along the

Fig. 4. Palaeozoic peri-Atlantic orogens in the Permian (from Matte 1991).

5

border-zone to Asia (Fig. 4). The relationship of these orogens to plate movement is treated by Cocks & Torsvik, who use palaeontological, palaeomagnetic, and other lines of evidence to reconstruct the relationship between Baltica and other continents, in particular the dominating supercontinent Gondwana. For about 100 million years, from the end of the Vendian through the Cambrian and Ordovician, Baltica existed as an independent continent surrounded by oceanic domains and largely covered by platform successions. By the Early Ordovician, subduction systems were closing the oceans and collision, with accretion of microcontinents (e.g. Avalonia) started in some areas in the Late Ordovician. Caledonian, Variscan and Uralian orogenesis followed though the mid and late Palaeozoic and much of the basic architecture of today's Europe was assembled, first as part of Laurussia and thereafter Pangaea. Only along the southern margin, as mentioned above, did subsequent Tethyan tectonics and Alpine collision of Africa with Eurasia substantially change the geology of the European plate. The Caledonides of northwestern Europe (Dewey & Strachan 2005; Gee 2005) define the collision zone of Laurentia with Baltica and the suturing of Laurussia. Subduction of the Baltoscandian margin of Baltica started in Late Cambrian-earliest Ordovician times and the Iapetus Ocean, separating these two continents, closed during the Ordovician, with collisional orogeny (Scandian) lasting from the Early Silurian into and through the Early Devonian. The sedimentary basin record of these events on the northwestern EEC is described by Sliaupa et al. Contractional orogeny had ceased by the Mid-Devonian and most of the EEC was undergoing extension, accompanied, in the later Devonian by widespread mafic volcanism (cf. Stephenson et al.). A SE-trending branch off the main Caledonide Orogen has been defined along the southwestern margin of the EEC in Denmark, northern Germany, and Poland (Katzung et al. 1993) with thrusting to the NE, well defined by seismic surveys and drilling. This deformation zone occurs in the footwall to overthrust Avalonian terranes and forms part of the Trans-European Suture Zone (TESZ), the latter defining the broad boundary between Phanerozoic western Europe and the EEC (Pharaoh et al.). Suturing of Gondwana-derived terranes to Baltica had begun by the end of the Ordovician and continued through the Palaeozoic (Fig. 4). Much of the TESZ structure is obscured by younger cover (cf. McCann et al.). It follows that deciphering the Palaeozoic history has depended on comprehensive integration of geophysical techniques for defining the structure of the deeper crust and mantle. Winchester et al. draw attention to the problems involved in defining the accretion of Neoproterozoic terranes along the margin of the EEC. Some of the former, in southern Poland, the Czech Republic, and probably Romania (Moesia), were apparently accreted before the Ordovician (perhaps early in the Cambrian) and then transgressed by Baltica's platform successions. These Neoproterozoic terranes are inferred to have existed as a promontory later in the Palaeozoic when Gondwana-derived Avalonian continental fragments docked along the Baltica margin. The central European Variscides have been treated extensively in Franke et al. (2000) and Winchester et al. (2002). In this volume, Franke looks at Palaeozoic plate kinematics and divides the Variscan tectono-stratigraphic evolution into major Early and Late Palaeozoic episodes, focusing on the latter. Relationships between Baltica and Gondwana during the development of the Variscides, from oceanic separation to the close proximity and the establishment of Pangea, are discussed. Franke also draws attention to evidence of widespread Mid-Devonian extension and mafic magmatism in the central Variscides; he discusses possible genetic affinities with the intraplate extension occurring at about the same time within the adjacent EEC, specifically in the Dniepr-Donets Basin, described by Stephenson et al. The Variscide Orogen also dominates the geology of the Iberian Peninsula. Simancas et al. focus on a southern transect where a

6

D.G. GEE & R. A. STEPHENSON

deep reflection seismic profile was recently acquired. The profile runs from the southernmost terrane in Iberia--the South Portuguese Zone--northwards into central part of the peninsula, crossing a major zone of transcurrent faulting that dominates this part of the Variscides. The paper relates surface geology to deep crust and mantle structure in one of Europe's outstanding examples of transpressional orogenesis. The major differences in the character of the lithosphere across the TESZ, described by Pharaoh et al., so conspicuous in compilations of the gravity and magnetic fields, are equally apparent at sub-Moho levels (Fig. 5) and deep into the mantle (e.g. Zielhuis & Nolet 1994). Major long-range, wide-angle reflection/refraction seismic surveys such as 'Polonaise' of the 1990s and Celebration-2000 (Janik et al. 2005), that ran from the EEC westward to the Variscan terranes of central Europe, were complemented by regional tomography studies, such as TOR (Gregersen et al.). High angle, non-symmetrical features extend deep into the mantle in the vicinity of the Tornquist Zone, displacing the SW-tapeting margin of the Craton (Grad et al. 2002). The Palaeozoic orogen of easternmost Europe, exposed in the Ural Mountains from the Aral to the Kara seas, marks the edge of the EEC from 48 ~ to 60 ~ N. Further north, the Urals truncate the grain of the NW-trending Timanides to 68 ~ N and then swing northwestwards into the Pai-Khoi-Vaigach-Novaya Zemlya fold and thrust belt. The Uralide Orogen, treated in four papers in this volume, is renowned for its preservation of ophiolites (Saveleva & Nesbitt 1996), arc and back-arc volcanic rocks (Brown et al. 2000), and associated mineralization. Footwall blueschists and eclogites are well exposed from the southern to the polar Urals, crystallizing in subducted Baltica-margin protoliths. For much of the Palaeozoic, from the Late Cambrian to the Carboniferous, the Uralian edge of Baltica developed as a passive margin (cf. Saintot et al. [b]; Sliaupa et al.); outboard, subduction-related complexes that formed in the Uralian ocean did not influence the off-shelf, slope-rise facies of the eastern edge of the EEC until the Late Devonian to Early Carboniferous. Not until the Late Carboniferous did the shelf collapse and flysch followed by Permian molasse herald Uralian orogenesis. This apparently occurred somewhat earlier in the south than in the far north, where folding and thrusting did not influence Novaya Zemlya until the middle Triassic. Comprehensive geophysical investigations of the Urals during the 1990s focused on profiles through its southern and central parts; a polar profile is now in progress. These investigations

Fig. 5. Shear wave velocity variations (in %) beneath Europe at a depth of 80 km with low velocities beneath Phanerozoic Europe in the west and high velocitiesbeneath the East European Craton (based on Zielhuis & Nolet, 1994).

provided the foundation for the integrated geological-geophysical studies reported in this volume. Brown et al. focus on the southern part of the mountain belt, where a 500 km long transect across the orogen was investigated in the mid-1990s by a combination of nearvertical seismic profiling, wide-angle reflection and refraction, and potential field methods (Berzin et al. 1996; Knapp et al. 1996). Brown et al. draw attention to the bivergent character of the orogen, with evidence for volcanic-arc collision with the Kazakhstan continent in the east, prior to closure of the ocean. They also discuss the evolution of intra-oceanic arcs, which occurred prior to Baltica-Kazakhstan collision, and the western foreland fold and thrust belt with its evidence of only limited shortening. Matte also concentrates on the southern Urals, emphasizing the contrast between the Uralides and Variscides, from the existence of the deep Moho beneath the former to the remarkable preservation of the middle Palaeozoic volcanic complexes; their dense root is inferred to account for the thicker crust. Accretion of oceanic and microcontinental terranes was not achieved by orthogonal collision and extreme overthrusting, as in the Scandinavian Caledonides; instead, transpression appears to have dominated the remarkably linear Uralide Orogen. The middle Urals transect is presented by Kashubin et al.; this is based on more than a dozen years of integrated geophysical and geological investigations. These authors summarize the evidence for relating the surface geology to the deep structure defined by CDP profiling and present a synthesis of the orogenic evolution. The crustal roots beneath the middle Urals reach 60 km depth and the truncation of the craton margin is abrupt. Interestingly, this part of the Baltica margin is marked by Timanian blueschists (Beckholmen & Glodny 2004) indicating that the late Neoproterozoic suturing coincides closely with that in the Late Palaeozoic. In the middle Urals, Palaeozoic terranes dominate the accretionary complex and microcontinents are apparently absent. The Ural mountain belt is famous for its mineralization, particularly volcanic-hosted massive sulphides (Koroteev et al. 1997; Allen et al. 2002). One paper in this volume, Bosch et al. directly concerns mineralization. Mafic-ultramafic bodies of the middle Urals, occurring within ocean-derived allochthons, are notable for their locally high platinum contents. Bosch et al. provide new isotope data indicating that these ophiolite-related allochthons are middle to Late Silurian in age, apparently substantially older than other similar Pt-bearing massifs of the Late Devonian. This range of ages coincides with those of related adjacent volcanic-arc complexes and testifies to the longevity of intra-oceanic Pt magmatism. The Uralide orogenic assemblage of ocean-derived allochthons (ophiolites and arc-volcanics), footwall high-pressure blueschists and eclogites and extensively developed late-orogenic granites, reaches from the far south of the mountain belt to the Kara Sea. Further north in Pai Khoi, Vaigach, and Novaya Zemlya only a fold and thrust belt is exposed, similar to that of the Uralian western foreland. Evidence for the continuation of the classical Uralide Orogen northwards into the Barents-Kara shelf and eastwards to Taimyr, as proposed by many authors (e.g. Bogdanov et al. 1996), is reconsidered by Gee et al.; the evidence is far from compelling. Likewise, the proposal (Cocks & Torsvik) that the Timanide margin of Baltica terminates immediately north of Novaya Zemlya is also in doubt. These interpretations are important for any attempt to reconstruct the evolution of the Arctic Basin; the northernmost shelf areas of Eurasia, with their vast hydrocarbon resources, will be the focus of many new studies in the coming years. Palaeozoic orogenesis, Caledonian and Hercynian along the margins of the EEC, was interrupted in the Mid-Late Devonian to Early Carboniferous by intracratonic tiffing and extensive basaltic magmatism. In a review of the Late Palaeozoic rift basins of the EEC, Stephenson et al. point out that the DnieprDonets Basin is a true intracratonic rift basin, cutting across the

INTRODUCTION Archaean-Palaeoproterozoic structural grain of its (Sarmatian) basement, whereas the East Barents-Pechora Basin (Gee et al. Kostyuchenko et al.) and the Peri-Caspian Basin (more speculatively) are pericratonic features, developed on reworked and juvenile crystalline basement accreted to the EEC during the Neo-(?Meso)proterozoic. It is speculated that some of this late Precambrian lithosphere may have rifted away from Lanrussia in the Late Devonian. Sliaupa et al. also note that Peri-Uralian basins developed as passive continental margin basins throughout the Early Palaeozoic. Sedimentation on the EEC was confined to the cratonic margins at this time with only limited intracratonic subsidence, in two distinct geodynamic settings, one where basins formed in response to continental break-up processes (break-up of Rodinia) and the other, where basins formed in response to the reassembly of continental lithosphere fragments and associated continental accretionary processes (Neoproterozoic and Caledonian). Fundamental differences in the thermo-compositional make-up of the Precambrian lithosphere of Europe played an important role in the development of the overlying Palaeozoic sedimentary basins described and discussed by Sliaupa et al., Stephenson et al. and Saintot et al. The influence of basement character on the distribution of Barents shelf hydrocarbon resources is noted by Gee et al. Artemieva et al. present a variety of lithospheric scale geophysical data indicating that differences in structure have both a compositional and a thermal origin and are a legacy of Precambrian terrane accretion and subduction as well as Phanerozoic rifting, volcanism, subduction, and continent-continent collision. With regard to the East European Craton, dominated by Archaean and Palaeoproterozoic terranes (Bogdanova et al.

7

2005), amalgamation of its three different segments (Fig. 6) is inferred to have occurred at c. 1800 Ma (Bogdanova 1993). The craton is largely covered by younger sedimentary rocks and is best exposed in the Fennoscandian Shield; it also crops out in the Ukrainian Shield and Voronezh Massif. Deep drilling and the analysis of invaluable drillcores, together with potential field and seismic surveys, has provided the foundation for current knowledge of the crustal rocks of the EEC--their igneous, sedimentary, and metamorphic histories and their evolution through the Archaean and Proterozoic. In this volume, aspects of the EEC are treated in six papers. One of these addresses the lithosphere as a whole; the seven others concern the surface geology and crustal structure of the Fennoscandian (Baltic) Shield, and the largely unexposed, but extensively drilled, regions of southernmost Fennoscandia and their relationships to Sarmatia. The Fennoscandian Shield is readily divisible into northern and southern regions separated, in eastern Finland and western Russia, by the Karelian block--an Archaean complex, overlain by Paleoproterozoic metasediments, but little influenced by tectonothermal activity of this age. Both to the north and to the south, evidence for Palaeoproterozoic orogenesis is widespread, in the north involving substantial Late Archaean terranes. Daly et al. present a tectonic synthesis of the northern part of the Shield, dominated by the Lapland-Kola Orogen. Here, extensive studies of structure, metamorphism, and geochemistry (including isotope age) together with deep reflection and wide-angle seismic profiling (Kostyuchenko et al.) have provided new insight into a Palaeoproterozoic belt of collision and accretion involving both Archaean terranes and younger juvenile crust. Slabunov et al. focus on the Archaean terranes of the northern part of the Shield, summarizing evidence of meso-Archaean protoliths

Fig. 6. Three-segmentconfiguration of the East European Craton (by courtesy of Svedana Bogdanova and Roland Gorbatschev).

8

D.G. GEE & R.A. STEPHENSON

and late Archaean orogeny. Both meta-ophiolites and high pressure eclogite-bearing assemblages are described, indicating the existence both of oceanic crust and deep underthrusting of continental crust during Archean orogenesis. Subsequent Palaeoproterozoic tectonothermal reworking and south-vergent thrusting emplaced these older complexes onto the northern margin of the Archaean Karelian Craton. To the south of the old Karelian core of Fennoscandia, a wide variety of Paleoproterozoic complexes occur that were thrust northwards and accreted to the Archean margin towards the end of the Palaeoproterozoic. Korja et al. integrate surface geology with a substantial geophysical database in Finland and northern Scandinavia. They present a model of the Svecofennian Orogeny involving several late Palaeoproterozoic pulses of accretion, terminating with gravitational collapse at about the same time (1.78 Ga) as orogenesis ceased along the northern side of the Karelian block. Further south in the EEC, the southern parts of the Fennoscandian segment and their relationships to Sarmatia have been investigated by a large multinational group of geologists and geophysicists. Bogdanova et al. describe a wide range of largely juvenile Palaeoproterozoic terranes, beneath the Neo-(? Meso)proterozoic cover of the craton. Relationships to Sarmatia are recognized and correlated with outcrops of igneous and metamorphic complexes in the Ukrainian Shield. Potential field data and wide-angle refraction/reflection seismic profiling allows definition of the extent and geometry of the terranes and interpretation of the crustal and upper mantle structure. These provide a comprehensive foundation for definition of the suture zone between Fennoscandia and Sarmatia and tectonic modelling of accretional, followed by collisional, orogenesis. Sarmatia is the focus for Claesson et al. who present new isotope age data including evidence of the oldest protoliths in the EEC (c. 3.65 Ga; perhaps 3.75 Ga). Most, but not all, of the ancient terranes are late Archaean in age but Palaeoproterozoic reworking was widespread, particularly during collision with Fennoscandia. The deep lithosphere of specific regions of the EEC was the subject of two major studies described in this memoir. One focused on the Fennoscandian Shield and the other on the craton margin where it wedges out southeastwards beneath the Trans-European Suture Zone. In the former, Hjelt et al. summarize the results of experiments involving electromagnetic measurements and seismic tomography. The tomography indicated that the lithospheric mantle under Fennoscandian Shield part of the craton extends down to at least 300 km; no boundary zone to the asthenosphere was detected. With regard to crust-mantle relationships, no expression of the Archaean-Palaeoproterozoic suture zone along the southern margin of the Karelian Craton was inferred. Both the electrical and seismic investigations indicated considerable lateral heterogeneity in the upper mantle. In contrast to this evidence from the internal parts of the EEC, seismic tomography (Gregersen et al.) across the Trans-European Suture Zone margin clearly indicates that the thin lithosphere of westem Europe's Phanerozoic terranes thickens rapidly into the craton. The westward thinning EEC margin is seen to be displaced vertically by at least two major zones of offset, one of which coincides with the previously well-known Tornquist Zone (Blundell et al. 1992). The shallow (c. 120 km depth) asthenosphere of western Europe is seen to deepen gradually northwards beneath the craton to more that 300 krn. Artemieva et al. have inte~ated these results with many others to present a series of lithospheric thickness maps of Europe as a whole, including the EEC, and discuss the origins of the regional scale lithosphere heterogeneities evident in Europe that have played and continue to play such as important role in its tectonic and geodynamic evolution. We thank Olga K. Bogolepova and Nina Lebedeva-Ivanovafor help with the text and diagrams and Johathan Turner for comments on rearranging them.

References ALLEN, R. I., WEIHED, P. & THE GLOBAL VHMS RESEARCH PROJECT TEAM 2002. Global comparisons of volcanic-hosted massive sulphide districts. In: BLUNDELL, D. J., NEUBAUER, F. & YON QUADT, A. (eds) The Timing and Location of Major Ore Deposits in an Evolving Orogen. Geological Society, London, Special Publications, 204, 13-39. BECKHOLMEN,M. & GLODNY,J. 2004. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia. In: GEE, D. & PEASE,V. (eds.) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125-134. BERZIN, R., ONCKEN,O., KNAPP,J. H., PEREZ-ESTAUN,A., HISMATULIN, T., YUNUSOV, N. & LlelLIN, A. 1996. Orogenic evolution of the Urals Mountains: Results from an integrated seismic experiment. Science, 274, 220- 221. BLUNDELL, D., FREEMAN, R. & MUELLER, S. (eds) 1992. A continent revealed: the European Geotraverse. Cambridge University Press, European Science Foundation. BOGDANOV, N. A., KHAIN, W. E. BOGATSKY, V. I., KOSTYUCHENKO, S. L., SENIN,B. V., SHIPmOV,E. V. & SO13OLEV,S. F. 1996. Tectonic map of the Barents Sea region and the northern part of the European Russia. Institute of the Lithosphere, Russian Academy of Sciences, Moscow [in Russian]. BOGDANOVA,S. V. 1993. Segments of the East European Craton. In: GEE, D. G. & BECKHOLMEN, M. (eds) EUROPROBE Symposium in Jablonna 1991. Polish Academy of Sciences and European Science Foundation, A-20(255), 33-38. BOGDANOVA, S. V., GORBATSCHEV,R. & GARETSKY, R. G. 2005. The East European Craton. In: SELLEY, R. C., COCKS, L. R. M. & Pt.IMER, I. R. (eds) Encyclopedia of Geology. Elsevier, 2, 34-49. BROWN, D., JUHLIN,C., & PUCHKOV,V. (eds) 2000. Mountain building in the Uralides: Pangea to the present. AGU Geophysical Monographs, 132. DEWEY, J. F. & STRACHAN, R. A. 2005. Caledonides of Britain and Ireland. In: SELLEY, R. C., CocKs, L. R. M. & PLIMER, I. R. (eds) Encyclopedia of Geology. Elsevier, 2, 56-63. GEE, D. G. 2005. Scandinavian Caledonides (with Greenland). In: SELLEr R. C., COCKS R. L. M. & PEnMEn, J. R. (eds) Encyclopedia of Geology. Elsevier, 2, 64-74. GEE, D. G. & ARTEMIEVA, I. (eds) 2001. EUROPROBE 1992-2001. Uppsala University, Sweden. GEE, D. G. & ZEYEN, H. (eds) 1996. EUROPROBE 1996--Lithosphere Dynamics. Origin and Evolution of Continents. EUROPROBE Secretariat, Uppsala University. GRAD, M., GUTERCH, A., & MAZUR, S. 2002. Seismic refraction evidence of crustal structure in the central part of the TransEuropean Suture Zone in Poland. In: WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) Palaeozoic Amalgamation of Central Europe. Geological Society, London, Special Publications, 201, 295-309. FRANKE, W, HAAK, V, ONCKEN, O. & TANNER, D. (eds) 2000. Orogenic Processes: Quantification and Modelling in the Variscan Belt. Geological Society, London, Special Publications, 179. JANIK, T., GRAD, M., GUTERCH, A., RADLEZ, D., YLINIEMIJ., TIIRA, T., KELLER, G. R., GACZYNSKIE. & CELEBRATION 2000 WORKING GROUP 2005. Lithospheric structure of the Trans-European Suture Zone along the TTZ CEL03 seismic transect (from NW to SE Poland), Tectonophysics, 411, 129-156. KATZUNG,G., GIESE, U, WALTER, R. & YON WINTERFELD,C. 1993. The Rugen Caledonides, northeast Germany. Geological Magazine, 130, 725 -730. KNAPP,J., STEER, D., BROWN,L., BERZIN,R., SULEIMANOV,A., STILLER, M., SHEN, E. L., BROWN, D., BULGAKOV, R., KASHUBIN, S. & RYBALKA,A. 1996. Lithosphere-scale seismic image of the southern Urals from explosion-source reflection profiling. Science, 274, 226228. KOROTEEV, V. A., DE BOORDER,H., NETCHEUKHIN,V. M. & SAZONOV, V. N. 1997. Geodynamic setting of the mineral deposits of the Urals. Tectonophysics, 276, 291-300.

INTRODUCTION

MATTE, P. 1991. Accretionary history and crustal evolution of the Vailscan belt in Western Europe. Tectonophysics, 196, 309-337. SAVELIEVA, G. N. & NESBITT, R. W. 1996. A synthesis of the stratigraphic and tectonic setting of the Uralian ophiolites. Journal of the Geological Society, London, 153, 525-537. ZIELHIUS, A. & NOLET, G., 1994. Deep Seismic Expression of an Ancient Plate Boundary in Europe. Science, 265, 79-81.

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WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) 2002. Palaeozoic Amalgamation of Central Europe. Geological Society, London, Special Publications, 201. WYBRANIC, S, ZHOU, S, THYBO, H., FORSBERG, R., PERCHUC, E., LEE, M., DEMIANOV, G. D. & STRAKHOV, V. N., 1998. New map compiled of Europe's gravity field. EOS, 79(37), 437-442.

Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga IRINA M. A R T E M I E V A t'2, HANS THYBO 2 & M I K H A I L K. K A B A N 3

1US Geological Survey, Menlo Park, CA 94025, USA (e-mail: [email protected]) 2Geological Institute, University of Copenhagen, Copenhagen, Denmark DK-1350 3GFZ, Potsdam, Germany D-14473

Abstract" We present a summaryof geophysical models of the subcrustal lithosphere of Europe. This includes the results from seismic (reflection and refraction profiles, P- and S-wave tomography, mantle anisotropy), gravity, thermal, electromagnetic, elastic and petrological studies of the lithospheric mantle. We discuss major tectonic processes as reflected in the lithospheric structure of Europe, from Precambrian terrane accretion and subduction to Phanerozoic rifting, volcanism, subduction and continent-continent collision. The differences in the lithospheric structure of Precambrian and Phanerozoic Europe, as illustrated by a comparative analysis of different geophysical data, are shown to have both a compositional and a thermal origin. We propose an integrated model of physical properties of the European subcrustal lithosphere, with emphasis on the depth intervals around 150 and 250 km. At these depths, seismic velocity models, constrained by body- and surface-wavecontinent-scale tomography, are compared with mantle temperatures and mantle gravity anomalies. This comparison provides a frameworkfor discussion of the physical or chemical origin of the major lithospheric anomalies and their relation to large-scale tectonic processes, which have formed the present lithosphere of Europe.

'Evidence obtained under different experimental conditions cannot be comprehended within a single picture, but must be regarded as complementary in the sense that only the totality of the phenomena exhausts the possible information about the objects. ' Niels Bohr 'One cannot embrace the non-embraceable.' Kozma Prutkov The European continent comprises tectonic structures ranging in age from Archaean to Cenozoic. A great variety of past and present tectonic regimes within the European continent provides a unique opportunity to analyse the effects of processes related to plate tectonics (e.g. continent-continent or continent-ocean collisions, leading to formation of continental orogens and subduction zones) and mantle dynamics (manifesting itself in magmatism, continental tiffing and formation of large sedimentary basins) on lithospheric structure. The Precambrian part of the continent is formed by the East European craton (EEC) that crops out in the Baltic and Ukrainian shields and underlies the Archaean-early Proterozoic East European Platform (EEP) (Fig. 1). The EEP is crossed by a cratonscale system of mid-late Proterozoic rifts in its central part (Gorbatschev & Bogdanova 1993) and Palaeozoic rifts in its southern parts, perhaps of plume origin (Lobkovsky et al. 1996). A unique feature of the EEP is the existence of a thick (typically c. 2 - 4 km, although locally 20 km thick) sedimentary cover over most of the platform (e.g. Nalivkin 1976; Khain 1985). Rapid subsidence of the EEP in the Palaeozoic was associated with subduction during the formation of the Uralides orogen (Mitrovica et al. 1996). The fundamental lithospheric boundary in Europe, the Trans-European Suture Zone (TESZ), which was first discovered from geological, palaeontological and magnetic data by W. K. de Teisseyre and A. J. H. Tornquist (Teisseyre 1903; Tornquist 1908), separates the Precambrian lithosphere of the EEC from the Phanerozoic lithosphere of Western Europe. Recent seismic reflection/refraction and tomography studies show a dramatic change in all lithospheric properties across the TESZ (e.g. Zielhuis & Nolet 1994; Arlitt 1999; Sroda et al. 1999; Villasefior et al. 2001). The Phanerozoic part of Europe includes a mosaic of tectonic structures, such as Caledonian, Hercynian (Variscan) and Uralides Palaeozoic orogens, Mesozoic rifts, areas of Cenozoic rifting and tectonomagmatic activity (the Central European Rift System), and Cenozoic collisional

orogens often associated with subducting lithospheric slabs (e.g. the Alps, the Pyrenees, the Carpathians). The goal of this paper is to present a comparative overview of lithospheric structure of the major tectonic provinces of Europe, in an attempt to distinguish the effects of the tectonic evolution of the continent from the Archaean to the present. The results of numerous recent multi-disciplinary international projects in European Earth sciences, the largest of which are the European Geotraverse (EGT) (Blundell et al. 1992) and the EUROPROBE programme (Gee & Zeyen 1996; Gee & Artemieva 2001), form the basis of this paper. The extensive set of geophysical information available for Europe does not permit even simple listing of the key publications. With the goal of summarizing the present knowledge on the European lithosphere on a continent scale, we have deliberately omitted local details. The comprehensive analysis of various geophysical data accumulated by the EUROPROBE research during the past decade is presented in the subsequent papers in this book. With rare exceptions, the lithospheric mantle is inaccessible for direct studies. Images of the upper mantle structure provided by remote geophysical sampling are non-unique, and different techniques measure variations in different properties of the mantle (e.g. density, elastic moduli and conductivity, which are related to variations in composition, structure, mineral alignment, and fluid and thermal regime). Geophysical data obtained by different methods are, to some degree, complementary, such that integrated interpretations of different data types may provide a comprehensive picture of the physical properties of the lithospheric mantle. We combine the highlights of recent achievements in different disciplines of geosciences to provide the reader with comparative and diverse information on the upper mantle structure of the major tectonic structures of the continent. Numerous recent seismological surveys of the deep European lithosphere include a set of continent-scale seismic tomography models. Comparison of these models with thermal and gravity models for Europe permits us to constrain an integrated model of the European lithospheric mantle, which reflects diversity in both its structure and composition.

Precambrian lithosphere of Europe The oldest crust within the European continent (in the Ukrainian Shield, Stepanyuk et al. 1998) is c. 3.6 Ga old and thus is one of

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 11-41. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Baltic Shield

Fig. 1. Simplified tectonic map of Europe. TESZ, Trans-European Suture Zone.

the oldest known on the planet. The oldest crust of the Baltic Shield and the EEP is younger, 3.0-3.1 Ga and 1.8-2.1 Ga, respectively (Fig. 1). The basement of the EEP is buried under a thick cover of Proterozoic and Phanerozoic sediments, which complicates dating of the basement rocks. Petrological studies of mantle xenoliths from Precambrian cratons of the world suggest that the crust and the entire lithospheric mantle of the cratons were formed simultaneously and remained attached ever since (Carlson et al. 1994; Pearson et al. 1999). Therefore, one may expect that the lithospheric mantle of a large part of the continent, from the Urals in the east to the TESZ in the west, also has Archaean-Proterozoic ages. Knowledge of the ages of the subcrustal lithosphere is important for interpretations of seismic and gravity data, as petrological studies of mantle xenoliths indicate that cratonic lithosphere has a unique composition, depleted in basaltic components. The highest depletion is found globally in the Archaean roots and it decreases in Proterozoic and Phanerozoic lithosphere (Griffin et al. 1998). Low iron content in the Archaean lithospheric mantle has important geophysical consequences: it implies higher (by 3-5%) seismic velocities and lower (by c. 1.5%) density than in the Phanerozoic mantle (Jordan 1988; Poudjom. Djomani et al. 1999, 2001; Deschamps et al. 2002). On the other hand, Archaean cratons have the lowest average values of surface heat flow measured on the continents (Nyblade & Pollack 1993). Low temperatures in Archaean lithospheric roots (Pollack & Chapman 1977; Artemieva & Mooney 2001) essentially compensate for the effect of the depleted composition on densities (Jordan 1988) and thus mask gravity anomalies produced by compositional variations in the mantle. However, low temperatures in cratonic lithosphere enhance the effect of depletion on seismic velocities. High mantle velocities, as observed in the EEC, are often interpreted in terms of 'hot' or 'cold' regions, but their origin can be both compositional and thermal. For example, a 1% velocity increase can be caused either by 4% Fe depletion or by 100-150~ temperature decrease in the mantle (Nolet & Zielhuis 1994; Deschamps et al. 2002). We present seismic and gravity models for Precambrian Europe and compare them with thermal models to distinguish structural and compositional variations in the lithospheric mantle.

S e i s m i c data. Most of the data on the lithospheric structure of the EEC come from the Baltic Shield, for which interpretations of seismic reflection/refraction profiles, regional upper mantle seismic tomography, electromagnetic, xenolith, thermal and elastic data became available over recent decades. This extensive dataset provides important information on the lithospheric evolution of the Baltic Shield since the Archaean and reveals the presence of a thick lithospheric keel beneath it. A 180-230 km thick lithosphere has been interpreted from explosion P-wave data along the long-range refraction FENNOLORA profile in the northern part of the Baltic Shield (Guggisberg & Berthelsen 1987). The existence of a high-velocity upper mantle down to 200-250 km beneath most of the EEC, including the Baltic Shield, is supported by regional dispersion analysis of long-period Rayleigh waves and by large-scale P- and S-wave seismic tomography models (Calcagnile 1982, 1991" Bijwaard & Spakman 2000; Shapiro & Ritzwoller 2002; Boschi et al. 2004) (Fig. 2). However, most surface-wave models lose resolution at depths below c. 200250 km and cannot provide reliable constraints on mantle structure below this depth (e.g. Panza et al. 1986). Some regional high-resolution P-wave tomography models have been interpreted as indicators of the existence of high seismic velocities (+2% anomaly compared with the global continental model iasp91, Kennett & Engdahl 1991) down to 250 __+50 km under the Baltic Shield of Finland (Bock et al. 2001; Sandoval et al. 2004). The region with the thickest lithospheric keel is located at the suture between the Archaean and early Proterozoic provinces, and spatially coincides with the anomalously thick crust that has formed during Palaeoproterozoic accretion of Svecofennian terranes to the Archaean Karelian block (Korja et al. 1993). The small size of the region (c. 200km x 300km), where both the crust and the lithosphere have anomalous thicknesses, suggests that both crustal and lithospheric roots could have been formed during the same tectonic event and may represent a unique preserved remnant of an ancient subduction zone. This hypothesis is supported by xenolith data that indicate a compositionally stratified mantle in the region (Peltonen et al. 1999), and by an eastward-dipping high-velocity anomaly in the mantle beneath the Archaean-Proterozoic suture (Sandoval et al. 2004). The geographical distribution of mid-Proterozoic rapakivi granite intrusions at the western and southern sides of the anomalous region of thick lithosphere suggests a deflection of ascending magmas by the pre-existing lithospheric keel. This deflection of mantle heat and magma could have assisted the survival of this thick keel during the mid-Proterozoic tectonothermal activity in the region, which 'embraces' the anomalous region of thick lithosphere and led to the formation of the Baltic-Bothnian Sea basin. A layer with reduced seismic velocities (c. 8.1 km s-1 for the mean model) has been identified at the depth range of 100-160 km within the high-velocity (8.6 km s -~ at 100 km depth) lithospheric mantle of the Baltic Shield (Perchuc & Thybo 1996). Similar seismic velocity structure has been revealed for the Archaean part of the Karelian province in a recent surfacewave based seismic tomography survey (Bruneton et al. 2004), similar to recent results from the Canadian Shield and Greenland (Darbyshire 2005). Tomographic inversion for velocities in the upper mantle in the Baltic Shield, based on the FENNOLORA data, suggests that the 100-160 km depth interval is also characterized by very small S-wave velocities, corresponding to a much more pronounced reduction in velocity for S waves than for P waves (Abramovitz et al. 2002). The nature of the reducedvelocity zone is still debated. Alternative interpretations include (1) regional metasomatism (Bruneton et al. 2004); (2) the presence of pockets of small-percentage melting or fluids (Perchuc & Thybo 1996), probably associated with ancient subduction zones

DEEP EUROPE TODAY (although the layer may be at supersolidus temperatures; Abramovitz et al. 2002); (3) petrological heterogeneities in the lithosphere (e.g. a compositional boundary from a highly depleted upper lithosphere to a less depleted lower lithosphere can produce a seismic pattern similar to the top of a low-velocity zone; Artemieva 2003). However, neither the existing seismic models nor petrographic data on mantle xenoliths (Kukkonen & Peltonen 1999) require the presence of asthenospheric material in the upper 250300 km beneath the Archaean-early Proterozoic part of the Baltic Shield. This conclusion is supported by electromagnetic studies in the region (Korja 1990), in which no highly conductive asthenospheric layer has been identified beneath the Finnish part of the Baltic Shield. Earlier interpretations of a high-conductivity layer below 100-130km depth (e.g. Jones 1982, 1984)should be considered with caution, as they did not account for high-latitude ( > 6 0 ~ distortions of the magnetic field (Osipova et al. 1989). Seismic evidence f o r Precambrian plate tectonics. At present, Precambrian plate tectonic processes are reliably identified only from deep mantle reflectors and associated structures in active seismic reflection surveys. Teleseismic tomography cannot resolve small velocity contrasts (e.g. < 1%) in the lithospheric mantle beneath Archaean and Proterozoic terranes (e.g. Poupinet et al. 1997; Sandoval et al. 2004). With the exception of the Archaean-Proterozoic suture in the Baltic Shield (as discussed in the previous section) and the Southern Baltic Sea (Abramovitz et al. 1997), neither the anomalous crustal structure typical for modern collisional orogens, nor a linear high-velocity seismic anomaly in the mantle (which might indicate the presence of a subducting slab) is documented for Proterozoic collisional structures. The only robust dipping high-velocity 'slab' anomaly in a cratonic root has been distinguished recently in P- and S-seismic tomography studies along the Western Superior Transect (Canada) down to c. 660km depth (Sol et al. 2002). Otherwise, the oldest slab of subducted lithosphere individually recognized in the mantle from teleseismic tomographic data is Jurassic in age (van der Voo et al. 1999). Welldocumented evidence for Precambrian plate tectonic processes was first presented by the BABEL Working Group (1989) for the Baltic Shield. Older relict (2.7-2.8 Ga) subduction has been imaged in seismic reflection studies by the Canadian LITHOPROBE programme in the Superior province (e.g. Calvert et al. 1995; Clowes et al. 1996) and in the Slave craton (Bostok 1998; Cook et al. 1998, 1999; Aulbach et al. 2001). Analogy between the observed reflection geometries and modern subduction zones allows interpretations of seismic images as ancient subduction of former oceanic crust (van der Velden & Cook 1999). Dipping mantle reflectors are of a particular importance, as they are interpreted as relict subduction zones. Two large-scale high-resolution marine seismic reflection experiments in the Baltic Shield (BABEL in the Bothnian Gulf and 'Mobil Search' in the Skagerrak between Norway and Denmark) have found evidence for sets of dipping mantle reflectors, which provide new insights into Precambrian tectonic processes. Distinct, dipping sub-Moho reflections have been identified at 40-110 km depths (BABEL Working Group 1990, 1993; Lie et al. 1990). Dipping at a 15-35 ~ angle, these reflections can be traced laterally over distances of up to 100 km, and in two out of three occurrences they are accompanied by a sharp 5 - 7 km offset of Moho. By analogy between the reflectivity patterns in the Baltic Shield and Cenozoic (e.g. the Alps and the Pyrenees) and Palaeozoic (the Caledonides and the Appalachians) orogens, these mantle reflectors are interpreted as relics of Proterozoic (0.9-1.2 Ga and 1.8-1.9 Ga) tectonic processes related to Svecofennian and Sveconorwegian plate convergence, subduction and accretion of terranes onto the Archaean nucleus of the Baltic Shield (BABEL Working Group 1990, 1993b).

13

This tectonic interpretation is supported by S m - N d isotopic data from the exposed volcanic arc complex in the Baltic Shield (Ohlander et al. 1993). Recent analysis of lithospheric-scale seismic data from 1.90-1.85 Ga subduction zones at the Slave and Baltic cratonic margins (Snyder 2002) reveals strong similarity between them and modern tectonic analogues. Thermal and xenolith data. Surface heat-flow values within the Baltic Shield are close to the global average for Precambrian cratons, 30-50 m W m -2 (Nyblade & Pollack 1993), although extremely low values (20-30 mW m -2) have been reported for the southern part of the Finnish-Karelian province (Bailing 1995; Kukkonen & Joeleht 1996) (Fig. 3). Several thermal models for the upper mantle of the Baltic Shield indicate that variations in the surface heat flow largely result from heterogeneous heat production in the crust (Pinet & Jaupart 1987; Kukkonen 1998). Estimates of Moho temperatures vary from 350 ~ to 600~ (Bailing 1995; Kukkonen & Joeleht 1996; Pasquale et al. 2001; Artemieva 2003); large scatter comes not only from different model constraints but also from a highly heterogeneous crustal structure, varying in thickness from c. 30 km in the Caledonides to c. 60 km at the Archaean-Proterozoic suture in southern Finland. Thermal models suggest that in the Archaean-early Proterozoic part of the Baltic Shield the thickness of the thermal boundary layer with predominantly conductive heat transfer (thermal lithosphere) is in the range from 200 to 280 km (Pasquale et al. 2001; Artemieva 2003). These values are in agreement with regional seismic tomography models, in which no lowvelocity layer has been found down to a 250-300 km depth (Fig. 4). However, a direct quantitative comparison of lithospheric thickness constrained by diverse techniques is inadequate, as they measure different physical properties of the upper mantle (Artemieva & Mooney 2002). For example, the difference between 'seismic' lithosphere (defined as the seismic highvelocity region on the top of the mantle) and 'thermal' lithosphere (defined as the depth at which the geotherm intersects the mantle adiabat or becomes supersolidus) can be up to several tens of kilometres (Jaupart & Mareschal 1999); this difference approximately corresponds to the thickness of the transition zone between purely conductive and purely convective heat transfer. In tomography studies, where seismic lithosphere is considered as the layer above the convecting mantle, its base is defined either as a zone of high velocity gradient or the bottom of a layer with positive velocity anomalies. However, seismic tomography and seismic refraction models would not necessarily indicate the same depth to the base of the lithosphere. In seismic reflection surveys, strong mantle reflectors are often interpreted as the base of the seismic lithosphere, as it is assumed that they originate at the transition from the lithosphere to a zone of partial melt (Lie et al. 1990). Furthermore, the base of the seismic lithosphere should be a diffuse boundary if the decrease of the seismic velocities associated with the lithospheric base is caused by high-temperature relaxation or by partial melting (Anderson 1989). Xenolith geotherms for mantle-derived peridotites from kimberlite pipes of the Finnish part of the Baltic Shield and the Arkhangelsk region confirm low mantle temperatures (Kukkonen & Peltonen 1999; Kukkonen et al. 2003; Malkovets et al. 2003) (see Fig. 6). Peridotites from Finnish xenoliths suggest that lithospheric mantle extends down to at least 240 km depth (the depth from which the deepest xenoliths originated) (Kukkonen & Peltonen 1999) as the peridotites show no variations in texture or composition that could be interpreted as indicators of the transition zone from conductive to convective heat transfer. For example, high-temperature sheared peridotites are absent even in the deepest sampled part of the lithospheric column.

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I.M. ARTEMIEVA ETAL.

Fig. 2. Cross-section of the European lithosphere at depths of 150 km and 250 km. Most of the Precambrian part of the continent has high seismic velocities and low attenuation, at least partly caused by low mantle temperatures. In contrast, Phanerozoic Europe is characterized by low seismic velocities, high attenuation and high temperatures. (a) P-wave velocity perturbations with respect to the ak135 model (based on the tomography model of Bijwaard & Spakman (2000), smoothed by Gaussian filtering). The lateral resolution of the model is very uneven. High resolution (c. 100 km) is achieved for regions with a good coverage of events and stations (Southern and Western Europe). For the EEP the lateral resolution is very low (500-1000 km) and this region is shown white. The vertical resolution of P-wave tomography models is poor, as body waves sample the entire mantle with almost vertical propagation. Most of the anomalies seen in the map propagate to deeper levels (see. (c)). (b) Rayleigh-wave phase velocities (based on the global model of Shapiro & Ritzwoller 2002). The vertical resolution is 50-100 km for the upper 250 km and coverage disappears at deeper levels; the lateral resolution does not exceed 500-1000 km. (c) As (a) for 265 km depth (based on the model of Bijwaard & Spakman 2000). The low lateral resolution for the eastern Baltic Shield and EEP, should be noted. (d) As (b) for 250 km depth (based on the global model of Shapiro & Ritzwoller 2002). The surface wave inversion loses resolution below depths of c. 250 km.

DEEP EUROPE TODAY

15

Fig. 2. Continued. (e) P-wave velocity perturbations with respect to the sp6 reference model (based on the tomography model of Piromallo & Morelli (2003), defined over the equi-spaced nodes with 0.5 ~ spacing). The model has been smoothed by Gaussian filtering. Vertical resolution is low compared with surface-wave tomography. The model resolves similar features in the upper mantle as the model of Bijwaard & Spakman (2000). (f) Mantle temperatures (in ~ at 150 km depth (Artemieva 2003, complemented by new data for Western Europe). Temperatures for the EEC are constrained by surface heat flow for steady-state conductive heat transfer; geotherms for Western Europe are constrained by lithospheric thickness data derived from different seismic models and assuming that 1300 ~ is reached at the lithospheric base. The uncertainty in temperatures is c. 10-15%, but for western Europe can be locally larger. Lateral resolution is c. 50-500 km. (g) Rayleigh-wave tomography for velocity model at 150 km depth (based on the model of Billien et al. 2000). The model is constrained effectively to 12th-degree spherical harmonics with a vertical resolution of c. 5 0 - 8 0 km at 150 km depth. (h) Rayleigh-wave tomography for inverse attenuation at 150 km depth (based on the model of Billien et al. 2000). The model is constrained effectively to 12th-degree spherical harmonics with a vertical resolution of c. 5 0 - 8 0 km at 150 km depth.

16

I.M. ARTEMIEVAET AL.

Fig. 3. Surface heat flow in Europe (after Pollack et al. 1993, updated for new heat-flow data); a low-pass filter has been applied to remove short-wavelength anomaliescaused by shallow effects (e.g. heterogeneitiesin crustal heat production and conductivity). Stars show locations of mantle xenoliths discussed in the text.

East European Platform Seismic data. The lithospheric mantle of the EEP is not studied as extensively as the upper mantle of the Baltic Shield. Continent-scale seismic tomography models (Fig. 2), especially for body waves, have insufficient resolution for the northeastern parts of the EEP as there are few seismic events and the distribution of stations is sparse. Regional electromagnetic models are limited to models of crustal conductivity. With rare exceptions, seismic reflection or refraction profiles do not image the lithosphere deeper than 5 0 - 6 0 k m (Vinnik & Ryaboy 1981; Garetskii et al. 1990; Grad & Tripolsky 1995; Kostyuchenko et al. 1999; EUROBRIDGE Working Group & EUROBRIDGE'95 2001; Grad et al. 2002; Thybo et al. 2003). Weak mantle reflectivity along the profiles, which image the lithosphere of the EEP to a significant depth, suggests either that the entire cratonic root was formed in a fast thermal event in the Precambrian, or that pre-existing reflectivity has been erased by later tectonic processes. However, the lack of significant tectonic activity in most of the EEC since the Precambrian rules out the latter hypothesis. Recent P- and S-wave tomography of the upper mantle of the entire EEP has demonstrated that it is characterized by constant shear velocities (4.65 km s-1) in the depth range 1 0 0 - 2 5 0 k m and radial anisotropy (c. 5%) down to a depth of 200-250 km, where the anisotropy decreases sharply to c. 2% (Matzel & Grand 2004). The depth of 250 km is interpreted as a transition from dislocation deformation to diffusion creep and thus may be considered as a rheological base of the EEP lithosphere. Seismic refraction data indicate that the lithosphere of the northem EEP (along the Peaceful Nuclear Explosion (PNE) profile Quartz) is c. 200 km thick (Mechie et al. 1993; Ryberg et al. 1996); the base of the lithosphere is likely to have a transitional character as no sharp velocity contrast was found at the inferred lithospheric base. Waveform inversion for the upper mantle

structure in the western part of the EEP along the 30~ meridian revealed similar values of lithospheric thickness, c. 200 km (Paulssen et al. 1999). These estimates of the seismic base of the lithosphere are, on the whole, in agreement with thermal estimates of the lithospheric thickness of the EEP, c. 170-200 km with small regional variations within the accuracy of the model (Artemieva 2003; Fig. 4c). Similar to the Baltic Shield, a pronounced reduced-velocity channel at a depth of 105-130 km has been identified within the lithospheric mantle of the northeastern EEP along the PNE profile Quartz (Ryberg et al. 1996). According to travel-time inversion of seismic data along the PNE profiles Quartz and Kraton, this feature extends eastwards as a continuous layer for at least 3000 km into the West Siberian Basin and the Siberian Shield (Nielsen et al. 1999). Similar reduced-velocity layers have been reported earlier for other cratonic regions of the world (Grand & Helmberger 1984; LeFevre & Helmberger 1989; Pavlenkova et al. 1996; Darbyshire 2005) and suggest that it may be a global characteristic of Precambrian lithosphere (Thybo & Perchuc 1997; Thybo 2006). The proposed models for such a layer, with a relatively low seismic velocity within high-velocity cratonic root, include the presence of fluids, partial melts (or temperature close to the solidus), metasomatism, or compositional variations. For example, in North America, a low-velocity zone was found in an S-wave model but was not observed in a P-wave model, which suggests that it is an indicator of a partially molten zone (Rodgers & Bhattacharyya 2001). Thermal data. The EEP is characterized by relatively homogeneous values of the surface heat flow (35-45 m W m -z, Fig. 3), that are within the range of the global average for the Archaean-early Proterozoic cratons of the world (Nyblade & Pollack 1993). Slightly higher values (40-55 m W m -2) have been measured in the southern parts of the platform although, locally, thermal anomalies can reach values as high as 7 0 - 9 0 m W m - 2 (i.e. in the Pripyat Trough). The transition to the Phanerozoic lithosphere of Western Europe is marked by a sharp step-like increase in surface heat flow by c. 20 mW m - 2 (Fig. 3). The thickness of the thermal lithosphere within the EEP has been estimated to be 1 7 0 - 2 0 0 k m (Cermak 1982; Artemieva 2003; Majorowicz et al. 2003) (Fig. 4c). Surprisingly, the Ukrainian Shield, which is the oldest part of the European continent, has similar lithospheric thickness, 1 8 0 - 2 2 0 k m (Kutas 1979). Such values have also been reported for the Archaean lithosphere of South Africa and Australia (Jaupart & Mareschal 1999; Artemieva & Mooney 2001). These cratons are among the oldest on the Earth: the major crust-forming events in the Kaapvaal, Zimbabwe, Indian and Pilbara cratons and the Greenland Shield occurred at c. 3.0-3.5 Ga, whereas in the East European, Siberian and North American cratons the major crust-forming events occurred significantly later, c. 1.8-2.5 Ga (Goodwin 1996). The large difference in lithospheric thickness of Precambrian regions, which were assembled into cratons at different times (Artemieva 2006), poses the question of whether different tectonic and/or mantle processes have operated in the early and late Archaean and led to the formation of cratons with significantly different lithospheric structures (Artemieva & Mooney 2002; Artemieva et al. 2002). As R e - O s isotope studies indicate similar geological ages (i.e. approximately the ages of crustal differentiation; Richardson et al. 1993) for all of the Archaean cratons, it is likely that anomalously thick lithospheric roots could have formed by different intensities of tectonic modification of pre-existing terranes during the cratonization stage and not as a result of different differentiation processes within the deep mantle. Precambrian rifts within the EEP. Mantle processes have played an

important role in the evolution of the continental lithosphere since

DEEP EUROPE TODAY

17

Fig. 4. Five models of lithospheric thickness in Europe. (For (a)-(c) see caption to Fig. 2 for more details.) (a) Lithospheric base defined by a 1% P-wave velocity perturbation (based on the model of Bijwaard & Spakman 2000, interpolated with a low-pass filter) with respect to the ak135 model. (b) Lithospheric base defined by a 2% S-wave velocity perturbation (based on the model of Shapiro & Ritzwoller 2002, interpolated with a low-pass filter) with respect to the global continental model iaspei91 (Kennett & Engdahl 1991). (c) Thermal lithosphere defined by the intersection of the geotherm with a 1300 ~ mantle adiabat (the model of Artemieva 2003). (d) Lithospheric thickness in Europe based on electromagnetic surveys (compilation of Hjelt & Korja 1993, interpolated with a low-pass filter). Dark blue corresponds to regions where depth to the highly conductive layer exceeds 200 km, or where electrical asthenosphere was not detected. (e) Lithospheric thickness calculated from P-residuals (Babu~ka et al. 1988) under the following assumptions: (1) variations in lithospheric thickness are proportional to P-residuals; (2) lateral variations in average lithospheric velocities (as a result of temperature or compositional variations) are ignored; (3) a homogeneous crustal thickness of 33 km is assumed for the entire Western European region; (4) the results are scaled by data from surface-wave dispersion analysis (Panza et al. 1986) on lithospheric thickness in the Western/kips (220 kin) and the Belgo-Dutch platform (50 km).

18

I.M. ARTEMIEVA ET AL.

its formation. Giant mafic dyke swarms (the oldest known, in SW Greenland, is c. 3.25 Ga old), continental rifting (the oldest known, in the Kaapvaal and Slave cratons, is c. 3.0-3.3 Ga old), and break-up of supercontinents (the oldest known is c. 2.5-2.7 Ga old) are believed to be surface manifestations of ancient plume-lithosphere interactions (Nelson 1992). The ages of the known large-scale mantle-lithosphere interaction events within the EEC are much younger than in other cratons (Khain 1985). In the Baltic Shield, the Riphean (1.35-1.05 Ga) tiring affected the Baltic Sea region with the emplacement of rapakivi granites and a subsequent subsidence of the basin (Ga~il & Gorbatschev 1987). Within the EEP, the fundamental trans-cratonic Central Russia Rift System (CRRS) formed at c. 1.3-1.0 Ga either by a large-scale tiring event or by amalgamation of three large terranes into the EEC (Gorbatschev & Bogdanova 1993) (Fig. 1). This process was followed by intensive intraplate volcanism at c. 1.0 G a - 6 5 0 Ma (Nikishin et al. 1996). However, there is otherwise little evidence for Precambrian rifting in the present-day structure of the deep lithosphere of the EEC, although this may be due to the sparse high-resolution geophysical data coverage of the upper mantle in this region (Figs 2 - 4 ) ; much of the knowledge comes from geological data. Nevertheless, joint interpretations of different geophysical datasets indicate significant compositional variations in the lithospheric mantle of the EEP, which may be related to Precambrian (as well as Phanerozoic) tectonomagmatic activity (see discussion below). Gravity data. Density inhomogeneities in the upper mantle, related

to variations in temperature and mineral composition, can provide significant driving forces of both vertical and horizontal motions of lithospheric blocks. As the gravity field contains effects of all density heterogeneities of the Earth, it is necessary to subtract all signals that do not originate from the mantle to extract the mantle component of the gravity field. These signals include the gravity effect of the crust, which is the largest, but can be approximated from independent a p r i o r i data. The resulting residual gravity anomalies reflect density anomalies in the mantle within the accuracy of the crustal model. Although attempts to calculate mantle gravity anomalies were made since the first seismic sections became available, a reliable 3D gravity model of the lithosphere of most of Europe (Artemjev et al. 1993, 1994) could not be constructed until sufficient data on the crustal structure had been accumulated. The new model of mantle residual Bouguer gravity anomalies, based on updated data on the crustal structure of Europe (Fig. 5), shows a sharp change in the sign of anomalies across the TESZ, from positive values over the EEC to negative values over Western Europe. A strong positive anomaly over the Caucasus implies the presence of a subducting slab, which, so far, has not been resolved in tomographic models (Fig. 2). Near-zero values of mantle gravity anomalies over the Baltic Shield are in agreement with the isopycnic hypothesis (Jordan 1988) and suggest that low lithospheric densities caused by Fe depletion of the cratonic keel are well compensated by low mantle temperatures. The positive anomalies of the EEP suggest that compositional density anomalies in the lithospheric mantle of the EEP are not compensated by temperatures as a result of either a more fertile composition or very low mantle temperatures (Fig. 6). However, a strong positive anomaly in the southern part of the EEP, which has been affected by Palaeozoic tiring, rules out a temperature origin of the gravity anomaly. Spatial correlation of the strongest positive residual gravity anomaly with the position of the Central Russia Rift System (Fig. 5) also suggests a compositional rather than a thermal origin of the anomaly. Furthermore, this conclusion is supported by high average crustal velocities in the CRRS (Fig. 7), which may be caused by magmatic underplating; it implies that infiltration of basaltic magmas into the lithosphere played an important role in the tectonic evolution of the CRRS.

Fig. 5. Mantle residual gravity anomalies, which are a part of a 3D global model (Kaban et al. 1999, 2003; Kaban & Schwintzer 2001), supplemented by higher-resolution regional data (Kaban 2001). The anomalies reflect density variations produced by compositional or temperature variations, presumably in the upper 40-60 km of the subcrustal lithosphere. The model is calculated by subtracting: (1) the anomalous gravity field of the sedimentary cover and water; (2) the anomalies related to the Moho depth variations; (3) density variations within the crystalline crust from the observed gravity field (Bouguer anomalies on land and free-airanomalies offshore). The results depend critically on seismic data on the crustal structure, because during calculations seismic velocities are converted to densities. The predictions of the present model are higher by c. 50 mGal than residual gravity anomalies for the European continent based on older data on the crustal structure (Yegorova & Starostenko 2002), although the general pattern of the anomalies remains similar. Density excess in the mantle is typical for Precambrian terranes and regions of Phanerozoic subduction. Density deficit in the Phanerozoic mantle may be caused by high temperatures and partial melt.

Contrast in lithospheric properties across the Trans-European Suture Zone (TESZ) The TESZ is a fundamental tectonic boundary within the European continent. It is formed by a broad complex zone of Palaeozoic terranes accreted to the southwestern margin of the East European Craton and marks the transition from the Precambrian cratonic lithosphere to the Neoproterozoic-Palaeozoic lithosphere of Western and Central Europe. The deep structure of the TESZ is characterized by a sharp change in lithospheric properties, well established by different geophysical methods (Thybo et al. 1999, 2002). The transition from the cratonic to the Phanerozoic lithosphere is characterized by the following features. (1) Crustal thickness changes sharply from 3 5 - 4 5 km in the EEP, to 4 0 - 5 5 km in the Teisseyre-Tornquist Zone, and to 2 8 - 3 2 km with a surprisingly flat Moho beneath the mosaics of Variscan and Caledonian terranes of Westem and Central Europe (Guterch et al. 1986; Abramovitz et al. 1998; Grad et al. 2002) (Fig. 7). Furthermore, the magnetization of the crust of Central Europe is extremely weak compared with the upper and middle crust of the EEC (Banka et al. 2002). Thin crust with a flat Moho and a lack of seismic signature in the lithospheric mantle of the European Caledonides and Variscides suggests that a large portion of the lower crust and the lithospheric

DEEP EUROPE TODAY

Fig. 6. Typical geotherms in different tectonic structures of Europe. For stable parts of the EEC the geotherms are constrainedby surface heat-flow data assuming steady-state conductive regime (Artemieva2003). Models of heat production distributionin the crust were constrained taking into account: (1) wavelength of surface heat-flow variations; (2) regional seismic models for the crustal velocity structure; (3) regional and global petrological models on the bounds on bulk crustal heat production (see details in Artemieva & Mooney 2001). For tectonicallyactive regions of Western Europe, mantle temperatures are based on a nonsteady-state conductive model constrained by data on Cenozoic magmatism(Artemieva 1993) and on the conversion of regional seismic tomography models into temperatures (Sobolev et al. 1996). For comparison, P-T data on mantle xenoliths are shown (Coisy & Nicolas 1978; Seck & Wedepohl 1983; Nicolas et al. 1987; Werling & Altherr 1997; Kukkonen & Peltonen 1999; Malkovets et al. 2003). Ar-ePt, Archaean-Early Proterozoic.

mantle could have been delaminated as a result of the Palaeozoic orogenies (Ziegler et al. 2004). (2) A pronounced and sharp decrease in seismic velocities (by 2 - 3 % ) down to the depth of 100-200 km is observed at the transition from fast cratonic lithosphere to Palaeozoic upper mantle (Zielhuis & Nolet 1994; Poupinet et al. 1997; Masson et al. 1999; Villasefior et al. 2001; Cotte et al. 2002) (Fig. 2). This velocity contrast is caused by differences in lithospheric composition and mantle temperatures. Part of the velocity anomaly may possibly be attributed to palaeosubduction along the cratonic margin, which increased the fluid content in the upper mantle (Nolet & Zielhuis 1994). (3) The transition zone between the lithospheric terranes of Precambrian and Palaeozoic ages dips at a steep angle to the vertical (c. 13-20 ~ in the Irish Caledonides and the Uralides, based on teleseismic studies (Masson et al. 1999; Poupinet et al. 1997). In comparison, the dip of the transition boundary across the Caledonian Deformation Front in the southern part of the Baltic Shield is shallow (c. 15-20 ~ to the horizontal with a SW dip based on a seismic normal-incidence reflection profile) (MONA LISA Working Group 1997). A subhorizontal boundary between the cratonic and Phanerozoic lithospheres implies that high-velocity lower crust, or a part of the subcrustal lithosphere of

19

Fennoscandia, may extend far to the south (i.e. to the E l b e Oder line), underlying Phanerozoic structures of Northern Europe (Thybo 1990; Cotte et al. 2002). This conclusion is supported by the results of a joint interpretation of seismic, gravity and magnetic data (Thybo 2001; Bayer et aL 2002) and by a likely compositional origin of the velocity anomalies observed in the TOR tomography experiment (see discussion below). A similar pattern of a non-vertical transition from Archaean to Proterozoic lithosphere has been documented by LITHOPROBE data at the margins of the Canadian Shield (Bostok 1999; Ludden & Hynes 2000). (4) A strong subhorizontal upper mantle reflectivity has been documented beneath the Variscides and Caledonides at the depth range of 5 0 - 1 0 0 km (Masson et al. 1999; Abramovitz & Thybo 2000; Grad et al. 2002), as compared with a weak mantle reflectivity in the cratonic lithosphere of the EEC, where only one significant mantle reflector was found at c. 10 km below Moho (BABEL Working Group 1993; Grad et al. 2002). (5) Surface heat flow changes abruptly by 2 0 - 3 0 m W m -2 from cratonic to younger Europe (Fig. 3), and is accompanied by a significant rise in lithospheric temperatures (Cermak 1993; Artemieva 2003, 2006). (6) Lithospheric thickness sharply changes from 150-200 km in the EEC to 8 0 - 1 2 0 km in Phanerozoic Europe (Figs 2, 4, 7 and 8, and Table 2) (e.g. Panza et al. 1986; Babugka et al. 1988; Zielhuis & Nolet 1994; D u e t al. 1998; Artemieva & Mooney 2001). (7) An abrupt change in the upper mantle density structure is reflected in a transition from near-zero or weakly positive isostatic gravity anomalies in the cratonic part to strongly negative anomalies in Western Europe (Fig. 5). Strong negative residual mantle anomalies suggest the presence of low-density masses within the upper mantle and provide indirect evidence for high mantle temperatures. Near-zero isostatic gravity anomalies in the cratonic part of the continent imply that the expected density increase caused by depleted composition of the cratonic lithosphere is entirely compensated by the density increase caused by low mantle temperatures, in agreement with the isopycnic hypothesis (Jordan 1988).

Palaeozoic structures of Europe Palaeozoic orogens of Europe include the Uralides at the eastern margin of the EEP and the Caledonian and Variscan (Hercynian) structures in the western part of the continent (Fig. 1). The crustal structure of European Palaeozoic orogens has been studied in detail by numerous seismic profiles (including normal incidence and wide-angle reflection seismic profiles) in the North Sea (BIRPS, MONA LISA), Germany (DEKORP BASIN 96), France (ECORS), Poland (POLONAISE), Ireland (VARNET- 96), Spain (IBERSEIS, ILIHA, NARS), and in the Urals (ESRU, URSEIS). However, data on the properties of the mantle lithosphere of European Palaeozoic orogens still remain limited (Blundell et al. 1992) and, in the case of the Caledonides, are restricted mainly to the transitional regions from the cratonic to post-cratonic lithosphere (i.e. across the Caledonian Deformation Front) (Masson et al. 1999; Roberts 2003). The Caledonides (named after Caledonia, the Latin name for Britain) and Variscides were formed during orogenic events involving a triple plate collision (Baltica, Laurentia and Avalonia) associated with the closure of the Iapetus Ocean and Tornquist Sea, and subsequent amalgamation of a series of terranes (Dewey 1969; McKerrow & Cocks 1976). Radiometric data on abundant granitoids and metamorphic rocks provide the ages of these Palaeozoic tectonic events, which included deformation, magmatism and metamorphism, as 500-400 Ma in the Caledonides and 4 3 0 - 3 0 0 Ma (possibly as late as 280 Ma) in the Variscan belt (e.g. Stille 1951; Emmermann 1977;

20

I.M. ARTEMIEVAET AL.

Fig. 7. Ranges of (a) average Vp seismic velocities in the crust, (b) crustal thickness, and (c) lithospheric thickness in different tectonic structures of Europe (based on Table 1). CRRS, Central Russia Rift System; CERS, Central European Rift System; PDDR, Pripyat-Dnieper-Donets rift; EEP, East European Platform. Ar-ePt, Archaean-Early proterozoic; Pt, Proterozoic; Pz, Phanerozoic; Mz-Cz, Mesozoic- Cenozoic.

Matte 1986). Opening of the North Atlantic Ocean disrupted the Caledonian orogenic belt into the European (Svalbard, Norwegian, Irish-British and Danish-Polish Caledonides) and the North American (the Appalachians and East Greenland) parts (Dewey 1969). The Uralides orogen, a well-preserved arc-continent collision zone composed of a series of late Proterozoic-Palaeozoic fold belts formed at c. 4 0 0 - 2 5 0 Ma, following the closure of the Uralian palaeo-ocean at c. 4 7 0 - 4 0 0 Ma and the accretion of the Kazakh terrane at the eastern passive margin of the EEC at c. 4 0 0 - 3 2 0 Ma (Edwards & Wasserburg 1985; Savelieva 1987; Sengfr et al. 1993; Bea et al. 1997; Puchkov 1997; Brown et al. 1998). This orogen is partly exposed in the Urals mountains, Severnaya Zemlya and the Taymyr Peninsula, whereas its

eastern part is buried under the West Siberian Basin. Further collisions of the EEC with the Siberian craton resulted in the formation of the Timan Ridge in Triassic-early Jurassic time. Compared with other Palaeozoic orogens, which have been essentially reworked during the late Palaeozoic and Meso-Cenozoic tectonomagmatic processes, the Uralides have remained intact since the Palaeozoic.

E u r o p e a n CaIedonides

A thin crust (Fig. 7), in places with a seismically laminated lower crust and a sharp subhorizontal Moho, that crosses pre-existing terrane boundaries, is typical of the Caledonides, Variscides and

DEEP EUROPE T O D A Y

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DEEP EUROPE TODAY

23

Fig. 8. Three profiles through the lithosphere of Europe: (a) north-south profile from the Baltic Shield to Corsica through the Alps; this profile follows the EGT profile (Blundell et al. 1992); (b) north-south profile from the Baltic Shield to Crete through the Pannonian Basin; (c) S W - N E profile from Ibefia to the Urals through the Central European Rift System, the Carpathians, and the East European Platform. The pronounced differences in lithosphefic thickness along the profiles should be noted; these are only partly coupled to variations in crustal thickness. The difference in wavelengths in crustal and lithospheric thickness variations may be caused by depth-dependent differences in resolution. Deep, normal-incidence reflection seismic data show traces of palaeo-subduction for all tectonic ages, independent of the lithosphefic thickness. A reduced velocity zone, identified beneath some cratonic terranes (see the section on the Baltic Shield), has absolute seismic velocities slightly lower than in the surrounding high-velocity layers in the cratonic mantle, but still c. 1% higher than in global continental reference models (ak135 or iaspei). The range of possible lithospheric thickness values is based on different methods (Table 2); the uncertainty is c. 50 km. M, Moho; STZ, Sorgenfrei-Tornquist Zone; TIB, Trans-Igneous Belt; LAB, lithosphere-asthenosphere boundary.

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DEEP EUROPE TODAY the northern Appalachians (Behr & Heinrichs 1987; Nelson 1992; Meissner 1996). This has long been believed to be typical for all Palaeozoic orogens. This crustal structure is often interpreted as an indication that a large part of the lower crust, and probably of the lithospheric mantle, has been delaminated during the Palaeozoic orogenies. However, seismic data from eastern East Avalonia shows no sign of lower crustal reflectivity (MONA LISA Working Group 1997). Nelson (1992) postulated another scenario of crustal modification during Palaeozoic orogenic events that includes: (1) post-compressional delamination of eclogitized lower crust and the uppermost mantle lithosphere resulting in crustal thinning (however, Abramovitz et al. (1998) interpreted low sub-Moho velocities at the northern edge of the former Caledonian orogeny in Denmark as being associated with the presence of lower crustal rocks of eclogite facies); (2) decompressional melting in upwelling asthenosphere tending to replace the foundering lithosphere; (3) ponding of mafic sills within the lower crust and at the crustal base, producing a sharp Moho and a laminated lower crust. As these processes took place after the main compressional events, the present crustal structure does not necessarily show any simple relationship to pre-existing terrane boundaries. Estimates of lithospheric thickness in the Norwegian and Danish-Polish Caledonides, based on surface-wave dispersion analysis, S-wave seismic tomography (Calcagnile 1982, 1991; Panza et al. 1980; Pedersen & van der Beek 1994) and thermal modelling (Cermfik 1994; Balling 1995; Zeyen et al. 2002), give values in the range of 90-130 km (see also Fig. 4). The MONA LISA Working Group (1997) detected subhorizontal seismic reflections at a depth of c. 80 km in the North Sea area, which can be interpreted as being close to the lithospheric base. In one case such reflectors are observed on two crossing profiles, thus ruling out side-swipes and other artefacts. Nevertheless, S-wave models may not have sufficient lateral resolution, such that an apparent lithospheric thinning in the Caledonides of Norway may result from smearing of a strong offshore low-velocity anomaly (e.g. Fig. 2d). Little is known about the structure of the subcrustal lithosphere of the British and Irish Caledonides; most upper mantle studies are restricted to the Iapetus Suture separating the Laurentian and Avalonian continents. Across the Caledonian Deformation Front, P-wave seismic velocities in the upper mantle increase by c. 0.26 km s-1 (Masson et al. 1999), and surface heat flow increases from 4 5 - 6 0 mW m - 2 in the cratonic lithosphere of Laurentia to 70-80 mW m - 2 in the Caledonides (Fig. 3). The latter values are significantly higher than in the Norwegian Caledonides (4555 mW m-2); it is, however, unclear if high heat flow values in the British and Irish Caledonides are caused by reduced lithosphere thickness or by shallow effects (e.g. high crustal heat production, groundwater circulation).

The Variscides Tectonics. The Variscan (Hercynian) orogeny has affected most of Central and Western Europe and forms a 700-1000 km wide and c. 3000 km long belt, extending from Poland and SE England to western Iberia (Franke 1986; Ziegler 1986; Fig. 1). The major tectonic features of the European Variscides are three N E SW-striking subparallel sutures (e.g. Neugebauer 1989), often interpreted as related to oceanic closure. However, plate tectonic interpretations of the origin of the Variscan orogen remain controversial, mainly because of the lack of evidence for the position of an ocean inside the Variscides (e.g. Ziegler 1986; Neugebauer 1989; Ziegler et al. 2004). Some workers (e.g. Behr et al. 1984) have . . . . . . . a o,. . . . . . . . . , ~,,,,,h,,,,,.a a;,,,.,;,,, o,,~.A,,ot;..n in the entire Variscan Europe. Others (e.g. Lorenz & Nicholas 1984; Matte 1986) favoured two-sided, north- and south-dipping, subduction caused by the closure of two Palaeozoic oceans,

25

followed by obduction and collision of Europe and Africa. The total crustal shortening during the Variscan orogeny exceeds 600 km; the terranes of Proterozoic to Carboniferous ages (e.g. Armorican, Ardennes, Iberian, Bohemian, French Massif Central) were deformed and partly metamorphosed, and large volumes of granitoids were emplaced between 370 and 280 Ma (Matte 1986). A large part of the Variscides has been later reworked by Mesozoic-Cenozoic events, related to tectonomagmatic activity in the Central European Rift System and large relative movements of the Eurasian and African plates.

Seismic models. Seismic studies of Hercynian Europe indicate that,

despite the strongly heterogeneous tectonic structures of the Variscan belt, the seismic velocity structure of the subcrustal lithosphere is rather uniform, with dominating subhorizontal wideangle reflectors in the upper 90 km (Him et al. 1973; Faber & Bamford 1979; ILIHA DSS Group 1993). These data imply that the Hercynian structures in the European lithosphere have not been preserved since the Palaeozoic formation of the orogen. However, one should bear in mind that the resolution in these studies is relatively low because of the > 3 km intervals between the seismic stations along the refraction profiles. Hence, it cannot be excluded that dipping orogenic structures, which could be ascribed to the Variscan orogeny, could exist in higherresolution, normal-incidence reflection seismic sections. A layered structure of the Variscan lithospheric mantle with a horizontal foliation of the upper layer and a vertical (or steeply dipping) layering in the lithospheric mantle below c. 45 km depth is supported by recent studies of spinel lherzolite xenoliths from the Bohemian Massif, which sample the Variscan lithosphere down to a depth of c. 70 km (Christensen et al. 2001). Data on Pn anisotropy and SKS shear-wave splitting provide further support for this conclusion (Fuchs & Wedepohl 1983). Christensen et al. (2001) argued that a horizontal olivine a-axis in the lower layer, with an approximately east-west strike, parallel to the observed fast shear-wave direction, has been inherited from the Variscan convergence. Strong seismic anisotropy (6.5-15% for P-wave velocities; Babugka & Plomerov~i 1992) in the lithospheric mantle of the Variscides provides evidence for palaeosubduction zones associated with the closure of the oceanic domains and the consequent Hercynian orogeny. By the pattern of seismic anisotropy, the Variscides can be subdivided into two domains with NW- or SE-dipping anisotropic structures in the lithospheric mantle (Babugka & Plomerov~i 1992). The general S W - N E orientation of the suture between the lithospheric domains with different anisotropy patterns differs from the north-south trend suggested by Panza et al. (1986). The depth range of seismic anisotropy in the lithospheric mantle is largely unknown. However, the boundary between the two domains approximately corresponds to the suture between the Saxothuringian and Moldanubian terranes and correlates with two features: (1) a pronounced step in lithospheric thickness, which increases southeastwards from 80-100 km to 120-140 km over a distance of c. 150 km (Fig. 4d; Babugka & Plomerovfi 1992); (2) a dip of a highly conductive layer in the mantle (Praus et al. 1990). Based on P-wave residuals (Fig. 4d), the typical thickness of the Variscan lithosphere is estimated to be 80-120km, with small values ( 6 0 - 8 0 k m ) in the Cenozoic Central European Rift system (see below), and large values (120-140 kin) beneath the Proterozoic-early Palaeozoic terranes (e.g. the NE part of the Massif Central and the Bohemian Massif). Because the variation in the P-wave residuals in Central Europe does not correlate with the present stress field (Muller et al. 1992), NW- and SE-dipping anisotropic structures in the lithospheric mantle of the Variscides are interpreted as traces of two divergent systems of palaeosubduction zones with olivine orientations inherited from subducted ancient lithosphere (Babugka & Plomerovfi 1992).

26

I.M. ARTEMIEVAET AL.

Similarly, two distinct patterns of upper mantle S-wave seismic anisotropy have been distinguished in the Armorican massif; upper mantle of the southern domain exhibits orogen-related anisotropy with N W - S E orientation of Pn and SKS fast directions, parallel to the strike of the South Armorican shear zone, whereas in the northern domain SKS fast directions do not follow the strike of major Hercynian shear zones. Furthermore, at 9 0 - 1 5 0 km depth the upper mantle has + 3 % P-wave velocity anomaly in the southern domain and - 3 % P-wave velocity anomaly in the northern domain (Judenherc et al. 2002). This seismic pattern is interpreted as evidence for a pre-Hercynian subduction process, which welded together two parts of the Armorican massif. Surface-wave tomography of Central Europe indicates low mantle velocities at depths below 150 km (Fig. 2b; Shapiro & Ritzwoller 2002); earlier estimations of lithospheric thickness, based on surface-wave dispersion analysis, are in the range of 7 0 - 1 0 0 k m (Panza et al. 1986; Du et al. 1998). In the Iberian peninsula, mantle velocities in surface-wave tomography models reach asthenospheric values between 80 and 180km depth (Badal et al. 1996). P-wave tomography, which has a much weaker vertical resolution (compare Fig. 2a and 2c), indicates that lithospheric thickness in Central Europe is less than 100 km (Bijwaard & Spakman 2000; Piromallo & Morelli 2003), except for the Armorican Massif, where lithospheric thickness may be as large as c. 150-200 km (Fig. 4a and d). A linear belt of large lithospheric thickness beneath SE Iberia, resolved by P-wave velocity models, is probably associated with a Cenozoic subduction zone (Blanco & Spakman 1993). Similar linear velocity anomalies are seen beneath other Cenozoic subduction systems (the Alps, the Hellenic arc; see Fig. 4a and d, and discussion below); but surprisingly, there is no seismic sign of a subducting slab beneath the Caucasus, despite the presence of a strong positive gravity anomaly (Fig. 5). Thermal models. Surface heat flow in the Variscides is high, c. 7 0 -

100 mW m - 2 (Fig. 3), and locally it significantly exceeds these values (Cermak 1995). Strong negative isostatic gravity anomalies ( - 4 0 to - 6 0 reGal; Fig. 5) indirectly imply high temperatures in the mantle of Hercynian Europe. However, the highly heterogeneous crustal structure as well as the transient thermal regime of the mantle induced by recent tectonic activity in many parts of the Variscan belt prevent reliable estimation of mantle geotherms from surface heat-flow data. Some attempts have been made by Cermak & Bodri (1995), who argued for a uniform lithospheric thermal thickness ( 7 0 - 8 0 km) in Hercynian Europe along the European Geotraverse with a slight southward decrease in thickness. Within the frame of this model, temperatures at 50 km depth were estimated to be in the range 700-900 ~ (Cermak 1995). For the Bohemian Massif, a steady-state thermal model of the mantle interpreted jointly with gravity data (Pasquale et al. 1990; Zeyen et al. 2002) has led to the conclusion that the thermal lithosphere beneath this terrane is c. 9 0 - 1 2 0 km thick. A melilite-nephelinite composition of magmas, typical for early stages of Cenozoic magmatism in the Massif Central and Rhenish Massif, implies that the thickness of the Hercynian lithosphere was at least 80-100 km in the Tertiary (Artemieva 1993). In comparison, based on analysis of Hercynian mafic magmas, Lorenz & Nicholas (1984) argued that the regional lithospheric thickness during the Variscan orogeny was probably between 40 and 50 km, implying a c. 4 0 - 5 0 km growth of the lithosphere by thermal cooling over a period of 200-300 Ma. Regional P- and S-wave tomographic models have been recently used to assess upper mantle temperatures in Western Europe (Goes et al. 2000). At present this work gives, probably, the best available constraints on the thermal regime of the European mantle, despite a significantly different lateral and vertical resolution of the two tomography models and inevitable weakly constrained assumptions on mantle composition and its fluid

regime. According to these estimates, mantle temperatures in Hercynian Europe along a 10~ profile may exceed 1000 ~ at a depth of 100kin, whereas the lithospheric thermal thickness, defined as the depth to an isotherm of 1300 ~ is expected to be c. 120-140 km. These values are close to thermal estimates for Palaeozoic rifts within the EEP (Fig. 6) (Artemieva 1993), such that, within the accuracy of model constraints, the range of mantle temperatures should be similar for most of the tectonic structures of Europe with Palaeozoic tectonothermal ages.

The Uralides Tectonics. The Uralian orogen, which is composed of a series of accreted island arcs, volcanic complexes and fold belts, is an unusual Palaeozoic orogen, as it has remained intact within the continental interior since its formation. Surface geology (in particular, the presence of ophiolite complexes), plate tectonic reconstructions and palaeomagnetic data have been used to argue that the formation of the Uralides began some time between the Early Ordovician and Early Carboniferous by accretion of late Proterozoic-Palaeozoic microcontinental fragments and island arcs formed at the active margin of the Kazakhstan plate to a passive continental margin of the EEC (Savelieva 1987; Zonenshain et al. 1990; Seng6r et al. 1993). The Main Uralian Fault, a 20 km wide zone of sheared schists with a deformation age of 450-385 Ma, is a well-preserved plate boundary, which separates the former passive continental margin zone of the EEC in the west from the accreted Asian island arc, oceanic and continental terranes to the east. It appears in normal-incidence reflection seismic profiles as a 40 ~ east-dipping reflectivity zone extending to a depth of at least 15 km (Knapp et al. 1998), and has been interpreted as an Ordovician subduction zone dipping beneath the Kazakhstan continent (Hamilton 1970). In SilurianEarly Devonian times (the ages of the oldest island-arc complexes of the Tagil and West Magnitogorsk zones), the eastern margin of the EEC could already have become an active continental margin, with a west-dipping subduction zone existing in the Devonian (Hamilton 1970; Degtiarev 2001). The formation of a subduction zone dipping beneath the EEC could have a strong influence on the Devonian tectonics of the EEP. Models of mantle convection that take into account the dynamic effect of a subducting slab provide a good explanation for a peak in sedimentation in the eastern part of the EEP, associated with a Devonian west-dipping subduction at the Urals (Mitrovica et al. 1996). At the final stages of the collision of the EEC and the Siberian-Kazakhstan plate (at c. 3 2 0 - 2 5 0 M a ) the remaining oceanic plate between the two cratons was subducted eastwards underneath the Kazakhstan continent, and the Urals fold belt was developed. However, the modern topography of the Urals came into existence only during the Tertiary-Quaternary (Lider 1976; Morozov 2001) and the recent uplift of the Urals is as enigmatic as Cenozoic uplift of the Caledonides of Norway and Greenland (Japsen & Chalmers 2000). Seismic data. The Urals orogen has a well-preserved, more than

50 km thick, crustal root, reaching a depth of about 65 km in the Polar Urals and under the Tagil-Magnitogorsk block (Druzhinin et al. 1990; Egorkin & Mikhaltsev 1990; Carbonell et al. 1996), very high average crustal velocities as a result of magmatic intrusions, and a 175-200 km thick lithosphere (Mechie et al. 1993; Ryberg et al. 1996; Knapp et al. 1996; Fig. 7). The most recent summary of geochemical and seismic data on the crustal structure along the length of the orogen, as well as new tectonic and geodynamic constraints on the subduction-related and orogenic processes, have been presented by Brown et al. (2002). However, data on the subcrustal lithosphere of the Uralides remain limited.

DEEP E U R O P E T O D A Y

The results of teleseismic tomography across the Middle Urals (Poupinet et al. 1997) show that, down to 100 km depth, the subcrustal lithosphere beneath the Western Urals has seismic velocities 2 - 3 % higher than beneath the accreted island arc complex to the east of the Main Uralian Fault. This result suggests that the fast lithosphere of the EEC dips underneath the low-velocity lithosphere of the Uralides. These results are consistent with seismic refraction interpretations along the PNE profile Quartz (Mechie et al. 1993; Ryberg et al. 1996), which show that the Urals are underlain by an eastward-dipping high-velocity block with compressional velocities of c. 8.7 km s -~ down to a 100km depth. Such high velocities may correspond to the palaeosubduction-related preferred mineral orientation in the underthrust lithosphere of the East European continental margin. However, modern tectonic models reject the idea that the Uralides are entirely underlain by lithosphere of EEC affinity (Morozov 2001). Along the URALSEIS seismic profile in the Southern Urals, the cratonic lithosphere down to depths of 6 0 - 2 2 0 km extends no further than 200-250 km to the east of the 'geological' edge of the EEC (Savelyev et al. 2001). Correlation of the seismic structure of the upper mantle down to 100- 200 km depth with the surface geology in the Urals suggests that orogenic processes have affected most of the lithosphere and that their signature has been preserved in the upper mantle for hundreds of millions of years. Seismic models of the crustal structure along the ESRU profile in the Middle Urals indicate that the Uralides extend beneath the sedimentary cover of the West Siberian Basin (Friberg et al. 2001). Based on an analysis of magnetic anomalies, Hamilton (1970) placed the eastern margin of the Uralides beneath the central part of the West Siberian Basin. This is consistent with seismic models of the upper mantle of northern Eurasia based on refraction data along the PNE profile Quartz (Ryberg et al. 1996), which show that the lithospheric thickness changes from c. 200 km, typical for the EEP and probably for the Uralides, to c. 150 km at a distance of 500 km eastwards from the Urals. Thus, it is likely that the highvelocity block beneath the western part of the West Siberian Basin is the extension of the Uralides. Similar to the northern EEP, a pronounced reduced-velocity zone is observed beneath the Uralides along the Quartz profile in the depth interval of 1 0 5 - 1 3 0 k m (Ryberg et al. 1996; Morozova et al. 2000; Fig. 8c). This highly reflective layer with reduced seismic velocities extends for 3000 km further eastwards (Thybo & Perchuc 1997) and is underlain by a high-velocity layer at c. 2 0 0 - 2 5 0 k m depth (Nielsen et al. 1999; Kuzin 2001). Seismic reflection profiling of the Southern Urals (Knapp et al. 1996) revealed mantle reflections at depths of c. 80 km and 175 km; the lower reflector was interpreted as possibly imaging the base of the lithosphere. Thermal data. The lithospheric thermal thickness at the eastern margin of the EEC, adjacent to the Ural mountains, is similar to estimates based on seismic interpretations for the Urals, c. 170-200 km (Artemieva & Mooney 2001). However, there is no reliable constraint of lithospheric temperatures beneath the Uralides, as anomalously low heat-flow values have been reported for the Southern Urals (Salnikov 1984; Kukkonen et al. 1997): c. 25 m W m - 2 in the 1500 km long Magnitogorsk block, compared with 4 0 - 5 0 m W m -2 in the EEP and in the eastern part of the Southern Urals (Fig. 3). Possible explanations for this thermal anomaly include palaeoclimatic variations, low crustal heat production, lateral groundwater heat transfer, or anomalously low mantle heat flow beneath the central part of the Southern Urals, perhaps associated with Palaeozoic subduction zones. For models with a low crustal heat production in island arc complexes of the crust, Moho temperatures (at a depth of c. 60 km) are estimated to be c. 550-600 ~ (Kukkonen et al. 1997). Downward continuation of this conductive geotherm would imply a lithospheric thermal thickness of c. 200 km.

27

Gravity data. T h e short wavelength of gravity anomalies in the Uralides (less than 100-200 km) suggests their crustal origin. Gravity studies across the Middle and Southern Urals show a + 5 0 mGal linear high of Bouguer anomalies above the Magnitogorsk block flanked by two negative gravity anomalies spatially limited to the area of the Pre-Uralian Foredeep, and the Western and Central Uralian zones ( - 7 5 to - 5 0 m G a l ) to the west from the Main Uralian Fault and to the Eastern Uralian Zone ( - 6 5 to - 4 0 mGal) in the Eastern Urals. The negative Bouguer anomaly in the Pre-Uralian Foredeep is attributed to thick sediments at the edge of the EEC; as the positive free-air anomaly in the Western and Central Uralian Zones is well correlated with the topography, the Bouguer gravity minimum in these tectonic zones is well explained by a superposition of low-density sediments and the nearby crustal root beneath the Tagil-Magnitogorsk block (D6ring et al. 1997). Similarly, the negative anomaly in the Eastern Zone has been explained by a joint effect of intruded granites and the nearby crustal root. Surprisingly, the crustal root beneath the Tagil-Magnitogorsk block is not reflected in the topography and produces a positive Bouguer gravity anomaly. 2D gravity modelling shows that gravity maximum can be explained by the joint effect of a subsurface load of mafic-ultramafic material superimposed on the negative gravity effect of a crustal root (D6ring et al. 1997). Seismic modelling supports this conclusion and indicates the presence of the crustal high-velocity body within the island arc material of the Magnitogorsk Zone (Carbonell et al. 2000).

P a l a e o z o i c rifts

The Precambrian part of Europe comprises extensional structures, the development of which may have involved deep mantle processes. The most important (and the most well-studied) Palaeozoic rifts include the Oslo rift in the southern part of the Baltic Shield (considered as a classical example of a 'passive rift') and the Pripyat-Dniepr-Donets rift in the southern part of the EEP (which is considered to be an 'active rift'). However, the amount of data on the structure of their subcrustal lithosphere is limited. rift (PDDR). Geophysical models of the lithosphere of the PDDR and the adjacent structures have been the goal of the GEORIFT project of EUROPROBE (Stephenson 2004), in the frame of which new regional gravity models of mantle anomalies (Yegorova et al. 1999) and geodynamic models of tectonic evolution of the region (Kusznir et al. 1996; Starostenko et al. 1999) were developed. However, seismic data on the deep lithospheric structure of the Palaeozoic rifts within the EEP are not available, as the deepest reaching reflection and refraction data of the DOBRE experiments provide seismic images to depths of only a few kilometres below Moho (DOBREfraction' 99 Working Group 2003). Geodynamic models of the formation of continental rifts are traditionally divided into models of 'passive' and 'active' rifting (Seng6r & Burke 1978); however, the validity of this approach is debated, as rifting activity is probably also governed by forces related to plate tectonics and thus many active continental rifts can be caused by stress-induced lithospheric extension (Ziegler & Cloetingh 2004). Traditionally, active models are based on the hypothesis that crustal extension results from a (plume-related?) thermal anomaly in the upper mantle. In these models, an uplift of hot mantle material to lithospheric depths (sometimes up to the crust) produces lithospheric extension and thinning. Indirect evidence for the presence of mantle plumes beneath some of the rift zones is provided by isotope data and the large volumes of magmas generated simultaneously with rifting. In particular, the model of active rifting is proposed for the Palaeozoic rifts in the southern part of the EEP

Pripyat-Dniepr-Donets

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I.M. ARTEMIEVA E T A L .

(Chekunov et al. 1992) based on a large volume of Devonian magmas (with a peak at c. 350 Ma) in the PDDR (Lyashkevitch 1987) and on geochemical data for the Dniepr graben (Wilson & Lyashkevitch 1996). A gravity maximum over the PDDR is interpreted to be caused by a large volume (c. 60%) of high-density mantle intrusive rocks in the crust (Yegorova et al. 1999), although a similar effect perhaps can be produced by eclogitization of the lower crust. The thermal regime of the lithosphere of the PDDR can be constrained from surface heat-flow data as the lithosphere has relaxed to a stationary thermal regime since the Devonian rifting. The PDDR is characterized by a linear, c. 200 km wide, anomaly of a slightly elevated surface heat flow (4555 m W m -z, reaching locally 70-90 mW m -2 in the Pripyat Depression), which separates the Ukrainian Shield (254 0 m W m -2) and the Voronezh Massif (Fig. 3). However, typical heat-flow values within the PDDR are similar to the values measured within most of the EEP, and a relatively short wavelength of the zone with higher heat flow suggests a chiefly shallow origin for heat-flow variations. Steady-state thermal models (i.e. Kutas 1979; Artemieva 2003) imply that the lithospheric thermal thickness in the southern part of the EEP, including the PDDR, is c. 120-150km, which, within the model accuracy, is similar to estimates for the Palaeozoic structures of Western and Central Europe (the Armorican and Bohemian massifs, in particular; see above and Fig. 6). It implies that the lower part of the cratonic lithosphere (c. 50-100 km) could have been thermally eroded or delaminated during the Devonian rifting. Alternatively, models of the transient thermal evolution since the impact of a presumed mantle plume (at 369 Ma) (Galushkin & Kutas 1995; Starostenko et al. 1999) result in lithospheric temperatures significantly lower than in steady-state models. In these interpretations, geotherms are similar to the EEP geotherms, implying a lithospheric thermal thickness of c. 180-200 km as in other Archaean-early Proterozoic cratons of the world (Jaupart & Mareschal 1999; Artemieva & Mooney 2001).

Because of the relatively small size of the Oslo rift, the structure of its lithospheric mantle cannot be resolved in large-scale geophysical models. Dispersion analysis of long-period Rayleigh waves implies that the thickness of the seismic lithosphere in southern Fennoscandia is c. l l 0 - 1 2 0 k m (Calcagnile 1982). Despite a low lateral and insufficient vertical (50-100 km) resolution of this model, these estimates agree with the depth at which strong, almost horizontal reflectors are continuously seen in the upper mantle; that is, 80-100 km over distances of 5 - 2 0 km (Lie et al. 1990). By analogy with lower crustal reflectors, they are interpreted as a transition from brittle to plastic deformation and thus can be considered to be the base of the rheological lithosphere. Similar estimates of lithospheric thickness in the southern part of the Baltic Shield in the vicinity of the Oslo rift were obtained by regional P-wave seismic tomography (Plomerovfi et al. 2001). The Oslo rift is characterized by positive Bouguer anomalies (0 to +50 mGal) compared with negative anomalies (less than - 5 0 reGal) in the adjacent southern Fennoscandia (Ramberg 1976). Despite the inherent non-uniqueness of gravity models, most researchers interpret positive anomalies to indicate large volumes of mantle intrusions in the crust (e.g. Neumann et al. 1995). Surface heat flow in the Oslo rift is similar to the values measured in the Proterozoic terranes of Fennoscandia (40-50 mW m-2), suggesting that a stationary thermal regime has been re-established in the rift zone. Short-wavelength, slightly increased heat-flow values along the rift axis are likely to be produced by higher crustal heat production in the areas of Palaeozoic magmatism. Estimates of Moho temperatures (at a depth of c. 29-34 km; Kinck et al. 1991) differ strongly: P - T petrological estimates give values of 250-350 ~ (Neumann et al. 1995), whereas lithospheric geotherms constrained by surface heat flow suggest temperatures of 550-650~ (Balling 1995; Cermak & Bodri 1995). Values of 450-550 ~ as for other Palaeozoic structures of Europe (Fig. 6), probably provide the most conservative estimate.

Oslo rift. The Oslo rift, which includes a chain of rift structures and grabens, extending from southern Norway to the TTZ or the Caledonian suture over a distance of c. 400-600 km, is considered to be a classical example of a passive rift (Pedersen & van der Beek 1994). Models of passive rifting assume that lithospheric extension is caused by tensional stresses at plate boundaries. If the stress is high (or the lithosphere is hot and thin), stress-induced lithosphere extension may cause rifting (Kuznir & Park 1984), accompanied by a passive upwelling of mantle material along weak lithospheric zones and its adiabatic melting. Because in this case the source of magmas is within the upper mantle, geochemical methods cannot reliably distinguish the models of passive from active rifting caused by small-scale mantle convection. Despite a large volume of basaltic magmas emplaced at c. 240-300 Ma (Neumann et al. 1995), the P - T analysis of their composition indicates that the magmatism was not caused by a high-temperature anomaly in the mantle (Neumann 1994). Numerical modelling of thermo-mechanical processes of rifting has shown that a step-like increase in lithospheric thickness at the eastern margin of the rift could have led to a passive diapirism and consequent rifting (Pascal et al. 2002). This explanation is close to the model by King & Anderson (1995) for the formation of large igneous provinces at cratonic margins by small-scale convection initiated by a step-like change in lithospheric thickness at the transition from a thick cratonic root to a thin younger lithosphere. Alternatively, based on analyses of the lateral distribution of seismic crustal velocities over the whole area to the south of the Oslo rift, Thybo (1997) proposed that the primary driving force for formation of the rift structures throughout the area could be related to deformation caused by far-field forces from the distant Variscan orogeny.

Lithosphere of Mesozoic-Cenozoic structures of Europe Most of the Hercynian orogen has been significantly reworked and overprinted as the result of plate tectonic processes related to the collision of the Eurasian and the African lithospheric plates, as well as by tectonomagmatic events associated with the formation and development of the Central European Rift System.

R e g i o n s o f Cenozoic s u b d u c t i o n a n d A l p i n e o r o g e n y Tectonics of the region. The huge volume of geological-geophysical information on the tectonic evolution and lithospheric structure of the Alps and the Mediterranean prevents even a simple listing of major results within the framework of the present review. For detailed information the reader is addressed to other publications (e.g. Mueller 1989, 1997; Blundell et al. 1992; Kissling & Spakman 1996; Pfiffner et al. 1997; Cavazza et al. 2004). The convergence of the Eurasian and African plates began at c. 120 Ma. It resulted in plate collision and subduction at c. 65 Ma and uplift of the Alpine orogenic belt after c. 23 Ma (Schmid et al. 1996; Castellarin & Cantelli 2000). The present convergence velocity is c. 9 mm a-1 (De Mets et al. 1994). These tectonic processes have led to the formation of a highly complex and heterogeneous structure of the crust (Him et al. 1980; Giese 1985; Pfiffner 1990; Ye et aL 1995; Bleibinhaus & TRANSALP Working Group 2001; TRANSALP Working Group 2001, 2002) and the upper mantle of the region (Him et al. 1984; Panza et al. 1986; Pfiffner et al. 1988; Kissling 1993; Lippitsch et al. 2003). Numerical models of mantle convection indicate that subduction of a lithospheric plate beneath continental lithosphere

DEEP EUROPETODAY causes a dynamic down-flexure of the lithospheric plate as a result of the down-pull by the dense cold subducting slab, leading to fast basement subsidence and basin formation (Gurnis 1992; Stern & Holt 1994; Pysklywec & Mitrovica 1998). This mechanism was used to explain the formation of the Po basin as the result of subduction beneath the Alps (Bott 1990), and can explain (at least partly) the formation of the Tyrrhenian, Aegean and Pannonian basins. It is likely that subduction-induced basin subsidence can explain one of the stages in the formation of the Northern Caucasus foredeep as the result of subduction of the Arabic (Turkish) plate under the Scythian plate. However, the existing geodynamic models attribute the formation of this basin chiefly to crustal processes (e.g. eclogitization or viscous flow in the lower crust) (Artyushkov 1993; Mikhailov et al. 1999; Ershov et al. 2003). Geophysical models f o r the Alps and the Mediterranean. Regional P-wave (Him et al. 1984; Spakman 1986, 1990; Blanco & Spakman 1993; Souriau & Granet 1995; Kissling & Spakman 1996; Piromallo et al. 2001; Lippitsch et al. 2003) and S-wave (Panza et al. 1986; Snieder 1988; Pasyanos & Walter 2002) refraction and tomography models provide the bulk of the available information on the structure of the crust and upper mantle of the Alps and the Mediterranean. They indicate the presence of several subduction zones in the region and pronounced lithospheric thickening associated with them, especially underneath the Alps (e.g. Figs 2a, e, g and 4a, d). The maximal crustal thickness (crustal root) beneath the western and central Alps is found in a block where high upper mantle velocities extend to a depth of 200-250 km (Cavazza et al. 2004), interpreted as a lithospheric plate (presumably continental European lower lithosphere) steeply subducting southeastwards beneath the Adriatic microplate (Lippitsch et al. 2003; Fig. 8). This high-resolution teleseismic P-wave tomography of the Alps further suggests the existence of the second NE-dipping subduction zone in the eastern Alps, interpreted as the continental Adriatic lower lithosphere subducting beneath the European plate (Lippitsch et al. 2003). Similarly, P-wave residuals models for Southern Europe (Babugka et al. 1990) advocate the existence of two regions, beneath the western and central Alps and beneath the eastern Alps, with high values of lithospheric thickness (>200km) with a sharp decrease in lithospheric thickness to c. 60 km beneath the Po basin (Fig. 4d). Similar lithospheric structures, with localized high-velocity blocks in the upper mantle interpreted as subducting slabs, have been identified in seismic tomography models for the Ligurian-Tuscany region of Italy (Panza et al. 1986) and southern Spain, where a detached subducted slab is identified in the regional tomographic images of the upper mantle (Spakman 1991; Blanco & Spakman 1993). Regional P-wave tomography models indicate the existence of a 30 km wide block with 2% lower velocities extending to a depth of c. 80-100 km beneath the central and eastern Pyrenees (Souriau & Granet 1995). This velocity anomaly has been interpreted as lower crust of Iberia subducted as the result of convergence of the Eurasian and the African plates (Vacher & Souriau 2001). By analogy to a model proposed earlier for the Alps (Austrheim 1991), weak negative residual gravity anomalies calculated for the Pyrenees are explained by eclogitization of the lower crust during its subduction (Vacher & Souriau 2001). Other zones of Cenozoic subduction (including the Hellenic arc, the Carpathians and the Caucasus) are characterized by linear belts of positive residual gravity anomalies (Fig. 5), ascribed to cold dense subducting lithospheric slabs in the underlying mantle. These gravity anomalies spatially correlate with linear high-velocity upper mantle structures resolved in regional P-wave seismic tomography models. Similarly, the presence of an ancient subducting slab beneath the western margin of the EEP as indicated by a regional S-wave tomography model (Nolet & Zielhuis 1994; Zielhuis & Nolet 1994) is supported by a linear belt of positive residual gravity anomalies along the TESZ (Fig. 5).

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Estimates of mantle temperatures for the tectonically active regions of Europe are scarce, as steady-state models constrained by surface heat-flow measurements (e.g. Della Vedova et al. 1990; Cermfik 1994; Zeyen et al. 2002) are not applicable. Thermo-kinematic models (e.g. Werner 1981; Royden et al. 1983b; Davy & Gillet 1986; Zeyen & Fernandez 1994; Bousquet et al. 1997) require detailed information on dynamic processes in the mantle, which are usually not completely understood, and, as a result, such models are poorly constrained. An advanced 2D thermo-mechanical model of the lithosphere of the Alps takes into account the processes of crustal shortening and formation of crustal and lithospheric roots during subduction (Okaya et al. 1996). According to this model, the Moho is an almost isothermal boundary with a temperature of c. 500-600 ~ although crustal thickness across the orogen changes from c. 30 km beneath the Variscides in the north to c. 55-60 km beneath the Alps and to c. 30-34 km beneath the Po basin in the south (Giese & Buness 1992; Pfiffner et al. 1997; Waldhauser et al. 1998; TRANSALP Working Group 2002); lithospheric thermal thickness gradually increases from north to south from c. 80 km beneath the Variscides to c. 120-150 km beneath the southern Alps-northern Apennines (Okaya et al. 1996). Steady-state thermal models for the lithosphere of Southern Europe give overestimated values of mantle temperatures and, thus, lithospheric thicknesses that are too small (70-80 km) (Della Vedova et al. 1990; Cermfik 1993). Although regional magnetotelluric (MT) studies indicate the presence of a highly conducting upper mantle layer at a depth of >90 __ 10 km (EREGT Group 1990; Fig. 4e), its origin can be ascribed not only to the presence of melt, but also to fluids or graphite (although the presence of fluids would cause the dissolution of the pyroxenes of the rocks into partial melt as interpreted in some places of the EEP and in central France; Thybo & Perchuc 1997). The Carpathians and the P a n n o n i a n Basin. A large number of geodynamic models for the Cenozoic evolution of the Pannonian Basin propose either an 'active' (e.g. Bergerat 1989) or a 'passive' role (Royden et al. 1983a, b; Le Pichon & Alvarez 1984; Horvath 1993; Huismans et al. 2001; Huismans & Bertotti 2002; Sperner et al. 2002) of the asthenospheric mantle in its formation and tectonic evolution. The large variety of passive models is probably due to a lack of detailed information on the interaction of the subducting slab with the asthenosphere-lithosphere system at different stages of subduction, especially when the continuous formation of the Alps affects the stress regime in the adjacent tectonic regions (Cloetingh et al. 2004). Seismic models based on P-wave residuals (Babu~ka et al. 1988) (Fig. 4d), MT and electromagnetic studies (Adam et al. 1982; Adam 1996; Adam & Bielik 1998), and geothermal (mostly steady-state) models (Bielik et al. 1991; Cermfik 1994; Cranganu & Deming 1996; Bojar et al. 1998; Andreescu et al. 2002; Zeyen et al. 2002) reveal an anomalously thin (60-80 km) lithosphere of the Pannonian Basin, with local values as small as c. 40 km (Posgay et al. 1995). Negative residual isostatic anomalies (Fig. 5 and Yegorova et al. 1998) indicate the presence of anomalous low-density asthenospheric material and support the hypothesis that an earlier passive stage of basin formation may have been replaced at present by an active mantle (Huismans et al. 2001). Low values of lithospheric thickness beneath the Pannonian Basin contrast with a thick lithosphere beneath the western Carpathians, where the thickness has been estimated to be 150 km by MT studies (Praus et al. 1990; Fig. 4e), 130-150 km by joint interpretation of surface heat-flow and gravity data (Zeyen et al. 2002) and seismic and MT data (Horvfith 1993), and c. 100km by steady-state thermal modelling (Cerm~ik 1994), although the steady-state thermal models are physically inadequate for Cenozoic tectonic structures. The thick lithosphere beneath the Carpathians is ascribed to westward subduction of the Eurasian slab (Wortel & Spakman 2000). The existence of a

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subduction zone beneath the southern Carpathians is well established from seismic data, with the main seismicity localized in the depth range 60-180 km along a steeply dipping plane of the Vrancea zone. Regions o f M e s o z o i c - C e n o z o i c tectonomagmatic activity Rift system of the North Sea (RSNS). The rift system of the North

Sea, deeply buried under thick Tertiary sediments, is one the most prominent Mesozoic rifts of Europe and includes the Viking Graben in the north and the Central Graben in the south. Although its formation probably began during the late stages of the Caledonian orogeny, the major phase was related to Mesozoic tiffing at the Atlantic passive continental margin; some researchers consider the RSNS as an unopened ocean or a failed arm of a broad Mesozoic rifting along the North Atlantic margins (Bott 1995). Mesozoic rifting started in Triassic-Early Jurassic times, continued for an unusually long time (c. 175 Ma; Bott 1995), and may have been affected by a mantle plume. The subsequent post-rift thermal subsidence occurred during the Tertiary (Ziegler 1992), and may be partly ascribed to delayed thermal reactions caused by late metamorphic reactions in the uppermost mantle (Vejbaek 1990). Despite a huge geological-geophysical database on the crustal structure of the RSNS, data on its upper mantle structure are very limited. Regional S-wave seismic tomography models (Fig. 2), which have better vertical resolution than P-wave tomography models, show high velocities in the mantle down to 100150 km depth. As it is unlikely that mantle temperatures in the Mesozoic rift are low, it is possible that high mantle velocities originate from compositional anomalies. Furthermore, residual gravity anomalies have strong negative values in the North Sea region (Fig. 5), implying a low-density (hot?) upper mantle beneath the RSNS, in agreement with a strong attenuation anomaly at a depth of 150 km (Fig. 2h). As seismic velocity and gravity models for the RSNS apparently contradict each other, the origin of the anomaly remains unclear. Central European Rift System (CERS). The CERS is formed by a continuous chain of Cenozoic rift structures that extend from the Atlas Mountains in northern Africa to the North Sea. Various geodynamic models, including plume-related active rifting, passive rifting in response to collisional processes in the Alps and Pyrenees, back-arc rifting, or slab pull associated with Alpine subduction, have been proposed to explain geological and geological data available for the CERS: a thin crust, high surface heat flow, weak seismicity, Cenozoic magmatism and anomalous properties of the upper mantle (for reviews see Ziegler 1992; Prodehl et al. 1995; Merle & Michon 2001; Dezes et al. 2004; Michon & Merle 2005). However, because of the narrow structures of the CERS, one cannot expect to resolve upper mantle anomalies in large-scale geophysical models (e.g. Figs 2, 4 and 6). Below we discuss in detail the lithospheric structure of three major tectonic provinces within the CERS: the Rhine Graben, the Rhenish Massif, and the French Massif Central. Rhine Graben. Intensive magmatism in the Rhine Graben began at 80 Ma and continued until 7-15 Ma; however, tiffing began only at 45 Ma in the southern part of the Rhine Graben, from where it gradually extended northwards. The crustal structure of the Rhine Graben is well known, although data on the properties of the upper mantle are non-unique. The surface expression of the rift zone does not exceed 36 km, whereas the width of the lithospheric zone with anomalous properties is estimated to be 200 km (Prodehl et al. 1995). Recent teleseismic surface-wave studies indicate that the region with low mantle velocities is localized to the Rhine Graben itself, whereas the regional value of lithospheric thickness is c. 80 km (Glahn et al. 1993). Absolute P-wave velocities estimated by tomography models do not reveal a low-

velocity anomaly in the upper mantle beneath the Rhine Graben down to a depth of c. 280 km (Achauer & Masson 2002). Furthermore, regional P-wave tomography indicates high mantle velocities beneath the Rhine Graben (Ansorge et al. 1979; Spakman 1986; Babugka et al. 1988). The Rhine Graben is characterized by weak negative Bouguer anomalies (less than - 30 mGal). They are explained either by an anomalous crustal structure without any significant thermal anomaly in the mantle (Grosse et al. 1990) or by the presence of anomalously low-density material in the upper mantle as required by strong negative mantle residual anomalies ( - 1 5 0 to - 2 0 0 mGal) (Yegorova et al. 1998). However, the latter conclusion is not supported by thermal data. High values of surface heat flow in the Rhine Graben (ranging from 70 to 140 mW m - 2 with an average around 1 0 0 m W m -2) were measured in shallow boreholes (Cermak 1995). They have a strong shortwavelength component, which implies that a large part of the heatflow anomaly has a shallow origin and is probably caused by groundwater circulation. Thus, geophysical data on the upper mantle structure do not provide evidence for a presence of a 'babyplume' beneath the Rhine Graben, but favour a passive mechanism of rifting, caused by lithospheric extension, which resulted from a complex stress field associated with the convergence of the Eurasian and the African plates. Rhenish Massif(RM). The intensive volcanism of the RM began in the Eocene with an eruption of nephelinitic magma (Wilson et al. 2004), which implies a lithospheric thickness of at least 80-100 km. At c. 25 Ma the composition of magmas changed to basalts and trachytes with a depth of generation < 6 0 - 8 0 km. The youngest volcanic areas in the western part of the RM have an age of c. 700 ka (Lippolt 1983). Uplift of the RM began in the late Oligocene and still continues. The upper mantle structure beneath the RM is asymmetric according to different geophysical data. Contrasting Bouguer anomalies with weakly negative values ( - 10 to - 20 mGal) to the west of the Rhine and weakly positive anomalies (+10 to + 2 0 mGal) in the eastern part are well explained by a heterogeneous crustal structure (Jacoby et al. 1983). However, low velocities in the upper mantle of the RM were found both in P-wave and S-wave models (Panza et al. 1986; Spakman 1986; Babugka et al. 1988; Ritter et al. 2001). Teleseismic studies of the RM reveal a zone with a 3 - 5 % low-velocity anomaly at a depth of 50-200 km, which is shallowest in the western part of the RM (Raikes & Bonjer 1983). A recent P-wave tomography experiment in the Eifel area supports earlier interpretations and shows a narrow (with a radius of about 100 km) low P-velocity anomaly in the upper mantle down to at least 400 km depth (Ritter et al. 2001). A lateral velocity contrast of up to 2% (with respect to the iasp91 model) within this columnar velocity anomaly can be explained by about 150-200 K excess temperature, which was attributed to the Rhenish plume. Nevertheless, the origin of Cenozoic tectonic and magmatic activity in the RM is still debated. The RM has high values of surface heat flow (c. 80 m W m - z ) with slightly higher values in its eastern part. Downward continuation of geotherms, constrained by upper mantle xenoliths from the RM (Seck & Wedepohl 1983), gives lithospheric thermal thickness estimates of c. 80-90 km (Fig. 6). Shallowing of the mantle transition zone beneath the western part of the RM is interpreted as an indicator of a possible upper mantle plume (Grunewald et al. 2001). Alternatively, partial melting in the upper mantle beneath the RM may be caused by passive adiabatic decompression (Schmincke et al. 1983) as a result of lithospheric extension during rifting (Ziegler & Cloetingh 2004). Massif Central (MC). Volcanic activity in the MC began in the Oligocene and was accompanied by uplift of the entire massif. The main phase of volcanism was at 2 - 5 Ma; but there is no correlation between the age of volcanism and its geographical distribution (Werling & Altherr 1997). Similar to the RM, P - T analysis of lower crustal and mantle xenoliths of different ages

DEEP EUROPE TODAY and from different locations (Coisy & Nicolas 1978; Werling & Altherr 1997) indicates that all of them approximately follow the 8 5 - 9 0 mW m -2 reference geotherm of Pollack & Chapman (1977; Fig. 6), implying a lithospheric thermal thickness of c. 7 0 - 8 0 km at the time of eruption. Three-dimensional regional P-wave tomography models reveal a low-velocity zone in the upper mantle of the MC at a depth of 6 0 - 1 0 0 km (Granet et al. 1995a), which is interpreted as the top of the mantle upwelling (plume?) (Granet et al. 1995b). Estimates of lithospheric thickness from P-wave residuals (Fig. 4d; Babugka et al. 1988, 1992) and surface waves (Souriau et al. 1980) also give a depth of c. 6 0 - 1 0 0 km. The region with a 3% velocity decrease in the upper mantle correlates spatially with both the area of recent volcanism and a local long-wavelength minimum of Bouguer anomalies ( - 4 5 mGal, Autran et al. 1976). However, the entire MC is characterized by the same range of residual mantle gravity anomalies ( - 5 0 to - 1 5 0 mGal) as other terranes of Proterozoic to early Palaeozoic ages within the Variscides (e.g. the Bohemian and the Armorican massifs; Fig. 5). P-wave tomography models for the MC (Granet et al. 1995a) have been used to constrain density and temperature of the upper mantle. Both gravity (Stoll et al. 1994) and temperature (Sobolev et al. 1996) models do not require the presence of large percentages of melt in the upper mantle of the MC, although the latter model assumes the presence of a mantle plume beneath the MC as responsible for a regional (50-70 km wide) lithospheric thinning to 70 km depth. Lucazeau et al. (1984) have modelled the thermal anomaly beneath the MC (where surface heat-flow values are 105 __ 13 m W m -2) by upwelling of a 40 km wide mantle diapir, and concluded that c. 50% of the anomaly can be attributed, to the crustal heat production and the rest should be ascribed to the combined effect of the mantle diapir and the Hercynian orogeny. Petrological studies of mantle xenoliths from the MC have revealed a significant difference in the upper mantle properties beneath its southern and northern blocks (Lenoir et al. 2000). Mantle peridotites from the northern domain have geochemical signatures similar to peridotites from Archaean cratons (though with low Mg#; Mg# -- MgO/(MgO + FeO)). Such difference in the composition of mantle peridotites may reflect a block structure of the Hercynian lithosphere formed by Palaeozoic accretion of continental terranes of different ages. Heterogeneous lithospheric structure of accreted terranes could have favoured the location of the Cenozoic mantle thermal anomaly beneath the young and thin lithosphere of the southern block of the MC. The existence of a hidden Hercynian suture zone in the lithosphere of the MC is indicated by seismic anisotropy models (Babugka et al. 2002), which suggest the existence of a Cenozoic asthenospheric flow from the Western Mediterranean to beneath the MC, channelled along a boundary between different lithospheric blocks (Barruol & Granet 2002). This model does not require the presence of a mantle plume (or diapir) to explain the mantle thermal anomaly beneath the part of the MC where the strongest seismic velocity anomaly is observed in tomography models.

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determined by different geophysical techniques may approximately correspond to the transition from the lithosphere to a zone of partial melt (see above for a detailed discussion). A diffuse character of the base of the seismic lithosphere together with a substantial thickness of the transition zone between purely conductive and purely convective heat transfer limits vertical resolution of any integrated model of lithospheric thickness to 50 km (Fig. 9). The integrated model of the lithospheric thickness in Europe (Fig. 9) is based on P-wave seismic tomography models (Spakman 1990; Bijwaard & Spakman 2000; Piromallo & Morelli 2003), surface-wave tomography models (Panza et al. 1986; D u e t al. 1998; Shapiro & Ritzwoller 2002), P-wave residuals (Babugka & Plomerovfi 1992), thermal models (Balling 1995; Cermak & Bodri 1995; Artemieva 2003), and P - T data for mantle xenoliths (Coisy & Nicolas 1978; Seck & Wedepohl 1983; Nicolas et al. 1987; Werling & Altherr 1997; Kukkonen & Peltonen 1999; Malkovets et al. 2003). Taking the limitations of different interpretation techniques into account, we compare and combine these models into a consistent map, to identify the bulk features of the lithospheric structure of Europe. Inevitably, the model smears some small-scale details; they can be found in corresponding publications of regional surveys (e.g. see the subsequent papers of this book). Our interpretation reveals continent-scale differences in both thickness (Table 2) and composition of the lithospheric mantle. These major differences reflect the tectonic history of the continent over c. 3.5 Ga and the effects of mantle processes on lithosphere modification. Thus, this integrated model provides a reference frame for comparing tectonic structures of Europe and their world analogues, and it forms the basis for a better understanding of geodynamic evolution of the European continent in space and time.

Synthesis: an integrated model of the European upper mantle structure and compositional variations Comparison of different seismic models of the upper mantle of the continent (including P- and S-wave tomography, P-wave residuals, reflection and refraction profiles) with MT, electromagnetic, thermal and gravity models and mantle xenolith data is used here to constrain an integrated model of the lithosphere of Europe. A change in physical properties of the upper mantle at the lithospheric base, as reflected in different geophysical models, is temperature dependent and may be caused by high-temperature relaxation or by partial melting. The lithospheric base as

Fig. 9. Integrated model of lithosphericthickness in Europe, based on seismic, thermal, MT, electromagneticand gravity interpretations. In general, a direct comparison of lithosphericthicknessvalues, constrainedby differenttechniques, is not valid, as they are based on measurementsof diverse physical parameters. The differencebetweenthe thicknessesof 'seismic' and 'thermal' lithospherecan be up to 40-50 kxn (Jaupart & Mareschal 1999), which approximately corresponds to the thickness of the transition zone between pure conductive and pure convectiveheat transfer. For this reason the isolines are drawn with a 50 km interval. North Africa, Central Asia and regions with the oceanic crust are excluded.

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Compositional variations within European lithospheric mantle Compositional variations within the cratonic roots (as a result of depletion in basaltic components) result in density and seismic velocity anomalies, which may be significantly masked by temperature variations in the upper mantle. As the Vp/Vs ratio is thought to be more sensitive to variation in composition than temperature (e.g. Lee 2003), we constrain maps of Vp/Vs ratio from smoothed and filtered P- and S-wave tomography models (Bijwaard & Spakman 2000; Shapiro & Ritzwoller 2002) at depths of 150 km and 250 km and interpret them as reflecting compositional variations in the subcrustal lithosphere of Europe (Fig. 10). Teleseismic P-wave tomography has the best lateral resolution, but poorly resolves the vertical extent of velocity anomalies (see Fig. 2a and c, where the shape of the velocity anomalies has basically the same pattern at all depths in the interval 100-265 km), whereas surface-wave tomography has the best vertical resolution. These differences reduce the obtainable resolution from straightforward comparison of S-wave and P-wave tomography results. It is, however, obvious that the cratonic and Phanerozoic parts of the European mantle at depths to 150250 km have significantly different composition. The lack of resolution in the P-wave tomography models for the northeastern part of the EEP does not permit interpretation of this part of the craton. However, a pronounced anomaly is detectable over the EEC at c. 250 km depth, which suggests that the lithosphere extends at least to this depth in the Finnish part of the Baltic Shield and the central-western part of the EEP. We supplement the data on variations in Vp/Vs ratio by data on the lateral variation of mantle residual gravity anomalies, which have a good lateral resolution and almost no vertical resolution. To separate the effects of temperature and composition on density anomalies, mantle residual gravity anomalies (Fig. 5) were corrected for thermal expansion using data on lithospheric temperatures (Figs 4c and 5) and following the approach of Kaban et al. (2003). The gravity effect of temperature variations in the upper mantle was estimated down to 225 km depth and removed from the total mantle gravity field; the resulting 'compositional' density variations are shown in Figure 1 l a. Another approach to separate the contributions of temperature from composition is based on independent free-board constraints (Fig. l lb; Artemieva 2003). There is a striking similarity between the two maps of density heterogeneities constrained by gravity and buoyancy (Fig. 11). However, both density maps lose resolution in the Caledonides (as a result of smearing of offshore gravity anomalies and unaccounted dynamic topography in free-board constraints). The strongest low-density anomalies, probably caused by a highly depleted lithospheric composition, are observed in the upper mantle of the Baltic Shield. A gradual increase of average (i.e. integral for the entire lithospheric column) lithospheric density in the EEP from north to south as a result of lateral variations of the composition is evident in both maps. The average density of the lithospheric mantle of the southern parts of the EEP is similar to the density of the Phanerozoic mantle of Western Europe. This density increase in the cratonic root can be related to metasomatic reworking of the cratonic lithosphere during large-scale intensive Devonian rift-related magmatism, when infiltration of Fe-enriched basaltic magmas may have increased the average lithospheric density (Artemieva 2003). Subduction zones of the Mediterranean and the Caucasus are marked by pronounced high-density anomalies (Fig. 11 a). Because gravity anomalies do not provide constraints on the depth distribution of anomalous masses in the upper mantle, a comparison of Figure 11 with maps of Vp/Vs at different depths (Fig. 10) permits us to speculate on their vertical distribution. There is a general overall agreement between the mantle density anomalies and the seismic compositional anomalies at

150-250km depth. In agreement with mantle xenolith data from craton and off-craton settings (e.g. Griffin et al. 1998), at these depths the transition from Archaean-early Proterozoic lithosphere of the Baltic Shield and the East European Platform to younger upper mantle of the Variscides, Caledonides and the Sveco-Norwegian province of the Baltic Shield is clearly seen in compositional variations (Figs 10 and 11). This finding supports the conclusion that, except for the subduction zones beneath the Western and Eastern Mediterranean, the Alps and the Carpathians, the lithosphere of Phanerozoic Europe does not reach 150 km depth. The high Vp/Vs ratio most likely results from the presence of partial melts at this depth in the upper mantle.

Compositional origin of velocity contrast in the TOR tomography The transition from depleted to non-depleted cratonic composition is clearly imaged in the TOR seismic tomography interpretations (e.g. Arlitt 1999; Gregersen et al. 2002; Shomali & Roberts 2002). As thermal models do not indicate any significant change in mantle temperatures across the transition zone from the Baltic Shield to the Danish Caledonides (Balling 1995; Cermak & Bodri 1995; Artemieva 2003), the sharp P-wave velocity contrast in the TOR tomography images across the Teisseyre-Tornquist Zone (TTZ) should be attributed to a purely compositional change. Moreover, if the entire velocity anomaly observed in the TOR models is caused by compositional variations in the upper mantle, it provides additional support to an earlier hypothesis that the lower crustuppermost mantle of Fennoscandia extends much further south than the geological boundary between the Baltic Shield and Danish Caledonides (Thybo 1990, 2001; Bayer et al. 2002). Interpretations of the TOR tomography model suggest a Vp contrast between the cratonic lithosphere of the Baltic Shield and the Caledonian lithosphere as large as c. 3% (6Vp c. +1% beneath the Sveconorwegian province and 6Vp c. - 2 % in the Phanerozoic mantle; e.g. Arlitt 1999). Experimental studies indicate that Vp is more sensitive to temperature variations than is V~, which is more sensitive to variations in composition (primarily, to the iron content) (e.g. Lee 2003). As a 1% Vs anomaly can be explained by a c. 4% anomaly in Fe content (e.g. Deschamps et al. 2002), probably most of the ~Vp anomaly beneath the Sveconorwegian province can be attributed to Fe depletion, although the required degree of depletion is about twice that expected for Proterozoic terranes (Griffin et al. 1998). The negative seismic velocity anomaly beneath Phanerozoic Europe cannot be explained in terms of iron-content variations and requires the presence of fluids or a strong mineralogical/compositional anomaly. The presence of fluids along the accreted cratonic margin, probably associated with ancient subduction zones, has been proposed earlier for the central segment of the TESZ (Nolet & Zielhuis 1994) and cannot be ruled out as a cause of a negative velocity anomaly on the Phanerozoic side of the TOR profile.

Summary Integrated analysis of the available geophysical, petrological and tectonic data for Europe reveals the major characteristics of its lithospheric structure and tectonic evolution. (1) Precambrian areas of Europe have a thick lithosphere, typically 150-220 km. Lithospheric thickness in the mid- and late Proterozoic provinces of the Baltic Shield is c. 120-180km. There is no obvious correlation between lithospheric thickness and the geological age of the crust (i.e. the absolute age of the oldest rocks determined from R e - O s isotope data) or the tectonic age (i.e. the age of the last major thermo-tectonic event) as proposed earlier (e.g. Poudjom Djomani et al. 1999).

DEEP EUROPE TODAY

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Fig. 10. Compositional anomalies in the lithosphere of Europe, represented by anomalies of Vp/V~ ratio at depths of 150 km (a) and 250 km (b) calculated from smoothed and filtered P-wave tomography model by Bijwaard & Spakman (2000) recalculated to absolute velocity by scaling by ak135 model values and S-wave tomography model by Shapiro & Ritzwoller (2002). Vp/V~ ratio is thought to be more sensitive to compositional than temperature variations (e.g. Lee 2003). Low lateral resolution for the northeastern parts of the maps is due to low resolution of the P-wave tomography model (compare with Fig. 2a and c).

Fig. 11. Density anomalies in the upper mantle of Eurasia of a non-thermal origin. (a) Mantle residual gravity anomalies (Fig. 5) corrected for temperature (Figs 2 g and 5). The resolution of this map is limited to approximately 3 ~ x 3 ~ which corresponds to a homogeneous resolution of thermal data in the study area. Conservative estimates of possible uncertainties of the residual anomalies are up to 75-100 mGal (Kaban et al. 2003). Amplitudes of the residual compositional anomaly significantly exceed this level (c. 600 mGal). (b) Density deficit in the subcrustal lithosphere calculated on a 5 ~ x 5 ~ grid from buoyancy (using data on the topography, crustal structure, lithospheric thickness and mantle temperatures) (from Artemieva 2003). A low-density anomaly over the Caledonides may result from a non-accounted dynamic topography. The general agreement of the zero-contour of gravity anomalies (a) and 0.8% contour of density anomalies from buoyancy (h) should be noted. The maps suggest a high degree of density deficit of a non-thermal origin in the northern parts of the EEP and the Baltic Shield. This anomaly can probably be associated with an Iron depletion of the cratonic lithospheric root. The pronounced difference in the gravity field from high (west) to low (east) across the TESZ correlates with the change in Vp/V~ (Fig. 10) from low values in Western Europe to high values in the Precambrian part. In contrast, the high densities in Southern Europe (corresponding to the subduction systems in the Eastern Mediterranean Sea) correspond to very high Vp/V~ ratios.

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An exceptionally thick lithospheric root is revealed by seismic, thermal and xenolith data for the Karelian part of the Baltic Shield, where it locally reaches a depth of c. 250-300 km. The region of thick lithosphere correlates spatially with the region of locally thick crust ( > 60 km) that formed during the Proterozoic orogenic event (Korja et al. 1993). Mantle xenoliths from the same region (at the edge of the Archaean terrane), brought to the surface by late Proterozoic (c. 600 Ma) kimberlite magmatism from depths as great as 240 kin, sample a fluid-free mantle; this conclusion is supported by regional MT data (Kukkonen et al. 2003). Gravity and buoyancy constraints on mantle density (Artemieva 2003; Kaban et al. 2003) reveal a strong density anomaly of compositional origin in this part of the Baltic Shield, which can be attributed to a highly depleted lithosphere. We speculate that local thickening of the lithosphere could have been produced during the same tectonic (orogenic) event as the formation of the crustal root; and that the depleted and devolatized composition of a thick cratonic root prevented its later destruction by mantle convection (Ballard & Pollack 1987). Thus, the lithospheric structure of the Karelian province may preserve evidence of tectonic processes that operated during the Proterozoic. Moreover, dipping and subhorizontal seismic reflectors at depths of 4 0 110 km at the margins of the Svecofennian and Sveconorwegian provinces, which are traced over distances of up to 100 km and correlate with a 5 - 7 km step on the Moho, are interpreted as evidence for Proterozoic subduction. Geophysical data reveal that the lithospheric structure of the Ukrainian Shield, which was formed by amalgamation of several Archaean-early Proterozoic terranes, is highly heterogeneous and different from that of the Baltic Shield: crustal thickness varies from 38 to 58 km and lithospheric thickness is in the range of 170-220 km. Similar values of lithospheric thickness are also typical for the north-central parts of the EEP. Southern parts of the EEP, affected by Palaeozoic rifting, have thin lithosphere (100-150 km), and it is likely that the cratonic lithospheric root has been thermally eroded (and/or delaminated) and metasomatized during the Devonian rifting. Seismic interpretations of refraction profiles (i.e. the PNE profile Quartz and FENNOLORA) and regional tomography models (i.e. SVEKALAPKO) suggest the existence of a layer at depths of 100-150 km with 1-2% lower seismic velocities than in the surrounding high-velocity cratonic upper mantle. It is important to note that seismic velocities in this reduced velocity layer within the cratonic lithosphere are c. 1% higher than average seismic velocities in the global continental models ak135 or iaspei. The nature of the reduced velocities is debated. Alternative models suggest high subsolidus temperatures (with a possible presence of small pockets of a partially molten material), the presence of fluids or compositional anomalies (i.e. a transition from a depleted upper layer to a non-depleted lower layer within the lithospheric root). (2) The Palaeozoic Variscan and Caledonian orogens of Western Europe were significantly reworked and overprinted by late Palaeozoic and Mesozoic-Cenozoic tectonic processes associated with the convergence of the Eurasian and the African plates (Ziegler & Drzes 2006). They have a uniform crustal thickness (typically 2 8 - 3 2 km), and the lithospheric thickness is in the range of 8 0 - 1 4 0 km, with the larger values beneath the Proterozoic-early Palaeozoic terranes (the Armorican, Bohemian and Brabant massifs, and the northern part of the Massif Central). The subcrustal lithosphere has a subhorizontal layering in the upper 90 km, revealed by seismic refraction studies and mantle xenolith data. Zones of strong seismic anisotropy in the upper mantle of the Variscides are interpreted as relict subduction zones. Compared with the Palaeozoic orogens of Western Europe, the Uralides, which remained intact within the continental interior and have not been reworked by later tectonic processes, have an

atypical structure of the crust (50-55 km thick with local roots reaching c. 65 km) and of the lithosphere (probably 170200 km thick). Palaeozoic rifts within the Precambrian part of Europe (the Oslo rift and the Pripyat-Dnieper-Donets rift) have lithospheric thickness, similar to the Variscan belt, of 100140 km. (3) The lithospheric structure of tectonically active parts of Western Europe is highly heterogeneous. Several Cenozoic orogens formed during closure of Tethyan ocean domains and subsequent continental subduction (the Alps, Carpathians, Caucasus, Apennines), followed by the development of back-arc basins (e.g. the Tyrrhenian, Aegean and Pannonian depressions). Crustal thickness in these orogens locally reaches 6 0 - 6 5 km in the convergence zone of lithospheric plates, where lithospheric thickness can exceed 150-200 km. The back-arc basins have thin crust (2530 km) and thin lithosphere (60-80 km). In the Central European Rift System, lithospheric thickness is similar to that of the adjacent Palaeozoic Variscan structures (80-120 km), although in some parts it can be as thin as 7 0 80 km. Available geophysical data do not provide distinctive evidence for a plume-related origin of the CERS. Instead, they suggest a passive mechanism of rifting, so that most of tectonomagmatic activity within the CERS was caused by a complicated stress regime associated with the convergence of the Eurasian and the African lithospheric plates. The authors are grateful to W. Spakman, C. Piromallo, N. Shapiro, M. Ritzwoller, G. Panza, Zhijun Du, M. Granet, M. Billien, J. Trampert, T. Yegorova and V. Starostenko for kindly providing their seismic tomography and gravity models and for permitting us to use them in this review. Special thanks are due to M. Cara, J. Ritsema, J. Trampert and E. Debayle for valuable discussions on the resolution of seismic tomography models. We are grateful to J. Ansorge, P. Ziegler and K. Fuchs for thoughtful, helpful and constructive reviews. The manuscript benefited from their valuable suggestions; the text, however, reflects the point of view of the authors, which does not always agree with that of the reviewers. The comments of M. Coble are appreciated. The research of I.M.A. is funded by a personal grant from Carlsbergfondet, Denmark, which is gratefully acknowledged. Economical support from the Danish Natural Science Research Council and the Carlsberg Foundation to H.T. is acknowledged.

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Crustal evolution of Western and Central Europe P. A. ZIEGLER & P. Dl~ZES Geological-Palaeontological Institute, Department o f Geosciences, University o f Basel, Bernoullistrasse 32, 4056 Basel, Switzerland (e-mail: [email protected])

Abstract: A new Moho depth map has been assembled for Western and Central Europe and the Western Mediterranean area that is exclusively based on published regional Moho depth maps. Tectonic overlays summarize Caledonian and Variscan tectonic units, Permo-Carboniferous fault systems and magmatic provinces, Mesozoic and Cenozoic rift-wrench systems, areas of intraplate compression, the outlines of Alpine orogens and the distribution of oceanic crust. Based on a comparison of these overlays with the Moho depth map we assess processes that controlled the evolution of the crust in the various parts of Europe through time. The presentday crustal configuration of Western and Central Europe results from polyphase Late Palaeozoic to recent lithospheric deformation that overprinted the margin of the Proterozoic East European Craton and particularly the Caledonian and Variscan crustal domains. Following consolidation of the Caledonides, their crustal roots were destroyed in conjunction with Devonian wrench tectonics and back-arc rifting. During the Permo-Carboniferous tectonomagmatic cycle, wrench faulting disrupted the crust of the Variscan Orogen and its foreland and the lithosphere of these regions was thermally destabilized. Late Permian and Mesozoic re-equilibration of the lithosphere-asthenosphere system was interrupted by the development of the Arctic-North Atlantic, Tethyan and associated rift systems. During the Alpine orogenic cycle, intraplate compressional stresses controlled basin inversion-related crustal thickening and lithospheric folding, as well as the evolution of the Rhine-Rhrne rift system. Variably deep crustal roots characterize the Alpine orogenic chains. Neogene back-arc extension disrupted the eastern Pyrenees, Betic-Balearic, Apennine and Dinarides orogens.

The depth map of the Moho discontinuity presented in Figure 1 was constructed by digitally scanning, scaling and assembling published regional Moho maps (see Bibliography and DOzes & Ziegler 2002) and by redrawing the depth contours as vectorized polygons. Most of the maps used for this compilation became available after the publication of earlier Moho compilation maps covering large parts of Europe by Meissner et al. (1987), Ziegler (1990) and Ansorge et al. (1992). The map resulting from our efforts (Fig. 1) gives the depth of the Moho discontinuity for Western and Central Europe and adjacent oceanic domains, but not the true thickness of the crust, as no corrections were applied for the thickness of its sedimentary cover nor for surface topography in elevated areas or for water depths in offshore areas. The objective of this compilation was to obtain an impression of the present-day crustal configuration of all of Western and Central Europe, including the Western Mediterranean area, and to develop a basis for the analysis of processes that through time contributed to the evolution of the crust in the various parts of Europe. To this end, a set of overlays was constructed, summarizing the main tectonic elements of the Caledonian and Variscan orogens (Fig. 2), the Stephanian-Early Permian fault systems and magmatic provinces (Fig. 3), the Mesozoic rift and wrench systems (Fig. 4), areas of intraplate compression, the outlines of the Alpine orogens, and Cenozoic rift and wrench systems (Fig. 5). In Figures 2 - 5 , which also show the present-day distribution of oceanic crust in the Atlantic and Mediterranean domains, these overlays are reproduced together with the Moho depth contour map. A comparison of Figures 1 and 2 clearly shows that the stable parts of the Proterozoic Fennoscandian-East European Craton are characterized by Moho depths as great as 48 km (within the frame of our map) whereas in more mobile Phanerozoic Europe Moho depths vary between 24 and 38 km and no longer bear any relation to the Caledonian and Variscan orogens. On the other hand, Moho depths of 2 0 - 2 6 km characterize the Proterozoic Hebridean craton and reflect strong modification of this crustal domain during Mesozoic rifting cycles. In contrast, the Alpine chains, such as the Western and Central Alps, the Carpathians, Apennines and Dinarides, as well as the Betic Cordillera and the Pyrenees are characterized by more or less distinct crustal roots reaching depths as great as 60 km. Inferring an Alpine crustal model (Stampfli et al. 1998) for the continent-continent collisional Caledonian and Variscan orogens, the present crustal configuration of extra-Alpine Phanerozoic Europe implies

post-orogenic destruction of their crustal roots and that their crust was repeatedly modified during Mesozoic and Cenozoic phases of rifting and intraplate compression. In this context it should be kept in mind that the present depth of the Moho discontinuity is controlled not only by the thickness and composition of the crust but also by the thickness of its sedimentary cover and, in offshore areas, by water depths. So far, we have not yet been able to construct a regional thickness map of the crystalline continental crust for the extra-Alpine domains, because in areas north of the Varsican deformation front the thickness of prePermian Palaeozoic sediments is still poorly constrained.

Processes controlling depth of the crust-mantle boundary During orogenic processes, the Moho discontinuity can be depressed to depths of 6 0 - 7 5 km in conjunction with subduction of continental crust, its imbrication and the stacking of basementcored nappes, as is evident, for example, in the Western and Central Alps and the Pyrenees (Roure et al. 1996; Schmid et al. 1996, 2004; Waldhauser et al. 1998). During nappe emplacement, the foreland lithosphere is deflected in response to its thrust- and slab-loading, accounting for the development of foreland basins and a corresponding depression of the crust-mantle boundary (Ziegler et al. 2002). Depending on convergence rates and the thermal state of the foreland lithosphere, underthrust continental crust is eclogitized at depths of 5 5 - 7 5 km and assumes densities in the range of p = 3.06-3.56, comparable with those of the average mantle (p = 3.35; Bousquet et al. 1997; Henry et al. 1997). In the process of this, the P-wave velocity of crustal material increases, depending on its composition, to 8.0-8.4 km s -1, and thus is transferred across the geophysically defined Moho discontinuity (break-over from Vp _< 7.8 to 8.0-8.2 km s -1) into the lithospheric mantle (Ziegler et al. 1998). This limits depth of orogenic crustal roots as defined by seismic velocities. In this respect, it should be kept in mind that the petrological and seismic crust-mantle boundary does not always coincide. Synorogenic thickening of the crest, involving subduction of continental lithospheric mantle and lower crustal material and stacking of upper crustal nappes (see, e.g. DOzes et al. 2004; Schmid et al. 2004), is accompanied by widespread metamorphism of crustal and sedimentary rocks, their metasomatic reactivation, and plutonic

From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 43-56. 0435-4052/06/$15.00 © The Geological Society of London 2006.

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52

54

56

58 Ki~on"~tres

ProjeCtion: Lambert Azimuthal Equal Area; Centre: 04~.00"/48"L00"; Region : W/E/N/S = 350°/28°/62°I34~'; Ellipsoide wgs-84

activity, as seen in the Variscan and Caledonian orogens. During post-orogenic times, erosional and tectonic unroofing of an orogen, often to former mid-crustal levels, results in the exposure of a newly formed crystalline basement complex. The post-orogenic re-equilibration of orogenically destabilized lithosphere with the asthenosphere can involve such processes as detachment of subducted lithospheric slabs and passive upwelling of the asthenosphere, thermal thinning and/or partial delamination of the mantle-lithosphere, interaction of mantle-derived melts with a felsic lower crust, and possibly also retrograde metamorphism of eclogitized crustal roots. All these processes contribute towards uplift and erosional unroofing of an orogen, and a corresponding shallowing of the crust-mantle boundary. Moreover, the postorogenic collapse of an orogen can be accelerated by its wrenchand/or extension-induced tectonic unroofing in response to a reorientation of the regional stress field (e.g. Devonian collapse of the Arctic-North Atlantic Caledonides, Permo-Carboniferous collapse of the Variscides). In such conditions, the crust of an

Fig. 1. Crustal thicknessmap of Western and Central Europe (after D~zes & Ziegler 2002). A full list of referencesis given in the Bibliography.

orogen can be thinned to 3 0 - 3 5 km within some 10-20 Ma after crustal shortening has terminated (e.g. Variscides). Furthermore, depending on the degree of post-orogenic thinning of the lithospheric mantle and the thickness of the crust, long-term thermal subsidence of the lithosphere of former orogens can account for the subsidence of intracratonic basins, involving a gradual depression of the crust-mantle boundary (Ziegler et al. 2004, 2006). During rifting and wrench faulting the continental crust is thinned by mechanical stretching (McKenzie 1978). The magnitude and mode of lithospheric stretching (pure or simple shear), the depth of the lithospheric necking level, whether or not magmatic processes contributed toward crustal thinning or crustal thickening by underplating, and the degree of thermal thinning of the lithospheric mantle, control the magnitude of synrift subsidence of extensional basins, the amount of uplift of their rift flanks and the position of the crust-mantle boundary. Similarly, these processes have a bearing on the magnitude of post-rift

CRUSTAL EVOLUTIONOF EUROPE subsidence of extensional basins and, thus, on the ultimate depth of the Moho discontinuity at the end of the post-rift re-equilibration of the lithosphere-asthenosphere system (Ziegler & Cloetingh 2004). During phases of collision-related intraplate compression, inversion of extensional basins and upthrusting of Rocky Mountain-type arrays of basement blocks can lead to crustal thickening and depression of the Moho discontinuity (Ziegler et al. 1998, 2002). Moreover, intraplate compressional stresses can cause large-scale folding of the lithosphere and, depending on its polarity, depression or uplift of the crust-mantle boundary (Cloetingh et al. 1999).

Caledonian crustal domain The Caledonides of the British Isles and Scandinavia are thought to have collapsed shortly after their earliest Devonian consolidation in response to orogen-parallel extension, reflecting the activation of the Arctic-North Atlantic megashear (Ziegler 1989; Braathen et al. 2002). Wrench faulting and rifting controlled the Devonian development of the Orcadian pull-apart basin and uplift of core complexes, and the subsidence of the Midland Valley and Dublin-Northumberland grabens, respectively. The Dublin-Northumberland graben is superimposed on the Iapetus suture (Fig. 2). During Late Devonian to Namurian times (370-315 Ma), the Midland Valley and Dublin-Northumberland grabens were sites of crustal extension. However, crustal thinning related to the development of these basins, which were partly inverted during the Late Westphalian (310-305 Ma), is not clearly reflected by the present-day depth of the Moho discontinuity (Ziegler 1989, 1990). The Mid-European Caledonides, which are exposed in the Ardennes and mark the Rheic suture between the Gondwanaderived East Avalonia and the composite ArmoricanSaxo-Thuringian terranes (Armorican Terrane Assembly; Pharaoh 1999; Winchester & the PACE TMR Network Team 2002; Verniers et al. 2002), were disrupted during the Early Devonian by back-arc extension, controlling the opening of the Rheno-Hercynian Basin and limited sea-floor spreading in its Lizard and Giessen-Harz sub-basins. In contrast, there is only limited evidence for Devonian tensional reactivation of the North German-Polish Caledonides, which are associated with the SW- to south-dipping Thor-Tornquist suture along which East Avalonia was welded to the East European Craton (Banka et al. 2002; Krawczyk et al. 2002). Following opening of the oceanic Lizard-Giessen-Harz Basin, its northern passive margin was transgressed and developed into the broad Rheno-Hercynian Shelf that was dominated by carbonate platforms, particularly during the Mid-Devonian and Early Carboniferous (Ziegler 1990). This reflects rapid degradation of the North German-Polish Caledonides and of the central North Sea area where they merge into the Scottish-Norwegian Caledonides. Midand Late Devonian development of a shallow sea arm, which extended from the Rheno-Hercynian Shelf into the central North Sea, may reflect mild tensional reactivation of the northwestern segment of the Thor-Tornquist suture (Ziegler 1990; Williamson et al. 2002). By Early Carboniferous times, the thickness of the continental crystalline crust underlying the Rheno-Hercynian Shelf may have been as great as 35 km in its northern parts and some 38 km in the area of the London-Brabant Massif, tapering to zero along the northern margins of the oceanic Lizard and GiessenHarz basins.

Variscan crustal domain The Variscan Orogen is delimited to the north by its external Rheno-Hercynian thrust belt (Fig. 2). This thrust belt evolved by imbrication of the crust and sedimentary cover of the Rheno-Hercynian Shelf, which, following Early Carboniferous closure of the oceanic Lizard-Giessen-Harz Basin was converted

45

into a flexural foreland basin that subsided in response to thrustand slab-loading (Oncken et al. 1999, 2000). The internal parts of the Variscan Orogen include a number of Gondwana-derived continental terranes, such as the Armorican, Saxo-Thuringian, Bohemian, Moldanubian and Aquitaine-Cantabrian blocks (Ziegler 1989; Pharaoh 1999; Franke et al. 2000). Sutures marking the location of subducted oceanic basins delimit these terranes. During the Late Devonian and Carboniferous main phases of the Variscan orogeny, major crustal shortening and subduction of continental lithospheric material was accompanied by widespread high-pressure metamorphism and associated magmatism (Franke et al. 2000; Ziegler et al. 2004). Intraplate compressional stresses, which were exerted on the Variscan foreland, are considered to have caused the Mid- to Late Carboniferous tensional reactivation of the Arctic-North Atlantic megashear and the onset of rifting in the NorwegianGreenland Sea area, as well as the Westphalian partial inversion of Carboniferous rifts on the British Isles (Ziegler 1989, 1990; Ziegler et al. 2002). By the end of Westphalian times (305 Ma), when crustal shortening in the Variscan orogen ended, its internal parts were probably characterized by a thermally destabilized lithosphere and 45-60 km deep crustal roots. Deep-reaching subducted lithospheric slabs were probably still attached to the Variscan lithosphere at the Bohemian-Moldanubian suture and the suture between the Armorican and Aquitaine-Cantabrian terranes, the latter being associated with the nappe systems on the Arverno-Vosgian and Ligerian zones (Fig. 2). On the other hand, the Rheno-Hercynian zone was underlain by a thermally stabilized foreland lithosphere that extended as a subduction slab some 200 km beneath the RhenoHercynian-Saxo-Thuringian suture. The oceanic parts of this slab, corresponding to the Lizard and Giessen-Harz basins, had already been detached from the foreland lithosphere during the Early Carboniferous (Ziegler et al. 2004, 2006).

Stephanian-Early Permian tectonomagmatic cycle At the end of the Westphalian (305 Ma), oblique collision of Gondwana and Laurussia gave way to their dextral translation. Stephanian-Early Permian (305-269 Ma) continued crustal shortening in the Appalachian and Scythian orogens was paralleled by the wrench-induced collapse of the Variscan Orogen. Continental-scale dextral shears, such as the Tornquist-Teisseyre, Bay of Biscay, Gibraltar-Minas and Agadir fractures zones, were linked by secondary sinistral and dextral shear systems (Fig. 3). Together, these overprinted and partly disrupted the Variscan Orogen and its northern foreland (Arthaud & Matte 1977; Ziegler 1989, 1990; Coward 1993; Ziegler & Stampfli 2001). Significantly, wrench tectonics, both of a transtensional and a transpressional nature, as well as associated magmatic activity, abated in the Variscan domain and its foreland during the late Early Permian (285-269 Ma), in tandem with the consolidation of the Appalachian Orogen (Ziegler 1989, 1990; Marx et aL 1995; Ziegler et al. 2004). Stephanian-Early Permian wrench-induced disruption of the rheologically weak Variscan Orogen and of its rheologically much stronger northern foreland was accompanied by regional uplift, widespread extrusive and intrusive magmatic activity, peaking during the Early Permian, and the subsidence of a multidirectional array of transtensional trapdoor and pull-apart basins in which continental clastic deposits accumulated (Fig. 3). Basins developing during this time span show a complex, polyphase structural evolution, including a late phase of transpressional deformation controlling their partial inversion (Ziegler 1990). Although Stephanian-Early Permian wrench deformation locally gave rise to uplift of extensional core complexes (Vanderhaeghe & Teyssier 2001), crustal stretching factors were, on a regional scale, relatively low, as seen in the Southern Permian Basin, which is

46

P.A. ZIEGLER & P. DEZES

17Z[[ZZ-]-7 L LZ_LJ

Lower Carboniferous rifts

Palaeozoic Suture

Variscan deformation front

Alpine deformation front

Caledonian deformation front

Oceanic basins

located in the Variscan foreland and encroaches in its eastern parts on the Rheno-Hercynian thrust belt (Ziegler 1990; van Wees et al. 2000). Stephanian-Early Permian wrench deformation of the Western and Central European lithosphere apparently caused detachment of the subducted Variscan lithospheric slabs and a general reorganization of the mantle convection system, involving the activation of a system of not very active mantle plumes. Upwelling of the asthenosphere induced partial delamination and thermal thinning of the mantle-lithosphere and magmatic inflation of the remnant lithosphere. This was accompanied by the interaction of mantle-derived partial melts with the felsic lower crust. These processes accounted for regional uplift and the destruction of the Variscan orogenic roots. By the end of Early Permian times, the Variscan crust was thinned to 28-35 km on a regional scale, mainly by magmatic processes and its erosional unroofing and

Fig. 2. Caledonian and Variscan structural elements superimposed on the crustal thickness map.

only locally by its mechanical stretching. Quantitative subsidence curves, derived from the Late Permian and Mesozoic record of intracratonic sedimentary basins, and their modelling suggest that, at the end of the Early Permian, the thickness of the remnant lithospheric mantle ranged between 10 and 50 km in the area of the Southern Permian and Paris basins, the Hessian Depression and the Franconian Platform (Ziegler et al. 2004, 2006). Development of a major magmatic province in northern Germany and Poland, which extends to the SW into the area of the Saar-Nahe Basin, may be a direct consequence of detachment of the Rheno-Hercynian slab. Related thermal thinning of the lithospheric mantle and magmatic thinning of the crust provided the driving mechanism for the Late Permian and Mesozoic subsidence of the Southern Permian Basin (van Wees et al. 2000). In NE Germany, the thickness of the

CRUSTAL EVOLUTIONOF EUROPE

47

6"0'

48 °

44 °

40 °

36"

[~iii~{

Alpine deformation front

~

Volcanics

~__~2-_~ Variscan deformation front

Wrench-induced sedimentary basins

~--~

Oceanic crust

~..j

Fault system Dykes

~

Sills

crystalline crust decreases from 32 km beneath the northern flank of the Southern Permian Basin to 22 km under its axial parts (Bayer et al. 1999). As this part of this basin was only mildly affected by Early Devonian rifting during the opening of the Rheno-Hercynian back-arc basin (Krawczyk et al. 2002), and is neither underlain by major extensional PermoCarboniferous basins nor overprinted by Mesozoic rifting, the observed crustal thinning has to be attributed to PermoCarboniferous magmatic destabilization of the crust-mantle boundary (Ziegler et al. 2004). Similarly, crustal thinning across the Oslo Graben, along the North Danish Basin (Sorgenfrei Line) and probably also along the Polish Trough must be largely attributed to PermoCarboniferous tectonic and magmatic processes related to the

Fig. 3. Permo-Carboniferous fault systems and magmatic fields superimposed on the crustal thickness map.

activation of the Tornquist-Teisseyre-Sorgenfrei Line. In Poland this line reflects wrench-induced reactivation of the suture between the Polish Caledonides and the Proterozoic East European Craton. However, to the NW, wrench faulting propagated along the Sorgenfrei Line through the Neoproterozoic Dalslandian crust of southern Scandinavia and terminated in the highly volcanic pull-apart Oslo-Skagerrak Graben (Ziegler 1990; Banka et al. 2002). Across this graben, upper crustal extension by faulting amounts to about 10-20 km, whereas its crustal configuration suggests 40-60 km of extension. This indicates that interaction of mantle-derived melts with the lower crust caused destabilization of the Moho, and thus contributed significantly to the observed crustal thinning (Ro & Faleide 1992; Ziegler & Cloetingh 2004).

48

P.A. ZIEGLER & P. DI~ZES

[,i. ........... L______J

l

...........

J

Mesozoic rifts & wrench faults Alpine deformation front Oceanic basins

The occurrence of extensive Permo-Carboniferous dyke swarms and sills in Scotland, and of Stephanian-Early Permian basins and volcanic rocks in the Irish Sea area and on the Western Shelves (Heeremans et al. 2004), shows that the British Isles were also destabilized during the PermoCarboniferous tectonomagmatic cycle, possibly contributing to thinning of their crust. By analogy with the Southern Permian Basin, subsidence of the Northern Permian Basin, which occupies much of the central North Sea and extends into northern Denmark, probably corresponds also to a zone of thermally driven PermoCarboniferous thinning of the lithospheric mantle and crust (Ziegler 1990). Of special interest is the major magnetic anomaly that transects the Paris Basin along the trace of the Seine-Loire wrench-fault system in the prolongation of the Sillion Houillier shear zone of

Fig. 4. Mesozoic rift and wrench fault systems superimposed on the crustal thickness map.

the Massif Central (Cavellier et al. 1980; Banka et al. 2002). This anomaly probably reflects emplacement of PermoCarboniferous basic magmas in the lower crust of the stable Cadomian Armorican block. However, in the area of the Sillion Houillier, which transects the Variscan nappe systems and associated intrusive bodies of the Arverno-Vosgian and Ligerian zones (Ledru et al. 2001), this anomaly is not evident. In contrast, the Sillion Houillier part of this deep crustal wrench zone is characterized by distinct crustal thinning that must be attributed to PermoCarboniferous magmatic destabilization of the Moho, as this zone was not overprinted by Mesozoic or Cenozoic rifting. Similarly, the NW-trending axis of crustal thinning that underlies the NE flank of the Aquitaine Basin and the Limousin must be attributed to Permo-Carboniferous destabilization of the crust-mantle boundary.

CRUSTAL EVOLUTION OF EUROPE

8oo

49

.60'

se°

. 5 0 ~'

° 52 ~

48

°.

44

°,

40

°,

36

~,

.4B

°

.44, °

.40

°

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co

~:~::2:_;.~:! F

!

Alpine deformation front

[2~-~]

Volcanics in sub-surface

Inverted grabens

Areas of Late Cret.-Tert. inversion

Rift system

Oceanic crust

We conclude that the Permo-Carboniferous tectonomagmatic pulse had a major impact on the crustal configuration of Western and Central Europe. Late Early Permian to Early Cretaceous thermal sag basins and rifts In the Variscan domain and its northern foreland, magmatic activity gradually abated during the late Early Permian (285-269 Ma) and decay of thermal anomalies, introduced during the StephanianEarly Permian tectonomagmatic cycle, commenced. This reflects that after the Permo-Carboniferous thermal surge (300-280 Ma) the temperature of the asthenosphere had retumed to ambient levels (1300 °C). During the Late Permian and Mesozoic, a new, ever-expanding system of intracratonic thermal sag basins

Fig. 5. Oligocene and younger rift and wrench systems and magmatism, superimposed on the crustal thickness map, showing areas of latest Cretaceous and Cenozoic compressional intraplate deformation.

developed, nucleating from the Southern and Northern Permian basins (Ziegler 1990; Ziegler et al. 2004). During the development of this basin system, the formerly elevated crust-mantle boundary was gradually depressed. For instance, in the axial parts of the Southern Permian Basin, the crust subsided during Late Permian to Cenozoic times by as much as 8 km in response to cooling and sedimentary loading of the lithosphere (Scheck & Bayer 1999). In large parts of Western and Central Europe, post-Early Permian thermal subsidence of the lithosphere was, however, overprinted and partly interrupted by the Late Permian-Early Triassic onset of a new rifting cycle, which preceded and accompanied the step-wise break-up of Pangaea. Major elements of this break-up system were the southward propagating Arctic-North Atlantic and the westward propagating Neotethys rift systems (Ziegler & Stampfli 2001). During the Triassic, a multi-directional rift system developed in Western and Central Europe, major

50

P.A. ZIEGLER& P. DI~ZES

constituents of which are the North Sea rift, the North DanishPolish Trough, the graben systems of the Atlantic shelves and the Bay of Biscay rift. Development of these grabens, partly involving tensional reactivation of Permo-Carboniferous fracture systems, persisted during the Jurassic and Early Cretaceous and was in some rifts accompanied by major crustal extension and commensurate thinning of the crust (Fig. 4; Ziegler et al. 2001). For instance, across the northern and central parts of the North Sea rift, upper crustal extension by faulting amounted to 3040 km (Ziegler 1990; Ziegler & Cloetingh 2004). The Norwegian-Greenland Sea rift propagated southward during the Late Permian into the northwestern shelves of the British Isles, and during the Triassic into the Central Atlantic domain, and the Neotethys rift systems propagated westward into the Bay of Biscay and NW Africa and linked up with the Atlantic rift system in the North Atlantic domain (Ziegler 1988, 1990; Ziegler & Stampfli 2001). This was accompanied by activation of the Central Iberian rift (Salas et al. 2001). During the Early Triassic, the North Sea rift, consisting of the Horda half-graben and the Viking, Murray Firth, Central and Horn grabens, was activated and transected the western parts of the Northern and Southern Permian basins whereas the North Danish-Polish Trough transected their eastern parts. Simultaneously, the rift systems of the Alpine domain, the Bay of Biscay and the Western Shelves were activated. The latter included the Porcupine, Celtic Sea and Western Approaches troughs. Crustal extension across the Celtic Sea and Western Approaches troughs was compensated, at their eastern termination, by reactivation of Permo-Carboniferous shear systems controlling the subsidence of the Channel and Wessex basins and intermittent destabilization of the Paris thermal sag basin (Ziegler 1990; Boldy 1995). Following late Early Jurassic crustal separation in the Central Atlantic (190-180 Ma) and Mid-Jurassic (177-160 Ma) crustal separation in the Alpine Tethys, the evolution of the Western and Central European rifts was dominated by northward propagation of the Atlantic rift system (Ziegler 1988; Ziegler et al. 2001; Stampfli & Borel 2004). During the Late Jurassic and earliest Cretaceous, accelerated rifting activity is evident in the Western Approaches, Celtic Sea and Porcupine, RockallFaeroe troughs and the Bay of Biscay. At the same time, rifting accelerated in the North Sea, focusing on its axial Viking and Central grabens. This was accompanied by the development of sinistral shear systems at the southern termination of the North Sea rift, controlling the subsidence of the transtensional Sole Pit, Broad Fourteens, West Netherlands, Lower Saxony, Sub-Hercynian and Altmark-Brandenburg basins. In contrast, crustal extension across the North Danish-Polish Trough apparently waned at the Jurassic-Cretaceous transition (Ziegler 1990; Kutek 2001). In the North Atlantic, crustal separation progressed gradually northwards during the Late Jurassic and Early Cretaceous and by Mid-Aptian times (±110 Ma) crustal separation was achieved in the Bay of Biscay (Ziegler et al. 2001). With this, the grabens on the Western Shelves became inactive and began to subside thermally. Following the Late Jurassic-Early Cretaceous rifting pulse, tectonic activity gradually abated also in the North Sea rift system and crustal extension focused on the zone of future crustal separation between Europe and Greenland (Ziegler 1988, 1990; Osmundsen et al. 2002). Post-rift thermal subsidence of the North Sea Basin then began.

Late Cretaceous and Palaeocene rifting and intraplate compression During the Late Cretaceous and Palaeocene, rifting activity was centred on the Rockall-Faeroe Trough and the area between the

Rockall-Hatton-Faeroe Bank and Greenland. During the Cenomanian-Santonian (98-84 Ma), limited sea-floor spreading may have occurred in the southern parts of the Rockall Trough. During the Campanian-Maastrichtian (84-65 Ma), the Iceland plume impinged on the North Atlantic-Greenland Sea rift system, giving rise to the Palaeocene (65-55 Ma) development of the Thulean flood basalt province, which had a radius of more than 1000kin (Morton & Parson 1988; Ziegler 1988; Larsen et al. 1999). At the same time, the northern parts of the British Isles and the Rockall-Hatton-Faeroe Bank were thermally uplifted and subjected to erosional unroofing. Mantlederived melts, underplating and intruding the crust, probably contributed to crustal thinning in the area of the Hebrides Shelf and the Rockall-Hatton-Faeroe Bank. At the Palaeocene-Eocene transition (55 Ma), crustal separation was achieved between Greenland and Europe to the west of the Rockall-HattonFaeroe Bank and in the Norwegian-Greenland Sea (Mosar et al. 2002). Volcanic activity then terminated on the conjugate margins and became centred on the evolving sea-floor spreading axes and on Iceland (Ziegler 1988, 1990). During the Turonian-Santonian (93.5-83.5 Ma), Africa began to converge with Europe in a counter-clockwise rotational mode (Rosenbaum et al. 2002). Resulting space constraints within the Tethyan belt caused activation of new subduction zones that controlled the gradual closure of the Alpine Tethys and the Bay of Biscay (Stampfli et al. 2001). Commencing in the late Turonian (_+90Ma), compressional stresses were exerted on the northern Tethyan shelves of the Eastern Alps and Carpathians, inducing inversion of Mesozoic tensional basins and upthrusting of basement blocks by reactivation of pre-existing crustal discontinuities. The Senonian pulse of intraplate compression (89-70Ma), which affected the North Danish-Polish Trough, the Bohemian Massif, the Brandenburg-Altmark, Sub-Hercynian, Lower Saxony, West Netherlands and Sole Pit basins, as well as the southern parts of the North Sea rift, can be related to compressional stresses that were projected from the Alpine-Carpathian orogenic wedge through the oceanic lithosphere of the Alpine Tethys into the lithosphere of Western and Central Europe. The more intense Palaeocene phase of intraplate compression (65-55Ma), which affected these areas, as well as the Tethyan shelves of the Western and Central Alps, the Paris Basin and the Channel area, probably marked the collision of the Alpine orogenic wedge with its East Alpine-Carpathian foreland and with the Brianqonnais terrane in the West and Central Alpine domain (Ziegler et al. 1998; D6zes et al. 2004). This phase of foreland compression, during which a Rocky Mountain-type array of basement blocks was upthrust in the Bohemian Massif and the Polish Trough was deeply inverted, involved also broad lithospheric folding and accelerated subsidence of the North Sea Basin. Strong inversion of the Polish Trough caused thickening of its crust and the development of an up to 50 km deep Moho keel that is offset to the NE with respect to the inversion axis as defined at supracrustal levels. Similarly, imbrication of the basement of the Bohemian Massif entailed crustal thickening (Fig. 5; Ziegler 1990; Ziegler et al. 1998, 2002; Bayer et al. 1999; Jensen et al. 2002). Convergence rates between the Africa-Arabian and European plates decreased sharply from as much as 20 mm a -a during the Late Cretaceous to practically zero during the late Maastrichtian and Palaeocene (70-55 Ma) (Rosenbaum et al. 2002). Presumably this resulted from strong collisional coupling of the Africa-Arabian and European plates across the Alpine-Mediterranean orogen. This gave rise to the Palaeocene pulses of intense intraplate compression, which affected not only Western and Central Europe but also the East European Craton and North Africa (Nikishin et al. 2001; Ziegler et al. 2001).

CRUSTAL EVOLUTIONOF EUROPE

Opening of North Atlantic, Cenozoic rifting and interaction of the Alpine Orogen with the European foreland With the early Eocene onset of sea-floor spreading between Greenland and Europe (55.9-53.3 Ma; Mosar et al. 2002), post-rift thermal subsidence of the Rockall-Hatton-Faeroe Bank, the Rockall-Faeroe Trough and the shelves of NW Ireland and Scotland commenced. However, during the late Eocene and Oligocene reorganization of sea-floor spreading axes in the NorwegianGreenland Sea, the Atlantic shelves of the British Isles were destabilized by minor wrench faulting in the prolongation of the Iceland ridge and the Charlie Gibbs fracture zone, causing the development of inversion structures and the subsidence of small pull-apart basins in the Irish Sea area (Fig. 5; Ziegler 1990; Boldreel & Andersen 1998; Mosar et al. 2002). Convergence rates between Africa and Europe gradually increased during the Eocene and Oligocene (55-23.8 Ma), but decreased again during the early Miocene (Rosenbaum et al. 2002). Thrust-loaded deflection of the foreland of the Western, Central and Eastern Alps, as well as of the Carpathians, commenced during the Eocene. Eocene to mid-Miocene emplacement of the East Alpine and Carpathian nappe systems was not accompanied by further intraplate compressional deformation of their foreland, thus reflecting mechanical decoupling of these orogens from their forelands. In contrast, late Eocene-early Oligocene and late Oligocene-early Miocene inversion pulses evident in the Celtic Sea, Western Approaches, Channel, Wessex, Paris, Sole Pit, Broad Fourteens and West Netherlands basins reflect transmission of compressional stresses from the evolving West and Central Alpine Orogen into its foreland and thus their mechanical coupling (Ziegler 1990; Ziegler et al. 2002; D6zes et al. 2004). The present deep crustal roots of the Central and Eastern Alps evolved in response to continued underthrusting of the foreland after detachment of the subducted oceanic Alpine-Tethys lithospheric slab towards the end of the Eocene (Schmid et al. 1996; Stampfli et al. 1998). In contrast, the subducted slab of the Western Alps remained attached to the lithosphere until early Pliocene times (D~zes et al. 2004; Schmid et al. 2004). Oligocene and later underthrusting and subduction of little attenuated foreland lithosphere, combined with the development of an upper plate mantle back-stop, accounted for increasing mechanical coupling of the West and Central Alpine orogenic wedge with its foreland at crustal and lithospheric mantle levels (Ziegler & Route 1996). Evolution of the Pyrenees commenced during the Campanian (80 Ma) and lasted until the early Miocene (+_20 Ma), involving northward subduction of the Iberian lithosphere under Europe and southward subduction of the oceanic crust of the Bay of Biscay under Iberia (Verg6z & Garcia-Senez 2001). During the Palaeocene and Eocene, foreland compression controlled the evolution of the Languedoc-Provenqal fold and thrust belt, and thrust-loaded subsidence of the Aquitaine and Ebro foreland basins commenced. During the late Eocene and Oligocene, the Ebro foreland basin became isolated in response to inversion of the Mesozoic Central Iberian and Catalan Coast Range rifted basins (Fig. 5; Salas et al. 2001; Ziegler et al. 2002). Development of the European Cenozoic rift system (ECRIS), which extends from the Dutch North Sea coast into the Western Mediterranean, commenced during the late Eocene (Fig. 5). Its southern elements are the Valencia Trough, the graben systems of the Gulf of Lions, and the north-striking Valence, Limagne and Bresse grabens; the latter two are superimposed on the Massif Central and its eastern flank, respectively. The Burgundy Transfer Zone links these grabens with the southern end of the north-striking Upper Rhine Graben. A further, although more diffuse transform fault system links the northern ends of the Limagne and Upper Rhine grabens and crosses the eastern parts

51

of the Paris Basin. Northward, the Upper Rhine Graben bifurcates into the NW-trending Roer Graben and the north-trending Hessian grabens, which transect the Rhenish Massif. The NE-striking Ohre (Eger) Graben, which cuts across the Bohemian Massif, forms an integral part of the ECRIS (Ziegler 1994; D6zes et al. 2004). Tensional reactivation of Permo-Carboniferous and Mesozoic shear systems played an important role in the localization of the ECRIS. The onshore parts of the ECRIS are associated with a distinct and broad shallowing of the crust-mantle boundary, which can be only partly attributed to Cenozoic rifting as upper crustal extension across the Upper Rhine Graben and the grabens of the Massif Central does not exceed 7 km (Fig. 5; D~zes et al. 2004). Evolution of the ECRIS was accompanied by the development of major volcanic centres in Iberia, on the Massif Central, the Rhenish Massif and the Bohemian Massif, particularly during Miocene and Plio-Pleistocene times (Wilson & Bianchini 1999). Mantle tomography reveals a system of upper asthenospheric lowvelocity anomalies beneath the ECRIS, interpreted as plume heads that have spread out above the 410 km discontinuity (Goes et al. 1999; Sibuet et al. 2004; Spakman & Wortel 2004). From these anomalies secondary, relatively weak plumes at present rise beneath the Eifel (Ritter et al. 2001) and Massif Central (Granet et al. 1995), but not beneath the Vosges-Black Forest arch (Achauer & Masson 2002). These upper asthenospheric anomalies presumably developed during the Palaeocene, following activation of the NE Atlantic and Iceland mantle plumes that rise from the core-mantle boundary (Hoernle et al. 1995; Bijwaard & Spakman 1999), and subsequently evolved further. This is compatible with volcanic activity in the ECRIS area that commenced during the Palaeocene and persisted into the Quaternary (D~zes et al. 2004). Because a shift with time in areas of major volcanic activity can be observed, it is likely that the supply of partial melts through secondary upper mantle plumes was not steady but pulsated and shifted in their location. In the ECRIS area, this plume activity caused thermal weakening of the lithosphere, thus rendering it prone to deformation, but was not the driving mechanism of tiffing. The ECRIS is generally considered to have evolved in response to passive rifting that was mainly controlled by compressional stresses originating in the Alpine and Pyrenean collision zones (D~zes et al. 2004). During the late Eocene, the Limagne, Valence, Bresse, Upper Rhine and Hessian grabens began to subside in response to northdirected compressional stresses that reflect collisional interaction of the Pyrenees and the Alps with their foreland (Merle & Michon 2001; Schumacher 2002; D6zes et al. 2004). During their Oligocene major extensional phase, these originally separated rifted basins coalesced and the Roer and Ohre grabens came into existence. During the late Oligocene, rifting propagated southward across the Pyrenean Orogen into the Gulf of Lions and along coastal Spain in response to back-arc extension that was controlled by eastward roll-back of the Alpine-Tethys subduction slab, which dipped beneath the Corsica-Sardinia-Balearic-Betic arc system. By late Aquitanian times (21.5 Ma), crustal separation was achieved in the Western Mediterranean Basin, the oceanic Provenqal-Ligurian Basin began to open, and the grabens of southern France and the Massif Central became inactive (S6ranne 1999; Roca 2001). In contrast, the Upper Rhine and Roer Valley grabens remained tectonically active until the present under a NW-directed compressional stress field that developed during the Miocene (D~zes et al. 2004). By the end of Oligocene time, the area of the triple junction of the Upper Rhine, Roer and Hessian grabens was uplifted and magmatic activity on the Rhenish Shield increased, probably accompanied by plume-induced thermal thinning of the mantle-lithosphere. By mid-Miocene times (+18 Ma) the Massif Central, the Vosges-Black Forest Arch and, slightly later, the Bohemian Massif were uplifted. This was accompanied by increased mantle-derived volcanic activity (Ziegler 1994; Merle & Michon 200l). At the Moho level, a

52

P.A. ZIEGLER & P. DEZES

broad anticlinal feature extends from the Massif Central via the Burgundy Transfer Zone and the Vosges-Black Forest into the Bohemian Massif. Development of this arch, which was paralleled by imbrication of the External Massifs of the Alps, can be attributed to folding of the lithosphere in response to the build-up of collisionrelated compressional stresses at mantle-lithospheric levels in the Alpine foreland (D~zes et al. 2004). This concept is compatible with the lack of lithospheric thinning beneath the Vosges-Black Forest Arch (Achauer & Masson 2002) and the lithospheric configuration of the Bohemian Massif (Babuska & Plomerova 2001). Uplift of this lithospheric fold entailed partial erosional isolation of the Paris Basin (Ziegler et al. 2002). Under the present NW-directed stress regime, which had intensified during the Pliocene, the Upper Rhine Graben is subjected to sinistral shear whereas the Roer Graben is extending nearly orthogonally (Dtzes et al. 2004). Moreover, the North Sea Basin is experiencing a Plio-Pleistocene phase of accelerated subsidence and a related depression of the Moho that can be attributed to stress-induced downward deflection of the lithosphere (van Wees & Cloetingh 1996). Similarly, lithospheric folding probably contributes to the continued uplift of the Fennoscandian Shield (Cloetingh et al. 2005). The present stress field reflects a combination of forces related to collisional interaction of the Alpine Orogen with its foreland and Atlantic ridge push (Gtlke et al. 1996).

Alpine orogens and Western Mediterranean basins The Alpine orogenic belts are characterized by variably deep crustal roots. The deepest roots are associated with the Pyrenees (Choukroune et al. 1990; Vergts & Garcias-Senez 2001), the Alps (Waldhauser et al. 1998), the northern Apennines (Finetti et al. 2001) and the Dinarides (Skoko et al. 1987), reflecting insertion of continental foreland crust into the mantle and the development of mantle back-stops, involving an offset of the upper and lower plate crust-mantle boundaries (Route et al. 1996). In contrast, the Betic Cordillera of Spain, the North African TellianMaghrebian chain and the Carpathians are characterized by shallower crustal roots or their absence (Fig. 1; see also Cavazza et al. 2004). This can be variably attributed to slab detachment, back-arc extension and early stages of post-orogenic extensional collapse of these orogens. For instance, in the Central Alps, crustal shortening persisted after the late Eocene detachment of the subducted oceanic AlpineTethys slab, accounting for the insertion of a secondary, 120 km long slab, consisting of continental lower crust and lithospheric mantle, into the asthenospheric mantle (Schmid et al. 1996, 2004; D~zes et al. 2004). Because at depths of 55-60 km subducted crustal material entered the eclogite stability field, its P-wave velocity increased to velocities typical for the mantle, and thus, by crossing the Moho discontinuity, limited the seismic depth of the crustal roots (Bousquet et al. 1997; Stampfli et al. 1998). On the other hand, the Betic-Balearic-Corsica Orogen, which was activated during the Late Cretaceous ( + 8 5 - 8 0 Ma, Faccenna et al. 2001), was disrupted by late Oligocene-early Miocene back-arc extension, culminating in B urdigalian (_+ 18 Ma) detachment of the Kabylian arc from the orogen. This marked the onset of opening of the oceanic Algerian Basin (Roca 2001), which was compensated by progressive subduction of the Alpine Tethys (Doglioni et al. 1999a). Following Langhian collision of the Kabylian arc with the North African margin, the subducted Tethys slab was detached, as evidenced by widespread bimodal magmatism, whereas compressional deformation of the Maghrebian-Tellian systems persisted intermittently into the Pleistocene, involving inversion of the Mesozoic Atlas rift system and commensurate crustal thickening (Carminati et al. 1998; Vergts & Sabat 1999; Frizon de Lamotte et al. 2000; Ziegler et al. 2002; Spakman & Wortel 2004).

Following opening of the Provenqal-Ligurian Basin and collision of the Corsica-Sardinia accretionary wedge with the Apulian passive margin, commencing in the north during the mid-Oligocene and progressing in time southward, the internal parts of the evolving Apennine orogenic belt were disrupted from the late Miocene onward by back-arc extension governing opening of the Tyrrhenian Basin, which is partly floored by denuded mantle (shown in Fig. 1 as oceanic crust) (Mauffret & Contrucci 1999; Stranne 1999; Faccenna et al. 2001; see also TRANSMED Transect III, Cavazza et al. 2004). Controlling mechanisms were delamination, roll-back, deformation and partial detachment of the subducted Alpine-Tethys slab from the Apulian lithosphere that was accompanied by a high-K calc-alkaline to shoshonitic magmatism (Carminati et al. 1998; Doglioni et al. 1999b; Wilson & Bianchini 1999; Argnani& Savelli 2001; Faccenna et al. 2001; Lucente & Speranza 2001; Spakman & Wortel 2004). After the Eocene collisional main deformation phase of the Dinarides, continued northward movement of the Apulian block caused dextral transpressional reactivation of the Sava-Vardar suture during the early Oligocene, triggering detachment of the subducted lithospheric slab and extensive shoshonitic magmatism (Pamic 2002; Pamic et al. 2002). During the Miocene, eastward extrusion of the Alpine-Carpathian Block and roll-back of the Carpathian subduction system was accompanied by continued crustal shortening in the Carpathians and wrench deformation of the internal Dinarides and the Pannonian domain, controlling the subsidence of transtensional and pull-apart basins (Horv~th 1993; Frisch et al. 1998; Tari & Pamic 1998; Fodor et al. 1999). This was coupled with intense thinning of the orogenically destabilized crust and lithospheric mantle of Pannonian Basin, involving upwelling of the asthenosphere (Tari et al. 1999; Cloetingh & Lankreijer 2001).

Summary and conclusions Depending on convergence rates and the thermal state of the foreland lithosphere, crustal roots of active continent-continent collisional orogens can extend to depths of 55-60 km (Alps) or even to 75 km (Himalayas). At these depths the subducted continental crust becomes eclogitized, assumes densities and velocities comparable with those of the mantle and, thus, is transferred across the seismically defined Moho discontinuity, and seismically appears to form part of the lithospheric mantle (Bousquet et al. 1997; Henry et al. 1997). By analogy with modem orogens, the Caledonian and Varsican orogens were presumably characterized, prior to their post-orogenic collapse, by variably deep-reaching crustal roots. The crustal roots of the Irish-Scottish-Scandinavian, North German-Polish and Mid-European Caledonides were presumably destroyed in conjunction with Early Devonian post-orogenic wrench faulting and back-arc rifting. Rifting during Late Devonian and Carboniferous times further disrupted the Caledonides of the British Isles. The crustal roots of the Variscan Orogen were destroyed during the wrench-dominated PermoCarboniferous tectonomagmatic cycle, in the course of which the crust thinned to 28-35 km. Simultaneously, the lithosphere of the Variscan foreland was destabilized by wrench faulting and magmatic processes. In the North Sea area, as in the British Isles, crustal thicknesses were probably variable prior to the onset of Mesozoic rifting. Late Permian and Mesozoic tiffing, affecting the area of the Atlantic shelves, the North Sea, the Tethys shelves and to a lesser degree the Sorgenfrei-Tomquist-Teisseyre Line, caused significant crustal thinning that progressed to crustal separation and the Mid-Jurassic opening of the Alpine Tethys, the Early Cretaceous opening of the North Atlantic, and the Mid-Cretaceous opening of the Bay of Biscay and the Valais Trough. Impingement of the Iceland plume,

CRUSTAL EVOLUTION OF EUROPE

immediately preceding crustal separation between Greenland and Europe at the end of the Palaeocene, caused further crustal thinning by magmatic destabilization of the crust-mantle boundary. On the other hand, Eocene to recent development of the ECRIS was apparently associated with less intense crustal thinning in the domain of Rhine-Rh6ne rift system, but progressed by early Miocene times to crustal separation and limited sea-floor spreading in the ProvenqalLigurian Basin. Back-arc extension, controlled by eastward roll-back of the Alpine-Tethys slab, governed the opening of the oceanic Provenqal-Ligurian and Algerian basins and subsidence of the Tyrrhenian Basin, involving mantle denudation. The oceanic floor of the Ionian Sea represents a remnant of PermoTriassic Neotethys Ocean (Stampfli et al. 2001; Ziegler et al. 2001; Stampfli & Borel 2004). The frequently observed discrepancy between the magnitude of upper crustal extension by faulting and the amount of extension derived from the crustal configuration of a rift zone is evidence for synrift magmatic destabilization of the crust-mantle boundary (e.g. Oslo Graben) or for non-uniform pre-rift crustal thicknesses (e.g. North Sea Rift, Upper Rhine Graben; Ziegler & Cloetingh 2004). Regarding the present-day crustal configuration of extra-Alpine Europe, we would like to further comment on some of the salient features of Figure 1. Broad zones of crustal thinning, which characterize the Rockall-Hatton-Faeroe Bank, the Atlantic shelves of the British Isles, France and Iberia, as well as the North Sea, must be largely attributed to Mesozoic crustal extension, although crustal thicknesses were probably not uniform prior to the onset of Mesozoic firing. Moreover, in the area of the Hebrides Shelf and the Rockall-Hatton-Faeroe Bank Palaeocene plume-related thermal destabilization of the crust-mantle boundary, combined with erosional unroofing of the crust in response to its thermal doming, probably contributed to crustal thinning. However, Late Cretaceous and Cenozoic post-rift thermal subsidence of these Mesozoic extensional basins caused a gradual depression of the Moho, amounting, for instance, in the Central North Sea to over 4 km. On the other hand, thinning of the Neoproterozoic Dalslandian crust in the North Danish Trough resulted partly from Mesozoic crustal stretching and partly from Permo-Carboniferous destabilization of the crustmantle boundary. Crustal thinning in the Oslo Graben must be exclusively attributed to Permo-Carboniferous rifting and magmatic destabilization of the crust-mantle boundary. Similarly, the regional Moho uplift that is still associated with the area of the Southern Permian Basin, despite its Late Permian to Cenozoic subsidence by as much as 8 km, must be attributed to Early Permian magmatic destabilization of the crust-mantle boundary. The SW-NE-trending broad zone of Moho shallowing, which is associated with the intracontinental part of the ECRIS, probably reflects a combination of (1) Permo-Carboniferous (predominantly magmatic) crustal thinning, (2) Cenozoic crustal extension, (3) erosional unroofing of the crust in response to Neogene plume-induced doming of the Rhenish Massif and the Massif Central, and mid-Miocene-Pliocene lithospheric folding controlling uplift of the Vosges-Black Forest arch. Latest Cretaceous and Palaeocene intraplate compressional deformation probably controlled crustal thickening in the Polish Trough and in the Bohemian Massif, and to a lesser extent also in the inverted basins of Denmark, Germany and the Netherlands. Eocene and Oligocene inversion of the Central Iberian and Catalan Coast Ranges resulted in crustal thickening, whereas late Oligocene-early Miocene tiring of the Valencia Trough caused thinning of the previously thickened crust. The present-day crustal configuration of Western and Central Europe resulted from polyphase Late Palaeozoic to recent deformation of the lithosphere that overprinted the margin of the Proterozoic East European Craton and, particularly, the Caledonian and Variscan crustal domains. In an effort to explain the crustal configuration of a given area, the total sum

53

of processes that affected it through time must be taken into consideration. This paper is a contribution to EUROPROBE by the European EUCORURGENT Project (Upper Rhine Graben: Evolution and Neotectonics). P. D~zes acknowledges financial support by a University of Basel ELTEM grant. Critical and constructive comments by W. Franke and I. Artemieva on a draft of this paper are gratefully acknowledged. Special thanks go to D. Gee for enthusiastically leading the EUROPROBE Project, which was instrumental in integrating the Earth Science communities of Western and Eastern Europe.

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289, 281-294. TORNI~, M., BANDA, E. • FERNANDEZ, M. 1996. Valencia Trough, Spain. In: ZIEGLER, P. A. & HORVATH, F. (eds) Peri-Tethys Memoir 2, Structure and Prospects of Alpine Basins and Forelands. M~moires du Mus6um National d'Histoire Naturelle, 170, 103-128. WALDHAUSER, F., KISSLING, E., ANSORGE, J. & MOLLER, S. 1998. 3D interface modelling with 2D seismic data: the Alpine crust-mantle boundary. Geophysical Journal International, 135, 264-278. VAN WEES, J.-D., STEPHENSON,R. A., ZIEGLER, P. A. ETAL. 2000. On the origin of the Southern Permian Basin, Central Europe. Marine and Petroleum Geology, 17, 43-59. ZEYEN, H., NOVAK, P., LANDES, M., PRODEHL, C., DRIAD, L. & HIRN, A. 1997. Refraction-seismic investigations of the northern Massif Central (France). Tectonophysics, 275, 99-117. ZIEGLER, P. A. 1990. Geological Atlas of Western and Central Europe, 2nd edn. Shell International Petroleum (distributed by Geological Society, London).

Europe from the Variscan to the Alpine cycles GI~RARD M. STAMPFLI l & HEINZ W. KOZUR 2 lInstitut de Gdologie et Paldontologie, Universiti de Lausanne, BFSH2, CH 1015 Lausanne, Switzerland (e-mail: Gerard.Stampfli @unil.ch ) 2Rdzsii u. 83, H-1029 Budapest, Hungary

Abstract: The time span between the Variscan and Alpine cycles is not devoid of any major tectonic activity, and corresponds to the Cimmeriancycle. Between the Early Permian and Late Triassic,the Eocimmeriancycle was markedby the closureof Palaeotethys and opening of Neotethys and of an array of south Eurasian back-arc basins. This was followed by the break-up of Pangaea and the Early Jurassic openingof the central Atlanticand AlpineTethys. However, in the area of the Eocimmeriancollision,the geodynamicevolution is relatively uninfluencedby this event, and a new cycle of Cimmeriandeformation affected the Hellenides,Dinarides,Balkans and Pontides in Jurassic-Early Cretaceous times. The anti-clockwiserotation of Africa during the Late Cretaceous heralded the onset of Alpine orogenic processes, characterized first by major east-west shortening, and opening and closure of younger oceanic basins of back-arc type.

A set of reconstructions covering the time period between the Variscan and Alpine cycles is presented in this paper (Figs 1-9). These reconstructions mainly delineate oceanic and continental domains, and most of them do not show the epicontinental seas (except Fig. 2); for those, readers can consult, for example, the results of the peri-Tethys group (Dercourt et al. 1986, 2000), those from IGCP 369 (Stampfli et al. 2001a; Ziegler et al. 2001), Golonka (2000), and the many references therein. These reconstructions have been developed following a new method of plate reconstruction (Stampfli & Borel 2002), which represents a distinct departure from classical continental drift models. These new plate tectonic models for the Palaeozoic and Mesozoic (Ordovician to Palaeogene) integrate dynamic plate boundaries, plate buoyancy factors, ocean spreading rates, subsidence patterns, new stratigraphic results, palaeobiogeographical and palaeomagnetic data, as well as major tectonic and magmatic events. Plates have been constructed through time by adding or removing oceanic material, symbolized by synthetic isochrons, to major continents and terranes. This approach offers a good control on plate kinematics, providing new constraints for plate tectonic scenarios, which are still numerous for the Tethyan realm. The relationship between the Variscan and the Cimmerian cycles in the Mediterranean-Alpine regions is illustrated in the first part of the paper by a set of detailed reconstructions. In a second part, a thorough account of the stratigraphy and geodynamic evolution of the main Tethyan oceanic realms is presented.

Part h geodynamic evolution Review of the pre-Variscan and Variscan cycle (Fig. 1) The well-known Variscan basement areas of Europe contain relict terranes characterized by a pre-Variscan geodynamic evolution testifying to their peri-Gondwanan origin. A Neoproterozoic active margin setting with volcanic arcs is observed along the entire length of the future European microcontinents formerly located at the Gondwanan border (von Raumer et al. 2002, and references therein). The evolution of this active Gondwana margin was guided by the diachronous subduction of a Prototethyan oceanic ridge under different segments of the margin. This subduction triggered the emplacement of magmatic bodies and the formation of back-arc rifts (yon Raumer et al. 2002), some of them becoming major oceanic realms (Rheic, Palaeotethys). One of the major blocks that rifted away from Gondwana in Early Palaeozoic time was Avalonia, accompanied by the opening of the Rheic ocean. A short Ordovician orogenic event

(von Raumer et al. 2002) was followed, after the Silurian, by the drifting of the Hun superterrane, accompanied by the opening of Palaeotethys. The slab roll-back of the Rheic ocean is viewed as the major mechanism for the drifting of the European Hun terranes towards Eurasia. The subduction of the Rheic ocean, in turn, generated a large slab-pull force responsible for the opening of rift zones within the passive Eurasian margin to which Avalonia had been accreted. The rifts evolved to the opening of peri-Eurasia oceanic domains (Rheno-Hercynian, Paphlagonian, Mugodzhar oceans). Therefore, the first mid-Devonian Variscan orogenic event, characterized by H P - L T metamorphism, is viewed as the result of a collision between terranes detached from Gondwana (the Hun superterrane) and terranes detached from Eurasia (Hanseatic terranes) (Stampfli et al. 2002b, and references therein) as well as collision with arcs derived from the Asiatic ocean. Subsequently, the amalgamated terranes collided with Eurasia in a second Variscan orogenic event in Vis~an times, accompanied by large-scale lateral escape and/or oroclinal bending of major parts of the accreted margin and the establishment of a Variscan cordillera. Final collision of Gondwana with Laurussia did not take place before the Late Carboniferous and was responsible for the Alleghanian orogeny, but the collision did not affect the eastern Alpine and Mediterranean part of the Variscan cordillera, which remained active for a much longer time, as shown by widespread arc-related plutonic activity in Late Carboniferous to earliest Permian (e.g. Calabria, Acquafeddra et al. 1994; Southern Alps, Oberh~insli et al. 1985; Dinarides, Pamic et al. 1996; Hellenides, Vavassis et al. 2000; see also Stampfli 1996, and references therein). This magmatic belt was thereafter the focus of Early Permian extension. In the Urals, a Late Carboniferous to Early Permian flexural (molasse) basin developed on the western side of the growing orogen. Pelagic sedimentation there lasted until the earliest Kungurian (Fokin et al. 2001); this basin connected the Tethyan realm and the Arctic. Final locking of Laurussia and the Kazakkstan plate can be placed at that time, whereas locking with the Siberian block is slightly younger (Mid-Permian; Russia-ITLP 1997).

Post-Variscan evolution (Figs 2 - 5 ) Late Palaeozoic suturing of Laurussia and Gondwana (Fig. 2) was accompanied and followed by a major plate boundary reorganization that involved subduction progradation from the Variscan suture in the interior of Pangaea to its peripheries and detachment of the Cimmerian composite terrane from the non-collisional

From: G~E, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 57-82. 0435-4052/06/$15.00 9 The GeologicalSociety of London 2006.

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G.M. STAMPFLI & H. W. KOZUR

Fig. 1. Drift history of Gondwana-derived basements between the Late Silurian and the Late Carboniferous, modified from Stampfli et al. (2002b). 1, Oceanic area and spreading ridge; 2, subduction zone; 3, suture; 4, transcurrent limit; 5, passive margin; 6, ocean. Abbreviations of key localities: AA, Austro-Alpine; Ab, Alboran; Ad, Adria s.s.; Ae, Abadeh; Af, Northern-Afghanistan, Band-e-Turkestan; Ag, Alada~-Bolkarda~; Ah, Agh-Darband; Aj, Ajat; A1, Alborz; Am, Armorica; An, Antalya, lower nappes; Ap, Apulia s.s.; Aq, Aquitaine; AP, Aspromonte, Peloritani; Ar, Arna accretionary complex; As, Apuseni, south, ophiolites; At, Attika; Au, Asterousia; Av, Arvi; Ay, Antalya, upper nappes; Ba, Balkanides, external; Bb, Band e Bayan; Bc, Biscay, Gascogne; Bd, Bey Da~lan; Be, Betic; Bf, Baft ophiolite; BH, Baer-Bassitand Hatay ophiolites; Bh, Bihar; Bi, Ba'id; Bj, Birjan ophiolite; Bk, Bozda~-Konya forearc; B1, Bitlis massif; BM, Bela, Muslim-bagh ophiolite; Bn, Bernina; Bo, Bolkarda~; Br, Brian~onnais; Bs, Bisitoun seamount; BS, Bator-Szarvasko ophiolites; Bt, Batain; Bu, Bucovinian; Bti, Btikk; Bv, Budva; BV, Bruno-Vistulian; By, Bey~ehir; Bz, Beykoz basin; Ca, Calabria autochthon; cA, central Afghanistan, Hazarajat; cB, central Bosnia; cD, central Dinarides ophiolites; Ce, Cetic; Cg, Chagai arc; (Continued)

VARISCAN-ALPINE

northern margin of Gondwana (Fig. 3), causing the Permo-Triassic opening of Neotethys (Stampfli et al. 2001a). Collapse of the Variscan cordillera was accompanied by regional uplift, the subsidence of an array of transtensional basins and widespread magmatism that can be related to the detachment of subducted slabs and lithospheric thinning. With the Mid-Permian consolidation of the Alleghenian orogen, tectonic and magmatic activity abated in the Variscan domain (from Iberia to west of the Tornquist line). The Late Permian and Triassic evolution was dominated by thermal relaxation of the lithosphere, southward propagation of the Arctic-North Atlantic rift system and westward propagation of the Tethys rift system (Ziegler & Stampfli 2001). Back-arc rifting associated with the Palaeotethys subduction zone caused Permo-Triassic opening of the M a l i a c - M e l i a t a Kfire-Svanetia system of marginal basins along the still active Eurasian margin (Figs 3 and 4). These were closed during the Late Triassic Eocimmerian (e.g. ~eng6r 1979) to Late Jurassic Cimmerian orogenic cycle (e.g. Nikishin et al. 1997), involving the collision of Cimmerian terranes with the Palaeotethys a r c trench system in Late Triassic times (Fig. 5). This Eocimmerian collision extends from Apulia to Thailand (~eng6r 1979; ~eng6r & Hsfi 1984). In most cases it consisted of a docking of terranes, therefore, no major orogen resulted from the collision; however, locally, relatively large molasse or foreland basins developed, such as the Kasimlar-t~aylr basin in the Taurus (Dumont 1976; Gutnic et al. 1979), and the Shemshak basin in Iran (Corsin & Stampfli 1977; Davoudzadeh & Schmidt 1981). The pefi-Tethys rift system provided avenues for Late Permian and Triassic transgressions of the Tethys seas onto Europe (Ziegler 1988; Stampfli et al. 2001a). The Norwegian-Greenland Sea rift system paved the way for the Late Permian Zechstein Sea transgression into Western and Central Europe (Ziegler 1989, 1990) (Fig. 3).

The A l p i n e cycle (Figs 5 - 9 )

There is a fundamental difference between the Alpine orogen (Alps and Carpathians) and the Tethysides (Dinarides-Hellenides, the Middle East mountain belts and the Himalayas sensu lato (s.l.). The Neotethys ocean, whose closure was responsible for the

59

formation of the Tethysides orogenic system, does not directly interfere with the Alpine domain sensu stricto (s.s.), and the Alpine Tethys should be regarded more as an extension of the central Atlantic Ocean into the Tethyan realm rather than a branch of the large and older Neotethys ocean (Bernoulli & Jenkins 1974). In that sense, the onset of the Alpine cycle could be placed in the Carnian, a period corresponding to the final closure of the Palaeotethys in the Mediterranean and Middle East regions (Fig. 5) (Kozur 1999; Stampfli et al. 2003) and to the onset of rifting in the central Atlantic-Alpine domain. Spreading in these regions did not start before Early to Mid-Jurassic time (Fig. 6) (e.g. Stampfli & Marthaler 1990; Manatschal 1995; Froitzheim & Manatschal 1996; Bill et al. 1997; Steiner et al. 1998). Within the Alpine domain s.s., there is a fundamental difference between the Austroalpine-Carpathian and Western Alps systems. The former presents an evolution rooted in the dynamics of Triassic back-arc basins located south of it (Meliata-Maliac domain). These back-arc basins were shortened as a result of the opening of the central Atlantic and rotation of Africa relative to Europe. Subsequent slab roll-back of the Maliac-Meliata-Ktire sea f o o r induced the opening of the Vardar suprasubduction zone (SSZ) and I z m i r - A n k a r a oceans, which, by Late Jurassic time, had completely replaced the pre-existing oceanic basins (Fig. 6). Continuing rotation of Africa provoked ridge failure in the Vardar and large-scale Late Jurassic ophiolitic obduction onto the Dinaride-Hellenide passive margin of the Pelagonian terrane (e.g. Laubscher & Bernoulli 1977; Baumgartner 1985; Dercourt et al. 1986). According to some workers (e.g. Smith et al. 1975; Jones & Robertson 1990; Clift & Dixon 1998) these obducted ophiolites came from the west (for discussion of this problem see Stampfli et al. (2003) and references therein). Roll-back of the Meliata-Maliac-Ktire slab generated an oroclinal bending of the Vardar upper plate, inducing collision on all its borders; thus, following the obduction of its western border, the Vardar shortening entailed a collision between its northeastern a r c trench system and the northern passive margin of Meliata, represented by the Northern Calcareous Alps (NCA) and Western Carpathians domain and the Rhodope, closing at the same time the Balkan rift system between Moesia and the Rhodope (Fig. 7). This event resulted in the NW Balkan orogen, accompanied by large-scale Early Cretaceous northward nappe

Ch, Channel; cI, central Iberia; Ci, Ciotat flysch; Ck, Chehel Kureh ophiolites; CL, Campania-Lucania; Co, Codru; Cn, Carnic-Julian; CP, Calabria-Peloritani; cR, circum-Rhodope; Ct, Cantabria-Asturia; Cv, Canavese; Da, Dacides; Db, Dent Blanche; DD, Dniepr-Donetz rift; Dg, Denizgrren ophiolite; DH, Dinarides-Hellenides; Di, Dizi accretionarycomplex; Dm, Domar; Do, Dobrogea; Dr, Drina-Ivanjica; Ds, Drimos, Samothrace ophiolites; Du, Durmitor; Dy, Derekry basin; eA, east Albanian ophiolites; E~, Eriq ophiolite; El, Elazig, Guleman ophiolites arc; eP, east Pontides; Er, Eratosthenes seamount; Es, Esfandareh ophiolites; Fa, Fatric; Fc, Flemish cap; FM, Fanuj, Maskutan ophiolite; Fr, Farah basin; GB, Grand Banks; gC, Great Caucasus; Gd, Geyda~-Anamas-Akseki; Gi, Giessen; Ge, Gemeric; GS, Gory-Sovie; GT, Gavrovo-Tripolitza; Gt, Getic; Gii, Gfimfishane-Kelkit; hA, High Atlas; Ha, Hadim; He, Helvetic rim basin; Hg, Hu~lu-Boyalitepe; HK, Hindu-Kush; hK, high karst; HM, Hu~lu-Mersin; Hr, Hronicum; Hy, Hydra; Hz, Harz; IA, Izmir-Ankara ocean; iA, intra-alpine terrane; Ib, Iberia, NW allochthon; Ig, Igal trough; Io, Ionian; It, Iranshar ophiolite; Is, Istanbul; Ja, Jadar; Jr, Jeffara rift; Jo, Jolfa; Jv, Juvavic; Ka, Kalnic; Kb, Karaburun; Kd, Kopet-Dagh; Ke, Kotel flysch.; Kg, Karabogaz Gol; Ki, Kir~ehir; Kk, Karakaya forearc; K1, Kabul block; Ko, Korab; KQ, Kunlun-Qaidam; Kr, Kermanshah; KS, Kotel-Stranja rift; KT, Karakum-Turan; Ku, Kura; Kfi, Kiire ocean; KW, Khost, Waziristan ophiolites; Ky, Kabylies; La, Lagonegro; 1A, lower Austroalpine; Lb, Longobucco; Le, Lesbos ophiolites; Lg, Ligerian; Li, Ligurian; LM, Lysogory-Malopolska; Lo, Lombardian; Ls, Lusitanian; LT, Lut-Tabas-Yazd; Lu, Lut; Ly, Lycian ophiolitic complex; Lz, Lizard ophiolitic complex; mA, middle Atlas; Ma, Mani; Mb, Magnitogorsk back-arc; Mc, Maliac rift or ocean; MD, Moldanubian; Me, Meliata rift or ocean; Mf, Misfah seamount; Mg, Magura; Mh, Mugodzhar ocean; Mi, Mirdita autochthon; Mk, Mangyshlak rift; M1, Meglenitsa ophiolite; Mm, Mamonia accretionary complex; MM, Meguma-Meseta; Mn, Menderes; Mo, Moesia; MP, Mersin, Pozanti ophiolites; Mr, Mrzlevodice forearc; MR, Masirah, Ra's Madrekah ophiolites; Ms, Meseta; MS, Margna-Sella; Mt, Monte Amiata forearc; Mz, Munzur Da~, Keban; nC, North Caspian; Ni, Niltifer seamount; Nk, Nakhlak; Nr, Neyriz seamount; Nn, Nain ophiolite; Ns, Niesen flysch; nT, north Tibet; Nt, Nish-Troyan trough; Ny, Neyriz seamount; OM, Ossa-Morena; Or, Ordenes ophiolites; Ot, Othrys-Evia-Argolis ophiolites; Oz, Otztal-Silvretta; Pa, Panormides; Pd, Pindos rift or ocean; Pe, Penninic; Pi, Piemontais; Pj, Panjao, Waser ocean; Pk, Paikon intra-oceanic arc; P1, Pelagonia; Pm, Palmyra rift; Pn, Pienniny rift; Pp, Paphlagonian ocean; Px, Paxi; Py, Pyrenean rift; Qa, Qamar; Rf, Rif, external; Rh, Rhodope; RH, Rheno-Hercynian ocean; Ri, Rif, internal; Rk, Ratuk ophiolite; Ru, Rustaq seamount; Rw, Ruwaydah seamount; Sa, Salum; sA, South Alpine; sB, sub-Betic rim basin; Sc, Scythian platform; sC, South Caspian basin; Sd, Srednogorie rift-arc; Se, Sesia; Sh, Shemshak molasse basin; Si, Sicanian; Sj, Strandja; Sk, Sakarya; sK, south Karawanken forearc; S1, Slavonia; Sm, Silicicum; SM, Serbo-Macedonian; sM, southern Mongolia; Sn, Sevan ophiolites; sP, south Portuguese; Sr, Severin ophiolites; SS, Sanandaj Sirjan; St, Sitia; Su, Sumeini; Sv, Svanetia rift; Sx, Saxo-Thuringian; Sz, Sabzevar ophiolite; Ta, Taurus, s.l.; Tb, Tabas; TB, Tirolic-Bavaric; tC, Transcaucasus; TD, Trans-Danubian; Tg, Tuzgrlti basin; Th, Thrace basin; Tk, Tuarkyr; Tin, Tarim; To, Talea Off; Tp, Troodos ophiolite; Tr, Turan; Tt, Tatric; Tu, Tuscan; Tv, Tavas + Tavas seamount; Ty, Tyros forearc; Tz, Tizia; uJ, upper Juvavic; UM, Umbria-Marches; Uy, Ust-Yurt; Va, Valais trough; Ve, Veporic; Vo, Vourinos (Pindos)-Mirdita ophiolites; wC, western Crete (Phyl-Qrtz) accretionary complex; Ya, Yazd; Z1, Zlatibar ophiolites; Zo, Zonguldak; Zt, Band e Ziarat ophiolites.

60

G.M. STAMPFLI & H. W. KOZUR

Fig. 2. WesternTethysreconstructionsfor the Late Carboniferousand Early Permian,showingthe extent of shallowseas, modifiedfrom Stampfli& Borel (2004). 1, Passivemargin;2, magneticanomalyor syntheticanomaly;3, seamount;4, intraoceanicsubduction-arc complex;5, spreadingridge; 6, subductionzone; 7, rift; 8, suture; 9, active thrust; 10, forelandbasin; 11, flexuralbulge; a, shallowmarineembayment;b, continentalbasin. (For abbreviationsof key localities,see Fig. 1.)

emplacement and metamorphism (Georgiev et al. 2001; Okay et al. 2001a; B o n e v & Stampfli 2003). This orogenic event started in the Late Jurassic and was sealed by Albian to Cenomanian molasse-type sediments. The latest deformation affecting that region took place in the Eocene after a renewed shortening phase starting in the latest Cretaceous-Palaeocene. Along the NCA segment, there was no real collision; elements of the Austro-Alpine microcontinent were scraped off, and incorporated into the accretionary wedge, to form the internal structural units of the Austro-Carpathian orogen (Kozur 1991b; Plagienka 1996; Faupl & Wagreich 1999). This event was accompanied by Early Cretaceous H P - L T metamorphism (e.g. Th6ni & Jagoutz

1992). Then, the enlarged accretionary wedge started to extend over the eastern segment of the Alpine Tethys (Figs 7 and 8) (Penninic-Vahic ocean), to finally collide with its northern border (Helvetic domain s.l., Magura rim basin) to give the present Eastern Alps and Carpathian orogen (e.g. Wortel & Spakman 1993). In this process, slab roll-back never really stopped, from its beginning in the Ktire domain during the Carnian until its propagation to the Eastern Carpathians in the Neogene. As a result of the Late Jurassic collision of the Vardar plate with the Rhodope and Balkans, eastward opening of the Alpine Tethys was stopped, and separation of Africa from Europe could not proceed following the previous Jurassic pattern. Then, the North

VARISCAN-ALPINE

61

Fig. 3. Western Tethys reconstructions for the Mid- and Late Permian, modified from Stampfli & Borel (2004). (For legend see Figs 1 and 2.)

Atlantic opening started, first between Iberia and Europe (Fig. 7). Therefore, the Western Alps domain has a totally different evolution from that of the Eastern Alps, marked by the drifting of the Iberic plate since Late Jurassic time (Fig. 6). It can be shown from the Atlantic magnetic anomalies that the drift of Iberia and Africa were similar and coeval during most of the Cretaceous (Stampfli & Borel 2002). Therefore, during that period, the African plate northern limit was between Iberia and Europe, detaching, together with Iberia, the Corsican-Brian~onnais microcontinent from southern France through the opening of the Pyrenean rift system to finally place them south of the Alpine domain (Frisch 1979; Stampfli 1993; Stampfli et al. 1998; Stampfli et al. 2002a). The northern margin of the Alpine Tethys was duplicated in that process, and the Brian~onnais eastern tip, following the eastward drift of Iberia, collided with the Austroalpine-Carpathian prism western border in the Late Cretaceous (Fig. 8),

creating an angular unconformity on which the Palaeocene Wildflysch of the Falknis nappe (eastern Brian~onnais) transgressed (Alleman 2002). This east-west shortening was followed by the onset of subduction of the Piedmont part of the Alpine Tethys in the latest Cretaceous, accompanied by H P - L T metamorphism of some elements of its former southern passive margin (e.g. Sesia Massif, Rubatto 1998). This subduction was located just north of a major transcurrent plate boundary (a palaeo-Insubric line), which allowed the AdriaTizia plate to be translated eastward and finally occupy a place to the south of its present location. These lateral displacements of Apulia s./.-Africa have been clearly documented in the Eastern Alps (Trtimpy 1988). This eastward movement was triggered by the subduction of the remnant Vardar ocean under Moesia and the development of the Srednogorie arc in Late Cretaceous times (Georgiev et al. 2001). By Late Eocene time the Western Alps

62

G.M. STAMPFLI& H. W. KOZUR

Fig, 4. Western Tethys reconstructions for the Early and Mid-Triassic, modified from Stampfli & Borel (2004). (For legend see Figs 1 and 2.)

orogenic wedge had included the Brianqonnais domain as an exotic terrane (Stampfli et al. 2002a) and most of the European passive margin was already subducting all along the Alpine Tethys from the Western/kips to the Eastern Carpathians. The remnant western Alpine Tethys, extending from Italy (Ligurian part) to Morocco (the Maghrebian Tethys), started subducting northward under Spain at that time (e.g. Puga et al. 1995), or possibly in Late Palaeocene time (e.g. Finetti et al. 2001) (Fig. 9). Slab roll-back of this remnant ocean gave birth to the Apenninic-Maghrebian orogenic wedge, detachment of the Corso-Sardinian-Kabylian block and opening of the Algero-Provenqal basin (Roca 2001; Cavazza et al. 2004). This orogenic wedge finally reached the Ionian basin (Neotethys westernmost tip), and roll-back proceeded southeastward in this narrow oceanic corridor, opening the Tyrrhenian back-arc and detaching the Calabrian block from Sardinia (Mantovani et al. 1994; Gueguen et al. 1998). Westward, slab roll-back

proceeded along another oceanic corridor (Gutscher et al. 2002), corresponding to the segment of the Alpine Tethys between Iberia and NW Africa (Maghrebian Tethys), giving rise to the Betics and Rif orogenic wedges, and to opening of the Alboran marginal basin (Spakman & Wortel 2004). C r e t a c e o u s e a s t - w e s t shortening in the Alpine-Mediterranean

domain

The east-west shortening in the Vardar region started in midCretaceous time, after the Balkan orogenic event. Then, subduction reversal took place as the Balkan domain, located on the lower plate during the Early Cretaceous orogenic event, was transformed into an upper plate with the development of the Late Cretaceous Srednogorie arc. The remnant Vardar slab started to subduct under the

VARISCAN-ALPINE

63

Fig. 5. Western Tethys reconstructions for the Late Triassic and Early Jurassic, modified from Stampfli & Bore1 (2004). (For legend see Figs 1 and 2.)

Balkan orogen after the Albian and further east under the western Pontides, where it was accompanied by H P - L T metamorphism dated to between 100 and 60 Ma (Okay et al. 1991; Okay & Tansel 1992). This subduction resulted in the Late Cretaceous Srednogorie-Pontides arc and, following roll-back of the Vardar slab, the western Black Sea back-arc opening in the Cenomanian (Robinson 1997; Kazmin et al. 2000) (Fig. 8). The closure of the Vardar ocean was diachronous, occurring first in the Dinaric region in the Maastrichtian-Palaeocene (Pamic 2002), and later in the Rhodope-Hellenic region, during the Palaeocene-Eocene (Fig. 9) (Yanev & Bardintzeff 1997). An oceanic space remained open longer along the Pontides (Okay & Ttiyztis 1999). There, the Late Cretaceous opening of the Lycian (SSZ) ocean followed the eastward slab retreat of the Izmir-Ankara ocean. Obduction of the Lycian s.l. ophiolites all along the Anatolides margin was nearly synchronous, as shown

by the development of the amphibolitic sole of most Tauric ophiolites at around 95 Ma (e.g. Dilek et al. 1999). Subduction-related processes lasted until the Eocene final closure of the remnant Lycian ocean all along the Pontides segment (Koqyigit 1991; Okay & Sahintiirk 1997; Kaymak~i et al. 2000). Whereas in the Turkish transect, north-south Tertiary shortening involved a young Cretaceous ocean obduction-subduction (Lycian domain), in the Hellenic transect, this time interval corresponds to the closure of the Late Triassic-Early Oligocene Pindos basin (Fig. 9) (Richter et al. 1993; Degnan & Robertson 1998). However, subduction of the eastern part of the Pindos-Antalya oceanic domain under the Tauride plate started in the Late Cretaceous as a result of the major east-west shortening movements. Remnants of this event are found in the Late Cretaceous metamorphic sequences of the Cyclades (Br6cker & Enders 1999) and the Asteroussia nappe of Crete (Seidel et al. 1976; Bonneau

64

G.M. STAMPFLI& H. W. KOZUR

Fig. 6. Western Tethys reconstructions for the Mid- and Late Jurassic, modified from Stampfli & Borel (2004). (For legend see Figs 1 and 2.)

1984). It is also recorded in the Late Cretaceous Pindos first flysch in Greece (Neumann & Zacher 1996; Wagreich 1996) and the emplacement of the Antalya ophiolites on a flexured Bey Da~lan platform in the Latest Cretaceous-Palaeocene (e.g. Gutnic et al. 1979). The Lycian nappe obduction on the TauricAnatolian plate (Robertson 2002, and references therein) should also be viewed in the frame of the Late Cretaceous east-west shortening and final emplacement of this plate to the south of the Pontides. The eastward subduction of the Pindos basin implies that elements located to the north (Pelagonia) and south of it (Greater Apulia, i.e. Bey Da~lan and lower Antalya domain) were imbricated on both sides of the Tauric-Anatolian plate, creating a duplication of older features such as the Palaeotethys suture zone (i.e. Tavas occurrences in the Lycian Taurus, Karaburun occurrences in the western Anatolides and Karakaya

complex in the Pontides; e.g. Kozur 1999b; Kozur & ~enel 1999; Rosselet & Stampfli 2003; Stampfli et al. 2003) (Fig. 9); this is certainly one of the main difficulties inherent in the geology of Turkey (for a review of this problem, see Stampfli & Borel 2004). Elements related to this Late Cretaceous juxtaposition correspond to the following features. (1) South of the Tauric-Anatolian plate is the Pamphylian (Antalya) suture (Monod 1977; Gutnic et al. 1979), which certainly extends westward under the Lycian nappes, to connect with the Cycladic domain. Besides the lower Antalya nappes of Pindos origin and their associated ophiolites, the upper Antalyan units (Marcoux 1987) present strong affinities with Hu~lu type sequences found in the Beysehir nappes and Mersin m61anges; like the latter they lack key index fossils (until the Cordevolian) that could relate them to the Neotethys basin to the south. A northern derivation was proposed by

VARISCAN-ALPINE

65

Fig. 7. WesternTethysreconstructionsfor the Early Cretaceous, modifiedfrom Stampfli& Borel (2004). (For legend see Figs 1 and 2.)

Marcoux (1987) and Marcoux et al. (1989), but refuted by many workers because the parauthochthonous sequences between the Bey~ehir and Antalya nappes present a continuous upper Cretaceous to Eocene sequence, preventing the passage of the nappe in that region. However, a derivation from the west is possible in the context of lateral Cretaceous shortening as proposed here. A derivation from the east (Mersin) can also be considered, in view of the necessary large clockwise rotation of the Tauric plate in the Cretaceous. The Antalya nappes are locally sealed by a Mid- to Late Eocene platform (Gutnic et al. 1979). (2) North of the Tauric-Anatolian plate is the Izmir-Ankara suture s.l., more precisely the Tavflenli and Ktitahya-Bolkarda~ zones (e.g. Okay & Ttiyztis 1999; Okay et al. 2001b), marked by blueschist metamorphism around the Campanian, and eastward, to the Kir~ehir massif, which is regarded as a core complex (Fayon et al. 2001). Both areas were subducted during

the obduction of the Lycian ocean (syncollisional (c. 95 Ma) and younger (c. 72 Ma) post-collisional granites of Kir~ehir; G6nctio~lu et al. 1997), then accreted to the Pontides in Eocene times during the final closure of the Lycian ocean. Ophiolitic remnants found in the Kir~ehir basement (Floyd et al. 2000) could be regarded as belonging to the Late Jurassic obducted Vardar ocean. In this context, the presence or absence of an inner Taurides suture is still an open question.

Part II: review of the oceanic realms involved in the history of the European post-Variscan lithosphere In the following we review, in some detail, key outcrops and stratigraphic data mainly from the oceanic areas involved in the western and central Tethys geological evolution from Late

66

G.M. STAMPFLI & H. W. KOZUR

Fig. 8. Western Tethys reconstructions for the Late Cretaceous, modifiedfrom Stampfli & Borel (2004). (For legendsee Figs 1 and 2.)

Palaeozoic to Cretaceous times presented above (Figs 2-8). Most of these areas have now completely disappeared, but they represent the main cornerstones of the proposed reconstructions. Also, some of the data presented here are new and are from recent investigations in Greece, Turkey and Iran. Before discussing the outcrops, there is a necessity to clarify, once more, the nomenclature of Tethyan oceans (see Stampfli 2000); we offer the reader herewith a more thorough historical review of the Tethys concept, which completes similar efforts made by ~eng6r (1985b, 1998).

Oceanic nomenclature for Palaeotethys and Neotethys concepts (Fig. 10) The terms Palaeotethys and Neotethys are used in very different manners, often in disagreement with the priority of the

Palaeotethys-(Neo)Tethys space-time concept of Kahler (1939). Therefore, the priority definition of the terms Tethys, Palaeotethys and Neotethys is discussed here. The term Tethys was established by Suess (1875) for a large Palaeozoic to Tertiary ocean between Gondwana in the south and Eurasia in the north ( ' w e . . . must concede the extinction of a great Palaeozoic, Mesozoic and Tertiary ocean in south-western Eurasia' (Suess 1893, p. 185)). As correctly stated by ~eng6r (1998), Suess even recognized in his papers that the Tethys contained several oceanic troughs, which opened and closed at different times. Kahler (1939) established the Palaeotethys as a wide ocean (in that sense departing from the geosyncline theory) and its shelves north of the Tethys s.s., the latter opening later than the Palaeotethys. The axis (suture zone) of the Tethys sensu Kahler (1939) was shown along the Zagros Zone and from there eastward to an area roughly between the Indian craton and Tibet. Later, Kahler (1974), referring to his earlier paper (Kahler 1939), produced a

VARISCAN-ALPINE

67

Fig. 9. Western Tethys reconstructions for the Palaeogene, modifiedfrom Stampfli & Borel (2004). (For legend see Figs 1 and 2.)

map with a Palaeotethys Ocean several thousand kilometres wide between Gondwana and Eurasia, a very modern picture indeed, which corresponds perfectly to the Palaeotethys s e n s u St6cklin (1974, 1977), Stampfli (1978) and Kozur (1999b). The Neotethys concept as an ocean located south of the Palaeotethys, which opened during the Late Permian (i.e. much later than the Palaeotethys) was also introduced by Kahler (1939), but the term Neotethys was not used; instead, this was simply designated as Tethys (s.s.), which opened, according to Kahler, during the Late Permian. Heritsch (1940) followed Kahler, but assumed that the Tethys s.s. ('zentrales Mittelmeer' or 'Mittelmeer' as he named the Tethys s.s.) opened in the 'Sosiostufe', which he correctly correlated with the upper Word 'formation' (today Middle Permian, Guadalupian, upper Wordian Stage). Also, he pointed out, as Kahler (1939) did, that his 'zentrales Mittelmeer' opened south of the Palaeotethys. To a large part, the 'zentrale Mittelmeer' s e n s u Heritsch (1940) corresponds to the Neotethys.

Without any doubt, the original Palaeotethys-(Neo)Tethys concept of Kahler (1939) and Heritsch (1940) was not only a timerelated concept of different opening and closure, but also a concept of two spatially clearly separated oceans, of which the (Neo)Tethys was the southernmost. Later attempts to reduce these concepts to a simple time-concept with a term Neotethys used inconsistently for some geodynamically different oceans, which opened at different Permian-Mesozoic times and were situated sometimes north of the Palaeotethys (e.g. Vardar ocean, Intrapontide Ocean, Izmir-Ankara or Lycian ocean(s) Fig. 10), and sometimes south of it (the peri-Arabian ocean, the priority Neotethys), basically violate the original priority Palaeotethys-Neotethys concept and lead to confusion. (e.g. the Hallstatt-Oman ocean of Kov~ics (1993) in which the Meliata-Hallstatt ocean was considered to be a continuation of the 'Neotethyan' V a r d a r - I z m i r Ankara ocean as Oman is Neotethyan; likewise, the Permian to Recent Palaeotethys of Csontos & V6r6s (2004), regarded as the

68

G . M . S T A M P F L I & H. W. K O Z U R

,..O

~z

.~

~o ~'-~o

,zz

,=~ o .~ .~ 9

9,,,.,

09 r~

...~

VARISCAN-ALPINE southernmost European ocean-east Mediterranean domain, violates the priority concept). The term Neotethys was first introduced by Stille (1944a,b), for a geosyncline that existed from the earliest Pennsylvanian to Tertiary time. This definition has only a historical value and cannot be used in a modern sense for any ocean, independent of the fact that there is no Tethyan ocean that existed through the PennsylvanianTertiary time interval. However, Stille (1951) attributed the priority for the term Neotethys to Heritsch (1940). He did not agree with Heritsch (1940) that the Neotethys opened within the Mid-Permian, but considered that the Neotethys opened at the Early-Late Carboniferous boundary. Thus, despite the fact that Heritsch did not mention the term Neotethys, Stille (1951) assigned the Palaeotethys-Neotethys concept to Heritsch (1940) and therefore to Kahler (1939), from whom Heritsch borrowed it. The Palaeotethys-Tethys s.s. concept of Kahler (1939) is basically the same as the Palaeotethys-Neotethys concept of St6cklin (1974) and Stampfli (1978), the Palaeotethys being located between the Variscan active margin and the Cimmerian blocks, and the Neotethys between the latter and Gondwana (Figs 3 and 10). As these workers were the first to use the term Neotethys in agreement with the priority Palaeotethys-(Neo)Tethys concept of Kahler (1939) and Heritsch (1940), the term Neotethys has to be used only in this priority-supported modern sense. In the field (Fig. 10), Neotethys-derived material is now found in the exotic blocks, under the peri-Arabian ophiolitic nappes (e.g. the Semail nappe in Oman; Glennie et al. 1974; Lippard 1983; Pillevuit et al. 1997) and in a similar relative position in the Himalayan suture (Bassoulet et al. 1980; Reuber et al. 1987). In Cretaceous times, the large Neotethys ocean was replaced by intra-oceanic (SSZ)-type domains (Stampfli & Borel 2002), and the Neotethysrelated sequences would most often be found as the most external (southernmost; Fig. 10), lower structural units of the Tethyan zone (e.g. within the peri-Arabian ophiolitic zone, Ricou 1983).

P a l a e o t e t h y s evolution

Remnants of the Palaeotethys are well known from China and Thailand, where its opening time was within the Early Devonian and its closure within the Mid- to Late Triassic (e.g. Yin & Nie 1996). Before that, a Silurian rift basin with graptolitic shales had existed (e.g. Ratanasthien et al. 1999), as it is the case in many other portion of the rift. In the central and western Tethyan realm, the rifting-opening timing of Palaeotethys was similar, but remnants of Palaeotethys were often not recognized. Palaeomagnetic, sedimentological and faunal data show no major separation of Armorica from Gondwana by Palaeotethys before the late part of Early Devonian time (Stampfli et al. 2002b, and references therein). The closure of Palaeotethys took place after the Vis6an in Morocco, close to the MississippianPennsylvanian boundary in southern France (e.g. Montagne Noire), further to the east (Sicily, Mrzle vodice in NW Croatia, Slovenian Southern Alps immediately south of the peri-Adriatic Line) during Roadian time (early Guadalupian, earliest MidPermian time), and in southern Greece, Turkey, Iran and eastern Tethys in Mid- to Late Carnian time. The Palaeotethys suture in the central Tethyan realm is well exposed in the Fariman area close to Mashhad (NE Iran) and extends to the Nakhlak area in Central Iran, where it was transported and duplicated during Alpine time, following the extrusion and rotation of the Lut block (Fig. 10) (Davoudzadeh et al. 1981). Whereas in the Alborz (northern [ran) only the southern shelf and slope of Palaeotethys is exposed (Stampfli 1978), in the Fariman area (Bagheri et al. 2003), an oceanic lower to upper Permian accretionary sequence is found, including a large seamount, associated with mafic tufts between pelagic limestones and red cherts with Early Kungurian conodonts (Kozur & Mostler 1991). Below this sequence, black pre-Kungurian cherts occur. This oceanic

69

sequence is overlain by middle to upper Permian shales, pelagic limestones and cherts without volcanic rocks and finally by a flysch of Triassic age, which is unconformably overlain by Liassic shallow-water coral and bivalve-beating limestones. This typical Palaeotethyan sequence of varying metamorphic grade can be traced until the Nakhlak area in Central Iran (Bagheri et al. 2003), accompanied by some ultramafic rocks and in the higher part of the sequence by intermediate to felsic volcanic rocks. The next well-known remnants of Palaeotethys are found in the Tavas units of the Lycian nappes in southern Turkey (de Graciansky 1972; Kozur et al. 1998; Kozur 1999b; Kozur & ~enel 1999). At Agiliovasi Yayla, several tectonic slices potentially belonging to an accretionary prism sequence contain different parts of the Palaeotethyan succession. The oldest flysch lies in the tectonic highest position and consists of siliciclastic turbidites and lydites and large blocks of lydites, pelagic, partly cherty limestones and mafic volcanic rocks and tufts. The matrix yielded radiolarians and conodonts of Vis6an and Serpukhovian ages; the blocks and clasts are mostly of Tournaisian age although some are of Vis6an or Serpukhovian age. This flysch is unconformably overlain by thin conglomerates, sandstones and a thick sequence of shallow-water Kungurian to Guadalupian fusulinid limestone. The next lower slice consists of mid-ocean ridge basalt (MORB) with a few intra-pillow fillings of red pelagic limestone, nearly unfossiliferous (few badly preserved Carboniferous radiolarians). The two upper slices indicate that the Palaeotethyan sea-floor accretion began either in the Vis6an, and probably during the Serpukhovian or, most probably, the Early Bashkirian. The spreading axis was probably subducted as the MORB have no oceanic sediment cover, and spreading was therefore probably active during the subduction under the Vis4an-Serpukhovian accretionary complex. The following thick lower tectonic slice contains a seamount sequence of late Moscovian and Kasimovian age, consisting of basalts, tufts and fusulinid limestones. The conodonts from these limestone clasts (which include the pelagic genus Gondolella) indicate the Moscovian-Kasimovian boundary level and are, therefore, contemporaneous with a part of the seamount sequence. Most interesting is the presence of quartz grains and small quartz pebbles in the Kasimovian part of the seamount sequence, which indicate that the seamount was entering the trench at that time. The lowest tectonic unit comprises Carnian shales above Middle Triassic to Carnian carbonates, in the upper part with long-ranging Carnian-Norian foraminifers (determination R. Rettori, Perugia). The oldest beds of that unit are shallow-water limestones of late Permian age. This sequence represents the southern shelf of Palaeotethys, overthrust during the Cimmerian event by oceanic Paleoeotethyan sequences. A Late Triassic red continental sequence is found unconformably on several of the described units (De Graciansky 1972). Thus, in the Tavas Nappe system, the onset of Palaeotethyan subduction (not later than Vis4an), the subduction of the spreading axis (Late Serpukhovian to Early Moscovian), the subduction of a seamount (Kasimovian) and the closure time of the Palaeotethyan ocean and nappe thrusting on its former southern passive margin (Mid- to Late Carnian or Early Norian) can be dated. These new data confirm previous data from Turkey (Monod & Akay 1984) and Iran (Stampfli 1978), and recent findings in Crete (Stampfli et al. 2003), regarding the Eocimmerian tectonic event as taking place in the Late Carnian-Early Norian.

P a l a e o t e t h y a n f o r e a r c basins

Immediately north of the Palaeotethys suture zone, forearc basins are known in a few areas of the Carnic Alps, Dinarides, Hellenides, Taurides (e.g. Mrzle vodice, Chios, Karaburun, Konya), Pontides (Karakaya basin, treated in the section on the periEuropean marginal ocean; see below) and in Iran (Figs 4 and 10). In most instances, no typical accretionary prism units have

70

G.M. STAMPFLI& H. W. KOZUR

been found so far; they would have been removed by tectonic erosion and/or thrust by the arc or forearc units during the Eocimmerian collision. These forearc basins developed above the former northern passive margin of Palaeotethys characterized by shallowwater (mainly reef) and pelagic Silurian to early Devonian limestones and shallow-water mid-Devonian limestones. These basins started immediately after the onset of Palaeotethys subduction during the Late Devonian, reworking pelagic material from the accretionary prism (e.g. lydites) and siliciclastic material from the hinterland and the arc. The Carnic Alps development (e.g. Sch6nlaub 1985; Venturini 1990) fits this scenario: the forearc Carboniferous flysch basin (Dimon and Hochwipfel formations with some volcanic rocks) is sealed by a shallow-water sequence of Pennsylvanian-early Permian age (Auernig-Rattendorf-Trogkofel groups). Late Permian rifting is marked by the deposition of the Val Gardena and Bellerophon formations, which extend to the Pelagonian domain (De Bono et al. 2001). Other remnants of potential forearc basins are found in Crete (Stampfli et aL 2003) and southern Peloponnesus, showing a continuous record of pelagic fauna from Bashkirian to Carnian time (Krahl et aL 1982, 1983, 1986, 1988; Kozur & Krahl 1984, 1987), in the Mrzle vodice area (NW Croatia, in northern Slovenia immediately south of the Periadriatic Line, Aljinovid & Kozur 2003), in the Monte Amiata area of southern Tuscany and in western Sicily (Catalano et al. 1991). In most instances these sequences are shallowing upward after pelagic conditions and cannot be interpreted as synrift. In Mrzle vodice (NW Croatia) the Permian forearc sequence consists of a Roadian flysch (ammonoid dated) overlain unconformably by shallow-water, molasse-type, mostly red sandy sediments. The clasts in the siliciclastic flysch consist of intermediate volcanic rocks, and mostly of deep-water sediments such as lydites (partly distally turbiditic) with thermally altered Visran conodonts (conodont alteration index, CAI = 5) and of pelagic limestone with a typical deep-water Albaillellaria radiolarian fauna typical for the Asselian and Late Tastubian (Early Sakmarian); some samples are of Late Pennsylvanian age (Aljinovid & Kozur 2003). In Crete (Trypali, Phyllite Quartzite group) and the southern Peloponnesus, the slope and outer shelf of the Palaeotethys southern margin are potentially exposed, with pelagic beds of Bashkirian to Mid-Carnian age; these could be regarded as one of the very few remnants of the Paleotethys accretionary prism. In an upper structural position, the Tyros Beds of eastem Crete represent the youngest known forearc sequence of Mid- to Late Triassic age (Stampfli et al. 2003). A pelagic Permian sequence is found under the Triassic back-arc and forearc sequence (Violet Schiefer, Agrilos schists); the geochemistry of these rocks shows felsic characteristics (Champod & Colliard 2003). These sequences were probably imbricated during the Eocimmerian deformation phase, sealed by a molasse sequence of early Norian age that grades into the Tripolitza carbonate platform. In central Crete, part of the flexural bulge related to the Eocimmerian phase is found in the Talea Ori parautochthonous sequence (Epting et al. 1972). The bulge shows erosion of most of the Triassic sequences down to Permian units. A shortlived Late Triassic platform covered the bulge during its downflexuring, then tectonic inversion took place, creating an angular unconformity covered by the Late Norian platform (Kock 2003). In the area between the Bey Da~lan and the Tauric plate, the flexural bulge must have been cut during the Mid-Camian (Figs 4 and 5), to link the Neotethys to the Pindos back-arc. Before that, and until the Cordevolian (Early Camian), the Eocimmerian flexural bulge separated faunistically (conodonts, holothurian sclerites, ostracodes, sponge spicules) this area from the Neotethys to the south (Kozur 1999b, 2000). In Chios (Greece) and Karaburun (western Turkey), pelagic Famennian limestones and upper Famennian to Tournaisian cherts derived from the Palaeotethys forearc or sea floor have been described (e.g. Kozur 1997a). The Mississippian siliciclastic

deep-water turbidites gradually changed during the latest Visran and Serpukhovian into pelagic limestones and finally shallowwater limestones, sandstones and conglomerates, which persisted to Bashkirian time (Garrasi & Weitschat 1968; Caridroit et al. 1997; Kozur 1997a, 1998a). Such a gradual change from deepwater turbidites to shallow-water sediments is typical for forearc basins. After a long gap, the Bashkirian shallow-water sequence was overlain by early Triassic shallow-water limestones grading into Spathian (Late Olenekian) to Cordevolian deep-water sediments (pelagic limestones, cherts) with many, mainly intermediate volcanic rock units. This synrift sequence of the Maliac back-arc basin (Rosselet & Stampfli 2003; Rosselet et al. 2003a,b; Stampfli et al. 2003), is overlain by a middle Carnian to Cretaceous shallow-water carbonate platform, which was finally fragmented to form the upper Cretaceous Bornova mrlange. In central Turkey, in the Konya region, a metamorphic Palaeozoic forearc type sequence is found (Kozur 1999a; Eren et al. 2004). Platform and pelagic limestones of Silurian-Devonian age are found in m~langes together with cherts and basalts (some of MORB type) potentially derived from the Palaeotethyan sea floor. The youngest blocks are Early Permian in age. The whole sequence (Slzma group) is cut by basaltic and andesitic dykes and dolerite bodies, pointing to a forearc setting evolving to an arc setting. The Late Permian-Triassic to Cretaceous Ard~qll group rests unconformably on older sequences and is regarded as representing the opening of a back-arc basin within the former active margin, during the Triassic as in Karaburun. Other metamorphic Palaeozoic forearc sequences, situated immediately north of the Palaeotethys suture, have recently been found in Central Iran, in the Anarak-Nakhlak-Jandaq area, which was previously mainly regarded as Precambrian (Fig. 10) (Bagheri et al. 2003). First studies yielded conodonts from a Late Devonian pelagic sequence, associated with younger coral-bearing shallow-water limestones of Late Carboniferous age, and volcanic rocks of arc affinity. A unit of early Palaeozoic to late Devonian age consisting mainly of metamorphosed rocks including ophiolitic rocks, pelagic sediments, flysch-like deposits and shallow-water limestones, occurs in the Anarak region. Southward it passes into Permian deposits associated with volcanic rocks of arc affinity, which are now juxtaposed with a complete Palaeozoic sequence of Cimmerian type, thus clearly marking the Palaeotethyan suture in that region.

Neotethys evolution

In the westernmost part of Variscan Europe the Palaeotethys closure had been completed by the Mississippian-Pennsylvanian boundary, but from the Slovenian part of the Southem Alps to southern Tuscany and western Sicily (Sicanian basin), closure of the Palaeotethys was in the Roadian (lower Guadalupian) (Catalano et al. 1991; Kozur 1999b). From Greece to Iran, the final closure took place in the Late Triassic (e.g. Crete: Krahl et al. 1996; Stampfli et al. 2003; Turkey: Kozur et al. 1998; Kozur & ~enel 1999; NE Iran: Bagheri et al. 2003). In China the Songpan ocean can be regarded as a back-arc derived from the Palaeotethys; the pelagic sedimentation there lasted until the Late Triassic and the Jurassic sequence is missing (e.g. Yin & Nie 1996; Yin & Harrison 2000). Subduction of the Palaeotethyan slab ended as Gondwana and Laurussia were locked together through the Alleghanian orogen. However, and despite the fact that Palaeotethyan slab detachment certainly took place in the orogenic areas (e.g. in Morocco, witnessed by the emplacement of post-collisional Early Permian granites; Amenzou & Badra 1996), slab-pull forces did not disappear eastward, and finally succeeded in detaching a continental ribbon from Gondwana, the Cimmerian terranes, accompanied by the opening of Neotethys. It is interesting that the youngest pelagic sediments found on Paleotethyan MORB in Fariman (Kungurian), before the ridge was finally subducted, have an age comparable with the

VARISCAN-ALPINE assumed age of earlier sea-floor spreading in the Neotethys (Kungurian-Roadian). Clearly the space lost in Palaeotethys was compensated by the opening of Neotethys. The westernmost occurrence of Neotethyan series is found in Sicily, in the Sicanian basin. There, the Late Roadian opening of Neotethys took place within the Late Artinskian to Early Roadian Palaeotethyan accretionary complex-foreland basin. Elsewhere, the Neotethys opened within the southern margin of Palaeotethys (Fig. 3). The opening took place generally within the Late Cisuralian (late Early Permian), indicated by palaeopsychrosphaeric ostracodes in the Roadian and Wordian (early to mid-Mid-Permian) of Oman, which require a wide connection to the world ocean. Thermal subsidence and flooding of the rift shoulders of the southern Neotethyan passive margin have been studied in detail, from Australia to Sicily (Stampfli 2000; Stampfli et al. 2001a; Borel & Stampfli 2002). This flooding took place during the Wordian in Oman (Pillevuit 1993; Pillevuit et al. 1997), allowing us to place the onset of sea-floor spreading within the Artinskian or just after. Only in the Australian sector of Neotethys did sea-floor spreading possibly begin earlier, during the earliest Permian or even Late Pennsylvanian (Yeates et al. 1987; Borel & Stampfli 2002), and it took place locally under an ice sheet. Relatively undisturbed northern Neotethyan margin sequences are known only in a few places. In Crete, in the Talea Ori massif, the Permian sequence (Krnig & Kuss 1980) presents a typical synrift evolution: rapid subsidence, significant clastic input and rapid flooding, followed by the progradation of a Midto Late Permian platform (Kock 2003). Other occurrences are found in Afghanistan (Vachard 1980): in the Central Mountains units (Blaise et al. 1977) and in the Kabul Block (Mennessier 1977). Again, typical Permian synrift sequences, marked by very rapid subsidence, angular unconformities and, locally, rapid appearance of pelagic facies have been described. The lower part of the Afghan sequence is also characterized by a coldclimate, peri-glacial fauna (Termier et aL 1973). The subduction of the Neotethys, along the Iranian transect, began between the Late Triassic and Liassic (Berberian & Berberian 1981). Closure occurred within the Late Cretaceous following the opening-obduction of the Semail-Troodos suprasubduction ocean (e.g. Robertson & Searle 1990; Robertson & Xenophontos 1993). To explain this opening, a failure of the Neotethys midocean ridge is invoked (Figs 6 and 7); again, this should be related to the onset of east-west shortening affecting the Tethyan area during the Cretaceous and major plate rearrangement at that time as a result of the break-up of Gondwana (Hauser et al. 2002; Stampfli & Borel 2002). Because of the systematic opening of intra-oceanic back-arc basins all along the northern margin of Neotethys in Cretaceous times, pure Neotethyan remnants are very rare; they are found, for example, under the peri-Arabic ophiolites from Oman to Cyprus (Fig. 10), as exotic material, dominated by Triassic seamounts. The detailed and systematic study of these exotic blocks in Oman allowed the reconstruction of the southern margin of Neotethys with some degree of confidence (Pillevuit 1993; Pillevuit et al. 1997). A remnant Semail ocean is still subducting under the Makran zone (Sea of Oman), and Neotethyan oceanic crust is still subducting under the Calabrian and Aegean arc today (Fig. 10).

Peri-European marginal oceans

Peri-European oceans developed on the border of Baltica and accreted Avalonia during the Variscan cycle (Fig. 1), and then developed on the border of accreted Variscan terranes during the Cimmerian and Alpine cycle (Figs 4 and 5). The R h e n o - H e r c y n i a n basin (Fig. 1) is characterized by significant volcanism since the Early Devonian (e.g. Walliser 1981;

71

Ziegler 1988), whose geochemical characteristics (Floyd 1995) show a purely ensialic extensional nature, and absence of any subduction-related signature. On the other hand, MORB have been found in some places (Lizard, Giessen, Harz) and point to sea-floor spreading, which probably started in the Emsian. From the Namurian onward, the basin northern margin became a flexural basin in the foreland of the advancing Variscan nappes; before this, the sedimentary records do not show evidence of any tectonic event. The term PaphIagonian Ocean was first used by Kozur (1999b) and Kozur et al. (1999). It corresponds to a Palaeozoic ocean in the middle Pontides (Fig. 10), which was formerly regarded as a potential Palaeotethyan remnant (~engrr & Yllmaz 1981; ~engrr 1984, 1985a; ~engrr et al. 1984; Gen~ & Yllmaz 1995; Yilmaz et al. 1997). The existence of this ocean is indicated by exotic olistoliths and pebbles in the Beykoz Formation, where they occur together with clasts from the Zonguldak Terrane, which was attached to stable Europe during the earliest Devonian (Fig. 1). This indicates a position of the Paphlagonian Ocean immediately adjacent to the Late Palaeozoic stable Europe; therefore, it was a continuation or equivalent of the Rheno-Hercynian ocean. This is confirmed by the occurrence of the same exotic clasts in the Eskiorda tectonic unit of SE Crimea (Fokin et al. 2001) separated in the present geological setting by the younger Black Sea (Fig. 10). In the middle Pontides, Mississippian and Pennsylvanian black radiolarite clasts, Late Pennsylvanian pelagic limestones with conodonts, and pelagic Permian limestones with ammonoids, pelagic ostracodes and conodonts, indicate an opening not later than Early Carboniferous (Kozur 1999b; Kozur et al. 1999, 2000), but probably in the Devonian, possibly at the same time as for the North Caspian basin (Fokin et al. 2001). In Crimea, only Permian pelagic limestones were found (Kotlyar et al. 1999), but there, microfaunas of pelagic limestones and radiolarites were not investigated. The youngest pelagic rocks are of Dorashamian age (latest Permian), indicating closure close to the Permian-Triassic boundary. Debris-flow blocks with Wordian to Dzhulfian matrix may indicate subduction or inversion until that time. This ocean was the connecting link between the Rheno-Hercynian and the Khanty-Mansi and Mugodzhar (Sakmarian basin) oceans east of Laurussia through the North Caspian basin (Russian-ITLP 1997; Fokin et al. 2001). After the closure of the Rheno-Hercynian ocean, connection with the Khanty-Mansi ocean (~engrr et al. 1993) and the Mugodzhar ocean remained open until the end of the Carboniferous (Fig. 2). Then, during the Permian, faunistic connection to the east occurred through the narrow Caucasian Dizi Basin (Figs 3 and 10) (Adamia & Kutelia 1987). This explains why palaeopsychrosphaeric ostracodes indicating deep-water conditions with free and broad connection to the world ocean (connected to cold bottom water currents) disappeared during the Mid- and Late Permian whereas pelagic deep-water ostracodes (free-swimming cypridinids and benthic ostracodes) were still common. The Kiire ocean comprises the largest part of the 'Palaeotethys' sensu ~engrr (1979, 1985a, b), but, as discussed above, it does not correspond to it, as it is located within the Variscan domain. The Ktire oceanic unit consists of ophiolites and a main sedimentary unit, the Akgrl Group (Ketin 1962), in which the dark middle Carnian to middle Jurassic siliciclastic turbidites and olistostromes represent a widespread flysch sequence characteristic of the active southern margin of the Ktire ocean. The middle Carnian to lower Norian part of the matrix can be dated by various Torlessia species (Kozur 1998b; Kozur et al. 2000). The olistoliths contain basalts and ultramafic rocks as well as shallow-water Scythian and pelagic, slope and shallow-water sediments (limestones and radiolarites) of Anisian and Ladinian age. The oldest pelagic rocks contain Chiosella timorensis, the conodont guide form of the basal Anisian. In the Neogondolella regalis Zone, the second conodont zone of the Anisian, the first palaeopsychrosphaeric deepwater ostracodes appear, which indicate a broad connection to

72

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the world ocean. The opening of the Ktire ocean is placed within the Late Olenekian, indicated by the first appearance of similar pelagic sediments in the Kotel Zone and in the Dobrudzha Trough (a western continuation of the Ktire ocean; Fig. 4), accompanied by the emplacement of Scythian enriched MORB (E-MORB) (Nicolae & Seghedi 1996; Seghedi 2001). The presence of siliciclastic flysch in the southern margin and of a passive margin with Late Triassic Hallstatt Limestone in the north at the same time (Kozur et al. 2000) indicates a southwarddirected subduction, as already assumed by Seng6r & Yllmaz (1981) and ~eng6r (1984). The northern slope of the Ktire ocean is represented by the ~alqa Unit, which is characterized by Pelsonian to uppermost Norian or lowermost Rhaetian Hallstatt Limestones that have the same facies succession as in the Hallstatt type area in Austria (Kozur et al. 2000). Through the Dobrudzha, north of the Moesian Platform, a trough extends to a proto-Pieniny basin (Figs 4 and 5), from which lower Anisian to Jurassic deep-water sediments and mafic volcanic rocks are known as exotic pebbles in conglomerates of Albian and Late Cretaceous age derived from the exotic Andrusov Ridge (Birkenmajer 1988), and on which the material of the protoPieniny basin was originally thrust as nappes (Birkenmajer et al. 1990). From the proto-Pieniny basin, pelagic fossils (ammonoids, conodonts) and articulate brachiopods spread to the Germanic basin during the Early Anisian (starting in the Aegean), through the Eastern Carpathian and Upper Silesian Gate, long before these fossils were present in the Western Carpathians, Northern Dinarides and Alps. In the Jurassic, the proto-Pieniny basin became part of the eastern Alpine Tethys northern margin; it was probably tectonically inverted in the Albian as a response to the Austrian orogenic event affecting the Austroalpine and Balkan domains. The Karakaya basin, often regarded as a long-lasting oceanic realm, is a potential Palaeotethys forearc basin, located in the Variscan Sakarya Terrane between the Palaeotethys and Ktire oceans, and related to the collapse of the Variscan cordillera and northward-directed subduction of Palaeotethys (Fig. 4). The onset of this narrow basin occurred in the Late Permian, in some places within the Early Scythian (~eng6r & Yllmaz 1981; ~eng6r & Hsu 1984; ~eng6r 1985a; Kozur & Kaya 1994; Kozur et al. 1996; Kozur 1999b). Upper Permian to Cordevolian series are characterized by pelagic limestones, red cherts, and mainly mafic volcanic rocks. The extensional regime in the Karakaya basin changed at the beginning of the Mid-Carnian to a compressional regime. The basin was so narrow that from the beginning of the Mid-Carnian, siliciclastic turbidites and olistostromes were deposited in the entire area. This sedimentation continued until the Mid-Norian, and in the Rhaetian-Liassic, the Karakaya sequence was overlain unconformably by shallow-water sediments. During the transpressional regime both subducted parts of the Palaeotethys ocean and older units of Variscan and preVariscan age were exhumed and juxtaposed to the Karakaya Unit. This led partly to the assumption that the Karakaya basin was the remnant of a long-lasting Palaeozoic-Triassic ocean, the Palaeotethys (e.g. Pickett et al. 1995; Usta6mer & Robertson 1995, 1997, 1999; Pickett & Robertson 1996). Variscan and older metamorphic rocks, non-metamorphic Silurian to lowermost Pennsylvanian (lower Bashkirian) pelagic limestones and radiolarites, and the Orhanlar Greywacke Unit (Variscan flysch, turbidites, debris flows and olistostromes, in the type area with lower Carboniferous to lower Bashkirian matrix and olistoliths; Okay & Mostler 1994) belong to the Variscan basement (Sakarya domain) in which the Karakaya basin developed. Partly Pennsylvanian, and Lower, Middle and Upper Permian shallow-water limestones belong to the cover of the Variscan basement. Non-metamorphic or slightly metamorphosed Late Permian to Cordevolian pelagic limestones, red cherts, tufts and mafic volcanic rocks belong to the extensional sequence, whereas middle Camian to middle Norian siliciclastic turbidites, sandstones, greywackes and olistostromes (mainly

Diskaya Unit; junior synonym: Hodul unit) belong to the tectonic inversion sequence of the Karakaya forearc. This change was caused by the closure of the Palaeotethys and contemporaneous onset of the southward-directed subduction of the Ktire ocean. The Niltifer Unit (Late Permian-Early Triassic blueschist: shallow-water, in the upper part, pelagic limestones, some black shales, many mafic metatuffs and metavolcanic rocks; Fig. 4) belongs to the exhumed subducted material from Palaeotethys, probably a former seamount and its surroundings (Usta6mer & Robertson 1999) that were possibly responsible for the Late Triassic closure of the basin (Okay 2000). The Niltifer Palaeotethyan H P - L T remnants were dated at 203-208 Ma (Okay & Moni6 1997; Okay & Ttiyztis 1999), which we interpret as exhumation ages, possibly related to the opening of the Izmir-Ankara rift. The whole Karakaya sequence is unconformably overlain by Rhaetian-Liassic shallow-water sediments. The M e l i a t a - H a l l s t a t t ocean of the Eastern and Western Carpathians and Northern Alps was located in a western prolongation of the Ktire ocean (Fig. 4). Former direct connections are apparently missing as a result of major lateral displacements around the Moesian promontory. However, a connection of the two basins through the south Moesian aborted rift (Georgiev et al. 2001) did exist (Kotel-Stranja rift: Chatalov 1991; Figs 4 and 5). The widespread occurrence of palaeopsychrosphaeric ostracodes in the Middle Triassic to Liassic units of the Meliata-Hallstatt ocean and its slope and outer shelves (Kozur 1991a) indicate a broad enough deep-water connection (below 500 m water depth) with the world ocean. The easternmost remnants of the Meliata-Hallstatt domain occur as exotic blocks of the Transylvanian nappes. There, one branch extended to the Pieniny basin (see above), and the other to the Meliata-Hallstatt ocean. As well as exotic blocks of the Meliata domain, other deepwater sediments are present among the exotic material of the Transylvanian nappes, including pelagic rocks of post-Oxfordian age, which do not occur in the Ktire or the Meliata ocean and were derived from the Vardar basin (see below). As in the Ktire ocean, the lower Anisian consists of pelagic limestone and pillow lava, but pelagic upper Olenekian rocks have not yet been found. In the typical Meliata domain of the Western Carpathians and Eastern Alps, a rifting phase occurred during the Late Permian (Fig. 3). In the Alps, thick hypersaline sequences, with some mafic volcanic rocks, yielded Dzhulfian ostracodes from dolomite intercalations. The Scythian and a large part of the lower Anisian sequences are characterized by shallow-water sediments. The earliest pelagic conditions began in the Late Bithynian. Pelagic limestones with some mafic volcanic rocks characterize the Pelsonian; sea-floor spreading began in the Illyrian. The Ladinian and Cordevolian are characterized by widespread pillow lava and red radiolarites. As in the Ktire ocean, sea-floor spreading suddenly stopped in the Mid-Carnian. Late Camian and Norian thermal subsidence caused the drowning of the outer shelves, characterized by a transition from shallowwater limestones to Hallstatt Limestones. The subduction of the Meliata ocean began during the Jurassic (Fig. 6) and its closure occurred within the Oxfordian, as in the Kfire ocean, where, however, the subduction began earlier (see above). The westernmost occurrence of the Meliata oceanic development is found in the eastern part of the Northern Calcareous Alps (Kozur 1991a, b; Mandl & Ondrejickovfi 1991; Kozur & Mostler 1992), from the Florianikogel Nappe in the east, including dismembered ophiolites and red radiolarite, to the Hallstatt area, where the oceanic part is subducted but the northern slope and outer shelf with Hallstatt limestones is preserved. The continental Late Permian rifting continued further to the west, to the area of Hall (east of Innsbruck). Further west the Partnach Basin (Lechtal Nappe) represents a Mid-Triassic restricted basin. Still further west, the only Ladinian-Cordevolian lavas of the Northern Alps, outside the Meliata domain, occur in the Arlberg Beds (southern Lechtal nappe). In the Jurassic, the Meliata subduction

VARISCAN-ALPINE zone in that area changed into a transform (strike-slip) zone marked by repeated formation of breccias (Lower and Middle Jurassic Eisenspitz breccia). In the M a l i a c o c e a n , sea-floor spreading started at the beginning of the mid-Carnian (Ferri~re 1974, 1976, 1977; De Bono 1998; De Bono et al. 1999), when it stopped in the K~re and MeliataHallstatt oceans (Kozur 1991a, b). This is locally connected with a pronounced shoulder uplift in the outer shelf of the Maliac domain. Deep-water Upper Olenekian to Cordevolian sediments (pelagic limestones and radiolarites connected with widespread intermediate volcanism) are overlain by middle Carnian shallowwater carbonates. This 'Maliac signal' is the opposite of the 'Meliata signal' (thermal subsidence of the slope and outer shelf). The Maliac ocean was subducted during the Jurassic, together with Meliata (Fig. 6). The northwestern evidence of Maliac remnants is found in Mts. Kalnik and Medvednica (northwestern Croatia; Halamid & Gorican 1995), but the 'Maliac signal' can also be observed in the eastern Drauzug. The 'Maliac signal' is also clearly recognizable in the Karaburun peninsula in Turkey (Stampfli et al. 2003), but its eastern continuation was nearly totally subducted in the IzmirAnkara Belt. Only very subordinate Late Carnian pillow lava and cherts of Maliac origin are preserved at the northern edge of the Tauride-Anatolide Platform, and were taken by G6ncfio~lu et al. (2001) and Tekin et al. (2002) as evidence for a Late Carnian opening of the Izmir-Ankara ocean. However, the latter opened only in the Jurassic (see below), and the Anatolide block was not the southern margin of this ocean; the two areas were juxtaposed during the Late Cretaceous (Fig. 6) (Stampfli & Borel 2004). T h e P i n d o s d o m a i n , located between the Gavrovo-Tripolitza (Greater Apulia) and the Pelagonian units (Fig. 5), was a deepwater basin from Late Triassic time. The Pindos-Olonos zone comprises a well-known sedimentary continuous sequence of pelagic facies of Late Triassic to Palaeocene age followed by a Palaeocene-Oligocene flysch (Fleury 1980; De Wever & OrigliaDevos 1982; De Wever & Cordey 1986; Richter & Mfiller 1993a,b; Richter et al. 1993). The basal part of the Pindos stratigraphic column, as shown by Aubouin et al. (1970) and Fleury (1980), is characterized by a formation, mainly of Carnian age, called 'D~tritique triasique'. Degnan & Robertson (1998) proposed the name Priolithos for this formation, and have shown the orogenic origin of the re-sedimented clasts. We attribute the main source of detritus to the uplifted continental basement located to the SW (Sitia microcontinent and Tyros forearc-future Tripolitza), corresponding to the most external part of the Variscan cordillera. The Mid- to Late Carnian age of this flyschoid formation indicates a possible mixed origin: syn-rift related to the Pindos back-arc opening and tectonic inversion related to the Eocimmerian event. Some remnant ophiolitic nappes of disputed origin are found in the Pindos realm. Triassic Pindos lavas were reported mainly in the Kerassies region (Robertson & Degnan 1992; Pe-Piper & Hatzipanagiotou 1993; Pe-Piper 1998), within the Avdella M~lange (Jones & Robertson 1990), in the 'Formation ?~blocs' (Pe-Piper & Piper 1991) and in the Alpine metamorphic belt of the Cyclades (Papanikolaou 1989). Generally speaking, these 'Jurassic or Triassic Pindos ophiolites' are allochthonous nappes (e.g. Pindos and Vourinos ophiolites) overthrust onto the Pindos and Pelagonian realms (Smith et al. 1975; Smith & Spray 1984; Robertson 1990; Robertson et al. 1990; Ross & Zimmerman 1996), first during the Eohellenic Late Jurassic phase (Vergely 1984) and later during the Alpine shortening, creating imbrication with Pindos slices through out-of-sequence thrusting and back-thrusting; these ophiolitic masses are rootless, and before the Alpine shortening they formed large ophiolitic klippen. There are enough sedimentological and stratigraphic data (Baumgartner 1985; Ferribre 1985; ThiSbault & C16ment 1992; Baumgartner et al. 1993; Thi6bault et al.

73

1994; De Bono 1998; De Bono et al. 2001) to demonstrate an eastern (Vardar) provenance for all the 'Jurassic Pindos ophiolites', an observation valid also for the Albanian (Collaku et al. 1993; Hoxha 2001) and Croatian ophiolites (Pamic 2002). The SSZ affinity of the 'Pindos' and Vardar ophiolites implies the subduction of an older ocean, by roll-back, for their formation. No Pindos exotic blocks have been found so far at the base of the 'Pindos ophiolites'; on the contrary, the exotic material is always of Maliac type (e.g. Ladinian Hallstatt type limestone or radiolarites). The disappeared ocean was not the Pindos domain that records everywhere a complete pelagic sequence from Late Carnian to the Palaeogene flysch, undisturbed by the Late Jurassic obduction of the ophiolites on the Pelagonian landmass, which is easily explained if one admits an eastern origin for all the Jurassic SSZ-type Greek and Albanian ophiolites, and their obduction on the eastern Pelagonian margin. Our own investigations (Payer 2001; Bellini 2002; Stampfli et al. 2003) have shown the presence of Early to Mid-Carnian basaltic lava flows and tuffites in the Pindos basal sequence (Priolithos formation) in eastern Crete (Kalos Potamos) (also present in the Pindos Mountains; Kozur & Mock, unpubl, data). REE distribution and other discriminative diagrams show a dominant volcanic arc affinity with a minor E-MORB signature, confirming a Carnian back-arc position for the Pindos ocean. No proof of spreading younger than Carnian has been found so far. Part of the Triassic arc was abandoned on the northern margin of the Pindos, in the Vardoussia region, characterized by sub-alkaline to calc-alkaline and shoshonitic (Mt. Ghiona) lavas with subductionrelated geochemical signatures (Pe-Piper & Mavronichi 1990; Lef6vre et al. 1993; Pe-Piper 1998). Pe-Piper (1998), following the model proposed by Robertson et al. (1991), abandoned the subduction hypothesis and proposed, on the basis of new isotopic data, an extension-related origin for the Triassic volcanic rocks, an interpretation that is also compatible with our model, in which the Palaeotethys subduction is accompanied by large-scale extension (1000-2000 kin) of the Eurasian upper plate (e.g. Fig. 11). Recent investigations in the Chamezi area of eastern Crete (Champod & Colliard 2003) have shown the presence of Early to Mid-Triassic rift-related series, potentially belonging to the southern margin of the Pindos domain. They can be correlated with the Tyros Mid-Triassic forearc sequences of Vai through a common Late Permian pelagic substratum and to the Pindos remnants of Kalos Potamos through a similar pelagic Late Triassic sequence. The Chamezi-Pindos rift opened within an Early to Mid-Triassic arc and was separated from the Vai forearc basin by a ridge composed of the Variscan basement units of the Sitia microcontinent. Eastward, the Pindos sequences can be extended to the Antalya domain, as proposed by Brunn et al. (1976), formerly located north of the Bey Da~lan Cimmerian platform, as discussed above (Figs 5 and 10). The lower and median Antalya nappes (e.g. AlaklrGay nappe) are characterized by Carnian deep-water facies (Pamphylian basin) rich in detritus and locally in volcanic rocks (Carnian pillows; Marcoux 1970; within-plate basalt (WPB) to MORB according to Robertson & Waldron 1990), similar to the lower Pindos sequences (Dumont et al. 1972; Gutnic et al. 1979; Marcoux 1987; Stampfli et al. 2003). One characterisitic of these sequences is the presence of Late Triassic clastic material (locally reworked granitic pebbles) marking the Eocimmerian event. They are followed by pelagic sequences locally lasting until the Late Cretaceous. These nappes are separated from the upper Antalyan units by a Cretaceous ophiolite that represents either the western end of the Semail-Troodos intra-Neotethys back-arc (SSZ) system or an independent late Cretaceous oceanic corridor related to the rotation of the Tauric plate. Part of the subducted Pindos domain is certainly also present in the Cyclades metamorphic domain, comprising the Dilek peninsula HP rocks of Western Turkey (Bozkurt & Oberhaensli 2001, and references therein). We also assign to an eastern continuation of the Pindos-Antalya domain (southern slope and shelf) the Hu~lu and Boyali Tepe units

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G.M. STAMPFLI & H. W. KOZUR

of the Central Taurus (Gutnic e t al. 1979), from the B e y s e h i r H o y r a n - H a d i m nappe (Andrew & Robertson 2002), assigned to so-called 'Neotethys northern branch' remnants by Andrew & Robertson. In the Hu~lu Unit, thick middle Carnian tufts and basalts with a few intercalations of cherty limestones are present, overlain by Upper Triassic cherty limestones and cherts with

Jurassic radiolarians (Kozur 1997b; Tekin 1999). An eastern continuation of the Hu~lu Unit was found in the exotic units at the base of the Mersin ophiolite (Parlak & Delaloye 2000; Parlak & Robertson 2004). Our own investigations (Masset & Moix 2004) have shown the following sequence: thin Hallstatt limestone and mafic tufts dated as Julian to early Tuvalian, overlying calci-turbidites

Fig. 11. Palinspastic scheme on a Carpathian-Libyan transect, from Carboniferous to Early Jurassic times. Location of transects are shown in Figures 2-6. The ages in the chronostratigraphic chart were taken from Gradstein et al. (2005) for the Carboniferous, Early Penrtian and Jurassic, and from Kozur (2003a, b) for the Mid- and Late Permian and Triassic. (For abbreviations, see Fig. 1.)

VARISCAN-ALPINE and debris flows of mid-Tuvalian to latest Norian age, followed by deep-water Jurassic sediments including Bathonian-Callovian radiolarites, pelagic cherty limestones of Early Cretaceous age, and Aptian to Cenomanian radiolarites covered by a wildflysch marking the flexure of the basin in front of the advancing ophiolite. The whole m~lange series rests now on a younger Late Cretaceous flexural series (Campanian) and m61ange units whose age extends to earliest the Maastrichtian; both were certainly re-displaced during the Palaeocene, after the ophiolite obduction that generally is sealed by a Late Maastrichtian platform nearly everywhere in the Taurus. The Gtilbahar Nappe of the Lycian domain with lower to middle Norian cherty limestones, cherts and tufts may also be derived from the southern margin of the Pindos ocean, but exposed Carnian rocks are not known from there. The western continuation of the Pindos is represented by the Budva domain of the external Dinarides (Cadet 1970; Gorican 1994). Upper Triassic rocks (pelagic Halobia limestone) are underlain by sandstones, marls, shales, limestones, cherts, tufts, basalts and intermediate volcanic rocks, for which a Mid-Triassic age is indicated, but without good faunal evidence. This lower sequence may correspond to the Palaeotethyan accretionary-forearc complex. This westernmost part of the Pindos basin could have been located on a major transcurrent fault, and locally inverted in Late Triassic times during the Eocimmerian phase and later during lateral displacement between Africa and Eurasia.

Alpine oceans (Figs 6 - 9 ) The Vardar ocean and I z m i r - A n k a r a - S o u t h

Caspian ocean

opened simultaneously in the Liassic (e.g. G6rtir et al. 1983). The Vardar opening was due to the Jurassic northward-directed roll-back of the Maliac ocean, and the Izmir-Ankara opening was related to the Ktire and Neotethys subduction. Both oceans finally closed within the Late Cretaceous or earliest Palaeocene, except for the South Caspian basin, the closure of which probably started in Late Miocene time and is still continuing. As noted above, we relate all Jurassic ophiolites of the Dinarides (Pamic et al. 2002) and Hellenides to the Vardar ocean, as they have very uniform fauna and geochemistry (SSZ), and similar Maliac-Meliata exotic material at their sole. The amphibolites found at the base of the ophiolites and generally used to determine their direction of emplacement are generally 20-30 Ma older (c. 160-170 Ma) than the final emplacement on the Pelagonian margin in latest Jurassic or Early Cretaceous time (150140 Ma). This amphibolitic metamorphism took place at the spreading centre of the SSZ ocean and subsequent rotation of large ophiolitic masses took place before final emplacement on the Pelagonian domain. We extend these Vardar attributions to the Southern Apuseni Mountains, Severin (Romania), Btikk Mountains (Hungary), Kalnik and Medvednica (Croatia; Halamid et al. 1999; Babic et al. 2002) ophiolites. All these ophiolites are covered by post-obduction middle to upper Cretaceous sediments, and together with the remnant Vardar ocean, they were accreted or subducted under the Serbo-Macedonian active margin to form the Sava-Vardar-Axios suture of Palaeocene age (Pamic 2002). This Tertiary accretionary process re-displaced ophiolites obducted in Late Jurassic-Early Cretaceous times, and, together with Alpine out-of-sequence thrusting and back-thrusting and lateral displacement, is at the origin of controversies regarding the number of ocean between Adria and the Balkan region. We consider here that the apparent palaeogeographical complexity should be related to complex structural patterns and that the former palaeogeography was simple, with a single Vardar ocean separating the two continental masses in Jurassic and Cretaceous times (Fig. 7). The Lycian ocean opening followed the east-directed roll-back of the Izmir-Ankara ocean in the general large-scale east-west shortening of the Tethyan area in the Cretaceous. It largely obducted its southern margin (Tauric-Anatolian plate) and,

75

locally, its northern margin (Pontides-Sakarya, e.g. Beccaletto 2004), producing numerous ophiolitic remnants of Cretaceous age with metamorphic sole ages centred around 95 Ma (Robertson 2002) and an SSZ geochemistry. Other ophiolites from southern Turkey (e.g. Baer-Bassit, Hatay, Guleman) are SSZ Neotethysrelated back-arc basin (part of the Semail ocean) obducted around the Arabian promontory and extending to Cyprus (Robertson 2002, and references therein). The Alpine Tethys opened in Early to Mid-Jurassic times, nearly contemporaneous with the Vardar opening, but related to the opening of the central Atlantic and break-up of Pangaea; pelagic sedimentation started in the Bajocian (Baumgartner 1987). The Alpine Tethys extends from the Maghreb through Italy (Ligurian-Piedmont) to the Penninic ocean. Spreading started in Late Toarcian to Aalenian times in the western part, and only in Aalenian to Bajocian time eastward (Bill et al. 1997; Stampfli et al. 1998, 2002a, and references therein). Further east, the Alpine Tethys is represented by the Vahic Ocean of Mahel' (1981), which was later overthrust by the Veporic nappe of the central Western Carpathians and higher nappes (e.g. Inovec Mountains: Plagienka et al. 1994; Kozur & Mock 1996; eastern Slovakia: Sotfik & Spisiak 1992; Kozur & Mock 1997). The Pyrenean r/ft opened around the M0 magnetic anomaly, when the central Atlantic started to extend between Iberia and Newfoundland and when spreading in the western half of the Alpine Tethys ended. The Pyrenean rift (e.g. Peybern~s & Souquet 1984) is characterized by the denudation of continental mantle in the French Pyrenees (Fabries et al. 1998), accompanied by a mid-Cretaceous thermal event . This opening is followed by the opening of the Biscay ocean and the formation of passive margins (Boillot 1984), related to the accelerated rotation of Spain-Africa (e.g. Olivet 1996) following the subduction of the Vardar ocean. This rotation entailed the onset of Alpine deformation along the European margin marked by large-scale tectonic inversion that started during the Turonian in the Pyrenees and Provence, and in the Maastrichtian in the Western Alps (Stampfli et al. 1998, 2002a). Pyrenean deformation is mainly sealed by Late Eocene deposits in the north, and Oligocene or younger deposits in the south. As seen above, this Cenomanian-Turonian accelerated rotation of Africa-Iberia heralds the onset of the Alpine orogenic cycle in many areas of the western Tethyan domain, from Spain to Turkey.

Conclusions The period between the Variscan and Alpine cycle corresponds to the final closing of Palaeotethys and opening of Neotethys in the Mediterranean and Middle East regions. This Cimmerian cycle affected Europe through pervasive extensional systems and re-equilibration of the Variscan lithosphere, mainly in Permian times. This was directly linked to the Variscan cordillera collapse along the still active Mediterranean Eurasian margin. There, Permian rifting was locally followed by Triassic back-arc openings, and the Palaeotethys arc-trench system started to drift away from Eurasia to collide with the Cimmerian terranes (Fig. 11). The latter drifted away from Gondwana from Early Permian time, accompanied by the opening of Neotethys. The soft collision of the two domains took place in Mid- to Late Permian time around the Apulian promontory. Further east, the collision was in the Late Triassic, and accompanied by large amounts of molasse-type deposits found in Turkey and Iran. The Eocimmerian cycle was followed by a second phase of ocean opening and closure, related to (1) the break-up of Pangaea (central Atlantic, Alpine Tethys) and (2) shortening affecting remnant Triassic back-arc basins (Meliata-Maliac-Ktire). The latter were affected by slab roll-back, which resulted in a Jurassic back-arc basin (SSZ Vardar; Izmir-Ankara). These younger oceanic domains had diverging evolutions: some were obducted in Late Jurassic time (Vardar), others (Izmir-Ankara) were affected by roll-back and

76

G.M. STAMPFL! & H. W. KOZUR

gave rise to a third generation of SSZ basins (Lycian). The final closure of these marginal basins took place in the Late Cretaceous-Palaeocene, accompanied by arc volcanism (Balkans, Pontides) and locally, again, by back-arc opening (Black Sea). The Late Cretaceous events heralded the Alpine cycle, marked by major e a s t - w e s t shortening and the onset of cordillera-building processes; however, in the Alps and the Hellenides, remnant oceanic basins (Alpine Tethys, Pindos, Fig. 9) were finally closed only in Oligocene to Early Miocene times, a period corresponding to the first Alpine orogenic phase. Orogenic processes lasted in the Alps for a longer time, whereas cordillera collapse and slab roll-back of remnant oceans (Maghrebian Tethys, East Mediterranean) affected most of the Mediterranean Alpine domain, from the Miocene onward, opening again a last generation of back-arc basins (West Mediterranean, Aegean basins). Large-scale displacement of terranes is responsible for the complex present-day tectonic framework of the Tethysides; in particular, suture duplications took place for most of the older oceans (Fig. 10). The authors have benefited during long years from many collaborations with geoscientists throughout the world whose knowledge of specific areas or material was indispensable in making up the models of Tethys evolution presented in this paper, and these colleagues are given here our sincerest thanks. We are grateful to J. Golonka, A. Saintot and an anonymous reviewer for their numerous comments and constructive criticisms that helped to produce the final version of this paper. Part of this research was supported by FNS grant 2000-068015.

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European geography in a global context from the Vendian to the end of the Palaeozoic L. R. M. COCKS 1 & T. H. T O R S V I K 2'3'4

tDepartment of Palaeontology, The Natural History Museum, Cromwell Road, London SW7 5BD, UK (e-mail: [email protected]) 2Center for Geodynamics, Geological Survey of Norway, Leif Eirikssons vei 39, N-7491 Trondheim, Norway 3Institute for Petroleum Technology and Applied Geophysics, Norwegian University of Science and Technology (NTNU), Trondheim N-7491, Norway 4School of Geosciences, Private Bag 3, University of Witwatersrand, WITS 2050, South Africa

Abstract: A succession of palaeogeographical reconstructions is presented, covering half the globe and the time interval from the latest Proterozoic (Vendian) at 550 Ma to the end of the Palaeozoic (latest Permian) at 250 Ma, mostly at 20 or 30 Ma intervals. The various terranes that today constitute Europe are defined and their margins discussed briefly; these are Gondwana, Avalonia, the RhenoHercynian Terrane, the Armorican Terrane Assemblage, Perunica, Apulia, Adria, the Hellenic Terrane (including Moesia), Laurentia, and Baltica. As time elapsed, many of these terranes combined to form first Laurussia and subsequently Pangaea. The further terranes of Siberia and Kara adjoined Europe and were relevant to its Palaeozoic development. Brief sections are included on the individual history and geography of the Vendian and the six Palaeozoic systems, with emphasis on their importance in the building of Europe.

During the 300 Ma from 550 to 250 Ma the geography of the Earth evolved greatly. At the beginning of this period there was only one superterrane, Gondwana, with a large number of other terranes at varying distances from each other, some separated by wide oceans. By the end of the Palaeozoic, at the end of the Permian, most of the other terranes had coalesced, and had also joined Gondwana, to form Pangaea, by far the largest superterrane in Phanerozoic history. Evolution of the biota over this huge time interval had progressed enormously, with great consequent diversity: at 550 Ma there were no animals or plants with substantial hard parts, and the colonization of the land by animals had not yet begun. In contrast, by the Permian there were probably millions of different animals and plants, not only in the marine habitats, but also over much of the land, a great part of which was covered with forests and jungles comparable in size with those known today. The end of the Permian also saw the largest faunal and floral turnovers and extinction event in the whole Phanerozoic. At no time in the 300 Ma period that we review here was Europe the geographical unity that it is today. It is the chief purpose of this paper to set a substantial part of Europe's geographical evolution in global context so that this book may be better appreciated by a wide audience. The Europe of today is made up of many terranes, some of which did not join the present continent until after the Alpine Orogeny during the Tertiary. Excellent and detailed reviews of the geology of Europe as it developed through time from the latest Silurian onward have been given by Ziegler (1989, 1990). We here present (Fig. 1) a simple flow diagram showing the break-up and amalgamations of the major European terranes and their associated orogenies from the Vendian to the end of the Palaeozoic. To reconstruct the successive ancient geographies we have pooled our different expertises of palaeontology and palaeomagnetism, and combined them with sedimentology and, to a lesser degree, structural evidence to produce a kinematically valid series of successive maps. We have already set out in detail elsewhere the criteria by which we work and have also reviewed much of the period in different time slices (Cocks 2000; Cocks & Torsvik 2002, 2004; Torsvik et al. 2002; Torsvik & Cocks 2004, 2005). The present shorter review both integrates some of our previous results and also focuses particularly on the terranes that make up modem Europe. It also depends heavily on the work of a great number of researchers who are not quoted in this brief paper: reference will be found to many of them in the landmark

publication edited by McKerrow & Scotese (1990) as well as in our own previous papers listed above. After a short review of each of the more important European and adjacent terranes, this paper presents a brief history of the events in Vendian and Palaeozoic Europe. The chief terranes are labelled in Figure 2.

Major terranes relevant to Palaeozoic Europe

Gondwana During the Early Palaeozoic this vast superterrane stretched from the South Pole to the Equator and beyond, and included at least South America, Africa, Madagascar, peninsular India, Antarctica and Australasia. None of what is termed 'core' Gondwana is today preserved in Europe; the nearest part of it is in northern Africa. However, at various times in the Palaeozoic, several terranes (termed peri-Gondwanan) that had originally formed integral parts of Gondwana separated and rifted from it. The principal ones that now form part of Europe are Avalonia, the Rheno-Hercynian Terrane, the Armorican Terrane Assemblage, Perunica, Apulia, Adria, the Hellenic Terrane and Moesia, and they will now be reviewed in turn, with a brief note on the Gondwana-derived Early Palaeozoic fragments caught up in the Cenozoic Alpine Orogeny.

Avalonia Avalonia today includes the eastem North America seaboard from Newfoundland as far south as Cape Cod, Massachusetts, and in Europe includes southern Ireland, Wales, England, Belgium, the Netherlands and parts of northern Germany, and its boundaries were described by Cocks et al. (1997). Its northern margin is defined by the closed Iapetus Ocean suture with Laurentia, its eastern margin by the closed Tomquist Ocean part of the TransEuropean Suture Zone (TESZ), rather than the Elbe Line as stated by Cocks et al. (1997), and its southern margin by the Rheic Ocean suture. Some workers use the terms West Avalonia and East Avalonia, but we believe that the terrane was a single entity during its relatively short Ordovician independent existence and do not consider that the two halves now separated by the

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 83-95. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Flow chart showing the break-up and amalgamations of the various European terranes with time from the Vendian to the end of the Palaeozoic. At the base is shown the longevities of the oceans whose sutures are included in modern Europe.

Fig. 2. Reconstruction of the Southern Hemisphere in the Late Vendian (550 Ma), modified from Hartz & Torsvik (2002) and Rehnstr6m et al. (2002). Figures 2 - 7 are Schmidt's Equal Area Projection, with projection centre at the South Pole. Spreading centres are shown as black lines, subduction zones as red lines with ticks, and transform faults as red lines with no extra ornament.The alternately dashed and dotted black line marks the limit of 'core' Gondwana. The terranes are also identified in Figure 9. In Figures 2-14, small terranes and island arcs are mostly omitted.

VENDIAN-PALAEOZOIC PALAEGEOGRAPHY

Atlantic Ocean were divided in the Palaeozoic. The terrane broke off from Gondwana before the Llanvirn Stage of the Ordovician; before that time a variety of palaeontological and sedimentological evidence suggests that it originally adjoined, and formed part of, core Gondwana and was probably adjacent to the northern part of South America (McKerrow et al. 1992).

The Rheno-Hercynian Terrane

The Rheno-Hercynian belt, which is chiefly within today's Germany, has been described by Franke (e.g. Franke 2000) and Stampfli and coworkers (e.g. Stampfli et al. 2002), and probably represents the opening and closing of the Rheno-Hercynian Ocean. From that it can be deduced that there was probably a separate Rheno-Hercynian Terrane, which was independent only from the Devonian (Emsian) to the early Carboniferous. The area today is largely tectonized and has few fossils that could be used to identify the terrane, and its limits consist entirely of postDevonian structures; we show it with an arbitrary shape in our Devonian reconstructions (see Figs 9 and 10). The Armorican Terrane Assemblage

This includes the Iberian Peninsula and most of France. The area is much tectonized and its geological history is contentious. However, the Cambrian to Devonian faunas and sediments clearly indicate that Armorica (as it is often called) remained an integral part of Gondwana at least until the end of the Silurian (Robardet 2003), and, as can be seen from the distributions of its higher latitude Early Palaeozoic 'Mediterranean Province' benthic faunas, was apparently located withrin Gondwana not far from its present location with respect to northern Africa. Robardet et al. (1990) have presented data that they interpret as indicating that Armorica remained part of Gondwana during the Devonian as well; however, in our 400 Ma Mid-Devonian map we show it as having left the superterrane with other terranes, following tiffing and the opening of the Palaeotethys Ocean of Stampfli et al. (2002). Some workers have included Perunica (Bohemia) as part of the Armorican Terrane Assemblage, but we believe the two were separate in the Early Palaeozoic. Perunica (Bohemia)

This area, most of which today forms the western part of the Czech Republic, has late Cambrian faunas and sediments that definitely link it to the Armorican and northern African part of Gondwana. However, palaeomagnetic (Tait et al. 1994) and faunal (Havlf6ek et al. 1994) evidence both demonstrate that Perunica left Gondwana in the early Ordovician and pursued an independent course, separate from the history of the Armorican Terrane Assemblage, across the Rheic Ocean before merging with what is today's northern Europe in the Variscan Orogeny. The faunal analysis by Havlf~ek et al. (1994) demonstrates that Perunica was at its most isolated in the late Ordovician (Caradoc). Some workers unite Perunica and the Rheno-Hercynian Terrane within a 'Saxo-Thuringian' Terrane, but because of the palaeomagnetic and fossil data we treat the two as separate. Apulia, Adria, Hellenic and Moesia terranes

These terranes make up most of the eastern part of southern Europe. We follow Stampfli et al. (1998, 2002) in the outlines and integrity of Apulia (southern Italy), Adria (the Adriatic Sea and adjacent areas) and the Hellenic Terrane (Greece and adjacent areas). However, there are no diagnostic palaeomagnetic or palaeontological data from any of these three regions before the

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Carboniferous. We have grouped Moesia with the Hellenic Terrane; Yanev (2000) reviewed the Palaeozoic faunal data from there, which, although sparse, indicate that Moesia probably had peri-Gondwanan rather than Baltic trilobites during the Ordovician. Stampfli et al. (1998, 2002) suggested that Adria, the Hellenic Terrane and Moesia may have left Gondwana at about the end of the Silurian with the opening of the Palaeotethys Ocean, but Apulia probably did not leave Gondwana until the Permian, as part of the opening of the Neotethys Ocean. Alpine fragments

Although these are not shown on our maps, there are today in the Alpine regions, particularly of Austria, a number of Palaeozoic fragments preserved in a variety of tectonic settings. Sch6nlaub (1997) and von Raumer (1998) have reviewed these and concluded that they represent peri-Gondwanan areas of unknown size and integrity. They chiefly consist of Cambro-Ordovician arc-related metavolcanics rocks and Ordovician granitoids; the oldest fossils there are mid-Ordovician brachiopods that are undoubtedly attributable to the higher latitude West Gondwanan (often termed Mediterranean) faunal province, which included Armorica, Perunica and northern Africa (Havlf~ek et al. 1994). Laurentia

This substantial terrane included most of North America; and, in Europe, Greenland, BjornCya, Svalbard, northwestern Ireland, Scotland and the upper parts of the nappes in the Scandian Caledonides (the Uppermost Allochthon). It was an independent entity from before the beginning of our study period, with the Iapetus Ocean to its east originally widening in the Proterozoic, at its maximum width in the latest Cambrian, starting to close in the earliest Ordovician, and continuing to close until the Silurian, when Laurentia collided with Avalonia-Baltica during the Caledonian (locally termed the Scandian) Orogeny to form the much larger terrane of Laurussia. At least two island arcs that also lay in the Iapetus Ocean were also involved in those orogenic events to eventually form the complex pattern of small terranes seen within the Iapetus Suture Zone today and illustrated by Armstrong & Owen (2001).

Baltica

Most of the northeastern part of modern Europe is attributable to this terrane, which includes the ancient and substantial Precambrian East European Craton of many workers (e.g. Bogdanova et al. 2001). Baltica is approximately triangular in modern outline, with its eastern limit defined by the Ural Mountains (extending northwards to include Novaya Zemlya), its northwestern edge defined by the British and Scandinavian Caledonides orogenic belt and its southwestern margin in general by the TESZ. Important exceptions are the Lysogory and Matopolska terranes exposed in the Holy Cross Mountains of Poland, which today lie south of the TESZ but which formed an integral part of Baltica in the Early Palaeozoic (Cocks 2002). Baltica was inverted in relation to its present-day orientation from the Neoproterozoic (Hartz & Torsvik 2002) until the mid-Cambrian, when it started to rotate anti-clockwise; a rotation that was largely completed by the end of the Ordovician. Baltica first collided obliquely and relatively softly with Avalonia at about 443 Ma, the end of the Ordovician (Torsvik & Rehnstr6m 2003), and then, fairly soon afterwards, in a more dynamic way with Laurentia to form Laurussia in the Caledonide Orogeny. The Uralian Orogeny, when the eastern part of Laurussia collided with Kazakhstania and intervening island arcs, took place in the late Carboniferous. The Baltica part of Laurussia did not merge with Siberia until the late

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Permian. The palaeogeography and history of Baltica during its Neoproterozoic to Silurian existence as a separate terrane has been described by Cocks & Torsvik (2005).

Siberia

No part of the old Siberian (sometimes termed Angaran) Terrane is today in Europe. As well as much of north-central modern Siberia, the terrane included the southern and central parts of the Taimyr Peninsula. It had been previously postulated by some workers (e.g. Cocks & Fortey 1998) that the central and southern parts of Taimyr formed part of Baltica in the Early Palaeozoic, but further palaeontological and palaeomagnetic work negated that hypothesis; the evidence has been summarized by Cocks & Torsvik (2002), and these two more southern parts of Taimyr are now seen as having formed integral parts of the Early Palaeozoic Siberia. For all of the Palaeozoic, Siberia was inverted relative to its present-day orientation; rotation to today's orientation began in the Late Devonian and ended in the Permian upon its collision with Pangaea (see Fig. 14). Today's northwestern margin of Siberia collided with Kazakhstania and intervening island arcs during the late Carboniferous, but the terrane was not finally accreted to Laurussia to become part of Pangaea until the Permian.

Kara

Severnaya Zemlya, the northern part of the Taimyr Peninsula and parts of the adjacent Arctic Ocean, despite being parts of Siberia today, for most of our study period formed the centre of an independent terrane (Torsvik & Rehnstrrm 2001), known as the Kara Terrane. Kara collided with Siberia at some time after 300 Ma. Zonenshain et al. (1990) used the term Arctida for a rather larger terrane area that they identified in the same vicinity.

Pangaea

Near the end of the Palaeozoic, from Late Carboniferous (330 Ma) times onwards, Laurussia and Gondwana merged to form the supercontinent of Pangaea, and they were subsequently joined by Kazakhstania and Siberia in turn. However, the Palaeotethys Ocean to the east of Pangaea separated the northern from the southern parts of today' s European collage during the later Palaeozoic (Stampfli e t al. 2002). Our reconstructions of Pangaea use the so-called Pangaea A configuration, which is the only one of several alternatives (some others being termed Pangaea B and C) that can resolve the tectonic, faunal and sedimentological evidence on the one hand and palaeomagnetic data on the other. Pangaea A can be true to the palaeomagnetic data only if it is assumed that the Earth's magnetic field had a 10-15% octupole component at the time (Torsvik & Van der Voo 2002; Torsvik & Cocks 2004), rather than entirely a dipole field, as presumed by many palaeomagnetists (e.g. Muttoni et al. 2003).

at that time. These arcs are represented today by rocks in largely tectonized areas in the eastern seaboard of North America (van Staal et al. 1998) and the British Isles (Armstrong & Owen 2001), and in the higher nappes of the Scandinavian Caledonides (Cocks & Torsvik 2005). The progressive accretion of the arcs to Laurentia, Avalonia and Baltica and their accretion to each other represent important phases in the widespread Ordovician to early Devonian Caledonian Orogeny. Most of the other side of the world, the northern hemisphere, which we do not illustrate here, was taken up by the vast Panthalassic Ocean, which was comparable in size with the Pacific today. That ocean included all the latitudes and longitudes now occupied by modern Europe. Latest P r o t e r o z o i c

The Vendian started at 600 Ma and continued until the beginning of the Cambrian at 543 Ma. However, we start our reconstructions here at 550 Ma (Fig. 2) because at that time the data concur that Gondwana, Laurentia and Baltica were certainly all separate continents. Before that, palaeomagnetic and some sedimentological and tectonic data have been used for tentative reconstructions (e.g. Torsvik et al. 1996), but the same degree of confidence is not present in those maps as in the ones for the Phanerozoic (Figs 3-14), and so we will not repeat those maps or discussions here. Knoll (2000) has reviewed both the basis for relative dating and also the glaciations in late Precambrian time; there is little evidence for glaciation within our study area in the late Neoproterozoic, except perhaps in Algeria. Figure 2 is modified from Hartz & Torsvik (2002) and Rehnstrrm et al. (2002). The opening of the southern Iapetus Ocean between Laurentia and the South American part of Gondwana was at an early phase at this time (550 Ma), as were the openings of the northern part of the Iapetus between Baltica and Laurentia, the Ran Ocean between Baltica and the Avalonian part of Gondwana; and the A~gir Sea between Baltica on the one hand and Kara and Siberia on the other. Avalonia, Armorica, the Rheno-Hercynian Terrane, Perunica and the other terranes now in southern Europe all formed part of Gondwana: they and North Africa were affected by the Cadomian Orogeny, which lasted throughout this period and finished in the early Cambrian at about 530 Ma. The Cadomian takes its name from NW France, but the scope and definition of that orogeny has been treated in different ways by different researchers and requires more stringent usage. For example, so-called 'Cadomian' rocks have allegedly been recorded from the Uralian margin of Baltica; these must surely represent an independent (although contemporaneous) orogeny that was perhaps linked with, or even formed part of, the Timanian Orogeny of northern Europe (Gee & Pease 2004). Although the Vendian Ediacaran Fauna soft-bodied fossils are well documented from this period, their distribution apparently occurs in a variety of terranes, and we have not found them relevant in determining terrane positions. During this period Laurentia moved from temperate and high latitudes into slightly warmer ones. From those early times to the midOrdovician, Baltica was independent and inverted in relation to its present-day orientation, with the fastest rotation occurring from the late Cambrian to the earliest Ordovician (references have been summarized by Torsvik & Rehnstrrm 2001).

Geological history Despite the fact that our maps (Figs 2 - 1 4 ) show considerably greater parts of the world, only the European sectors of them will be discussed here. We treat the geological history system by system. We have omitted from these reconstructions all the many island arcs and smaller terranes that were undoubtedly present at each period. For example, it is now well documented (e.g. Harper et al. 1996; van Staal et al. 1998) that in the Ordovician there were two subparallel island arcs in the Iapetus Ocean, which acted as stepping stones for the spread of benthic faunas

Cambrian

The duration of the Cambrian, from 543 to 490 Ma, was a substantial 47 Ma, and we present two maps (Figs 3 and 4) showing the palaeogeography at 535 and 500 Ma. The former has been constructed chiefly from palaeomagnetic data, with some tectonic input, and modified from Torsvik & Rehnstr6m (2001), and the latter from a combination of palaeomagnetic and faunal data (Cocks & Torsvik 2002, fig. 3). Figure 3 shows a wide and still

VENDIAN-PALAEOZOIC PALAEGEOGRAPHY

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Fig. 3. Early Cambrian (535 Ma) reconstruction, modified from Torsvik & Rehnstr6m (2001). Symbols as in Figure 2. The terranes are labelled in Figures 2 and 9. The Sibumasu, South China and Annamia terranes were then NE of Siberia: parts of them may or may not have overlapped into the Southern Hemisphere, although they are not shown here.

Fig. 4. Late Cambrian (500 Ma) reconstruction, modified from Cocks & Torsvik (2002, fig. 3). Symbols as in Figure 2; the terranes are labelled in Figures 2 and 9. The Annamia, South China and Tarim Terranes to the north of the Hellenic Terrane (labelled in Fig. 9) appear in our figures for the first time; also shown are the Sibumasu Terrane west of South China and the Precordillera Terrane to the then north of central South America. Two small triangles, representing idealized parts of what were later to accrete to form Kazakhstania, are shown to the then west of Tarim.

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Fig. 5. Early Ordovician (Arenig, 480 Ma) reconstruction, modified from Cocks & Torsvik (2002, fig. 4). Symbols as in Figure 2; terrane names as in Figures 2 and 9. Those intra-Iapetus islands with reliable palaeomagnetic data (Cocks & Torsvik 2002) are also shown.

Fig. 6. Mid-Ordovician (Caradoc, 460 Ma) reconstruction, modified from Cocks & Torsvik (2002, fig. 5). Symbols as in Figure 2, terrane names as in Figures 2 and 9.

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Fig. 7. Latest Ordovician-earliest Silurian (440 Ma) reconstruction, modifiedfrom Cocks & Torsvik (2002, fig. 6), showingthe latest Ordovician (Hirnantian) glacial deposits. Symbolsas in Figure 2, terrane names as in Figures 2 and 9.

spreading Iapetus Ocean between Laurentia on the one hand and Baltica and Gondwana on the other. In the earliest Cambrian (Tommotian), before its northward movement, Laurentia was not too far separated from Baltica, a position reinforced by the similarities of the archaeocyathids on both terranes at that time (Debrenne et al. 1999). There was a substantial Tornquist Ocean between Baltica and Gondwana. The evidence from the faunal distributions of terrane-dependant benthos, such as brachiopods and most trilobites, shows that Avalonia, Armorica and Perunica were all integral parts of Gondwana at this time; and all at temperate to high palaeolatitudes, with the then South Pole situated in the NW African part of Gondwana. After moving northwards, Laurentia reached the palaeoequator by the end of the period, where it remained for most of the Palaeozoic, with rich Cambrian and early Ordovician benthic faunas very different from those of the European parts of Gondwana, and generally different again from the faunas of Baltica. Siberia and Kara were clearly not too distant from Baltica at the time, because, although most of their faunas were distinct from Baltica and there were many endemic genera, particularly the articulated brachiopods, there were some key late Cambrian trilobites in common between the three terranes (Rushton et al. 2002). However, most of Baltica was inhabited by the widespread Olenid trilobite realm during the late Cambrian, whose distribution was controlled by relatively poor bottom-water circulation and consequent reduction in oxygen levels. That realm is known from several areas and is not terrane-specific, and its widespread late Cambrian distribution makes the identification and elucidation of Cambrian terrane-related faunal provinces much more difficult than at other times.

Ordovician

The start of the Ordovician is dated at 488 Ma and its end at 443 Ma, giving a duration of 45 Ma. Exactly when the Iapetus Ocean reached its widest point is uncertain in detail, but it was probably in about the late Tremadoc or early Arenig (c. 480 Ma), after which subduction and consequent closure started, with the ocean steadily narrowing through the rest of Ordovician time. Avalonia rifted off from the South American part of Gondwana in the earlier part of the Ordovician, and probably also in the Arenig, with a widening Rheic Ocean to its south. Whether or not that rifting and ocean-floor spreading was part of the same tectonic process as the Iapetus closure is uncertain. The Tornquist Ocean (Cocks & Fortey 1982) between Avalonia and Baltica was also narrowing throughout the Ordovician, with subduction beneath Avalonia, and finally closed at about Ordovician-Silurian boundary time with the soft oblique Avalonia-Baltica docking (Torsvik et al. 1996; Torsvik & Rehnstr6m 2003). Thus Avalonia was an independent terrane for less than the total duration of the Ordovician. Perunica also rifted from Gondwana during the early Ordovician and proceeded across the Rheic Ocean; Havlffiek et al. (1994) concluded that it had its highest proportion of endemic terrane-specific brachiopod and trilobite genera in the mid-Ordovician (Caradoc), indicating substantial oceanic separation around it. In contrast, southern Europe (Armorica and the terranes to the east of it) clearly remained part of Gondwana itself, as can be deduced from both the facies and faunas (Robardet 2003). That part of Gondwana was close to the South Pole, whose probable position was in modern Libya. We present three Ordovician reconstructions. The first map (Fig. 5), is for the Early Ordovician (Arenig, 480 Ma),

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Fig. 8. Late Silurian (Ludlow-P~doli, 420 Ma) reconstruction, modified from Cocks & Torsvik (2002, fig. 8). Symbols as in Figure 2; terrane names as in Figures 2 and 9. Figures 8 - 1 0 are Schmidt's Equal Area Projection, with projection centre at 30~

Fig. 9. Early mid-Devonian (Emsian, 400 Ma) reconstruction, modified from Torsvik & Cocks (2004, fig. 5), also showing the Old Red Sandstone continents in Laurussia and Gondwana. Symbols as in Figure 2; terrane names are labelled. The outline of the Rheno-Hercynian (RH) Terrane is arbitrary.

VENDIAN-PALAEOZOIC PALAEGEOGRAPHY

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Fig. 10. Late Devonian (Famennian, 370 Ma) reconstruction, modified from Torsvik & Cocks (2004, fig. 6). Symbols as in Figure 2: terrane names as in Figure 9. The Rheno-Hercynian (RH) Terrane was in the process of amalgamation with Laurussia at this time.

Fig. 11. Early Carboniferous (Tournaisian, 340 Ma) reconstruction, modified from Torsvik & Cocks (2004, fig. 8). Figures 11 and 12 are Schmidt's Equal Area Projection, with projection centre at 15~ Symbols as in Figure 2; terrane names as in Figure 9. The Sibumasu Terrane (unlabelled) is shown as largely attached to the NW Australian part of Gondwana.

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Fig. 12. Late Carboniferous (Westphalian, 310 Ma) reconstruction, modified from Torsvik & Cocks (2004, fig. 9). Extensive glacial deposits (not shown) covered much of the Southern Hemisphere. Symbols as in Figure 2; terrane names as in Figure 9.

Fig. 13. Early Permian (Asselian, 280 Ma) reconstruction of the Western Hemisphere, modified from Torsvik & Cocks (2004, fig. 10). Extensive glacial deposits (not shown) covered much of the Southern Hemisphere. Figures 13 and 14 are Schmidt's Equal Area Projection, with projection centre at the Equator. Symbols as in Figure 2; terrane names as in Figure 9. W, WrangelliaAlexander Terrane; S, Stikinia Terrane; EK, Eastern Klamath Terrane.

VENDIAN-PALAEOZOIC PALAEGEOGRAPHY

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Fig. 14. Permo-Triassicboundary (250 Ma)

reconstruction of the Western Hemisphere,modified from Torsvik& Cocks (2004, fig. 11), also showingthe flood basalts in Siberia and China. Symbols as in Figure 2; terrane names as in Figure 9.

when the different benthic trilobite faunal provinces indicate that the Iapetus was at its widest (about 5000 km), and with Baltica well separated from both Gondwana and Laurentia. The second map (Fig. 6), for the mid-Caradoc (460 Ma), shows an independent Avalonia between a closing Iapetus and an opening Rheic, and also the initial sea-floor spreading of the ocean, which was separating Bohemia and Annamia (Indochina) from Gondwana. Our third reconstruction, at 440 Ma (Fig. 7), although technically representing Silurian time (the Ordovician-Silurian boundary was at 443 Ma), has plotted on it the latest Ordovician glaciogenic deposits, which were laid down during the half-million-year glacial interval of the Hirnantian Stage of the latest Ashgill Series.

Silurian

The start of this system is dated at 443 Ma and the end at 416 Ma, giving a duration of 27 Ma; much shorter than the other systems that preceded and followed it in the Palaeozoic. The major tectonic event was the acme of the Caledonian (including Scandian) Orogeny of eastern North America, the British Isles and Scandinavia, caused by the collision of Laurentia with the combined Baltica-Avalonia. Our first map, for 440 Ma (Fig. 7), shows the Baltica-Avalonia docking completed but with a narrow Iapetus Ocean separating that combined terrane from Laurentia. Although that ocean may have contained islands, it was still an effective barrier to the migration of ostracodes (which have no planktonic larval stages) until near the end of the Silurian. In the early Ordovician, Avalonia had been at high latitudes and Baltica at intermediate latitudes, but they both drifted northwards after that and were in tropical palaeolatitudes during the Silurian, with substantial carbonate deposits, including the famous bioherms of Gotland, Sweden, and Wenlock Edge, England. It is from Wenlock age

beds in Britain that the first true land plants from anywhere are known (although trilete spores inferred to have come from land plants are known from as early as the mid-Ordovician). By the time of our second reconstruction, at 420 Ma (Fig. 8), Baltica, Avalonia and Laurentia had all coalesced to form the new superterrane of Laurussia, thus completing the jigsaw for the northern parts of today's Europe. The hydrothermal vent communities of Silurian age found in the central Urals (Little et al. 1997) provide a sure indication that that area was in a truly oceanic environment, and thus seaward of the eastern margin of Baltica. Zonenshain et al. (1990, Fig. 20) showed both the cratonic and the marginal facies that made up the eastern (Ural) margin of Baltica, as well as the fragmented Cambrian to Devonian island arc fragments. All of these were also caught up in the strike-slip movements of the late Palaeozoic Uralian Orogeny, which accounts for the unnaturally straight outcrop of the Urals today.

Devonian

The start of the system is at 416 Ma and the end at 359 Ma, giving a duration of 57 Ma. We present two reconstructions. The first (Fig. 9), at 400 Ma, the early mid-Devonian (Emsian), shows the substantial and largely desert Old Red Sandstone continents present in both Laurussia and Gondwana, although the latter were fringed by seas rich in marine benthos. The new Palaeotethys Ocean, which had probably started its opening in the latest Silurian, included the spreading centre which was by then separating most of Southern Europe (the Armorican Terrane Assemblage, Adria, the Pontides of Turkey, and the Hellenic Terrane including Moesia) from Gondwana (Stampfli et al. 2002). The second reconstruction (Fig. 10) is for the Late Devonian at 370 Ma (Famennian); by that time the Variscan Orogeny was at its maximum in

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the central part of Europe. Winchester et al. (2002) have documented the amalgamation of Central Europe. McKerrow et al. (2000) have summarized the evidence to show that the Rheic Ocean between Gondwana and Laurussia was not then wide enough to prevent the migration of key faunas during the Late Devonian. The period also saw the start of the relatively brief independent existence of the Rheno-Hercynian Terrane (Franke 2000).

Carboniferous

The start of this system was at 359 Ma and the end at 299 Ma, making it 60 Ma long, the longest in the Palaeozoic. Our two reconstructions are for the Tournaisian at 340 Ma (Fig. 11) and the Westphalian at 310 Ma (Fig. 12). The northern Europe part of Laurussia straddled the palaeoequator. The Palaeotethys Ocean had widened considerably at its eastern end by the start of the Carboniferous, so that Southern Europe (apart from Apulia) was at some distance from the north African part of Gondwana. However, at its western end, and away from the European region, subduction caused contact between the South American part of Gondwana and Laurussia, which formed the incipient superterrane of Pangaea for the first time. By the end of the Carboniferous (Fig. 12), more of the western part of the Palaeotethys had closed, and the Laurussian-Gondwanan collision zone had progressively stretched as far eastward as the Iberian Peninsula. The Kazakhstania Terrane, which had coalesced from numerous fragments over the previous 200 Ma (Sengor & Natalin 1996), and which continued to enlarge by accretion after that, collided with the Uralian margin of Laurussia to cause the Uralian Orogeny. The late Carboniferous and early Permian were marked by a significant glacial episode, which may have persisted for as long as 50 Ma, but the glacial deposits did not reach the lower latitudes of contemporary Europe. The late Carboniferous also saw the most extensive forests of the Phanerozoic, which are very well represented as coal in Britain, Belgium, Germany, Poland and elsewhere in Europe.

Permian

The Permian extended from 299 Ma at its start to 251 Ma at its end, a duration of 48 Ma. Figure 13 shows the situation in the Early Permian (Asselian) at 280 Ma, by which time most of Europe had moved north of the Equator and the coalescence of Pangaea was more complete than in the Carboniferous. Significant evaporite deposits were laid down within the New Red Sandstone deposits that covered much of Europe. Figure 14, for 250 Ma, shows the situation at the Permo-Triassic boundary times, when the vast flood basalts of the Siberian Traps and the Emeishan Traps of China poured out, a period that coincided with, and probably contributed to the causes of, the greatest biological extinction event in the whole of the Phanerozoic. The amalgamation of Siberia with Laurussia also began in the early Permian. Ziegler and coworkers (e.g. Ziegler et al. 1997) have provided excellent accounts of Permian geography and climates on a global basis. Glennie et al. (2003) have published detailed palaeogeographical maps of the Central European area, showing the progressive migration of the southern margin of the Boreal Ocean (locally termed the Zechstein Sea) over the North-Central European part of Pangaea.

Concluding remarks (1) By pooling our separate expertises of palaeontology and palaeomagnetism and combining the results from the numerous researchers in those fields with selected sedimentological data,

we have been able to make revised palaeogeographical reconstructions for the whole of the Palaeozoic. (2) By considering such a considerable length of geological time, from 550 to 250 Ma, we have been made constantly aware of the imperative need to maintain kinematic continuity in the movement of terranes over that 300 Ma period, which we have striven to implement in our successive reconstructions. (3) By treating Europe in its global context, rather than confining our reconstructions to its modern boundaries, we have become aware of the influence of the various terranes that do not today form part of Europe but whose presence nearby in the Palaeozoic directly affected the current European terrane patterns and geography. We acknowledge with pleasure all the colleagues with whom we have developed these reconstructions over many years, in particular R. A. Fortey, the late W. S. McKerrow and J. Mosar. We also thank E. Rehnstrrm for help in the preparation of the figures, VISTA and NGU for financial support, and The Natural History Museum for the provision of facilities. Much of the work was accomplished through participation in the Europrobe project funded by the European Science Foundation.

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Post-Variscan evolution of the lithosphere in the area of the European Cenozoic Rift System P. A. ZIEGLER l, M. E. S C H U M A C H E R 2, P. Dl~ZES a, J.-D. V A N WEES 3 & S. C L O E T I N G H 4

1Department of Earth Sciences, University of Basel, Bernoullistrasse 32, 4056 Basel, Switzerland (e-mail: paziegler@ magnet, ch) 2Unterer Zielweg 77, 4143 Dornach, Switzerland 3Netherlands Institute for Applied Geosciences TNO, Prins Hendriklaan 105, 2508 TA Utrecht, Netherlands 4Faculty of Earth & Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 H V Amsterdam, Netherlands

Abstract: The European Cenozoic Rift System (ECRIS) transects Variscan basement, Permo-Carboniferous troughs and Late Permian to Mesozoic series, deposited in thermal sag basins, which are exposed on rift-related arches. We have analysedprocesses governing the transformation of the orogenically destabilized Variscan lithosphere into end-Cretaceous stabilized cratonic lithosphere, prior to its renewed Cenozoic rift-related destabilization. In the ECRIS area, crustal and lithospheric thicknesses at present are in the range of 24-35 km and 60-120 km, respectively. The Variscan orogen was characterized at the time of its end-Westphalian consolidation by 45-60 km deep crustal roots, marking major sutures. During the Stephanian-Early Permian wrench-induced collapse of the Variscan orogen, subducted lithospheric slabs were detached, causing upwelling of the asthenosphere, thermal thinning and/or partial delamination of the lithospheric mantle, and regional uplift. By mid-Permian times, the crust was thinned to 28-35 km owing to its regional erosional unroofing, localized mechanical stretching and the interaction of mantle-derived melts with its basal parts. By mid-Permian times, when the temperature of the asthenosphere returned to ambient levels, thermal subsidence of the lithosphere commenced, controlling development of a system of Late Permian and Mesozoic intracratonic basins. These experienced repeated minor subsidence accelerations, related to the build-up of far-field stresses, which did not involve renewed lithospheric destabilization. Modelling of observed subsidence curves indicates that during the mid-Permian lithospheric thicknesses ranged in the ECRIS area between 40 and 80 kin, but had increased by the end of the Cretaceous to 100-120 km. Cenozoic rifting and mantle-plume activity caused renewed lithospheric thinning.

The European Cenozoic Rift System (ECRIS) extends over a distance of more than 1000 km from the shores of the North Sea to the Mediterranean and transects the essentially SW-NE-striking French and German parts of the Variscan orogen (Ziegler 1990, 1994; Dbzes et al. 2004). In the ECRIS area, the deeply degraded Variscan orogen was, and in part still is, covered by extensive Late Permian and Mesozoic sediments. During the evolution of ECRIS these sediments were disrupted in conjunction with rift-related uplift of the Rhenish Massif, the V o s g e s - B l a c k Forest Arch, the Massif Central and the Bohemian Massif, in which parts of the Variscan orogen are exposed, thus providing insight into its architecture. This offers a unique opportunity to evaluate processes that controlled the transformation of the orogenically destabilized Variscan lithosphere into end-Mesozoic thermally stabilized cratonic lithosphere, particularly as the ECRIS area was only marginally affected by Mesozoic rifting (Ziegler et al. 2004). The main constituents of ECRIS are the Lower Rhine (Roer Valley), Hessian, Upper Rhine, Limagne, Bresse and Eger (Ohre) grabens (Fig. 1). The Lower Rhine and Hessian grabens transect the external parts of the Variscan orogen, corresponding to the Rheno-Hercynian thrust belt (Oncken et al. 2000). The Upper Rhine, Bresse and Limagne grabens cross-cut the internal parts of the Variscan orogen (see Fig. 3), corresponding to the Mid-German Crystalline High and the Saxo-Thuringian, Bohemian-Armorican and MoldanubianArverno-Vosgian zones, all of which are characterized by basement-involving nappes and a widespread syn- and postorogenic magmatism. The Eger graben is superimposed on the eastern parts of the Saxo-Thuringian zone (Franke 1989, 1995, 2000; Eisbacher et al. 1989; Schreiber & Rotsch 1998; Pharaoh 1999). In the ECRIS area, the depth to the Moho varies at present between 24 and 30 km and increases away from it to 3 4 - 3 6 k m and more (Fig. 2; Prodehl et al. 1995; D~zes & Ziegler 2002). The thickness of the lithosphere decreases from

about 1 0 0 - 1 2 0 k i n in the Bohemian Massif and along the southern end of the Upper Rhine Graben to 6 0 - 7 0 km beneath the Rhenish Massif and Massif Central, and appears to increase to some 120km or more in the Western Netherlands and beneath the Paris Basin (Babuska & Plomerova 1992, 1993; Sobolev et al. 1997; Goes et al. 2000a,b). Beneath Western and Central Europe, anomalously low P- and S-wave velocities characterize the upper asthenosphere (Zielhuis & Nolet 1994; Goes et al. 2000a,b). Tomographic images suggest that lowvelocity structures, rising from the deep mantle, feed smaller upper-mantle plumes, the most important of which well up beneath the Rhenish Massif and the Massif Central (Granet et al. 1995; Goes et al. 1999; Ritter et al. 2001). The present crustal and lithospheric configuration of Western and Central Europe bears no relationship to the major structural units of the Variscan orogen, but shows strong affinities to the ECRIS (Ansorge et al. 1992; Mengel 1992; Ziegler & D~zes 2006). However, development of the Variscan orogen involved major crustal shortening and subduction of substantial amounts of supracrustal rocks, continental and oceanic crust and lithospheric mantle (Ziegler et al. 1995, 2004). By analogy with modern examples, such as the Alps (Schmid et al. 1996, 2004; Stampfli et al. 1998), the Variscan orogen must have been characterized at the time of its late Westphalian consolidation (305 Ma) by a significantly thickened crust and lithosphere. Therefore, its orogenically destabilized lithosphere must have re-equilibrated with the asthenosphere in post-Variscan times so that regional crustal and lithospheric thicknesses of about 2 8 - 3 5 km and 100-120 km, respectively, were achieved towards the end of the Mesozoic. Processes controlling the post-orogenic modification of the Variscan lithosphere have been variably attributed to slab detachment, delamination of the lithospheric mantle, crustal extension, and plume activity during the Stephanian-Early Permian phase of wrench faulting and magmatism that overprinted the Variscan orogen and its foreland (Lorenz & Nicholls 1984;

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 97-112. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Location map of the ECRIS in northern Alpine foreland, showing Cenozoic fault systems (continuous lines), rift-related sedimentary basins (light grey), Variscan massifs (dark grey) and Cenozoic volcanic fields (black). Continuous barbed line, Variscan deformation front; dashed barbed line, Alpine deformation front. BF, Black Forest; BG, Bresse Graben; EG, Eger (Ore) Graben; FP, Franconian Platform; HG, Hessian grabens; LG, Limagne Graben; LRG, Lower Rhine (Roer Valley) Graben; OW, Odenwald; TF, Thuringian Forest; URG, Upper Rhine Graben; VG, Vosges.

Fig. 2. Depth map of Moho discontinuity, contour interval 2 km (after Dbzes & Ziegler 2002) with superimposed ECRIS fault systems and volcanic centres (black fields). Continuous barbed line, Variscan deformation front; dashed barbed line, Alpine deformation front.

EVOLUTION OF VARISCANLITHOSPHERE Ziegler 1990; Henk 1993, 1999; Henk et al. 2000; Prijac et al. 2000; van Wees et al. 2000). In an attempt to evaluate the relative importance of processes contributing to the post-orogenic modification of the Variscan lithosphere to its configuration prior to the onset of Cenozoic rifting, we review the synorogenic evolution of the various Varsican units that are transected by the ECRIS, inspect available crustal seismic reflection profiles, develop quantitative subsidence curves for selected wells penetrating the sedimentary cover of the Variscan basement, and compare these with a theoretical thermal decay curve. During the last 310 Ma, the tectonic setting of the Variscan domain underwent repeated changes. Following the late Westphalian (305 Ma) consolidation of the Variscan orogen, its Stephanian-Early Permian collapse (305-280Ma) was controlled by wrench faulting and associated magmatic activity. During Late Permian to Cretaceous times, large parts of the Variscan domain were gradually incorporated into sedimentary basins that evolved in response to thermal contraction of the lithosphere during its re-equilibration with the asthenosphere. During the latest Cretaceous and Palaeocene, the Variscan domain was affected by an important pulse of intraplate compression that was related to early phases of the Alpine orogeny (Ziegler 1990; Ziegler et al. 1995, 1998, 2002). At the same time an array of mantle plumes impinged on the lithosphere of Western and Central Europe, including the NE Atlantic (Hoernle et al. 1995), Iceland (Bijwaard & Spakman 1999), Massif Central (Granet et al. 1995) and Rhenish plumes (Goes et al. 1999; Ritter et al. 2001). The resulting rise in the potential temperature of the asthenosphere caused a renewed destabilization of the lithosphere, as evidenced by the Palaeocene injection of olivine-melilite and olivine-nephelinite dykes in the Massif Central, Vosges-Black Forest and Bohemian Massif, reflecting low-degree partial melting of the lithospheric thermal boundary layer at depths of 60-100 km (Wilson et al. 1995; Adamovic & Coubal 1999; Ulrych et al. 1999; Michon & Merle 2001). Starting in late Eocene times, the ECRIS developed in the foreland of the evolving Alpine and Pyrenean orogens, with crustal extension and continued plume activity causing further destabilization of its lithosphere-asthenosphere system. Crustal discontinuities, which had developed during the Permo-Carboniferous phase of wrench faulting, played an important role in the localization and evolution of the ECRIS (Ziegler 1990; Schumacher 2002; D6zes et al. 2004). In terms of defining initial boundary conditions for modelling the post-orogenic evolution of the Variscan lithosphere, the stages that are of primary interest are its Late Carboniferous (305 Ma), late Early Permian (280 Ma), end-Cretaceous (65 Ma) and present-day configurations. The last of these is well constrained at crustal levels by geophysical data (Cazes & Toreilles 1988; Meissner & Bortfeld 1990; Blundell et al. 1992; B r u n e t al. 1992; Prodehl et al. 1995), whereas constraints on the thickness of the lithosphere are more controversial (Babuska & Plomerova 1992, 1993; Goes et al. 2000a). On the other hand, the Late Carboniferous, late Early Permian and end-Cretaceous-Palaeocene pre-rift lithospheric configurations must be inferred from circumstantial evidence. In this respect, vertical tectonic movements of the crust, derived from its Permian and younger sedimentary cover, provide important constraints to the post-orogenic evolution of the Variscan lithosphere. In the following, we summarize the Late Palaeozoic and Mesozoic evolution of those parts of the Variscan domain that are transected by the ECRIS, the results of quantitative and forward modelled subsidence analyses on selected wells, and a proposed model for the post-orogenic evolution of the lithosphere (for details see Ziegler et al. 2004). We applied the time scales of Menning et al. (2000) and Menning (1995) for the Carboniferous and Permo-Triassic, respectively, and for later times the scale of Gradstein & Ogg (1996).

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Variscan orogeny

Evolution of the Variscan orogen, which forms part of the Hercynian mega-suture along which Laurussia and Gondwana were welded together, involved the stepwise accretion of Gondwana-derived terranes to the southern margin of Laurussia and ultimately the collision of Africa with Europe (Ziegler 1989, 1990; Tait et al. 1997; Stampfli & Borel, 2004). During the Late Vis6an to Westphalian main phases of the Variscan orogeny (330-305 Ma), the collision front between Gondwana and Laurussia propagated eastward and southwestward in conjunction with the progressive closure of the Palaeotethys and Protoatlantic oceans, respectively. By Westphalian times, Gondwana had collided with North America, whereas, to the east, Palaeotethys remained open. Thus, the western parts of the Hercynian mega-suture were characterized by a Himalayan-type (continent-continent collision) setting, whereas its eastern parts remained in an Andean-type (continent-ocean collision) setting. It is noteworthy that the setting of the West European segment of the Variscan orogen was transitional between a Himalayanand an Andean-type (Stampfli 2000; Ziegler & Stampfli 2001; Stampfli & Borel 2004). Subduction of large amounts of oceanic and continental lithosphere (including sediments) along subduction zones associated with the Palaeotethys arc-trench system and the boundaries between the intra-Variscan Gondwana-derived terranes, as well as with the closure of Devonian back-arc basins, accounted for a Late Devonian to Early Carboniferous synorogenic calc-alkaline I-type (island-arc related) and S-type (syncollisional) magmatism (von Raumer 1998; Vigneresse 1999; Franke 2000; Henk 2000). Considering the sparse relics of Palaeozoic ophiolites (e.g. Lizard and Giessen-Harz nappes), this magmatism and its distribution provides important clues to the location of Variscan suture zones and their subduction polarity (Franke 2000). The ECRIS transects the suture between the external RhenoHercynian and the more internal Saxo-Thuringian zone, which is located in the domain of the Mid-German Crystalline Rise, as well as the sutures between the Saxo-Thuringian and Bohemian and the Bohemian and Moldanubian zones (Figs 1 and 3). The triple junction of the Upper Rhine, Lower Rhine and Hessian grabens is superimposed on the south-dipping RhenoHercynian-Saxo-Thuringian suture. The Upper Rhine Graben transects the south-dipping Saxo-Thuringian-Bohemian suture in the northern parts of the Vosges-Black Forest Arch (Lalaye-Lubin and Baden-Baden zone), and the north-dipping Bohemian-Moldanubian suture in the southern parts of the Black Forest (Badenweiler-Lenzkirch zone) (Eisbacher et al. 1989; Franke 2000; Hegner et al. 2001). The latter suture can be traced towards the SW into in the northeastern parts of the Massif Central (Mt. du Lyonnais suture; Lardeaux et al. 2001) where it is transected by the Bresse and Limagne grabens. The Rheno-Hercynian-Saxo-Thuringian suture marks the location of the Early Devonian oceanic Giessen-Harz back-arc basin that was gradually closed during the Late Devonian-Early Carboniferous by southward subduction beneath the arc system of the Mid-German Crystalline Rise, as evidenced by the occurrence of Early Carboniferous I-type intrusive rocks in the Odenwald. During the early VisEan, this magmatism changed to a syncollisional high-K to monzonitic type (340-335 Ma), that gradually gave way to a shoshonitic type (340-332 Ma; Altherr et al. 1999). This suggests that, following collision of the toe of the Rheno-Hercynian Shelf with the Mid-German Crystalline Rise arc, the subducted lithosphere of the Giessen Ocean (original width c. 250km) was detached from the Rheno-Hercynian lower plate during the mid-Vis~an. This slab detachment entailed unflexing of the Rheno-Hercynian plate and its increased collisional coupling with the Mid-German Crystalline Rise during their subsequent accelerated convergence. This controlled deformation of the Rheno-Hercynian pro-wedge thrust belt, which persisted

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Fig. 3. Variscan tectonic framework of the ECRIS area with superimposed ECRIS fault systems, showing the location of deep reflection lines referred to in the text. Continuous barbed lines, Variscan deformation front and boundaries between Varsican terranes; dashed barbed line, ALpinedeformation front. Interrupted part of Line D-2S corresponds to seismic section given in Figure 5. until the end of the Westphalian (305 Ma) (Ziegler et al. 1998, 2002). Balanced cross-sections through this thrust belt yield crustal shortening values of 180-200 kin during a time span of about 20 Ma (Behrmann et al. 1991; Oncken et al. 1999, 2000). At the end of the Variscan orogeny, the foreland crust extended essentially unbroken under the entire thrust belt with only its upper parts being involved in thrust sheets. Correspondingly, the subducted continental lithospheric slab, which dipped beneath the Mid-German Crystalline Rise, had a length of some 200 km. The depth to which this slab penetrated into the sub-lithospheric mantle is uncertain. Evolution of the Rheno-Hercynian-Saxo-Thuringian suture was coupled with the development of a south-verging, basement-involving retro-wedge thrust belt that flanks the MidGerman Crystalline Rise to the south, representing the conjugate to the Rheno-Hercynian pro-wedge thrust belt. Crustal shortening in this retro-wedge thrust belt, which was active during early Vis6an (340 Ma) to Westphalian (312-308 Ma) times, amounts to some 100 km (Sch~ifer et al. 2000), with stacking of crustal flakes accounting for considerable crustal thickening and synorogenic uplift (Seyferth & Henk 2000). By analogy with the Bohemian Massif, amalgamation of the western parts of the Saxo-Thuringian and Bohemian terranes involved Late Devonian southward subduction of the oceanic Saxo-Thuringian Basin that had separated them. By contrast, suturing of the Bohemian and Moldanubian terranes involved closure of an oceanic basin by northward subduction (Eisbacher et aL 1989; Franke 2000; Hegner et al. 2001; Lardeaux et al. 2001). During Early Vis~an post-collisional crustal shortening and nappe emplacement, continental lithospheric mantle was subducted to high- and ultrahigh pressure conditions (near diamond stability field; 340-335 Ma) (Altherr & Kalt 1996). This was followed by a major heat advection into the lithosphere, an important pulse of high-K calc-alkaline dioritic to granitic plutonism spanning late Vis6an and Westphalian times (336-310 Ma), regional

low-pressure-high-temperature metamorphism, and rapid extensional-transtensional exhumation of high-pressure metamorphic rocks (328-326 Ma) (Eisbacher et al. 1989; Rey et al. 1992; Altherr & Kalt 1996; Schaltegger et al. 1999; Franke & Stein 2000; O'Brien 2000). This sequence of events presumably reflects mid-Vis6an detachment of the subducted Saxo-Thuringian lithospheric slab (O'Brien 2000), upwelling of the asthenosphere and partial melting of the lithospheric mantle and later of crustal rocks (Altherr et al. 2000). The oldest post-kinematic granites, dated as 320 Ma in the northern and 335-325 Ma in the southern parts of the Black Forest, suggest that in this area crustal shortening gradually decreased during mid- to late Vis6an times (Henk 2000; Schaltegger 2000). Namurian and early Westphalian continental clastic deposits, resting on mid-crustal rocks, accumulated in transtensional basins (Eisbacher et al. 1989). However, along the Bohemian-Moldanubian suture, crustal shortening and wrench faulting persisted into the Late Carboniferous, as evidenced by the tectonostratigraphic record of the Southern Vosges (Maass 1988). The total amount of crustal shortening accommodated at the south-dipping Saxo-Thuringian-Bohemian suture is estimated at a minimum of 180 km. An additional 150km of shortening is thought to have occurred at the north-dipping Bohemian-Moldanubnian suture (Franke 2000). Total Carboniferous lithospheric shortening in the area transected by the ECRIS presumably exceeded 600 km. Of this, some 300 km can be attributed to late Vis6an to Westphalian shortening in the Rheno-Hercynian pro-wedge and the Saxo-Thuringian retro-wedge thrust belts. During the Namurian and Westphalian, deformation of the internal zones of the Variscan orogen was dominated by lateral escape of the relatively rigid Saxo-Thuringian and Bohemian-Moldanubian terranes and the subsidence of intramontane transtensional basins in which neo-autochthonous continental clastic deposits accumulated (e.g. Saar-Nahe, Saale, Pilzen basins: Ziegler 1990; Dallmeyer et al. 1995; Henk 1995, 1999; Oncken et al. 1999). Progressive uplift of the

EVOLUTION OF VARISCANLITHOSPHERE Saxo-Thuringian-Rheno-Hercynian collision zone commenced during the Namurian (327-316 Ma), in response to underplating by the foreland lithosphere and basal accretion of crustal flakes (Oncken et al. 2000). By end-Westphalian times (305 Ma), the continental foreland lithosphere underlying the Rheno-Hercynian zone dipped southward under the Rheno-Hercynian-Saxo-Thuringian suture, forming a subduction slab about 200 km long. The sedimentary cover and parts of the upper crust of the foreland lithosphere were incorporated into the Rheno-Hercynian thrust belt. The Rheno-Hercynian-Saxo-Thuringian suture was probably characterized by a thermally destabilized and orogenically thickened crust ( _+60 km). For the Saxo-Thuringian, Bohemian-Armorican and Moldanubian-Arverno-Vosgian zones a synorogenic thermally destabilized 45-60 km thick crust can be visualized, considering that late Visran (330 Ma) exhumation of high-pressure rocks in the Saxonian Granulite Massif involved synorogenic lower-crustal flow (Franke & Stein 2000; Henk 2000). Similarly, StephanianAutunian extensional uplift of core-complexes in the Massif Central (Malavieille et al. 1990; Brun & van den Driesschen 1994; Burg et al. 1994; Lardeaux et al. 2001) involved crustal flow. In this respect, it should be noted that extensional middle and lower crustal flow occurs only when thick, rheologically weak crust (>45 km, felsic composition, elevated temperatures, and/or presence of partial melts) is stretched at high strain rates (Burov & Cloetingh 1997; Bertotti et al. 2000; Vanderhaeghe & Teyssier 2001). Furthermore, outcropping plutons, with crystallization depths between 8 and 20 km (Altherr 1999; Vigneresse et al. 1999), indicate that substantial amounts of upper and middle crustal material were eroded during late and post-orogenic times. This implies considerably greater crustal thicknesses than the now observed 28-35 km (Seyferth & Henk 2000). The widespread Visran to Westphalian magmatic activity in the Saxo-Thuringian, Bohemian-Armorican and MoldanubianArverno-Vosgian zones indicates that their asthenospherelithosphere system was far from an equilibrium steady-state regime at the end of the Variscan orogeny. This magmatic activity can be related to late synorogenic slab detachment, upwelling of the asthenosphere, thermal thinning and/or partial delamination of the lithospheric mantle, and crustal melting that was accompanied by transtensional unroofing of the crust (Rey et al. 1992; Henk 2000; Ledru et al. 2001). In summary, we emphasize that in the area of the future ECRIS the configuration of the lithosphere at the end of the Variscan orogeny can only be loosely defined.

Stephanian-Early Permian disruption of the Varisean orogen End-Westphalian consolidation of the Variscan orogen was followed by its Stephanian-Early Permian wrench-induced collapse (305-280Ma). Continental-scale dextral shears, such as the Tornquist-Teisseyre and the Bay of Biscay fractures zones, were linked by secondary sinistral and dextral shear systems. Together, they overprinted and partly disrupted the Variscan orogen and its northern foreland. This deformation, which reflects a change in the Gondwana-Laurussia convergence from oblique collision to a dextral translation, was kinematically linked to Permo-Carboniferous crustal shortening in the Appalachian and the Scythian orogens (Arthaud & Matte 1977; Ziegler 1989, 1990; Coward 1993). Significantly, wrench tectonics and associated magmatic activity abated in the Variscan domain and its foreland at the transition to the Late Permian, in tandem with the consolidation of the Appalachian orogen (Ziegler 1989, 1990; Marx et al. 1995; Ziegler & Stampfli 2001; Ziegler & D~zes 2006).

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Stephanian-Early Permian wrench-induced disruption of the rheologically weak Variscan orogen was accompanied by regional uplift, widespread extrusive and intrusive magmatism peaking during the Early Permian, and the development of a multidirectional array of transtensional trapdoor and pull-apart basins in which continental clastic deposits accumulated (Fig. 4; Ziegler 1990). Microtectonic analyses in the Massif Central indicate that during Stephanian-Early Permian times the principal horizontal compressional stress axis progressively rotated from north-south to east-west (Bl~s et al. 1989). Basins, which developed during this time span, show a complex, polyphase structural evolution, including a late phase of transpressional deformation controlling their partial inversion (Ziegler 1990). Although Stephanian-Early Permian wrench deformation locally gave rise to uplift of core-complexes (e.g. Massif Central: Malavieille et al. 1990; Burg et al. 1994; Lardeaux et al. 2001; Montagne Noire: Brun & van den Driesschen 1994; see also Vanderhaeghe & Teyssier 2001), crustal stretching factors were relatively low on a regional scale. This is evidenced, for instance, in the area of the Southern Permian Basin that is located in the Variscan foreland and encroaches in its eastern parts on the Rheno-Hercynian thrust belt (Ziegler 1990; van Wees et al. 2000). Regional erosional and locally confined tectonic unroofing of the Variscan orogen, as well as the interaction of mantle-derived basic melts with the felsic lower crust, contributed to a re-equilibration of the Moho at depths of 28-35 km and, locally, less. By MidPermian times (_+ 280 Ma), some 25 Ma after consolidation of the Variscan orogen, its crustal roots had apparently been destroyed. Although the model of the Cenozoic Basin-and-Range Province has been repeatedly invoked for the Stephanian-Early Permian collapse of the Variscan orogen (Lorenz & Nicholls 1976, 1984; Jowett & Jarvis 1984; Mrnard & Molnar 1993; Malavieille 1993; Prijac et al. 2000), there are fundamental differences in the kinematics controlling the development of these two provinces. This concerns mainly the dominantly wrench-induced collapse of the Variscan domain versus the extension-dominated collapse of the Cordillera, as well as their megatectonic setting (Himalayan-type western Variscides, Andean-type eastern Variscides and Cordillera) (Ziegler 1990; Ziegler & Stampfli 2001). Permo-Carboniferous

magmatism and lithospheric

destabilization

The widespread Stephanian-Early Permian (305-285 Ma) alkaline intrusive and extrusive magmatism of the Variscan domain and its northern foreland is mantle derived and shows evidence of strong crustal contamination (Bonin 1990; Bonin et al. 1993; Neumann et al. 1995; Marx et al. 1995; Benek et al. 1996; Cortesongo et al. 1998; Breitkreuz & Kennedy 1999). Melt generation by partial melting of the uppermost asthenosphere and the lithospheric thermal boundary layer was probably triggered by a rise in the potential temperature of the asthenosphere and localized transtensional decompression. Upwelling of the asthenosphere was presumably induced by the detachment of deep-reaching subducted lithospheric slabs, causing a reorganization of the mantle convection system and the impingement of a system of not very active mantle plumes onto the base of the lithosphere, partly at considerable distances to the north of the Variscan orogen (e.g. Oslo Graben, northern British Isles; see Heeremans et al. 2004; Ziegler & D~zes 2006). Supporting evidence for a contribution from deeper mantle sources comes from the isotopic signature of the most primitive melts (Neumann et al. 2004). Crustal-scale fractures provided avenues for magma ascent to the surface. Although uplift and exhumation of the Variscan internides had commenced already during the Late Devonian and accelerated during the Late Visran and Namurian, regional uplift of the

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Fig. 4. Stephanian-Early Permian tectonic framework of ECRIS area, showing sedimentarybasins (vertical shading), major volcanic fields (horizontal shading)and

sills (cross-hatched), and fault systems (modifiedafter Ziegler 1990) with superimposedVariscan terrane boundaries (continuous barbed grey lines) and Alpine deformation front (dashed, barbed grey line). * Analysed wells (for names see Fig. 6). BU, Burgundy Trough; KT, KraichgauTrough; SB, SchrambergTrough; SN, Saar-Nahe Trough. entire orogen and its foreland began only after crustal shortening had ceased at the end-Westphalian. Stephanian-Early Permian uplift and erosional, as well as tectonic, unroofing of the Variscan orogen, in many areas to formerly mid-crustal levels (Burg et al. 1993; Vigneresse 1999; Seyferth & Henk 2000), can be attributed to a combination of wrench deformation, heating of crustal roots and related eclogite to granulite transformation (Bousquet et al. 1997; Le Pichon et al. 1997), detachment of subducted slabs, upwelling and partial melting of the asthenosphere causing thermal attenuation and partial delamination of the lithospheric mantle, and magmatic inflation of the lithosphere. Mantle-derived basic melts, which had ascended to the base of the crust, underplated it, inducing crustal anatexis, and the intrusion of fractionally crystallized granitic to granodioritic-tonalitic melts into the crust (Cortesogno et al. 1998; Breitkreuz & Kennedy 1999). This can be observed at the Ivrea Zone of the Southern Alps, where at basal crustal levels Early Permian asthenosphere-derived mafic intrusions caused partial melting of metasediments, triggering the ascent of granitic magmas into the middle and upper crust and volcanic activity in a contemporaneous wrench-induced sedimentary basin (Schmid 1993; Wittenberg et al. 2000). P e r m o - C a r b o n i f e r o u s crustal thinning in the Southern P e r m i a n B a s i n

The Southern Permian Basin, although largely located in the Variscan foreland, provides a model for Stephanian-Early Permian lithospheric destabilization that can also be applied to the post-orogenic development of the internal parts of the Variscan orogen. In the area of the Southern Permian Basin, up to 800 m thick Stephanian continental red beds were deposited in a broad successor basin to the Namurian-Westphalian Variscan foreland

basin. This Stephanian basin was disrupted during the late Stephanian-Early Permian by predominantly transpressive wrench tectonics, as evidenced by the deep truncation of Late Carboniferous series and the conspicuous absence of deep Early Permian sedimentary basins (Ziegler 1990; McCann 1999). This wrench deformation was accompanied by extensive volcanic activity, particularly in NE Germany and NW Poland (Ziegler 1990; Plein 1995; Breitkreuz & Kennedy 1999; Scheck & Bayer 1999). The crystalline crust of the Southern Permian Basin thins from 32 km beneath its northern flank to about 22 km under the basin centre (Bayer et al. 1999). This can be translated into a 'stretching' factor of about 1.45. As NE Germany was apparently only marginally affected by rifting during the Early Devonian opening of the oceanic Giessen-Harz Basin, and was not overprinted by Mesozoic rifting (Kossow & Krawcyk 2002), the observed crustal thinning must be attributed to late StephanianEarly Permian magmatic destabilization of the crust-mantle boundary that was paralleled by major thermal attenuation of the lithospheric mantle (van Wees et al. 2000). In the area of the Southern Permian Basin, the Moho is overlain by a 2 - 5 km thick band of subparallel, anastomosing high-amplitude reflectors (Bayer et al. 1999), which can be related to mantle-derived basaltic sills and possible metamorphic layering of the lower crust (Meissner 1986; Ziegler 1996; Meissner & Rabbel 1999; Ziegler & Cloetingh 2004). The observed crustal thinning, combined with major crustal contamination of extruded magmas (Benek et al. 1996; Breitkreuz & Kennedy 1999) and the presence of a thick high-density lower crust (Bayer et al. 1999), indicates magmatic destabilization of the Moho in the presence of a felsic lower crust. Subcrustal velocities of 8.0-8.1 km s -1 (Bayer et al. 1999) suggest that mafic melts and cumulates that had underplated the crust cooled in time into the eclogite stability field at depths of 30-33 km and at Moho temperatures around 600 ~ (Griffin et al. 1900).

EVOLUTION OF VARISCANLITHOSPHERE P e r m o - C a r b o n i f e r o u s evolution o f the E C R I S Z o n e

During the Stephanian-Early Permian, a system of essentially NE-SW-trending transtensional intramontane basins developed in the area of the Bresse, Limagne and Upper Rhine grabens, whereas NW-striking fracture systems transected the Rhenish Massif (Fig. 4). Development of these basins, which contain thick continental clastic deposits and volcanic units, involved reactivation of the Variscan structural grain, predominantly by dextral shear. The Saar-Nahe Trough is superimposed on the Rheno-Hercynian-Saxo-Thuringian and partly on the SaxoThuringian-Bohemian sutures. The Kraichgau Trough broadly reflects reactivation of the Saxo-Thuringian-Bohemian suture. The Schramberg, Burgundy and Jura troughs can be considered as being associated with the Bohemian-Moldanubian suture. Evolution of these basins was coupled with uplift and erosion of intervening highs and the development of NNE-SSW-trending sinistral shears, partly outlining the Upper Rhine and Hessian grabens (Boigk & Sch6neich 1970; Eisbacher et al. 1989; Weber 1995a; Schumacher 2002). Post-orogenic uplift of the Rheno-Hercynian thrust belt, prior to the deposition of Late Permian sediments, amounted along its northern margin to 2 - 3 km, increased southwards to 6 k m (Littke et al. 2000) and in the area adjacent to the Saar-Nahe Basin reached some 10 km (Oncken et al. 2000). Transtensional subsidence of the partly inverted Saar-Nahe Basin, which contains up to 5.6km of Permo-Carboniferous clastic deposits, accounts for a stretching factor of > 1.36. Contemporaneous extrusion of voluminous basalts and rhyolites in this basin, dated as 300-297 Ma, reflects profound destabilization of its lithospheric system (Henk 1993; Stollhofen & Stanistreet 1994; Korsch & Sch~ifer 1995; Weber 1995b). In the area of the Saxo-Thuringian retro-wedge thrust belt, late to post-kinematic granites range in age from 313 to 282 Ma, with increasing post-kinematic magmatism being attributed to crustal melting processes. Stephanian exhumation of granites and diorites (298-289 Ma), partly involving transtensional faulting (Thomson & Zeh 2000), reflects some 10 km of post-orogenic uplift of the Saxo-Thuringian retro-wedge thrust belt prior to the deposition of Late Permian sediments (Seyferth & Henk 2000). In view of the above, we postulate that under the Stephanian stress field the NE-trending German part of the Rheno-HercynianSaxo-Thuringian suture was transtensionally strongly reactivated, whereas its NW-trending French and British and its south-trending Polish parts were only mildly reactivated. Wrench-induced reactivation of the German part of this suture caused detachment of the subducted continental Rheno-Hercynian slab that had dipped beneath the Mid-German Crystalline Rise. Passive upwelling of the asthenosphere into the space formerly occupied by this slab triggered partial melting of the asthenosphere and remnant lithospheric mantle, ascent of melts to the base of the crust and anatexis of lower crustal rocks. In conjunction with a general reorganization of asthenospheric flow patterns, a not very active mantle plume apparently welled up to the base of the lithosphere to the north of the RhenoHercynian-Saxo-Thuringian suture in the area of the eastern parts of the future Southern Permian Basin, causing thermal attenuation of the lithospheric mantle and magmatic destabilization of the crust-mantle boundary. A branch of the South Permian Basin upwelling system extended southwestward into the area of the Saar-Nahe Trough. This is compatible with the occurrence of a distinct middle and lower crustal positive magnetic anomaly that can be traced from NE Germany into the area of the Saar-Nahe Trough (Hahn & Wonik 2002); this anomaly is attributed to the injection of mantle-derived basic melts into the lower and middle crust. On the other hand, the Rhenish Massif was little affected by magmatic processes but was subjected to regional uplift, reflecting unflexing of the lithosphere in response to slab detachment and possibly also retrograde eclogite to granulite metamorphism of lower crustal material.

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In the area of the Upper Rhine, Bresse and Limagne grabens, development of a system of NE-SW-trending Stephanian-Early Permian fault-bounded troughs (Boigk & Sch6neich 1970; Philippe et al. 1996), overlying high-grade crystalline rocks, reflects transtensional reactivation of the Variscan orogenic fabric. Subsidence of these basins and uplift of intervening highs in response to extensional unloading of the lithosphere was accompanied by the intrusion of abundant rhyolite and granite porphyry dykes and the extrusion of Permian rhyolites. This reflects the ascent of mantle-derived partial melts to the base of the crust, lower crustal anatexis and destabilization of the Moho in response to a major thermal surge (Eisbacher et al. 1989; von Raumer 1998; Prijac et al. 2000; Lardeaux et al. 2001). In zones of Permo-Carboniferous crustal extension or transtension, mechanical thinning of the crust may have significantly contributed to the transformation of the orogenically thickened Variscan crust to an end-Mesozoic average thickness of 2 8 35 kin. However, for intervening unextended areas this mechanism cannot be invoked, unless diffuse extension of the lower crust and lithospheric mantle is assumed (Henk 1999). On the other hand, in areas of core-complex development (Vanderhaeghe & Teyssier 2001), lower and middle crustal ductile flow towards extensional zones may have contributed to regional thinning of the crust. Deep seismic reflection lines show that in the Saxo-Thuringian and Bohemian-Moldanubian zones the orogenic fabric of the Variscan crust extends from upper crustal levels down to the Moho at which it either soles out or is truncated (DEKORP Research Group 1988, 1994; Vollbrecht et al. 1989; Bois et al. 1990; Meissner & Bortfeld 1990; Bankwitz & Bankwitz 1994). Moreover, seismic velocity analyses and the study of xenoliths contained in Cenozoic extrusive rocks indicate that the lower crust is characterized by a felsic composition and that basic material occurs only near the crust-mantle transition (Mengel 1992; Downes 1993; Wittenberg et al. 2000). As the geophysical and petrological Moho do not always appear to coincide (Wittenberg et al. 2000), it is likely that during orogenic processes crustal material was eclogitized, in the process of which, and depending on its composition, its P-wave velocity increased to 8 . 0 - 8 . 4 k m s -1, and thus was transferred across the Moho, defined as the break-over from Vp _< 7.8 to 8.0 km s -1, into the continental lithospheric mantle (Ziegler et al. 1985; Bousquet et al. 1997; Stampfli et al. 1998). Similar to the Southern Permian Basin, the lower crust of the Variscan internal zones is in many places characterized by a highly reflective band that parallels the Moho and overprints the crustal orogenic fabric (Eisbacher et al. 1989; Meissner & Rabbel 1999). In contrast to the Southern Permian Basin, this 'laminated' lower crust can attain in the Variscan internides thicknesses of 10-15 km (Fig. 5). Development of such a thick laminated lower crust may be related to the intrusion of mantle-derived sills into a crust that was already thermally destabilized by Namurian and Westphalian magmatism. However, a distinction must be made between 'pristine' Variscan crust (e.g. DEKORP line D-2S) and crust that was overprinted by Cenozoic rifting and magmatism (e.g. DEKORP K-8401, MVE-90 East, D-3A; for line location see Fig. 3). For instance, lines crossing the Rhenish Massif show that in their northern parts the lower crust is non-reflective, whereas southward lower crustal lamination exists in areas of Cenozoic volcanism and persists in their southern parts (e.g. lines D-1 and D-2N). On line K-8401, which runs along the axis of the Black Forest and parallels the Upper Rhine Graben, lower crustal lamination is well developed but decreases southward. On line MVE-90, which parallels the Cenozoic Eger graben, lower crustal lamination is very well expressed. Therefore, the question arises whether lower crustal lamination can be exclusively attributed to the PermoCarboniferous destabilization of the crust-mantle boundary or whether Cenozoic rift-related magmatic activity contributed to it

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P.A. ZIEGLERETAL.

Fig. 5. Segment of DEKORP 2S deep seismic reflectionline, crossing the FranconianPlatform and showing subhorizontal,high-amplitudelower crustal lamination in the two-way travel time (TWT) interval of 5.0-9.5 s, which overprints the orogenic fabric of the crust, as evident particularlyin the Saxo-Thuringianparts of this line. The Moho discontinuitycorresponds to the base of the laminated interval (after Meissner & Bortfeld 1990). This line segmentis located outside the confines of Permo-Carboniferoustroughs. (For location see Fig. 3, interrupted part of line D-2S.)

(Eisbacher et al. 1989) or was even exclusively responsible for its development. An alternative model proposes that the laminated lower crust of the Variscan internides developed in response to its deformation by distributed horizontal ductile shear during Permo-Carboniferous crustal extension, a mechanism that may have contributed, at least on a local scale, to post-orogenic crustal thinning (Rey 1993; Henk 1999). However, as PermoCarboniferous extension cannot be invoked for the Rhenish Massif, development of lower crustal lamination in this area probably has to be attributed to Cenozoic volcanism with a possible contribution from cryptic Permo-Carboniferous magmatic activity along its southern margin, adjacent to the Saar-Nahe Basin. Similarly, it is uncertain whether lower crustal lamination evident in the Black Forest area (Eisbacher et al. 1989) and along the margins of the Eger volcanic-tectonic zone can be exclusively attributed a Permo-Carboniferous age. However, lower crustal lamination evident in the Paris Basin (Cazes & Toreilles 1988) can only be attributed to the Stephanian-Early Permian tectonomagmatic cycle, which in this area involved only minor crustal extension, as indicated by the distribution of Stephanian-Early Permian basins and their close association with multi-directional wrench faults (Fig. 4). Similarly, lower crustal lamination evident on line D-2S (Fig. 5), which is clearly located outside the confines of Permo-Carboniferous troughs, developed in all likelihood during Permo-Carboniferous times. Invoking the Southern Permian Basin model, we prefer to relate development of lower crustal laminations in those parts of the Variscan internides that were not overprinted by Cenozoic rifting to PermoCarboniferous intrusion of mantle-derived basic sills, with lower crustal ductile shear playing a subordinate role. The widespread occurrence of a bimodal Permo-Carboniferous intrusive and extrusive magmatism in the Variscan internides, indicative of lower crustal melting (Henk 1999), combined with the regional distribution of a primarily Permo-Carboniferous laminated lower crust, is taken as evidence for magmatic destabilization of the crust-mantle boundary in response to a major Permo-Carboniferous thermal surge. The resulting upward displacement of the Moho contributed to thinning of the Variscan crust. Contemporaneous thermal thinning and possibly partial delamination of the lithospheric mantle, as well as magmatic inflation of the remnant lithospheric mantle, contributed to regional uplift and erosional unroofing of the Variscan crust and thus to its thinning. Extensional unroofing played an important role in crustal thinning only locally. Post-intrusion unroofing of Late Variscan granitic bodies amounted to 4 - 9 km with tectonic

processes perhaps playing a greater role than erosion in bringing high-grade metamorphic rocks to the surface (Vigneresse 1999; Sch~ifer et al. 2000; Lardeaux et al. 2001). For the Permo-Carboniferous evolution of the Variscan internides we visualize that the inferred major thermal surge, controlling thermal thinning of the lithospheric mantle, magmatic destabilization of the crust-mantle boundary and regional uplift, involved upwelling of the asthenosphere in response to rollback and detachment of the subducted lithospheric slabs, such as that associated with the Bohemian-Armorican to MoldanubianArverno-Vosgian suture (Matte 1991; Najoui et al. 2000; Lardeaux et al. 2001). By the end of Early Permian times, crustal thicknesses had been reduced on a regional scale to 28-35 km. Significantly, no mantle reflectors related to subducted crustal and lithospheric mantle material (Ziegler et al. 1998) could be detected in the Variscan Internides, despite dedicated surveys (Meissner & Rabbel 1999). This may be an effect of strong Permo-Carboniferous thermal thinning of the lithospheric mantle.

Late Permian and Mesozoic thermal subsidence and rifting By late Early Permian times ( _ 280 Ma), magmatic activity had abated and thermal anomalies introduced during the PermoCarboniferous tectonomagmatic cycle began to decay, controlling regional thermal subsidence of the lithosphere. In combination with progressive erosional degradation of the remnant topography and cyclically rising sea levels (Haq et al. 1988), increasingly larger areas subsided below the erosional base level and were incorporated into a new system of intracratonic basins. However, in large parts of Western and Central Europe thermal re-equilibration of the lithosphere-asthenosphere system was overprinted and partly interrupted by the Triassic onset of a new rifting cycle that preceded and accompanied the stepwise break-up of Pangaea. Major elements of this break-up system are the southward-propagating Arctic-North Atlantic and the westward-propagating Neotethys rift systems (Ziegler & Stampfli 2001). At the same time a multi-directional rift system developed in Western and Central Europe, comprising the North Sea rift, the Danish-Polish Trough and the graben systems of the Atlantic shelves. Stress fields controlling the evolution of this rift system changed repeatedly during late Mid-Jurassic and Early Cretaceous

EVOLUTION OF VARISCANLITHOSPHERE

105

Fig. 6. Isopach map of restored Triassic series, contour interval 500 m (after Ziegler 1990), showing location of analysed wells and Variscan (continuous barbed line) and Alpine (dashed barbed line) deformation fronts. Horizontal shading indicates area not mapped. BU, Burgundy Trough; FP, Franconian Platform; GG, Gltickstadt Graben; HD, Hessian Depression; KT, Kraichgau Trough; NP, Nancy-Pirmasens Trough; PB, Paris Basin; PT, Polish Trough; SP, Southern Permian Basin; TB, Trier Basin; WN, West Netherlands Basin. times prior to the Late Cretaceous concentration of tiffing activity on the Norwegian-Greenland Sea area (Ziegler 1990; Stampfli 1993; Ziegler et al. 2001; Ziegler & D~zes 2006). Although much of the ECRIS area was only marginally affected by Mesozoic rifting, minor diffuse crustal stretching probably contributed to the subsidence of the Kraichgau, Nancy-Pirmasens, Burgundy and Trier troughs (Fig. 6). Triassic and Jurassic reactivation of Permo-Carboniferous faults, controlling subtle lateral facies and thickness changes, is also evident in the Paris Basin (Bessereau et al. 1995; Goggin et al. 1997) and in the area of the Burgundy Trough (Wetzel et al. 2002). On the other hand, Mesozoic crustal extension played a more important role in the subsidence of the West Netherlands Basin and its prolongation into the area of the future Lower Rhine Graben, as well as in the lower Rh6ne Valley (see Ziegler & D~zes 2006). In an effort to quantify Late Permian and Mesozoic vertical movements of the lithosphere in the wider ECRIS area, we carried out subsidence analyses on selected wells from the Paris Basin, the Upper Rhine Graben and the Franconian Platform, applying the back-stripping method of Christie & Sclater (1980). Resulting tectonic subsidence curves, similar to those by Loup & Wildi (1994), Prijac et al. (2000) and van Wees et al. (2000), show that re-equilibration of the lithosphere with the asthenosphere commenced during the late Early Permian ( + 2 8 0 Ma) and continued throughout Mesozoic times. Detailed tectonic subsidence curves show that, superimposed on the long-term thermal subsidence trends, intermittent and generally local subsidence accelerations occurred during the Mesozoic (Fig. 7) (Ziegler et al. 2004). These reflect either tensional reactivation of Permo-Carboniferous fault systems or compressional deflection of the lithosphere (Cloetingh 1988) under stress fields related to far-field rifting and wrench activity. Temporal and spatial variations in these subsidence accelerations probably relate to differences in the orientation of pre-existing crustal

discontinuities and changes in the prevailing stress field. However, despite these anomalies, overall subsidence trends clearly reflect thermal re-equilibration of the lithosphere-asthenosphere system through time.

Tectonic subsidence modelling In an attempt to define the end-Early Permian configuration of the lithosphere, we compared the tectonic subsidence curves to a theoretical thermal decay curve, applying a numerical forwardbackward modelling technique which automatically finds the best-fit stretching parameters for the observed subsidence data (van Wees et al. 1996, 2000). Forward-backward modelling of tectonic subsidence is based on lithospheric stretching assumptions where (6 is the crustal stretching factor and /3 is the lithospheric mantle stretching factor) under which the lithosphere is represented by a plate with constant temperature boundary conditions, adopting a fixed basal temperature (McKenzie 1978; Jarvis & McKenzie 1980; Royden & Keen 1980). For thermal calculations, a 1D numerical finite-difference model was used, adopting parameters as given by van Wees et al. (2000), which allows for incorporation of finite and multiple stretching phases, as well as for crustal heat production effects and conductivity variations (van Wees et al. 1992, 1996, 2000; van Wees & Stephenson 1995). Differential stretching of the crust and lithospheric mantle can be applied to simulate thermal attenuation of the latter. Input parameters for forward-backward modelling of the observed subsidence curves include the pre-rift crustal and present lithospheric thickness, and for each stretching phase its timing, duration and mode of lithospheric extension (uniform 6 =/3, McKenzie 1978; twolayered 6
P.A. ZIEGLER ETAL.

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Fig. 7. Example of (a) an air-loaded tectonic subsidence curve and (b) a modelled subsidence curve (well Bourneville, Paris Basin; for locations see Fig. 6). Control points derived from the penetrated sedimentary sequence. The positive part of the modelled subsidence curve reflects uplift of the crust in response to thermal thinning and/or delamination of the lithospheric mantle, whereas its negative part reflects subsidence of the crust during thermal re-equilibration of the lithosphere-asthenosphere system.

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parameters are changed until a good fit is obtained between the observed and modelled subsidence curves. The best-fit stretching parameters thus determined give a measure of the PermoCarboniferous thermal perturbation of the lithosphere, as well as during subsequent tensional events that interfered with the Late Permian and Mesozoic re-equilibration of the asthenospherelithosphere system. Our modelling of the lithosphere evolution in the ECRIS area is based on the concept that, after the Permo-Carboniferous thermal surge ( 3 0 0 - 2 8 0 M a ) , the temperature of the asthenosphere returned rapidly to ambient levels (1300 ~ at which it remained until the end-Cretaceous renewed flare-up of plume activity. Therefore, we adopted in our forward-backward model relatively low lithospheric thicknesses of 1 0 0 - 1 2 0 km that, according to Babuska & Plomerova (1992, 1993), are representative for areas not affected by Cenozoic rifting. Furthermore, as most of the analysed wells are located outside Permo-Carboniferous troughs, initial crustal thicknesses were assumed to be close to the present values. The subsidence curves were modelled with PermoCarboniferous differential crustal and lithospheric mantle extension (attenuation), allowing /3 factors to attain significantly greater values than 6 factors. The high /3 factors represent the effects of delamination and thermal thinning of the lithospheric mantle. On the other hand, the temporary Mesozoic subsidence accelerations were successfully modelled with uniform lithospheric extension ( 6 = / 3 ) . The modelled subsidence curves (Fig. 8) demonstrate that, after an initial uplift phase between

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300 and 280 Ma, which gives a measure of Stephanian-Early Permian lithospheric thinning, the evolution of the lithosphere in the ECRIS area was governed from 280 Ma onward by the long-term decay of thermal anomalies that were introduced during the Permo-Carboniferous tectonomagmatic cycle. This is

EVOLUTION OF VARISCAN LITHOSPHERE

in accordance with the assumption that the temperature of the asthenosphere returned rapidly to ambient levels after the Permo-Carboniferous thermal surge (Ziegler et al. 2004). Good fits between observed and modelled tectonic subsidence curves were obtained, assuming initial crustal thicknesses of 30-35 kin, final lithospheric thicknesses of 100-120km, and a Permo-Carboniferous 'stretching' phase that spanned 3 0 0 - 2 8 0 M a and involved decoupled crustal extension and attenuation of the lithospheric mantle. These assumptions are compatible with the concept that during the Permo-Carboniferous re-equilibration of the crust-mantle boundary, crustal extension played a signifcant role only locally. In this respect it is noteworthy that the major Permo-Carboniferous troughs, which occur in the Massif Central and the Bohemian Massif and beneath the Franconian Platform, do not coincide with major Late Permian and Mesozoic depocentres, whereas no major Permo-Carboniferous basins are located under the Southern Permian Basin and Paris Basin depocentres (Ziegler 1990). This suggests that during the Permo-Carboniferous tectonomagmatic cycle uniform and/or depth-dependent lithospheric extension was, on a regional scale, only a contributing but not the dominant mechanism of crustal and lithospheric mantle thinning, as advocated for the Paris Basin by Prijac et al. (2000). In contrast, lithospheric stretching may have played a somewhat more important role in the evolution of the Hessian Depression, NancyPirmasens and Burgundy system of Late Permian and Mesozoic basins that is superimposed on a Basin and Range type array of Permo-Carboniferous troughs (Figs 4 and 5). Modelled subsidence curves reflect that during the PermoCarboniferous tectonomagmatic cycle the lithospheric mantle was significantly attenuated and that/3 factors attained values in the range of 1.8-10 (Table 1). As these values are subject to large lateral variations, they reflect that thinning of the lithospheric mantle was heterogeneous and generally more intense in areas that evolved into Mesozoic depocentres than in areas marginal to them. Similarly, in areas that remained positive features throughout most of Mesozoic times, such as the Bohemian and Armorican massifs, the lithospheric mantle was apparently not significantly thinned during the Permo-Carboniferous and retained a thickness of 70-100 kin, as well as an orogen (subduction)related anisotropy (Babuska & Plomerova 2001; Judenherc et al. 2002). Sensitivity studies indicate that the best fits between

Table 1. Input parameters and modelling results for Permo-Carboniferous thermal destabilization of the lithosphere (after Ziegler et al. 2004) Basin and well name

Input EL

Calculated IC

Paris Basin Lyon-la-For& 120 35 Bourneville 120 35 Champotran 120 35 Sennely 120 35 Lorraine and Rhine Graben area Trois-Fontaines 120 35 Wiesloch-Neibsheim 100 30 Freiburg 100 30 Otterbach 100 30 Franconian Platform Benken 100 30 Trochtelfingen 100 30 Aalen 100 30

End-Early Permian

/3

6

RLM

RC

RL

10.00 9.99 5.08 1.77

1.04 1.08 1.07 1.03

8.5 8.5 17.0 48.1

33.6 33.4 33.0 34.0

42.1 41.9 50.0 82.1

3.06 5.79 3.89 3.72

1.06 1.13 1.06 1.12

27.8 12.1 18.0 18.8

33.0 26.5 28.3 26.7

60.8 38.6 46.3 45.5

2.42 3.02 4.49

1.04 1.05 1.04

28.9 23.2 15.6

28.8 28.6 28.8

57.7 51.8 44.4

EL, equilibrated lithosphere thickness in krn; IC, initial crustal thickness in km; ]3, lithospheric mantle attenuation factor; 3, crustal stretching factor; RLM, remnant lithospheric mantle thickness in km; RC, remnant crustal thickness in km; RL, remnant lithosphere thickness in km.

107

observed and modelled subsidence curves were obtained when the thickness of the thermal lithosphere at its end-Mesozoic equilibration with the asthenosphere was set at 100 or 120 km. Yet, even at these values, Permo-Carboniferous /3 factors and the end-Early Permian remnant thickness of the lithospheric mantle (RLM) varied significantly (e.g. Trochtelfingen: 100 km lithosphere, /3 = 3.02, RLM 23.2 km; 120 km lithosphere, /3 = 1.96, RLM 43.3 km). For the Paris Basin, the best fit between observed and modelled subsidence curves was achieved with a lithosphere thickness of 120 km, whereas for the Upper Rhine Graben and the Franconian Platform best fits were obtained with a lithosphere thickness of 100 km. Whereas a 100 km lithosphere thickness is compatible with the Palaeocene plume-related segregation depth of olivine-melilitic partial melts in the Vosges, Black Forest and Bohemian Massif (Wilson et al. 1995), we have no explanation for the apparently greater lithosphere thickness beneath the Paris Basin. In view of the above, values given in Table 1 for the end-Early Permian thickness of the RLM should be regarded as rough approximations. Nevertheless, we conclude that substantial Permo-Carboniferous thinning of the lithospheric mantle provided the principal driving mechanism for the Late Permian and Mesozoic subsidence of thermal sag basins that developed in the ECRIS area. On a regional scale, modelled Permo-Carboniferous crustal extension was relatively low. Under the assumption of initial crustal thicknesses of 30-35 km, automated modelling yields 6 factors of 1.04-1.13 and crustal thicknesses close to the present values. The minor, intra-Mesozoic subsidence accelerations, which overprint the long-term thermal subsidence curves, were successfully modelled by uniform lithospheric extension with cumulative = / 3 values in the range of 1.01-1.07 (Ziegler et al. 2004). As corresponding extensional faulting is generally poorly documented, stress-induced deflections of the lithosphere (Cloetingh 1988) may have contributed to some of these subsidence anomalies.

Discussion and conclusions A conceptual model for the Late Carboniferous to endCretaceous evolution of the lithosphere along a transect that extends from the Lower Rhine Graben across the Rhenish Massif and the Upper Rhine Graben into the area of the Jura Mountains is shown in Figure 9. By end-Westphalian times, when crustal shortening had ceased in the Variscan orogen, the crustal and lithospheric configuration of the ECRIS area was heterogeneous. Whereas the Rheno-Hercynian zone was underlain by continental foreland lithosphere, the Saxo-Thuringian, Bohemian and Moldanubian zones were characterized by an orogenically thickened lithosphere that was thermally destabilized, as evidenced by widespread granitic magmatism. In the internal zones of the Variscan orogen, crustal thicknesses were probably in the range of 45-60 kin, with crustal roots marking the Rheno-Hercynian- Saxo-Thuringian, Saxo-ThuringianBohemian and Bohemian-Moldanubian sutures. An approximately 200 km long south-dipping subducted continental lithospheric slab extended from the northern foreland under the Rheno-Hercynian-Saxo-Thuringian suture. The oceanic part of this slab, corresponding to the Giessen-Harz Basin, had been already detached from it during the mid-Vis6an. A major, north-dipping subducted slab, consisting of oceanic and continental material, was probably still associated with the Bohemian-Moldanubian suture, whereas the south-dipping Saxo-Thuringian-Bohemian slab had already been detached from the lithosphere during mid-Vis6an times (Fig. 9a). During the Stephanian and Early Permian dextral translation of Gondwana and Laurussia, transtensional and transpressional

108

P.A. ZIEGLER ETAL.

Permian times the thickness of the remnant lithospheric mantle varied between 9 km and 5 0 k m in areas that subsequently evolved into Late Permian and Mesozoic thermal sag basins; (3) thinning of the lithospheric mantle was heterogeneous and more intense in areas that developed into Late Permian-Mesozoic depocentres (e.g. the Southern Permian and Paris basins) than beneath flanking areas and persisting highs; (4) there is no obvious relationship between the degree of lithospheric mantle thinning and the various Variscan tectonic units. We conclude that during the Permo-Carboniferous tectonomagmatic cycle mechanical stretching of the lithosphere played a subordinate role, whereas thermal thinning of the lithospheric mantle and magmatic and erosional thinning of the crust dominated, thus providing the principal driving mechanism for the Late Permian and Mesozoic subsidence of intracratonic basins. Spatially and temporally variable Mesozoic short-term subsidence accelerations, which are superimposed on the long-term thermal subsidence trend of the wider ECRIS area, reflect subtle reactivation of pre-existing crustal discontinuities by far-field stresses that are related to rift and wrench activity in the North Sea, on the Atlantic shelves and in the Tethys domain. These subsidence anomalies were not associated with significant destabilization of the lithosphere-asthenosphere system. However, the lithosphere-asthenosphere system of the ECRIS area became destabilized again at the transition from the Cretaceous to the Palaeocene in conjunction with a phase of major intraplate compression that was accompanied by the impingement of mantle plumes. With the late Eocene activation of the ECRIS, crustal extension, and particularly Neogene increased plume activity, caused further destabilization of its lithosphereasthenosphere system (Drzes et al. 2004). As a result, the present thickness of the lithosphere decreases from 100-120 km in areas flanking the ECRIS to 6 0 - 7 0 km beneath some parts of the Rhenish Massif and the Massif Central (Babuska & Plomerova 1993). Fig. 9. Conceptual model for the Late Carboniferous to end-Cretaceous evolution of the Variscan lithosphere in the ECRIS area along a transect that extends from the Lower Rhine Graben in the north to the Jura Mountains in the south (not to scale). (For discussion see text.)

wrench deformation of the Variscan orogen controlled the development of a multidirectional array of pull-apart and trapdoor basins, the detachment of subducted lithospheric slabs, upwelling of the asthenosphere and widespread mantle-derived magmatic activity that abated towards the end of the Early Permian. Partial delamination and thermal thinning of the lithospheric mantle, magmatic inflation of the remnant lithosphere and interaction of mantle-derived partial melts with the lower crust accounted for the destruction of the crustal and lithospheric roots of the Variscan orogen and its regional uplift. By the end of the Early Permian, the crust was regionally thinned to 2 7 - 3 5 km, mainly by magmatic processes and erosional unroofing and only locally by mechanical stretching (Fig. 9b). After the Permo-Carboniferous thermal surge, which probably reflects a reorganization of the asthenospheric flow patterns and perhaps minor plume activity or secondary convection, the temperature of the asthenosphere apparently returned rapidly to ambient levels. Re-equilibration of the lithosphere with the asthenosphere, commencing during the late Early Permian and persisting during the Mesozoic, accounted for the long-term thermal subsidence of the crust and the development of a system of intracratonic sedimentary basins that covered much of the ECRIS area (Fig. 9c). Modelling of observed subsidence curves, based on the comparison with a theoretical thermal decay curve, indicates that: (1) by end-Cretaceous times the lithosphere had equilibrated with the asthenosphere at depths of 100-120 km; (2) by end-Early

This paper is a contribution by the EUCOR-URGENT Project (Upper Rhine Graben: Evolution and Neotectonics) to EUROPROBE. M.E.S. and P.D. acknowledge financial support by a University of Basel ELTEM grant. Critical and constructive comments by W. Franke, A. Henk, M. Wilson and I. Artemieva on earlier versions of this paper are gratefully acknowledged.

References

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EVOLUTION OF VARISCAN LITHOSPHERE

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Thermo-mechanical controls on Alpine deformation of NW Europe S. CLOETINGH 1, P. A. ZIEGLER 2, F. B E E K M A N t, P. A. M. ANDRIESSEN 1, N. HARDEBOL 1, J. VAN WIJK 1'3 & P. DISZES 2

1Netherlands Research Centre for Integrated Solid Earth Science, Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 H V Amsterdam, Netherlands (e-mail: [email protected]) 2Geological-Palaeontological Institute, Department of Geosciences, University of Basel, Bernoullistrasse 32, 4056 Basel, Switzerland 3Institute of Geophysics and Planetary Physics, Scripps Institution of Oceanography, University of California San Diego, La Jolla, CA 92093-0225, USA

Abstract: The lithosphere of the Northern Alpine foreland has undergone a polyphase evolution with an intense interplay between upper mantle thermal perturbations and stress-inducedintraplate deformationthat points to the importance of lithosphericfolding of the thermally weakened lithosphere. In this paper we address relationshipsbetween deeper lithosphericprocesses, neotectonics and surface processes in the Northern Alpine foreland with special emphasis on tectonicallyinduced topography. We focus on lithosphere memory and neotectonics with special attention to the thermo-mechanical structure of the lithosphere, mechanisms of large-scale intraplate deformation, Late Neogene anomaliesin subsidenceand uplift, and links with surface processes and topography evolution.

The lithosphere of the Alpine foreland system has undergone repeated tectonic reactivation (Ziegler 1989a; Ziegler et al. 1995, 1998), expressed by significant intraplate seismicity, differential vertical motions and the development of dynamic topography at large distances from plate boundaries (Fig. 1). Increasing evidence is accumulating for widespread Neogene uplift and tectonics around the North Atlantic (e.g. Chalmers & Cloetingh 2000; Japsen & Chalmers 2000), occurring simultaneously with accelerations in subsidence and sedimentation (Cloetingh et al. 1990), and commencing too early to be attributed solely to the effects of climate change and post-glacial rebound of the North Atlantic borderlands. Results of geothermochronological studies have provided constraints on the timing and magnitude of uplift of the North Atlantic borderlands, demonstrating a polyphase record that commences as early as Oligocene time, and that is interpreted as the combined result of upper mantle perturbations, intraplate stresses and post-glacial rebound (Hendriks & Andriessen 2002). Over the last few years, seismicity studies and geomorphological evidence from Brittany (Bonnet et al. 2000), Normandy and the Channel and Dover Strait areas (Lagarde et al. 2000; Van Vliet-LanoE et al. 2000) and southern England (Preece et al. 1990), and work on the Ardennes-Eifel region (Demoulin et al. 1995; Meyer & Stets 1998; Van Balen et al. 2000), the Upper Rhine Graben (Nivi~re & Winter 2000) and the North German Basin (Ludwig 1995) demonstrate the important contribution of neotectonics to the topographic evolution of intraplate Europe (Fig. 1). Regional Neogene exhumation of the British Isles and the margins of the North Sea Basin (Japsen 1997) is documented by analyses of petroleum industry data. Combined studies of geomorphology and the record of sedimentary basins, integrating results of subsidence analyses and geothermal chronology (e.g. Ter Voorde et al. 2004), have also demonstrated that neotectonics has a strong bearing on topography and drainage patterns. The nature of intraplate stress fields in continental lithosphere and its relationship to plate tectonic driving forces has been the subject of a large number of observational (e.g. Van der Pluim et al. 1997; Marotta et al. 2001) and modelling studies (e.g. G61ke & Coblentz 1996; Bada et al. 1998). These studies have revealed the existence of consistently oriented first-order patterns of intraplate stress in, for example, the NW European platform (Fig. 2) and the North American craton. The effect of these stresses on vertical motions of the lithosphere, expressed in terms of, for example, apparent sea-level fluctuations (Cloetingh et al. 1985),

foreland bulges (Ziegler et al. 2002), basin inversion (Ziegler et al. 1995, 1998) and lithosphere folding (Martinod & Davy 1994; Cloetingh et al. 1999), has been demonstrated to be an important element in the dynamics of intraplate continental interiors (Cloetingh 1988; Van der Pluim et al. 1997). Stress propagation occurs in a lithosphere that can be significantly weakened by inherited structural discontinuities, but also by thermal perturbations in the upper mantle (e.g. Goes et al. 2000a,b). Below we present thermo-mechanical models for large-scale intraplate deformation, and we discuss constraints on these models inferred from studies carried out during the last few years on the rheology of the NW European foreland.

Structure of Europe's intraplate lithosphere European Cenozoic rift system Development of the still active European Cenozoic Rift System (ECRIS), which extends from the Dutch North Sea coast to the western Mediterranean, commenced during the late Eocene. Its southern elements are the Valencia Trough, the graben systems of the Gulf of Lions, and the northerly striking Valence, Limagne and Bresse grabens; the latter two are superimposed on the Massif Central and its eastern flank, respectively. These grabens are linked via the Burgundy transfer zone to the northerly striking Upper Rhine Graben, which bifurcates northwards into the NW-trending Roer Valley Graben and the NE-trending Hessian grabens, which transect the Rhenish Massif. The NE-striking Eger Graben, which transects the Bohemian Massif, forms an integral part of the ECRIS (Ziegler 1994). Localization of the ECRIS involved the reactivation of Permo-Carboniferous shear systems. Although the onshore parts of the ECRIS are characterized by relatively low crustal stretching factors, they are associated with a distinct uplift of the crust-mantle boundary. To what extent this feature must be attributed to Cenozoic rifting or whether PermoCarboniferous processes have contributed to it remains an open question (D~zes et al. 2004). Evolution of the ECRIS was accompanied by the development of major volcanic centres in Iberia, the Massif Central, the Rhenish Massif and the Bohemian Massif, particularly during Miocene and Plio-Pleistocene times. Seismic tomography indicates that mantle plumes well up beneath the Massif Central and the Rhenish Massif (Granet et al. 1995; Ritter et al. 2001; Ritter 2006), but not beneath the

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 113-127. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Topographic map of Europe with the distributionof seismicitysuperimposed(red dots), illustratingthe present-day active intraplate deformation.Also shown are intraplate areas of Late Neogene uplift (circles with plus symbols) and subsidence(circles with minus symbols).Background elevation image is extracted from the ETOPO2 dataset. Earthquake epicentres are obtained from the NEIC Data Center. Boxed areas indicate regions for which the relationshipbetween lithospheric therrno-mechanicalstructure and vertical motions is discussed in more detail in this paper. Vosges-Black Forest arch. Similar data are, however, not available for Iberia and the Bohemian Massif. Despite this, the evolution of the ECRIS is considered to be a clear case of passive rifting (Drzes et al. 2004). During the Late Eocene, the Limagne, Valence, Bresse, Upper Rhine and Hessian grabens began to subside in response to northerly directed compressional stresses that may be related to collisional interaction of the Pyrenees and the Alps with their foreland (Merle & Michon 2001; Schumacher 2002). During their Oligocene main extensional phase, these originally separated rifted basins coalesced, and the Roer Valley and Eger grabens came into being. During the Late Oligocene, rifting propagated southward across the Pyrenean orogen into the Gulf of Lions and along coastal Spain in response to back-arc extension that was controlled by eastward roll-back of the subducted Betic-B alearic slab. By Late Burdigalian times, crustal separation was achieved in the Gulf of Lions, the oceanic Provenqal Basin began to open and the grabens of southern France became inactive (Roca 2001). In contrast, the intra-continental parts of the ECRIS remained tectonically active until the present, although their subsidence was repeatedly interrupted, possibly in conjunction with stresses controlling far-field inversion tectonics. By end-Oligocene times, the area of the triple junction between the Upper Rhine, Roer Valley and Hessian grabens became uplifted, and magmatic activity on the Rhenish Shield increased. By mid-, Late Miocene times, the Massif Central, the Vosges-Black Forest arch and, slightly later, the Bohemian Massif were uplifted. This was accompanied by increased mantle-derived volcanic activity. Because at the level of the Moho a broad anticlinal feature extends from the Massif Central via the Burgundy transfer zone and the Vosges-Black

Forest into the Bohemian Massif, uplift of these arches may have involved folding of the lithosphere in response to increased collisional coupling of the Alpine orogen with its foreland. Uplift of the Burgundy transfer zone entailed partial erosional isolation of the Paris Basin. Under the present NW-directed stress regime (Fig. 2), which started during the Miocene and intensified during the Pliocene, the Upper Rhine Graben is subjected to sinistral shear, the Roer Valley Graben is under active extension (Drzes et al. 2004), whereas the North Sea Basin experiences a late phase of accelerated subsidence that can be related to stress-induced deflection of the lithosphere (Van Wees & Cloetingh 1996). Similarly, the continued uplift of Fennoscandia is thought to be controlled by folding of the lithosphere under the present stress field that reflects a combination of Atlantic ridge-push and collisional coupling of the Alpine orogen with its foreland (Grlke & Coblentz 1996).

C o n s t r a i n t s on c r u s t a l a n d u p p e r m a n t l e s t r u c t u r e

Ziegler & D~zes (2006) have compiled the results of crustal studies that were carried out since the publication of Moho depth maps by Meissner et al. (1987), Ziegler (1990) and Ansorge et al. (1992) to obtain a better understanding of the present-day crustal configuration of Western and Central Europe, and to analyse processes and their timing that controlled the evolution of the crust in the various parts of Europe. Spectacular improvements have been made in global travel time tomography. A new model parameterization technique and new 3D ray tracing algorithms (Bijwaard & Spakman 1999a)

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Fig. 2. Intraplate stress map for Western and Central Europe, displaying the present-day orientation of the maximum horizontal stress (SHmax). Symbols stand for stress indicators and the length of the lines represents the data quality, 'A' being of highest quality. Background shading indicates topographic elevation (darker is higher). Stress map is extracted from the World Stress Map database (http://www.world-stress-map.org/). NF, normal faulting; SS, strike-slip; TF, thrust faulting; U, unknown tectonic domain. resulted in global mantle models that, for the first time, exhibit regional-scale (60-100 km) detail (Bijwaard & Spakman 2000). Improved focusing on lower mantle structures led to the first evidence for a whole mantle plume below Iceland (Bijwaard & Spakman 1999b) and for upwelling of the lower mantle beneath Europe, which is proposed as underlying the Cenozoic Central European volcanism (Goes et al. 1999). Goes et al. (2000a) developed a new inversion strategy that takes advantage of the fact that most of the variation in seismic wave velocities in the upper mantle is caused by variations in temperature. State-of-the-art seismological models and experimental data on the physical properties of mantle rocks are inverted for upper mantle temperatures (e.g. beneath Europe). Inferred mantle temperatures beneath Europe agree reasonably well with independent estimates from heat flow and general geological considerations (e.g. Dunai & Baur 1995). Global seismic tomographic studies image beneath the ECRIS the presence of a system of upper asthenospheric low-velocity anomalies, interpreted as plume heads that have spread out above the 410km discontinuity (Goes et al. 1999; Spakman 2004). Local tomographic studies (Granet et al. 1995) show that beneath the Massif Central a secondary mantle plume with a diameter of 100-300km, involving material 100-200 ~ hotter than the ambient mantle, rises from such a deep-seated asthenospheric anomaly. These findings, and similar results for the Eifel region (Ritter et al. 2001; Ritter 2006), support diapiric upwelling of small-scale, finger-like convective instabilities from the base of the upper asthenosphere, which presumably act as the main source

for Tertiary-Quaternary volcanism of Western and Central Europe. This volcanism is spatially and temporally linked to the development of the ECRIS and to domal uplift of Variscan basement massifs (Wilson & Patterson 2001). However, the absence of an upper mantle plume beneath the Vosges-Black Forest arch (Achauer & Masson 2002) suggests that the Upper Rhine Graben developed as a passive rift under a complex regional stress field with inherited structures playing a controlling role in its localization (D~zes et al. 2004; Ziegler & D~zes 2006).

Strength and deformation mode of Europe's intraplate lithosphere Strength o f the lithosphere

The strength of continental lithosphere is controlled by its depthdependent rheological structure (Fig. 3), in which the thickness and composition of the crust, the thickness of the mantlelithosphere, the potential temperature of the asthenosphere, the presence or absence of fluids, and strain rates play a dominant role (e.g. Carter & Tsenn 1987; Kirby & Kronenberg 1987). In contrast, the strength of oceanic lithosphere depends on its thermal regime, which controls its essentially age-dependent thickness (Panza et al. 1980; Kusznir & Park 1987; Stephenson & Cloetingh 1991; Cloetingh & Burov 1996). Theoretical rheological models indicate that thermally stabilized continental lithosphere consists of the mechanically strong

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Fig. 3. From crustal thickness (top left) and thermal structure (top right) to lithosphericstrength (bottom); conceptual make-up of the thermal structure and composition of the lithosphere, adopted for the calculationof 3D strength models. upper crust, which is separated by a weak lower crustal layer from the strong upper part of the mantle-lithosphere, which in turn overlies the weak lower mantle-lithosphere. In contrast, oceanic lithosphere has a more homogeneous composition and is characterized by a much simpler rheological structure. Rheologically speaking, thermally stabilized oceanic lithosphere is considerably stronger than all types of continental lithosphere. Atlantic-type continental margins mark the transition from oceanic to continental lithosphere, and are the sites of thinned continental lithosphere that was extended and heated during continental break-up. This has caused a substantial lateral variation in the mechanical strength of such margins (e.g. Fig. 3, lower panel) that is controlled by complex variations in crustal thickness, composition of the lithospheric layers, and the thermal regime. The strength of continental crust depends largely on its composition, thermal regime and the presence of fluids, and also on the availability of pre-existing crustal discontinuities. Deep-reaching crustal discontinuities, such as thrust- and wrench-faults, cause significant weakening of the otherwise mechanically strong upper parts of the crust. As such discontinuities are apparently characterized by a reduced frictional angle, particularly in the presence of fluids, they are prone to reactivation at stress levels that are well below those required for the development of new faults. The strength of the continental upper mantle-lithosphere depends to a large extent on the thickness of the crust but also on its age and thermal regime. Extension of stabilized continental crustal segments precludes ductile flow of the lower crust and faults will be steep to listric and propagate towards the hanging wall, that is, towards the basin centre (Bertotti et al. 2000). Under these conditions, the lower crust will deform by distributing ductile shear in the brittle-ductile transition domain. This is compatible with the occurrence of earthquakes within the lower crust and even close to the Moho (e.g. Upper Rhine Graben: Bonjer

1997; East African rifts: Shudofsky et al. 1987). In young orogenic belts, which are characterized by a crustal thickness of up to 60 km and an elevated heat flow, the mechanically strong part of the crust is thin and the mantle-lithosphere is also weak. Extension of this type of lithosphere can involve ductile flow of the lower and middle crust along pressure gradients away from areas lacking upper crustal extension towards zone of major upper crustal extensional unroofing, involving the development of core complexes (Bertotti et al. 2000). Generally, the upper mantle of thermally stabilized, old cratonic lithosphere is considerably stronger than the strong part of its upper crust (e.g. Moisio et al. 2000). However, the occurrence of upper mantle reflectors, which generally dip in the same direction as the crustal fabric and are probably related to subducted oceanic and/or continental crustal material, suggests that the continental mantle-lithosphere is not necessarily homogeneous but can contain lithological discontinuities that enhance its mechanical anisotropy (Ziegler et al. 1995, 1998; Vauchez et al. 1998). Such discontinuities, consisting of eclogitized crustal material, can potentially weaken the strong upper part of the mantlelithosphere. These factors contribute to weakening of former mobile zones to the end that they present rheologically weak zones within a craton, as evidenced by their preferential reactivation during the break-up of Pangaea (Ziegler 1989b; Janssen et al. 1995; Ziegler et al. 2001). From a rheological point of view, the thermally destabilized lithosphere of tectonically active rifts, as well as of rifts and passive margins that have undergone only a relatively short postrift evolution (e.g. 25 Ma), is considerably weaker than that of thermally stabilized rifts and of unstretched lithosphere (Ziegler et al. 1998; Ziegler & Cloetingh 2004). In this respect, it must be realized that during rifting, progressive mechanical and thermal thinning of the mantle-lithosphere and its replacement

ALPINE DEFORMATIONOF NW EUROPE by the upwelling asthenosphere is accompanied by a rise in geotherms causing progressive weakening of the extended lithosphere. In addition, its permeation by fluids causes its further weakening. Upon decay of the rift-induced thermal anomaly, rift zones are rheologically considerably stronger than unstretched lithosphere. However, thermal blanketing through the accumulation of thick syn- and post-rift sedimentary sequences can cause a weakening of the strong parts of the crust and mantlelithosphere of rifted basins (Stephenson 1989). Moreover, as faults permanently weaken the crust of rifted basins, they are prone to tensional as well as compressional reactivation (Ziegler et al. 1995, 1998, 2001, 2002; Ziegler & Cloetingh 2004). In view of its rheological structure, the continental lithosphere can be regarded under certain conditions as a two-layered viscoelastic beam (Reston 1990; Ter Voorde et al. 1998). The response of such a system to the build-up of extensional and compressional stresses depends on the thickness, strength and spacing of the two competent layers, on stress magnitudes and strain rates, and on the thermal regime (Zeyen et al. 1997). As the structure of continental lithosphere is also areally heterogeneous, its weakest parts start to yield first once intraplate stress levels equate their strength (e.g. Brun 2002). The presence of crustal and mantle-lithospheric discontinuities can significantly reduce the strength of the lithosphere. In this, the orientation of such discontinuities with respect to the prevailing stress field plays an important role in terms of their reactivation potential (Ziegler et al. 1995; Brun & Nalpas 1996). On the other hand, oceanic lithosphere behaves as a single-layer beam that is thinner than the competent parts of thick cratonic continental lithosphere. However, in view of the high strength of mature oceanic lithosphere, its deformation requires considerably higher stress levels than the deformation of continental lithosphere (Cloetingh et al. 1989). This suggests that tectonic stresses transmitted through mature oceanic lithosphere can be large enough to cause failure of the continental lithosphere forming part of the same plate, without at the same time causing deformation of the oceanic lithosphere (Ziegler et al. 1998). Nevertheless, evidence for tensional reactivation of rifts that had been abandoned millions of years ago suggests that crustal-scale faults permanently weaken the lithosphere to the degree that rifts are prone to tensional and compressional reactivation (Ziegler et al. 1995, 2001, 2002). At the scale of the lithosphere, the strength of the mechanically strong upper part of the mantle-lithosphere, which depends on its thermal state and the thickness of the crust, plays an important role in the localization of rift zones (Ziegler & Cloetingh 2004). Moreover, lateral thickness heterogeneities of the lithosphere appear to play an important role in the localization of rifts (e.g. Oslo Graben: Pascal et al. 2002). At the scale of the crust, its composition, the thickness of its mechanically strong upper part and the availability of pre-existing crustal discontinuities, which can be tensional reactivated, play a dominant role in the width and deformation mode of an evolving rift. Strength profiles and effective elastic thicknesses have been calculated over the last few years for a number of locations in Europe (e.g. Cloetingh & Burov 1996). Most of these strength profiles and estimates of integrated strength were calculated along available deep seismic crustal cross-sections, for example, the European Geotraverse (Cloetingh & Banda 1992) and the TransAlp deep seismic profile (Willingshofer & Cloetingh 2003). So far, lithospheric strength maps have been calculated for restricted areas of Europe only, including the Pannonian Basin-Carpathian region (Lankreijer et al. 1999) and the Baltic Shield (Moisio et aL 2000), but are not available on a regional scale for intraplate Europe. Drawing on the newly compiled Moho map of Europe of Ziegler & Dtzes (2006) and on constraints on the thermal lithospheric structure from heat-flow studies and upper mantle seismic tomography, as well as estimates of the lithospheric thickness from seismological studies (Plomerova et al. 2002), we constructed a 3D strength map for the lithosphere of a large part of Europe.

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The underlying strength model is based on a 3D multilayer compositional model involving one upper mantle layer, two to three crustal layers and a sedimentary cover layer (Fig. 3) (Hardebol et al. 2003). The temperature structure of the lithosphere below Europe inferred from seismic tomography (Goes et al. 2000a,b) has only limited resolution in the mechanically strong part of the lithosphere. Therefore, in this study the temperatures in the lithosphere were calculated analytically, using Fourier's law for heat conduction. Thermal rock properties were taken from Cloetingh & Burov (1996), whereas the thermal boundary conditions were extracted from Babuska & Plomerova (1992, 1993) and Plomerova et al. (2002), or, where available, from higher quality regional or local studies. A comparison of the calculated thermal cube with temperature structures inferred from seismic tomography studies of the upper mantle below Europe (Goes et al. 2000a,b) shows a first-order overall agreement at depths corresponding to the lithosphere- asthenosphere boundary. Figure 4a shows the integrated compressional strength of the entire lithosphere of Western and Central Europe, whereas Figure 4b and c displays the integrated strength of the mantle and crustal parts of the lithosphere, respectively. As evident from Figure 4a, Europe's lithosphere is characterized by major spatial mechanical strength variations, with a pronounced contrast between the strong lithosphere of the East European Platform east of the Teisseyre-Tornquist line and the relatively weak lithosphere of Western Europe. A similar strength contrast occurs at the transition from strong oceanic lithosphere to the relatively weak continental lithosphere of Western Europe. Within the Alpine foreland, pronounced NW-SE-trending weak zones are recognized that coincide with major geological structures, such as the Rhine Rift System and the North Danish-Polish Trough, that are separated by the high-strength North German Basin. Moreover, a broad zone of weak lithosphere characterizes the Massif Central and surrounding areas. The presence of thickened crust in the area of the TeisseyreTornquist suture zone gives rise to a pronounced mechanical weakening of the lithosphere, particularly of its mantle part. Whereas the lithosphere of Fennoscandia is characterized by relatively high strengths, the North Sea rift system corresponds to a zone of weakened lithosphere. Other areas of pronounced lithospheric strength are the Bohemian Massif and the LondonBrabant Massif, both of which exhibit low seismicity (Fig. 1). A pronounced contrast in strength can also be noticed between the strong Adriatic indenter and the weak Pannonian Basin area. Comparing Figure 4a, b and c reveals that the lateral strength variations of Europe's intraplate lithosphere are primarily caused by variations in the mechanical strength of the mantle-lithosphere, whereas variations in crustal strength appear to be much more modest. The variations in mantle-lithospheric strength are primarily related to variations in the thermal structure of the lithosphere, reflecting thermal upper mantle perturbations imaged by seismic tomography, with lateral changes in crustal thickness playing a secondary role, apart from Alpine domains that are characterized by deep crustal roots. For instance, the strong lithosphere of the East European Platform, the Bohemian Massif, the LondonBrabant Massif and the Fennoscandian Shield can be explained by the presence of old, cold lithosphere, whereas the ECRIS corresponds to a major axis of weakened lithosphere in the NW European Platform. Weakening of the lithosphere of southern France can be attributed to the presence of tomographically imaged plumes rising under the Massif Central (Granet et al. 1995; Wilson & Patterson 2001).

L i t h o s p h e r i c f o l d i n g : important m o d e o f d e f o r m a t i o n

Folding of the lithosphere appears to play a more important role in the large-scale neotectonic deformation of Europe's intraplate

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Fig. 5. Schematic diagram of lithospheric or crustal folding and decoupling, and the consequences for vertical motions and erosion and sedimentation at the Earth's surface, v is horizontal shortening velocity; upper crust, lower crust and mande layers are defined via corresponding rheologies and physical properties. A typical brittle-ductile strength profile (in black) for decoupled crust and upper mantle lithosphere adopting a quartz-diorite-olivine rheology is shown for reference (after Cloetingh et al. 1999).

Fig. 4. Integrated strength maps for intraplate Europe. Adopted composition for upper crust, lower crust and mantle is based on a wet quartzite, diorite and dry olivine composition, respectively. Rheological rock parameters are from Carter & Tsenn (1987). The adopted bulk strain-rate is 10-i6 s-1. Contours represent integrated strength in compression for (a) total lithosphere, (b) mantle, and (e) crust. The main structural features of Europe are superimposed on the maps (Ziegler 1988; D~zes et al. 2004).

domain than hitherto realized (Cloetingh et al. 1999). The large wavelength of vertical motions associated with lithospheric folding (Fig. 5) necessitates integration of available data from relatively large areas (Elfrink 2001), often going beyond the scope of regional structural and geophysical studies that target specific structural provinces. Recent studies of the German Basin have revealed the importance of its structural reactivation by lithospheric folding (Marotta et al. 2000). Similarly, folding of the Variscan lithosphere has recently been documented for Brittany (Bonnet et al. 2000), the adjacent Paris Basin (Lefort & Agarwal 1996) and the Vosges-Black Forest arch (Dbzes et al. 2004). Lithospheric folding is a very effective mechanism for the propagation of tectonic deformation from active plate boundaries far into intraplate domains (e.g. Stephenson & Cloetingh 1991; Burov et al. 1993; Ziegler et al. 1995, 1998, 2002; Burov & Molnar 1998). An interesting analogue at the scale of a microcontinent that was affected by a succession of collisional events is provided by Iberia (Cloetingh et al. 2002), a well-documented natural laboratory for quantifying the interplay of neotectonics and surface processes. An important factor in favour of a lithosphere-folding scenario for Iberia is the compatibility of the wavelength of observed deformations, the thermotectonic age of the lithosphere and the total amount of shortening with well-documented examples of continental lithospheric folding coming from other cratonic areas (Fig. 6). A prominent example of lithospheric folding occurs in the Western Gobi area of Central Asia, involving a lithosphere with a thermo-tectonic age of 400Ma. In this area, mantle and crustal wavelengths are 360 km and 50 km, respectively, with a shortening rate of c. 10 mm a-1 and a total amount of shortening of 200-250 km during 10-15 Ma (Burov et al. 1993; Burov & Molnar 1998). Another case of recently detected Quaternary folding of Variscan lithosphere is the Armorican Massif of Brittany at the western margin of the Paris Basin (Bonnet et al. 2000). The wavelength of the folds is 250 kin, pointing to a lithospheric mantle control of the deformation. As pointed out by Bonnet et al. (2000), the spatial pattern and the timing of the uplift inferred from river incision studies of Brittany is incompatible with a glacio-eustatic origin. Bonnet et al. linked the observed patterns of vertical motions in N W France to the NW-SE-directed principal compressional axis of the present-day intraplate stress field of NW Europe. The stress-induced uplift pattern appears to control the amount of fluvial incision in the area, as well as the location of the main

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high-quality tomographic data (Goes e t al. 2000b) permitted imaging of the thermal structure of the sublithospheric mantle beneath NW Europe, revealing mantle plumes upwelling beneath the Massif Central (Granet et al. 1995) and the Rhenish Massif (Ritter et al. 2001). In this context, it is noteworthy that studies on the mechanical properties of Europe's lithosphere reveal a direct link between its thermo-tectonic age and bulk strength, whereas inferences from P- and S-wave tomography and thermomechanical modelling point to pronounced weakening of the lithosphere in the area of the Massif Central and Rhenish Massif owing to high upper mantle temperatures (Hardebol et al. 2003). Uplift of the Rhenish Massif by as much as 250 m during the last 0.8 Ma (Meyer & Stets 2002) can be directly attributed to the load of the impinging mantle plume and related thermal thinning of the lithosphere (Fig. 7) (Garcia-Castellanos et al. 2000; Drzes et al. 2004). Fig. 6. Comparison of observed (I) and modelled (9 wavelengths of folding in Iberia (Cloetingh et al. 2002) with theoretical predictions (Cloetingh et al. 1999) and other estimates (D) for wavelengths documented from geological and geophysical studies (Lambeck 1983; Stephenson & Cloetingh 1991; Nikishin et al. 1993; Ziegler et al. 1995; Bonnet et al. 2000). Wavelength is given as a function of thermo-tectonic age at the time of folding. Thermo-tectonic age corresponds to the time elapsed since the last major perturbation of the lithosphere prior to folding. It should be noted that neotectonic folding of Variscan lithosphere has recently also been documented for Brittany (Bonnet et al. 2000). Both Iberia and Central Asia are characterized by separate dominant wavelengths for crust and mantle folds, reflecting decoupled modes of lithosphere folding.

drainage divides. The area located at the western margin of the Paris Basin and along the rifted Atlantic margin of France has been subject to thermal rejuvenation during Mesozoic extension related to North Atlantic rifting (Robin et al. 2003; Ziegler & Drzes 2006) and subsequent compressional intraplate deformation (Ziegler et al. 1995), also affecting the Paris Basin (Lefort & Agarwal 1996). Levelling studies in this area (Lenotre et al. 1999) also point towards its continuing deformation. The inferred wavelengths of these neotectonic lithosphere folds are consistent with the general relationship that was established between the wavelength of lithospheric folds and the thermotectonic age of the lithosphere on the basis of an inventory of global examples of lithospheric folding (Fig. 6) (Cloetingh & Burov 1996). In a number of other areas of continental lithosphere folding, smaller-wavelength crustal folds have also been detected, for example in Central Asia (Burov et al. 1993; Cobbold et al. 1993; Nikishin et al. 1993).

Intraplate deformation in the Alpine foreland: role of plumes and neotectonic fault reactivation In conjunction with the World Stress Map project and the Task Force Origin of Sedimentary Basins, both sponsored by the International Lithosphere Programme (ILP), new databases were developed for the stress field of NW Europe and recent crustalscale vertical motions. The present-day stress field of NW Europe (Fig. 2) (Mtiller et al. 1997) could be successfully modelled by taking Alpine collisional coupling and Atlantic ridge-push forces into account (Grlke & Coblentz 1996; Goes et al. 2000a; Ziegler et al. 2002). These studies established a close link between the stress field, late Neogene to Quaternary intraplate deformation, earthquake distribution and topography (see also Fig. 1). There is increasing evidence that the European lithosphere responds to intraplate compressional stresses by lithospheric folding (Cloetingh et al. 1999), as evidenced for instance by the Plio-Pleistocene subsidence acceleration of the North Sea Basin and contemporaneous uplift of the Fennoscandian Shield (Van Wees & Cloetingh 1996). Furthermore, acquisition of

Fig. 7. (a) Elements of the 3D conceptual model adopted for a 2D numerical elastic-plastic plate model. Uplift is assumed to be a flexural response to a buoyant load acting in the base of the lithosphere. Both the amount of load and its extension are assumed to be related to the thermal anomaly at lithospheric depths. (b) Input temperature distribution. (c) Calculated uplift. (d) Stress distribution (lower panel) predicted in the model assuming a low compression regional setting (Fx = 1 TN m l). Positive stresses (in blue) mean extension. The black lines show the yield strength profiles (YSE) representative for both the unperturbed and the anomalous zones. It should be noted that the mantle has nearly no strength in the central area as a result of thermal weakening owing to mantle upwelling. The dashed line is the predicted elastic thickness. UC, upper crust; LC, lower crust; LM, lithospheric mantle. Bold lines indicate the base of the upper and lower crust.

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the upper crust can be attributed to collision-related stresses that are transmitted from the Alps above an incipient mid-crustal detachment level. By contrast, lower crustal extension may be related to folding of the mantle-lithosphere, controlling uplift of the Vosges-Black Forest arch, in response to collision-related stresses transmitted from the Alps through the mechanically strong parts of the mantle-lithosphere. The effects of Cenozoic rifting are still present in the lower crust, as is visualized by relative P-wave velocity imaging of the lower crust (Lopes Cardozo & Granet 2003; Lopes Cardozo et al. 2005). Also important in this context is the observation that earthquakes occur almost down to the Moho but are absent below it (Plenefisch & Bonjer 1997; Deichmann e t al. 2000). Moderate Pliocene and Quaternary extension across the Bresse and Upper Rhine grabens was presumably accompanied by sinistral movements along the seismically still active Burgundy transfer zone, which links them. Geodetic data indicate continuing, shortening rates of 3 - 4 m m a - 1 for the French Jura Mountains (Jouanne et al. 1998) and horizontal displacement rates across the Upper Rhine Graben of 0.8 mm a -1 (Rozsa et al. 2005). From about 4 Ma onward, compressional deformation of the Jura Mountains was no longer exclusively thin skinned, but also involved the basement, as indicated by intra-crustal earthquakes (Roure et al. 1994; Becker 2000).

North Atlantic continental margin: rheology and stress controls on basin configuration and vertical motions

Fig. 8. Topographyof the Netherlandsand surroundingsin a colour-coded relief map (data from GTOPO30). Red lines are reactivated Base Tertiary faults in the subsurface,based on data from NITG-TNO (Geluk et al. 1995; De Mulder et al. 2003). Also shown is the total seismicity(tectonic and man-induced) (data from ORFEUS data centre; see also Dirkzwageret al. 2000).

Research on the role played by inherited crustal weakness zones in the distribution of seismic activity in the Netherlands (Fig. 8) involved quantification of (1) the role of lithospheric rheology on basin reactivation and (2) the geomechanical control of border faults on basin reactivation, contributing to neotectonic deformation of the Roer Valley Graben (Dirkzwager et al. 2000). Integrated analysis of crustal-scale cross-sections and their comparison with isopach maps of Tertiary and Quaternary sequences demonstrated the strong control of Mesozoic faults on recent differential vertical motions in the Roer Valley Graben and coastal areas. Central in this context were the development of a revised Moho map and the construction of compaction trends for the Netherlands (Dirkzwager et al. 2000). Results of gravity 3D backstripping, carried out to isolate crustal and upper mantle sources, demonstrate a clear correlation between the main residual gravity features, characterized by a distinct positive anomaly, and the main structural trends of the Lower Rhine-Roer Valley Graben system. In the southern parts of the Upper Rhine Graben, PlioQuaternary tectonic activity is indicated by folding of the Pliocene Sundgau gravels along the Jura Mountains thrust front (Giamboni et al. 2004), by faults extending through Quaternary deposits of the graben fill, and by the seismicity of the area. Earthquake focal mechanisms indicate that deformation of the upper crust is controlled by a strike-slip to compressional stress regime whereas the lower crust is subjected to extension (Plenefisch & Bonjer 1997; Deichmann et al. 2000). Transpressional deformation of

The Atlantic margin of Mid-Norway (Fig. 9) is one of the best documented continental margins of the world owing to the availability of a wealth of high-quality industry data and intense research collaboration between industrial, governmental and academic institutions (Mosar 2003; Tom6 et al. 2003). For this reason the Mid-Norway margin serves as a natural laboratory for studying processes controlling the development of a passive margin, the crustal separation phase of which was accompanied by extensive magmatic activity.

Fig. 9. Sketch of the three rift zones that characterize the mid-Norwegian margin (after Van Wijk & Cloetingh 2002). COB, continent-ocean boundary.

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Fig. 10. Thermal evolution of the lithospherefor a migratingrift case (Vext= 6 mm a- 1). Times are in millionyears afterrifting started (after Van Wijk & Cloetingh 2002). Moreover, it permits us to analyse, for passive margins, enigmatic features such as its post-break-up partial inversion and erosional truncation in nearshore areas in conjunction with uplift of the Norwegian Caledonides, the controlling mechanisms of which have long been a matter of debate. Previous work on the North Atlantic margins has led to the development of a new generation of models for controls on continental break-up and the subsequent evolution of ocean-continent boundary zones. Van Wijk et al. (2001) developed dynamic models for the quantification of melting processes in volcanic rifted margins, as well as for the lateral migration of rifting activity (see also Van Wijk & Cloetingh 2002) and associated vertical motions. When extension is characterized by low strain rates, as typical for the rifting of the Norwegian margin, the thermal evolution of the extending lithosphere shows the development of a cold spot in the area that previously underwent extension (Fig. 10). As a result, temperatures in the lithosphere below the first-stage basin become lower than temperatures in the surrounding lithosphere. Whereas initially the first-stage basin is marked by the smallest values of lithospheric strength, in the later phases strength increases with time, ultimately resulting in a reversal in the spatial pattern of strength distribution, with the smallest values found on both sides of the central basin, which now has become the strongest part (Fig. l la). This thermal inversion process leads to distinct differential topography of the surface (Fig. 1 lb), and to temporal and spatial variations in tectonic subsidence and uplift (Fig. 1 lc). It appears that break-up processes, as manifest in the Norwegian margin, have set the stage for subsequent tectonic reactivation of the margin under a compressional stress regime and overprinting upper mantle thermal anomalies (Ziegler & Cloetingh 2004). Systematic quantification of surface uplift and erosion of marginal highlands by thermochronology and other techniques has revealed pronounced along-strike variations in the magnitude of their late-stage uplift and related geomorphological development. The Mid-Norway margin is a representative part of the Norwegian-Greenland Sea rift, along which crustal separation between NW Europe and Greenland was achieved in earliest Eocene times (Torsvik et al. 2001; Mosar et al. 2002). The Norwegian-Greenland Sea rift, which had remained intermittently

Fig. 11. (a) Temporal and lateral evolution of the integrated strength of the lithosphere for a migratingrift case (Vext= 6 mm a- 1). (b) Evolution of relative surface topography for this migratingrift case. Grey lines indicate the positions of the corresponding syntheticsubsidencecurves derived from this panel and shown in (c). (e) Synthetic subsidencecurves for three locations indicated in (b): in the first-stagebasin(rightpanel); outside this basin but in the new basin (left panel); and in the transition zone (middle panel) (after Van Wijk & Cloetingh 2002).

active for some 280 Ma from the Late Carboniferous until the end of the Palaeocene, is superimposed on the Arctic-North Atlantic Caledonides (Ziegler 1988; Ziegler & Cloetingh 2004). During the Devonian and Early Carboniferous, orogen-parallel extension controlled the collapse of the Caledonian orogen, the subsidence of pull-apart basins and uplift of core complexes (Braathen et al. 2002; Eide et al. 2002). During the subsequent tiffing stage, tensional reactivation of Caledonian and Devonian-Early Carboniferous crustal discontinuities played an important role in the structuring of the Mid-Norway margin (Mosar 2003). During the evolution of the Norwegian-Greenland Sea rift, initially a broad zone was affected by crustal extension. In time, rifting activity concentrated progressively on the zone of future crustal separation, with lateral elements becoming abandoned stepwise (Ziegler 1988, 1990; Mosar et al. 2002; Ziegler & Cloetingh 2004). On the Mid-Norway margin, synrift sediments attain thicknesses of up to 10 km with post-rift series reaching thicknesses of 2 - 3 km (Osmundsen et al. 2002). Whereas Late Carboniferous to early Late Cretaceous

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crustal extension was not accompanied by volcanic activity, volcanism started during the Campanian-Maastrichtian, peaked during the Palaeocene and terminated on the Mid-Norway margin upon earliest Eocene crustal separation, when it became centred on the newly developing system of sea-floor spreading axes. The Palaeocene development of the Thulean flood basalt province, which was centred on Iceland and had a radius of more than 1000 km, is attributed to the impingement of the Iceland plume on the Norwegian-Greenland Sea rift (Morton & Parson 1988; Ziegler 1988; Larsen et al. 1999; Skogseid et al. 2000). Tomographic data image beneath Iceland a mantle plume rising from near the core-mantle boundary (Bijwaard & Spakman 1999a). Correspondingly, the Palaeocene extrusion and intrusion of large volumes of basaltic rocks on the Mid-Norway margin is generally attributed to a plume-related temperature increase of the asthenosphere. However, the interplay between extension and magmatism during continental break-up is still debated. Following Early Eocene crustal separation, the Mid-Norway margin was partly inverted during the Late Eocene-Early Oligocene and in the Miocene in the prolongation of the Iceland and Jan Mayen fracture zones (Mosar et al. 2002; Dor~ & Lundin 1996). As theoretical models predict inversion of passive margins in response to the build-up of compressional ridge-push forces a few tens of million years after crustal separation, mechanisms controlling the development of inversion structures of the MidNorway margin remain speculative. Most of the shortening of the Mid-Norway margin was accommodated along pre-existing major fault zones (Gabrielsen et al. 1999; Pascal & Gabrielsen 2001). Compressional structures, such as long-wavelength arches and domes, strongly modified the architecture of the deep

Cretaceous basins and controlled the pattern of sedimentation during the Cenozoic (Bukovics & Ziegler 1985). From the Oligocene onward, the nearshore parts of the MidNorway margin were uplifted and deeply truncated (Holtedahl 1953; Dor~ & Jensen 1996). Mechanisms controlling the observed broad uplift of the inner shelf and the adjacent onshore areas, as evident all around Norway, remain enigmatic. However, the long wavelength of the uplifted area suggests the existence of mantle processes (Rohrman & Van der Beek 1996; Rohrman et al. 2002). Olesen et al. (2002) interpreted the long-wavelength component of the gravity field in terms of both Moho topography and large-scale intra-basement density contrasts. Comparing the Bouguer gravity field with the gravity responses from Airy roots at different depths for the northern Scandinavia mountains shows that the compensating masses are located at relatively shallow depths in the upper crust. Consequently, the gravity field with the northern Scandinavian mountains must originate from intracrustal low-density rocks in addition to Moho depth variations. These results are in contrast to the situation for southern Norway, where the mountains are probably supported by low-density rocks within the mantle. Southwestern Norway was uplifted by as much as 2 km during Neogene times (Rohrman et al. 1995), as clearly shown by increased Miocene sedimentation in the adjacent basins (Jordt et al. 1995). Uplift patterns and timing in the northern Scandes differ from those of southwestern Norway (Hendriks & Andriessen 2002). By Late Tertiary times, cold climatic conditions prevailed. Uplifted land areas adjacent to the coast were submitted to strong glacial erosion, which in turn enhanced uplift and increased sedimentation and acceleration of subsidence of the shelf basin (Fig. 12). Fission-track analyses along major onshore lineaments

Fig. 12. Tectonic subsidencecurves for the mid-Norwegianmargin. The simultaneouslyoccurring accelerationsin subsidence (coloured bars) in four sub-basinsand platforms of the mid-Norwegianmargin should be noted (a-d). The pre-Neogene accelerations reflect pronounced tectonic phases associated with the rifting and inversion of the margin, whereas the Late Neogene acceleration in subsidence(light yellow bar) coincideswith a major plate tectonic reorganizationin the northern Atlantic (after Reemst & Cloetingh 2000).

ALPINE DEFORMATIONOF NW EUROPE in southern Norway show that pre-existing structures, in particular major normal faults, ranging in age from the Permian to the present, played a significant role in controlling uplift patterns (Hendriks & Andriessen 2002). Under the currently prevailing NW-directed compressional stress field (see Fig. 2) the Mid-Norwegian margin and its adjacent highlands are seismically active (Grtinthal 1999), with some faults showing evidence for recent movement (M6rner 2004). Uplift of the South Norwegian highland continues whereas the North Sea Basin experiences a phase of accelerated subsidence that began during the Pliocene and that is attributed to stress-induced deflection of the lithosphere (Van Wees & Cloetingh 1996). Landscapes resulting from continental break-up are in some areas characterized by highly elevated margins and associated escarpments. Post-rift uplift patterns are a distinct feature of the North Atlantic margins (Holtedahl 1953; Japsen & Chalmers 2000; Lidmar-Bergst6m et al. 2000; Rohrman et al. 2002). Definition of the dominant mechanism causing these surface expressions at passive margins requires an understanding of the interaction between tectonics and landscape evolution and constraints on the timing and quantification of processes that control denudation and morphological development of such receding rift flanks. The North Atlantic margin of Norway experienced substantial vertical movements during the Cenozoic and Quaternary that were coupled with the subsidence of offshore basins (Fig. 12) and the emergence of their nearshore parts (Cloetingh et al. 1990; Reemst & Cloetingh 2000). Two phases of Tertiary uplift are now recognized to have affected the entire North Atlantic region. These are a Palaeogene phase that was associated with volcanism, related to the impingement of the Iceland mantle plume, and a Neogene phase characterized by substantial uplift of several regions in the North Atlantic domain that was associated with subsidence of adjacent basins. Although progress has been made in the timing and spatial resolution of this Neogene surface uplift, controlling mechanisms are still debated (see Olesen et al. 2002) and may include intraplate compression, mantle phase changes, magmatic underplating and small-scale asthenospheric convection (Rohrman et al. 2002).

Discussion and conclusions The thermo-mechanical structure of Europe's intraplate lithosphere is characterized by significant spatial variations in its mechanical strength. The Teisseyre-Tornquist suture marks the main strength contrast in intraplate Europe, with relatively strong lithosphere to the east and north and relatively weak lithosphere to the west and south. Within Western and Central Europe, further strength contrasts are evident between the strong Palaeozoic massifs, such as the London-Brabant Massif and the Bohemian Massif, and the weak lithosphere of the ECRIS (see also Cloetingh et al. 2005b) and the Pannonian Basin (see also Cloetingh et al. 2006). Predicted strength patterns presented in this paper are in agreement with the spatial distribution of effective elastic thickness values inferred from flexural studies in the Pannonian-Carpathian system (see Cloetingh et al. 2006). Another feature is the strong Adriatic indenter, which is surrounded by weak lithosphere of the Pannonian Basin to the north and the Tyrrhenian and Ligurian back-arc to the west. In general, these Neogene back-arc basins are characterized by an overall low-strength lithosphere, the spatial pattern of which points to a prime control by thermal perturbations in the underlying mantle. The same appears to apply also for the large lowstrength area of the Massif Central, which is associated with a mantle plume and recent volcanism (Granet et al. 1995). In contrast, other regional areas of weakness, such as the North Sea Basin, appear to be more localized and parallel main structural trends.

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These spatial differences may play an important role in the localization and expression of neotectonic deformation of Europe's lithosphere (see, e.g. Cloetingh et al. 2005a; Cloetingh & Van Wees 2005). Another important factor is the orientation and magnitude of the present-day stress field (Fig. 2). In general, back-arc basins, which are located close to the boundaries of the European plate and are surrounded by areas of high topography, such as the Pannonian Basin, the Alboran Sea and the Tyrrhenian Sea, are likely to be subjected to a higher stress level (see also Bada et al. 1998, 2001) than areas located in the far-field domain, such as the North Sea Basin. These features (low strength, high stresses) may explain why present-day compressional reactivation is more manifest (by, e.g. higher seismicity) in Europe's back-arc basins than in such areas as the North Sea Basin. In several areas of Europe the rheological structure of the lithosphere will deviate locally and/or regionally from our first-order 3D model, and consequently will affect the estimated strength to some degree. For instance, local variations in crustal composition and architecture (e.g. caused by faults offsetting parts of the crust) were not incorporated in our model, as these hardly affect the firstorder patterns of integrated strength. On the other hand, the orogenic zones of the Alps and Pyrenees, with a substantially thickened lithosphere and a complex crustal architecture, as well as areas close to plate boundaries, should not be included in any interpretation. Other second-order regional or local processes may influence strength estimates, such as the presence of water or serpentinite in the mantle-lithosphere, both of which reduce its strength (e.g. Bassi 1995; P6rez-Gussiny6 & Reston 2001), or the removal of melts, which strengthens it (Van Wijk & Cloetingh 2002). A depth-varying rheology, such as employed in this study, depends on several parameters of which the most important are the crustal thickness, composition and temperature. Spatial variations in these parameters across Europe are described by first-order models, as explained above, thus neglecting local, second-order scale deviations. The strength of the ductile and viscous layers in the lithosphere depends also on strain rates, for which we have adopted a value of 10 - 1 6 s - 1 that is characteristic for the long-term and first-order bulk deformation of intraplate Europe. However, as for the other rheological parameters, it is likely that local deformation mechanisms are better described by higher strain rates, such as, for instance, in the ECRIS areas. In these tectonically active areas, our strength predictions may underestimate the true strength of the lithosphere. Summing up, we want to emphasize that the 3D strength cube for the European intraplate lithosphere in its present state, as given in this paper, is based on first-order variations in the geometry, composition and temperature of the lithosphere. Thus, interpretations and conclusions reflect only first-order variations in (integrated) strength across Europe. Any interpretation of the estimated strength has to be first order and preferably qualitative, comparing different parts of intraplate Europe in terms of being either (much) stronger or weaker. Increased seismic activity is associated with the Upper Rhine and the Roer Valley grabens, the Armorican shear zone and the Massif Central, as well as with the Eger Graben. This cannot be directly attributed to the activity of mantle plumes impinging on the attenuated lithosphere of the Massif Central, the Rhenish Massif and the Bohemian Massif (Wilson & Patterson 2001), but rather to the reactivation of Cenozoic and older crustal-scale faults under the present compressional stress field of the Alpine foreland (Fig. 2). On a much broader scale, seismicity (Fig. 1; Griinthal et al. 1999) and stress indicator data (Mtiller et al. 1997; Tesauro et al. 2005, see also G61ke & Coblentz 1996) demonstrate that active compressional deformation also continues in the Alpine foreland outside the various segments of the ECRIS. Zones of concentrated seismic activity correspond to areas of crustal contrast between the Cenozoic rifts and their surrounding platform areas, as well as to areas of crustal contrast in the

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rifted northeastern Atlantic margins. In general, earthquakes are associated with pre-existing faults, such as those bounding the Bohemian Massif and transecting the Armorican Massif. Recent GPS data indicate that the highest deformation rates are associated with the ECRIS, with strain rates being of the order of 10-16 s-1 (Tesauro et al. 2005). The simultaneous occurrence of thrust-, normal- and strike-slip-faulting mechanisms in the Alpine foreland supports a stress distribution dominated by heterogeneous crustal structures, including weak zones therein (Handy & Brun 2004). Thermal thinning of the mantle-lithosphere, often associated with volcanism and doming, enhances lithospheric folding and appears to control the wavelength of folds. Substantial thermal weakening of the mantle is consistent with higher folding rates in the European foreland as compared with folding in Central Asia (Nikishin et al. 1993), which is marked by pronounced mantle strength (Cloetingh et al. 1999). The European lithosphere, characterized by large areas of relatively low strength, appears to be more affected by compression-induced large-scale folding than hitherto realized. Lithospheric folding has been documented for areas with weak lithosphere, such as the Pannonian Basin (Cloetingh et al. 2006), the ECRIS area (e.g. Drzes et al. 2004), the Paris Basin (Lefort & Agarwal 1996; Cloetingh et al. 1999), the North German Basin (Marotta et al. 2000), and Iberia (Cloetingh et al. 2002). Recent and present-day intraplate deformation of Europe's lithosphere is associated with distinct differential vertical motions of the Earth's surface occurring at different spatial scales (Cloetingh & Cornu 2005). These vertical motions have clear expressions in, for instance, geomorphology, drainage patterns, and sediment s o u r c e - s i n k relationships. Therefore, a better understanding of the thermo-mechanical structure and evolution of the lithosphere is required for an improved characterization of the intrinsic coupling between intra-lithosphere and surface processes. This research was funded through the European Union (grant ENTEC and EUROBASIN), the Netherlands Organization of Scientific Research (grant NEESDI), the EUCOR-URGENT program, and the Netherlands Research Centre for Integrated Solid Earth Science. Constructive discussions with F. Roure, J. Negendank and J. Mosar are gratefully appreciated.

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Lithosphere structure and tectonic evolution of the Alpine arc: new evidence from high-resolution teleseismic tomography E. K I S S L I N G 1, S. M. S C H M I D 2, R. L I P P I T S C H 3, J. A N S O R G E t & B. F U G E N S C H U H 2

1Institute of Geophysics, ETH Hoenggerberg, CH-8093 Ziirich, Switzerland (e-mail." kissling @tomo. ig.erdw, ethz. ch) 2Department of Earth Sciences, Bernoullistrasse 32, CH-4056 Basel, Switzerland 3Department of Geophysics, ZAMG-Central Institute for Meteorology and Geodynamics, Hohe Warte 38, A-1190, Wien, Austria

Abstract" Several continental and oceanic plates and/or terranes amalgamated during the formation of the tectonically complex Alpine

arc. Reliable knowledge of the present structure of the lithosphere-asthenosphere system throughout the Alpine arc from the Western through the Central to the Eastern Alps is crucial for understanding the evolution of this orogen and the current interaction of lithospheric blocks, and additionally, for assessing the amount and orientation of lithosphere subducted in the geological past. We have compiled results from earlier geophysical studies and reinterpretations of existing seismic and geological data for the Alpine crust and Moho. High-resolution teleseismic tomography was used to produce a detailed 3D seismic model of the lower lithosphere and asthenosphere. The combination of these techniques provides new images for the entire lithosphere-asthenosphere system, showing significant lateral variations to depths of 400 km. Over the years the crustal structure has been determined extensively by active seismic techniques (deep seismic sounding) with laterally variable coverage and resolution. For a closer view three international seismic campaigns, using mainly near-vertical reflection techniques in the Western, Central and Eastern Alps, were carried out to assess the crustal structure with the highest possible resolution. The synoptic reinterpretation of these data and an evaluation of existing interpretations have allowed us to construct four detailed deep crustal transects across the Alps along the ECORS-CROP, NFP-20/EGT and TRANSALP traverses. In addition, contour maps of the Moho for the wider Alpine region and of the top of the lower crust were compiled from existing seismic refraction, near-vertical and wide-angle reflection data. Substantial structural differences in the structure of the deep crust appear between the Western, Central and Eastern Alps: doubling of European lower crust in the west resulted from collision with the Ivrea body; indentation of lower Adriatic crust between European lower crust and Moho occurred in the Central Alps; and a narrow collision structure exists under the transitional area between the western and eastern subduction regime under the Tauern Window of the Eastern Alps, where the crustal structure resembles a large-scale flower structure. Most recently, high-resolution teleseismic tomography based on the a priori known 3D crustal structure and compilation of a high-quality teleseismic dataset was successfully developed and applied to derive reliable detailed images of the lower lithosphere. Along strike of the Alps a fast slab-like body is revealed which in the western part is subducted beneath the Adriatic microplate. In the Western Alps detachment of parts of the lower continental slab occurred, possibly induced by the Ivrea body, which acted as a buttress in the collision process of the European and Adriatic plates. The generally SE-directed subduction of the European continental lithosphere changes gradually from west to east to almost vertical under the westernmost part of the Eastern Alps (western Tauern Window and Giudicarie lineament). Unexpectedly, some 50 km further east the subducted continental lower lithosphere is now part of the Adriatic lithosphere and dips NE beneath the European plate. Our tomographic image documents clear bipolar slab geometries beneath the Alpine orogen. The depth extent of the subducted continental lithospheric slab agrees rather well with estimates of post-collisional crustal shortening for the Western and Central Alps. This kinematic control on amounts of lateral motion of the collision zone in the west also allows estimates of the subduction and collision process in the Eastern Alps. The new 3D lithospheric picture for the wider Alpine region to 400 km depth demonstrates the clear connection and interaction between the deep structure of the lithosphere-asthenosphere system and near-surface tectonic features as seen today. It provides new and unexpected evidence for the entire Alpine tectonic evolution, a process which obviously changes significantly from west to east.

T h e very successful E u r o p e a n Geotraverse (EGT) (Blundell et al. 1992) was based on the concept o f a continuous continental swath extending f r o m northern Scandinavia to Tunisia. It p r o v i d e d consistent present-day information regarding the lithospheric structure across each o f the tectonic provinces, ranging in age f r o m A r c h a e a n to recent. T h e present paper, however, is partly based on the rationale o f an equally important f o l l o w - u p project, n a m e l y E U R O P R O B E (Gee & Z e y e n 1996). This project elucidates tectonic phases and processes in various E u r o p e a n areas in time and space, one of these areas b e i n g the A l p i n e orogeny. In recent years data on the Alps w e r e collected as part of the E G T and E U R O P R O B E projects, during specific F r e n c h - I t a l i a n and Swiss crustal seismic reflection campaigns, E C O R S - C R O P (Roure et al. 1990) and N R P 2 0 (Pfiffner et al. 1997), respectively, as well as during earlier deep seismic sounding experiments. W e will d e m o n s t r a t e that this recent mosaic of structural information, w h i c h mostly pertains to the crust, eventually leads to a consistent 3D picture for the lithospheric-scale tectonic evolution of the Alps w h e n c o m b i n e d with the latest results f r o m n e w l y d e v e l o p e d m e t h o d s o f teleseismic tomography. Several continental and oceanic plates a n d / o r terranes a m a l g a m a t e d during the formation of the tectonically c o m p l e x present-day A l p i n e arc, w h i c h is characterized by very m a j o r

along-strike c h a n g e s in crustal structure f r o m the W e s t e r n ( F r e n c h - I t a l i a n ) to the Central ( S w i s s - I t a l i a n ) and Eastern (Aust r i a n - I t a l i a n ) Alps. T h e attribution o f the various tectonic units of the Alps to particular p a l a e o g e o g r a p h i c a l d o m a i n s (Fig. 1) is based on stratigraphical analysis and retro-deformation o f nappe stacking. T h e r e m n a n t s of the f o l l o w i n g m a j o r p a l a e o g e o g r a p h i c a l units, w h o s e present-day position is indicated in Figure 1, are f o u n d at present in the Alps (after F r o i t z h e i m et al. 1996; S c h m i d 2000; see S c h m i d et al. 2004a,b for a revised overall architecture o f the A l p i n e orogen): (1) T h e E u r o p e a n margin: external massifs and their cover, Helvetic cover nappes and their p r e s u m e d b a s e m e n t f o r m i n g the lowe r m o s t Penninic units (Sub-Penninic nappes) and almost reaching as far south as the I n s u b r i c - P e r i a d r i a t i c line. (2) The late J u r a s s i c - C r e t a c e o u s Valais o c e a n ( A l p i n e Tethys), w h i c h closed during the late E o c e n e collision: r e m n a n t s p r e d o m i nantly consist of Cretaceous-age Bfindnerschiefer and are f o u n d within the L o w e r P e n n i n i c nappes. (3) T h e P i e m o n t - L i g u r i a n o c e a n (Alpine Tethys) o f midJurassic to Early Cretaceous age: largely subducted b e l o w the Adriatic microplate since the onset o f Late Cretaceous (Eoalpine) to Tertiary o r o g e n y although slivers are preserved within the U p p e r Penninic nappes.

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 129-145.0435-4052/06/$15.00

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Fig. 1. Sketch map of palaeogeographicalunits in the Alps (modifiedafter Froitzheimet al. 1996) with locations of the four crustal geophysical-geological transects shown in Figures 3, 13 and 14 (see also Fig. 12 for location of the lithosphereprofiles). Bold dashed contour outlines the Ivrea geophysicalbody; fine dashed contours indicate extrapolated tectonic lineaments.Details of surveys on ECORS-CROP, NFP-20 WEST and EAST, TRANSALP have been given by Roure et al. (1990), Pfiffneret al. (1997) and TRANSALP Working Group (2002), respectively.

(4) The Brianqonnais microcontinent or terrane: originally situated between the two above-mentioned oceans and part of the European margin before being rifted off along the future Valais ocean, preserved as Middle Penninic cover and basement nappes, but only in the Western And central Alps. (5) The Margna-Sesia fragment: a small splinter of the Apulian margin rifted off during the opening of the Piemont-Ligurian ocean, later incorporated into the accretionary wedge forming at the active northern margin of Apulia, and forming a part of the Lower Austroalpine nappes. (6) The Apulian plate south of the Periadriatic line: part of the Apulian plate referred to as 'Adriatic plate' and including the Ivrea body at its western margin, consisting of rigid south Alpine lower crustal rocks that were exhumed already during Mesozoic rifting. (7) The Apulian plate north of the Insubric line: parts of the Apulian plate at present forming most of the Austroalpine nappes, that is thin crustal flakes floating on the remnants of the Piemont-Ligurian ocean. (8) The Neotethys and its distal passive margin: mostly only the remnants of the distal passive margin of the Apulian plate facing Neotethys are preserved in the Alps (the Triassic-age Meliata ocean, a second former ocean, originally located SE of the Alpine Tethys, was closed by early Cretaceous times) and form small parts of the Austroalpine nappe system (restricted to the Eastern Alps of Austria). To a large extent the above-mentioned palaeogeographical findings are based on the analysis of near-surface evidence. The concept of plate tectonics showed clearly that the evolution of the Alpine orogen could be understood only by the additional

assessment of the detailed structural image of Alpine crust and lower lithosphere. Active seismic experiments (controlled source seismology; CSS) began in 1956 in the Western Alps and have continued since then throughout the Alps with increasing resolution of crustal structure. In particular, near-vertical reflection surveys along several across-strike transects provided the necessary insight into mechanisms of collision between continental crustal units. Earlier, various other seismic methods (e.g. dispersion of seismic surface waves and initial analysis of travel times) were used to derive first images of the Alpine lower lithospheric structure. However, until now these structures could not be resolved precisely enough, although knowing their geometry is indispensable for the understanding of the evolution of the Alps. For a summary of earlier results and their significance the reader is referred to Kissling (1993) and Mueller (1997). Only travel-time tomography, based on regional and teleseismic earthquakes, provided the desired resolution at the decisive depth range. We are fully aware that seismic anisotropy represents additional information. However, currently detailed and reliable anisotropy information covering a significant part of the volume under study is missing. At present, the combination of the available techniques and data allows us to establish a 3D image of the lithosphere to depths of 400 km and to quantitatively unravel the evolution of the Alps.

Crustal structure Classical deep seismic sounding with refraction and wide-angle reflection surveys provides a basic overview of the Alpine

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Fig. 2. Crustalcross-section alongthe European Geotraverse (EGT) across the CentralAlps (NFP-20 East in Fig. 1) from the northern Alpine forelandto the LigurianSea derived from active refractionand wide-anglereflectiondata. A-F, shot points; IL, Insubricline; numbersgive P-wave velocities(in km s-1); M, crust-mantle boundary (after Ye et al. 1995). crustal structure to Moho depths. Good information is available from a rather dense network of profiles in the Western and Central Alps, but much less is known from the Eastern Alps (for a compilation see Waldhauser et al. 1998). Interpretation of these data along profiles with various orientations provides excellent information on the topography of the crust-mantle boundary and on the average P-wave velocities used as input for an average 3D crustal model as described in the next section. In addition, characteristic structural details allowing for tectonic inferences can be derived in areas with favourable location and orientation of surveys. The large number of along-strike profiles in the Central Alps, combined with a densely occupied refraction survey along the north-south-oriented EGT (NFP-20 East in Fig 1; Fig. 4), allowed the derivation of a reasonably detailed crustal image in terms of P-wave velocities (Fig. 2) extending across the Alps and the Po Plain from the northern foreland to the Ligurian Sea (Buness 1992; Ye et al. 1995; Kissling et al. 1997), also including a careful error assessment (Waldhauser et al. 1998). Underneath the very variable and complicated sedimentary cover and the Alpine nappes, the main features of the Alps are the deep-reaching autochthonous Aar Massif (Pfiffner & Hitz 1997), which exhibits little internal structure, a high-velocity layer in the middle to lower crust north of the Insubric line, and two clear Moho offsets underneath the Alps and the northern Apennines, where lower European and Adriatic lithosphere are subducted under the Adriatic and Ligurian plates, respectively. This velocity information is a prerequisite for a satisfactory interpretation of near-vertical reflection transects. Figure 1 shows the location of four 2D Alpine crustal reflection transects (western transect ECORS-CROP, Roure et al. 1990; central and eastern transects NRP20, Pfiffner et al. 1997; Eastern Alps transect TRANSALP, TRANSALP Working Group 2001, 2002; Gebrande et al. 2006). Depending on the data quality, which is largely determined by the difficult topographic and geological conditions across the Alps, these transects provide the best available structural resolution for the entire crust. Structural details at depth can directly be incorporated into the near-surface geological and tectonic structure. However, these transects are limited in number. Figure 3 (Schmid & Kissling 2000; Schmid et al. 2003, 2004a) shows unified interpretations of these available nearvertical reflection surveys, including evidence from seismic refraction surveys and geological data along the three western transects and a preliminary interpretation of the TRANSALP transect. Going from the ECORS-CROP to NFP20-EAST (the TRANSALP profile will be discussed later) major common and/ or contrasting features are (Schmid & Kissling 2000): (1) ESEto south-directed subduction of European lithosphere; (2) offset between European and Adriatic Moho, as also seen in Figure 2;

(3) duplication and back-thrusting of lower European crust in the Western Alps (Fig. 3a and b) and wedging of Adriatic lower crust into the European middle crust under the Central Alps (Fig. 3c), respectively, covered by a stack of piled up and refolded upper crustal flakes (the Alpine nappes) in all three transects; (4) Adriatic Moho rising towards the Alps in the west and descending Moho at the base of the lower crustal wedge under the Central Alps; (5) eastwards increasing amounts of back-thrusting in the vicinity of the Insubric line; (6) strong shortening within the southern Alps in a foreland fold and thrust belt above the Adriatic lower crust (Sch6nborn 1992), which is exposed in the Ivrea Zone (Handy & Zingg 1991; Schmid 1993).

Moho topography

As mentioned above, a wealth of CSS crustal profiles in the wider Alpine region provides ample information on the Moho topography, depicted in Figure 4 (for an overall compilation see Waldhauser et al. 1998; for the Western Alps see Him et al. 1989; Thouvenot et al. 1990). Wide-angle reflections from the crust-mantle boundary are the most reliable and clear signals on most of these profiles. Based on these data, a method was developed that assesses the quality of Moho reflections, depths and crustal velocities (Kissling et al. 1997). Based on this method a reproducible 3D crustal model was established, comprising mean crustal velocities and a Moho contour map with least roughness within the estimated error estimates (Waldhauser et al. 1998). This model (Fig. 4) serves several purposes. First, it provides a good and reliable overview for variations in crustal thickness and mean velocity, and the relative position of these features with respect to surface tectonics and other geophysical observations. Second, it reliably shows the location of offsets of the Moho where subduction does occur. Third, it can be used to correct crustal travel times for teleseismic tomography, as discussed below. Figure 4a shows the contours of the Alpine Moho in 2 km intervals as derived by interpolation of the migrated CSS travel-time data located in the shaded areas, and Figure 4b shows a perspective N E - S W view of these surfaces. The number and lateral distribution of shaded areas also provides a measure of the high information density, which is unique worldwide. The image of the Alpine crust-mantle boundary shows two offsets, resulting in three separate Moho interfaces, namely the European, Adriatic and Ligurian Moho. The European Moho features a continuous change from an eastward dip under the Western Alps to a southern dip under the Central Alps, as already seen in the detailed transects of Figure 3. The Adriatic Moho is best imaged near the EGT-NFP20 profile, where it is up-domed below the Po Plain between the European and the

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Fig. 3. Crustal geophysical-geological transects through the Alps (after Schmid et al. 2004a). (For locations see Figs 1 and 12.)

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Fig. 4. (a) Crust-mantle boundary in the

wider Alpine region contoured at 2 km intervals derived from smoothest interpolation of 3D migrated CSS data. Numbers indicate isoline depth values in kilometres. Broken bars indicate locations of wide-angle Moho reflection elements. (b) Perspective SW view on the European, Adriatic and Ligurian Moho. The Ivrea geophysical body is schematically indicated (afterWaldhauser et aL 1998).

Ligurian Moho. Near the southern rim it is overthrust by the Ligurian crust, and at the western margin of the Po Plain the Adriatic Moho merges into the structure of the Ivrea Zone (see also Fig. 3b). This Alpine Moho topography reflects the large-scale Alpine structure resulting from the latest stage of continental collision (Schmid & Kissling 2000). Moho offsets and gaps, including their location, play key roles in tectonic interpretations of 3D lithospheric structure. A gap in otherwise continuous seismic information (Fig. 5) (e.g. as seen along near-vertical reflection profiles) could be interpreted as a

zone of symmetric subduction of lithosphere, or so-called 'Verschluckungs-Zone', as proposed by Laubscher (1970). However, Valasek e t al. (1991) and Holliger & Kissling (1992) imaged the expected Moho structure in the same area where no nearvertical reflections were observed (Fig. 5). Those workers used wide-angle data from CSS cross profiles and wide-angle reflections along the EGT-NFP20, respectively. The results obtained by networked wide-angle and near-vertical profiling prove that the Moho interface exists everywhere under the wider Alpine region. However, there is clear evidence for Moho offsets (Figs 3 and 5),

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Fig. 5. Combination of crustal structure and Moho depth along NFP-20 EAST/EGT transect (Fig. 1). (a) Line drawings of migrated near-vertical reflections and generalized seismic crustal structure derived from orogen-parallel refraction profiles (Holliger & Kissling 1992). Continuous line indicates Moho (M), wavy line the top of lower crust (C), dotted line the base of Penninic and Helvetic nappes, fine line the Insubric line; RRL, Rhine-Rhrne line. AUC, Adriatic upper crust; ALC, Adriatic lower crest; EUC, European upper crust; ELC, European lower crust. (b) Normal incidence representation of wide-angle Moho reflections along the EGT refraction profile perpendicular to the orogen across the Central Alps. The gap in the reflectivity signal from the lower crust is clearly covered by wide-angle reflection data (Valaseket al. 1991; Valasek & Mueller 1997).

indicating asymmetric subduction geometries. Hence, the Moho is not a continuous interface. Laterally bounded Moho signals define these offsets. Provided their relative positions are clearly defined, the sense of subduction can be inferred.

Lower crustal wedge structures

The high-resolution transect images shown in Figure 3 a - c clearly show that there is no common crustal model that would be valid for the entire Alpine arc in terms of a simple collision or shortening mechanism. Special features are the lower crustal wedges as found in the western and central Alpine transects. The wedge in the lower crust of the central transect (Fig. 3c) consists of Adriatic lower crust with a P-wave velocity of 6.5-6.6 km s - t (Fig. 2). This wedge lies above lower European crust, as can be derived from the clear reflection seismic data visible in Figure 5a. It is bounded by the Adriatic upper to lower crustal interface C and by the Adriatic Moho M. Its shape also agrees with the interpretation of CSS seismic refraction and wide-angle reflection observations (Fig. 2) by Ye et al. (1995) and Schmid & Kissling (2000). According to palinspastic reconstructions (Schmid et al. 1996, 1997), the northward intrusion of the wedge occurred during mid- to late-Miocene times and was contemporaneous with the formation of a fold and thrust belt within the upper Adriatic crust of the southern Alps. Hence, wedging is a rather late and suddenly appearing feature during the Alpine collision. It should be noted that this wedging requires complete detachment near the interface between lower and upper crust, and probably also at the base of the Adriatic lower crust that directly overlies the European lower crust, although this lower interface of the wedge is not clearly identified. Figure 6 (courtesy of Sch6nborn, pers. comm., based on Sch6nborn 1992) gives a perspective N W - S E view of the Alpine crustal model along the EGT-NFP20 transect, revealing the late-collisional wedging of Adriatic lower crust into the subducting European plate with the uncovered detachment interface of upper to lower crust. Rectangular arrows indicate the Late Miocene to present-day

NNW-SSE-oriented maximum horizontal stress direction that produced significant lateral extrusion of the Eastern Alps to the east (Ratschbacher et al. 1991). Shortening in the two western transects, which are located close together (Fig. 3a and b), predominantly occurred within the external European and Brianqonnais realms. The overall geometry suggests south-directed subduction of the European plate, as can also be inferred from the central transect (Fig. 3c). It should be noted, however, that in the western transects lower crustal wedging occurs within the European plate, as was extensively discussed by Schmid & Kissling (2000). The clear identification of the top of the lower crust, based on the exact position of the detailed near-vertical reflection profiles, together with the location derived from refraction and wide-angle reflection profiles between them, allowed the compilation of a contour map for this internal crustal interface (Schmid & Kissling 2000). Figure 7 shows the topography of the Conrad discontinuity (top lower crust) and identifies the lateral extent of lower crustal wedges of different origin, situated in the hinge zone between the north-south-striking Western Alps and the east-west-striking Eastern Alps. The Adriatic lower crustal wedge under the Central Alps and the European crustal wedge under the Western Alps (Figs 3 and 8) meet at depth below the location where the Simplon fault zone branches off the Insubric line; that is, about halfway between NFP20 reflection profile segments W3 and C2 in Figure 7. This indicates a rather sharp transition from Western to Central Alps at depth. In accordance with recent findings in exhumed high-pressure rocks (Lund et al. 2004) the geometry of this wedging of Adriatic and European lower crust, as discussed so far, suggests that the bulk of the lower crust is made up of high-strength material, contrary to a widely held belief in a 'weak lower crust' by the geoscience community (e.g. Meissner & Kusznir 1987; Banda & Cloetingh 1992; Willingshofer & Cloetingh 2003). Such low-viscosity material, however, must be present within relatively thin layers forming the interfaces of the lower crustal wedges with the upper crust and the upper mantle, respectively, allowing for detachment near these interfaces bounding the wedges.

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Fig. 6. Perspective NW-SE view of the Alpine crustal model with the NFP-20/EGT transect at the western edge, showing the late collisional wedging of lower Adriatic crust into the subducting European plate with uncovered detachment of upper to lower crust interface. Rectangular arrows indicate the present-day NNW-SSE-oriented horizontal stresses; round arrows indicate lateral escape movement of the upper crustal units to the east and SW (courtesy of G. Sch6nborn, after Sch6nborn 1992).

A quantitative post-35 Ma kinematic reconstruction of S E - N W shortening along the ECORS-CROP and Central Alps transects is displayed in Figure 8. The total amount of shortening is a composite of the westward strike-slip component of the Adriatic plate relative to Europe and the amount of north-south shortening along the EGT-NFP20 transect (Schmid et al. 1996). As shown in Figure 7 the Adriatic Moho rises to shallower depth and merges into the structure of the Ivrea Zone (see also Fig. 3a and b) at the western margin of the Po Plain. The subvertical orientation of rigid lower crust and upper mantle material that rises to the surface is probably responsible for the presence of a backstop, causing the relatively young back-thrusting and doubling of the European lower crust under this part of the Western Alps.

E a s t e r n A l p i n e crustal structure

A fourth and latest crustal transect (Fig. 3d) TRANSALP was established across the Eastern Alps (Fig.l) by high-resolution

reflection and refraction seismic survey and other geophysical methods (Ebbing et al. 2001; TRANSALP Working Group 2001, 2002; Bleibinhaus 2003; Lueschen et al. 2003), following much earlier work by Miller et al. (1977). First interpretations of the data were presented by Lammerer & TRANSALP Working Group (2003) and Castellarin et al. (2003). Additional information on this transect has been given by Nicolich et al. (2003). The boundary between the Western and Eastern Alps coincides roughly with the north-south-striking western margin of the Austroalpine nappes (Fig. 1, western margin of Apulian plate north of the periadriatic line), which formed by top-to-the W N W suturing of the Austroalpine units with the Piemont-Ligurian ocean during a first orogenic cycle in the Cretaceous (Froitzheim et al. 1994). However, the more external Briangonnais microcontinent and Valais ocean, bordering the European margin, were not sutured to the Austroalpine units before the end of a second orogenic cycle in the latest Eocene (e.g. Schmid et al. 1996). Another important boundary running across strike and situated immediately east of the western end of the Eastern Alps is formed by the

Fig. 7. Contour map of top of European lower crust in the north, top of European lower crustal wedge in the west, top of Adriatic lower crust in the north-central region, and of Adriatic Moho in the south with position of main crustal transects (Fig. 3). The tentative depth extension of the Insubric line and location of high-velocity Ivrea material (after Schmid & Kissling 2000) should also be noted.

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lithosphere recently revealed by high-resolution tomography in an area situated immediately to the east (Lippitsch 2002; Lippitsch et al. 2003; Schmid et al. 2003). These data will be discussed below.

Lower lithosphere Early m o d e l s o f l o w e r lithospheric structure derived w i t h o u t crustal corrections

Fig. 8. A kinematic reconstruction of post-collisional(post 35 Ma) shortening along the western transect. The vector triangle (top right) illustrateshow the total amount of post-collisionalshortening (124 km) of vector AE (movementof Adriatic microplate relative to stable Europe) is composed of 100 km dextral strike slip between the Adriaticmicroplateand the CentralAlps (vector AC) onto the western transect; 71 km of vector CE results in north-south shorteningwithin the Central Alps parallel to the NFP-20/EGT transect (after Schmid & Kissling 2000).

sinistrally transpressive Giudicarie belt (Fig. 1). Sinistral shearing along the Giudicarie belt during the early Miocene caused northdirected indentation of the eastern part of the southern Alps and massive north-south shortening in the Tauern Window (Ftigenschuh et aL 1997; Stipp et al. 2004). This was accompanied by lateral extrusion of the Eastern Alps east of the Brenner normal fault (Ratschbacher et al. 1991) associated with dextral strike slip along the Periadriatic line (Fig. 6). The interpretation of the TRANSALP transect shown in Figure 3d only partly follows that proposed by the TRANSALP Working Group (2002) regarding the near-surface structures and it drastically differs in terms of the interpretation of the deep structure. We attribute the core of the Tauern Window to tectonic units of the European margin, this core having the appearance of a flower structure (Dobrin & Savit 1988). We agree with one of the two models presented by the TRANSALP Working Group (2002) regarding a continuation of south-dipping Tauern Window reflectors underneath the southern Alps south of the Periadriatic line inherited from an initial stage of south-directed subduction. However, we do not see enough evidence for an Adriatic lower crustal wedge indenting the European crust from the seismic data presented by the TRANSALP Working Group such as observed in the NFP20-EAST profile (see Fig. 3c) further west. In fact, these data are also not clear enough to indicate an Adriatic Moho descent northward under the European lithosphere as drawn in Figure 3d, either. However, when proposing the interpretation given in Figure 3d, we are strongly guided by data indicating a NE-oriented subduction of south Alpine

As clearly follows from the discussion above, the tectonic evolution and structure of the Alps cannot be understood or reconstructed without reliable and sufficient knowledge of the 3D structure of the lower lithosphere and the lithosphere-asthenosphere boundary at least in terms of P-wave velocity distribution. Any technique chosen has to resolve possible differences in the regional structure between the Western, Central and Eastern Alps to depths of at least the penetration of the lithospheric slabs. Therefore, over the years a considerable effort was made to gain more insight into the relevant upper mantle structure. Kissling (1993) compiled a critical summary of pre-1993 knowledge on lower lithospheric structure and concluded that there exists a thickened lower lithosphere beneath or near the Alps that indicates a southerly and southeasterly dip under the Central and Western Alps, respectively. These results were mainly based on surface-wave analysis and travel-time residual studies (Babuska et al. 1990; Suhadolc et al. 1990; Guyoton 1991; Viel et al. 1991; Ansorge et al. 1992) as well as on seismic tomography using datasets of limited accuracy or covering only a small area in the SW Alps (Cattaneo & Eva 1990; Spakman 1991; Spakman et al. 1993), respectively. However, methods and resolution were not sufficient to image the 3D structure in detail and, in particular, to properly resolve slab geometries. Since then, travel-time tomography has developed to a powerful tool for the resolution of global and large-scale regional structures in the upper mantle by mainly using data collected by international seismic bulletins such as those produced by the International Seismological Centre (ISC). Regarding the wider Alpine region we show two important cases of upper mantle structure recently derived by travel-time inversions (Bijwaard & Spakman 2000; Piromallo & Morelli 2003). Both examples use the same ISC data with different selection and resolution criteria in the inversion process. Piromallo & Morelli (2003) showed a continuous structure of high-velocity material underlying the Alps from west to east, as seen in a map view in terms of P-wave velocity variations at a depth of 150km. This structure is rather diffuse at shallow levels and disappears with increasing depth. Piromallo & Morelli (2003) applied corrections for global non-spherical velocity structure outside the model volume of the wider Mediterranean area but did not apply crustal corrections. Figure 9a depicts a vertical crosssection in the Eastern Alps through the model of Piromallo & Morelli (2003). A diffuse subvertical high-velocity body is seen under the Alps, reaching the 410 km discontinuity. The vertical section across the Central Alps (Fig. 9b) derived by Bijwaard & Spakman (2000) is again based on a large set of ISC first P-wave arrival times without crustal corrections. In this transect the high-velocity structure beneath the Alps appears to dip in a southerly direction. On a larger scale, the two models by Piromallo & Morelli (2003) and Bijwaard & Spakman (2000) agree regarding the existence of a high-velocity structure beneath the Alps. However, in the latter model this structure varies rather unsystematically in a horizontal direction, and it may even disappear further east. In summary, these tomographic mantle models were derived without crustal corrections. Furthermore, they were obtained by inversion of P-wave travel times determined with different picking routines from seismograms recorded on a variety of

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Fig. 9. P-wave velocities as derived by travel-time tornography. Crustal structure is not resolved in both cases. (a) NE-SW lithospheric cross-section from Central Europe to northern Africa across the Eastern Alps after Piromallo & Morelli (2003). Lower part shows section A-a/. The northerly vergence of subducted high-velocity material below 150 km depth under the Eastern Alps should be noted. (b) NW-SE cross-section through the Central Alps after Bijwaard & Spakman (2000). (For location see B-B 1in Fig. 12.) The SE vergence of subducted European lithosphere, and the high-velocity material in the upper mantle transition zone between 400 and 620 km depth should be noted.

seismographs and reported to ISC. The significant error in the ISC data and the uncorrected effects of the 3D crustal structure seriously limit the resolution of this kind of regional tomography. Hence, the detected structure in the lower lithosphere cannot be correlated clearly with independently determined geophysical a n d / o r tectonic evidence (Fig. 3), nor can it be used to discriminate between hypotheses for the tectonic evolution of the Alpine orogen.

L o w e r lithospheric structure based on high-resolution teleseismic tomography In spite of the difficulties regarding the desired resolution discussed above, teleseismic tomography is a very valuable tool to obtain reliable basic structural information in depths ranges that are hard to assess in detail with other methods. To further increase the resolution of teleseismic tomography we recently developed a new and different approach (Waldhauser et al. 1998, 2002; Arlitt et al. 1999), which can only be summarized here. This procedure uses: (1) a set of carefully selected teleseismic events with digital

signals transformed to the same standard recording response; (2) a uniform picking routine for seismic phases resulting in a highly consistent dataset; (3) a careful correction of observed travel times for 3D crustal contributions. Figure bOa schematically shows the ray paths from a single teleseismic event to an array of recording stations through the standard Earth model IASP91. These signals traverse the comparatively slow crust above the study area in a subvertical direction. The lack of crossing rays within the crust prevents a reliable resolution of lateral velocity variations in this crustal layer. Therefore, a representative evenly gridded 3D model (Fig. 10b) was compiled for the Alpine crust from all available seismic data obtained by active CSS methods in terms of mean crustal velocity structure, Moho topography, and sedimentary basins (see Fig. 4). This a priori known crustal velocity model allows us to correct the observed teleseismic travel times for the carefully calculated crustal contribution (Waldhauser et al. 2002), which may account for up to 50% of the total travel-time residuals. In a second step we merged selected evenly distributed teleseismic events (Fig. l lb) recorded at sufficiently dense Austrian, French, German, Italian, Slovenian and Swiss permanent seismic

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Fig. 10. (a) Schematic ray path diagram for teleseismic events recorded at a given station array. Observed travel times contain contributions from mantle, lower lithosphere and crust. Crustal contributions are determined independently for correction of observed total travel times. (b) Perspective NW view of 3D Alpine crustal P-velocity model to 70 km depth used for the calculation of crustal travel times, shown in a cut-away display. ME, European Moho; ML, Ligurian Moho; M A, Adriatic Moho.

networks in the wider Alpine region, and, additionally, at the passive seismic network of TRANSALP (Fig. l la) (Lippitsch 2002; Lippitsch et al. 2003). The dataset of this study consists of the travel times of 4698 manually picked first arrivals from 79 events with even azimuthal distribution. Absolute travel times from the selected events were used to determine an initial reference subcrustal velocity model for the investigated area. Sensitivity and resolution tests with synthetic data show that a combination of non-linear inversion, high-quality teleseismic data and the use of the a priori 3D crustal model allows us to reliably resolve structures of about 60 km linear length in the upper mantle in most areas of the investigated region.

Seismic structure of the Alpine lithosphere derived by integrating crustal and mantle structure In the following we present images of the lower lithosphere derived from high-resolution tomography, and we then assess their relation to the independently determined crustal transects discussed above to achieve a unified lithosphere model for the Alpine orogen. Figure 12 shows a detailed map view of the Alpine lithosphere-asthenosphere system for the 135-165 km depth interval. The continuous high-velocity structure beneath the Alps found by Piromallo & Morelli (2003) at the same depth range is now split into two separate slabs, situated in the Western to Central and Eastern Alps, respectively. Both again follow the strike of the orogen. As will be discussed below, such a clear separation is supported by strong differences in the structure of these two slabs which are clearly visible in the vertical transects. There are also indications for further differentiations in the structure of the western slab. The same horizontal section additionally reveals a pronounced negative velocity anomaly situated under the eastern

Po Plain. It remains unclear whether this represents a singular local feature, or alternatively, the northern part of an extended low-velocity structure in the Adriatic plate as imaged by Piromallo & Morelli (2003) and Di Stefano et al. (2006). The seismic images of the two major lithospheric constituents, crust and mantle lithosphere, were not derived simultaneously but separately and over many years. Hence, we cannot a priori expect that the locations of crustal transects, selected on the basis of surface tectonic or practical experimental criteria, are also the most suitable transects for optimally representing lower lithosphere structure and geometries. The profiles in Figure 12 indicate the positions of three lithosphere profiles, A - A ' , B - B ' and C - C ' , that exhibit very clear images of the newly derived 3D lower lithosphere slab geometries (Fig. 13) in blue. Figure 12 also shows the positions of four lithospheric profiles I - I V (Figs 13a and 14), that contain the well-imaged crustal structures discussed earlier with respect to Figure 3, in red. The tomographic images illuminate the upper mantle to depths of 400 km and reflect the current status of the complex processes that shape the Alpine orogen today. Transect B - B ' (Fig. 13b), situated at the transition from the Central to the Eastern Alps (Fig. 12) displays the clearest image of the present European-Adriatic collision structure and process. European and Adriatic Moho, as derived from CSS surveys, serve as guidelines for fixing the crustal structure that remains unresolved by mantle tomography, and they define the location of the suture at depth. The high-velocity volume subducted to the SE, outlined by a dashed line, clearly distinguishes the European lower lithosphere from the surrounding area exhibiting background velocities identifiable as asthenosphere in that depth range. The amount of more or less undeformed subducted lower lithosphere, when interpreted as continental, implies a shortening of c. 120 km since the onset of continental collision. For the

ALPINE LITHOSPHERE

Fig. 11. Seismic stations and teleseismicevents used for the tomography study. (a) Location of permanent and temporary seismic networks in the wider Alpine region and (b) location of selected earthquakes. These are evenly distributedand were recorded with high quality.

first time this provides an independent measure of post-collisional shortening, against which palinspastic reconstructions based on crustal structure can be checked, assuming that the transect runs parallel to the direction of subduction. Further high-velocity volumes visible between 350 km and 400 km are not yet reliably identified. They might be remnants of earlier subducted and detached oceanic lithosphere, as suggested by von Blanckenburg & Davies (1995) to explain magmatic intrusions found along the Periadriatic Lineament. Significant low-velocity bodies at shallow asthenospheric depth appear beneath the southern Rhine Graben and the Po Plain. Further information on the Rhine Graben rift system has been given by Prodehl et al. (1995) and Lopes Cardozo & Granet (2003). Transect A - A ' (Fig. 13a) in the Western Alps is chosen to coincide with crustal cross-section I (Figs 12 and 3a) provided by ECORS-CROP, now incorporated into the lithospheric image at proper scale. Dotted lines indicate the CSS-derived European and Adriatic Moho. In general, continental European lower lithosphere is subducted east to SE beneath the Adriatic microplate, with the high-velocity material reaching a depth of at least 400 kin. The variation of velocity pattern at about 300 km depth may indicate a change in origin of the subducted material at greater depth from continental to oceanic. The slope of the subducted lithosphere clearly varies with depth, being subvertical to 250 km with a slight tendency of rollover. At Moho depth the thrust wedge formed by European lower crust (Fig. 3a) ends against the adjacent high-velocity Ivrea body and bends into the subvertically oriented high-velocity lower lithosphere (Schmid & Kissling 2000). At about 100 km depth detachment of the deeper parts of the European continental lithosphere has occurred, possibly induced by the Ivrea body, which seems to have acted as a buttress in the collision process of the European and Adriatic plates, leading to the present subvertical

139

orientation. Sue et al. (1999) also have invoked slab detachment under the Western Alps to explain extensional earthquake mechanisms. The detachment may be accompanied by additional upwelling of asthenospheric material in the pronounced lowvelocity region immediately to the west of the detachment. A third transect, C - C ' (Fig. 13c), representative for the mantle structure in the Eastern Alps, lies east of the TRANSALP traverse (Fig. 12). It again reveals a rather obvious subduction pattern, although, very surprisingly, with a subduction polarity opposite to that found in the Western Alps, as the European lithosphere here represents the overriding plate. In this section through the 3D lithospheric model, the Adriatic lower lithosphere is found to have been subducted to the NE and underneath the European plate to a depth of 270 km. Using the suture between European and Adriatic Moho as reference, the total shortening since collision amounts to some 210 km, significantly more than observed for the SE-directed subduction along transect B - B ' situated further west. There is no indication of a detachment in the subducted lower lithosphere within the observed depth range. Transect C - C ' reaches the Po Plain low-velocity anomaly seen more clearly in Figure 12 near its SW end. The 3D tomographic model clearly documents that the dip direction of subduction flips from SE to NE between transects B - B ' and C - C ' . This flip occurs over the relatively short distance of about 8 0 k m and between the two separate high-velocity volumes shown in Figure 12. This important transition occurs roughly beneath the Giudicarie tectonic lineament and beneath the TRANSALP profile, and is illustrated in more detail in Figure 14. Figure 13 presents the most important and clearest features of the Alpine lower lithosphere and its lateral variations in the form of representative cross-sections. As the locations of these mantle sections, ideally chosen for depicting mantle structures, do not coincide with the available crustal transects depicted in Figure 3, additional sections are provided in Figure 14. This figure combines the lower lithosphere and crustal structures in the form of whole lithosphere transects that also incorporate the crustal images within the 3D tomographic model crustal images at their proper locations I - I V (Fig. 12). The combined sections A - A ' and I have already been shown and discussed above as part of Figure 13a. Crustal profile II (Fig. 14a) along the NFP20 WEST transect shows structural features that are nearly identical to those found in the combined profile along the ECORS-CROP transect shown in Figure 13a. It confirms the sense of subduction and also the important role of the Ivrea body, which probably caused the detachment of a lower European crustal wedge (Schmid & Kissling 2000). Probably, it is also responsible for the observed slab detachment of lower lithosphere visible under the crustal transect. Further to the south both the lithospheric transects A - A ' and the deep structure under crustal profile II cover nearly the same area and exhibit near-identical velocity patterns. The transect shown in Figure 14b follows crustal profile III (NFP 20 EAST/EGT) across the eastern margin of the Western Alps and it obliquely crosses transect B - B ' (Fig. 13b; see also Fig. 12). A careful analysis of the crustal data (Holliger & Kissling 1992; Schmid et al. 1996; Pfiffner et al. 1997; Valasek & Mueller 1997) led to the conclusion that a wedge of relatively rigid lower Adriatic crust did indent the European crust. Consequently the sheared-off European lower crust possibly was subducted together with its lower lithospheric substratum (Pfiffner et al. 1997). However, teleseismic tomography cannot resolve such a detailed feature within the subducted volume; yet, as seen in Figure 14b, the lower boundary of the European lower crust can most easily be extrapolated into the outlines of the high-velocity lower lithosphere subducted towards the SE. This could also suggest that only a small part of the lower continental European crust has been subducted (Burg et al. 2002). The difference in strength and rigidity between the European crust in the north and the Adriatic crust in the south, which includes the deep-reaching Ivrea body at its western end

140

E. KISSLING ETAL.

Fig. 12. P-wave velocity distribution between 135 and 165 km depth, with linear interpolation between inversion cells (from Lippitsch 2002; Lippitsch et al. 2003). Velocity variations are plotted relative to a 1D initial reference model determined from absolute travel times for the research area. Areas with no resolution are left grey; areas with critical resolution are displayed in pale colours. Thick black dashed lines indicate areas of high-velocity European and Adriatic lower lithosphere, which is subducted east to SE and north to NE under the Western and the Eastern Alps, respectively. Thick white dashed lines indicate (from west to east) the Insubric, Giudicarie and Periadriatic lineaments (PL) and the Tauern Window (TW) as part of the Eastern Alps. Red dashed lines I, II, III and IV mark locations of crustal geophysicalgeological transects (see Figs 1 and 3). Blue dashed lines A-A', B-B' and C-C' mark locations of lower lithospheric transects (Fig. 13). Dark red area indicates the Po Plain anomaly.

(Figs 13a and 14a) but normal continental crust further east, has probably caused the contrasting deep lithospheric structures between the westernmost Western Alps and the eastern part of the Western Alps traversed by the EGT profile. The necking of the high-velocity material, seen at around 300 km depth under transects A - A ' and II (Figs 13a and 14a), is confirmed by the narrow vertical separation of high-velocity material (Fig. 14b) at 250 km depth, as is seen under the southern extension of the obliquely crossing transect III. This may suggest an increasing separation of detached and subsiding dense high-velocity material from east to west. The transition between the opposite subduction regimes characteristic for the Western and the Eastern Alps, respectively, can be seen on the north-south-oriented transect IV that corresponds to the T R A N S A L P profile shown in Figure 14c. Features of lower crustal wedging and imbrications as observed further west are now replaced by a relatively narrow collision structure situated under the western end of the Tauern Window where the Giudicarie line joins the Periadriatic Lineament (Schmid e t al. 2003). Here the originally flat-lying piles of nappe structures, as observed in the west, is dramatically steepened, and the structure under the Tauern Window resembles a large-scale flower structure as commonly seen in sediments (Dobrin & Savit 1988). Judging from the 2D section alone, the polarity of subduction is not as obvious as it is either further east (Fig. 13c) or further west (Fig. 13b). The location of profile T R A N S A L P in the horizontal section provided for a depth of 150 kin (Fig. 12, profile IV), however, makes it clear that the lithospheric structure underneath this profile represents the transition zone between west and east. Hence, very probably it is already similar to that which predominates underneath the Eastern Alps, as is depicted in Fig. 13c and in another

lithosphere-scale transect, 'EASTERN ALPS', discussed elsewhere (Schmid et al. 2004b). A subvertically oriented structure characterized by moderately high velocity, without sharp boundaries, extends to 220 km depth into the surrounding asthenosphere. Parts of the diffuse geometrical outlines of the high-velocity material are caused by the east-to-west averaging of the grid elements over about 100 km, which has resulted in the inclusion of structural features from the western as well as the eastern subduction systems. Possibly, the derived tomographic image reflects only the northern end of the deep-reaching left-lateral Giudicarie shear zone along which no clear direction of subduction can be defined.

Discussion and conclusions This study compiles and combines the major features of the 3D structure of the crust with the 3D structure of the lower lithosphere in terms of its P-wave velocity distribution. Until recently, and for strictly methodological reasons, crust and lower lithosphere were investigated and interpreted separately. Thanks to the latest advancement of deep crustal exploration techniques with higher resolution, for example appropriate combined refraction, near-vertical and wide-angle reflection seismic surveys, together with high-resolution teleseismic tomography, particularly including crustal corrections for the lower lithosphere, these structural units can now be explored together. This allows for a better understanding of their unseparable roles in tectonic evolution. It also provides a much better understanding of the derived present-day 3D structure in time and space, even within comparatively small orogens such as the Alps. The careful analysis of the actively acquired seismic data for the crust indicated and proved the existence of different types and

ALPINE LITHOSPHERE

141

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mechanisms of wedging and indentation. These include features such as (1) back-thrusting or doubling of lower European crust as a consequence of the collision with the deep-reaching rigid Ivrea part of the Adriatic plate in the west, and (2) wedging of lower Adriatic into European crust, implying intense shearing at the

C

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Fig. 13. Three lower lithosphere transects that exhibit very clear images of the newly derived 3D lower lithosphere slab geometries: (a) A - A ' , Western Alps, with superimposed crustal transect ECORS-CROP as depicted in Figure 3a; (b) B-B', Central Alps, (c) C-C', Eastern Alps. (For location and display mode see Fig. 12.) Crustal layers are set to zero velocity deviation, with Moho topography taken from the 3D crustal model in Figure 10b. Bold dashed line indicates lithosphere- asthenosphere boundary (LAB). IL, Insubric line; AF, Alpine front.

interface between upper and lower crust and leaving the European lower crust closely connected to its lower lithospheric underpinnings. The fourth transect, that is TRANSALP through the Eastern Alps, positioned near the transpressive Giudicarie line, lacks wedging

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Fig. 14. Combined crustal and lower lithosphere transects along the high-resolution crustal geophysicalgeological profiles depicted in Figure 3 b - d (see Fig. 13a for the ECORS-CROP crustal transect depicted in Fig. 3a). Location and display mode are shown in Figure 12, profiles II, III and IV. (a) NFP-20 WEST; (b) NFP-20 EAST/EGT; (c) TRANSALP.

ALPINE LITHOSPHERE

within the lower crust. Kummerow et al. (2003, 2004) have derived an independent and alternative north-south cross-section that features the Moho along the TRANSALP transect based on receiver function analysis. In an area between the northern Molasse Basin and the centre of the orogen this cross-section agrees with our compilation, which is based on earlier CSS data and the new TRANSALP data. However, it differs significantly in the southern part. There, Kummerow et al. proposed a subhorizontal crust-mantle boundary at 40 km depth. It will be highly interesting to know which type of lower crustal collision scheme will be found further east, where the extensive ALP2002 active seismic experiment was carried out recently (Brueckel et al. 2003; Gebrande et al. 2006). The mechanisms by which large portions of lower crust and lithosphere disappeared during the latest continental collision process, as well as the location of the subducted remnants, have been a topic of discussion since the early hypothesis of 'Verschluckung' (Laubscher 1970). Mueller (1997) provided evidence for a lithospheric root of deeply south-dipping structures. South-dipping subduction in the Central Alps was also postulated and suggested based on the interpretation of active seismic experiments, geological data and seismic tomography (Fig. 9b) (Spakman et al. 1993; Pfiffner & Hitz 1997; Stampfli & Marchant 1997; Bijwaard & Spakman 2000). The present work, which is based on the construction of a representative 3D crustal model for crustal travel-time corrections, combined with the use of high-quality teleseismic travel-time data, provides a quantitative assessment of the length of lithospheric slabs, which varies significantly along the orogen (Fig. 13a-c). Based on the estimates of post-collisional crustal shortening along the EGT and ECORS-CROP transects (Schmid & Kissling 2000), we identify the slab beneath the Central Alps to represent lower continental lithosphere (Figs 13b and 14b). The detached slab beneath the inner arc of the Western Alps (Figs 13a and 14a), however, is likely to contain continental and oceanic lithosphere. This strongly supports the model of von Blanckenburg & Davies (1995), who postulated (oceanic) slab break-off early during collision at least for parts of the Alpine subduction zone. The increased resolution allows for the separation of two nearly opposing subduction regimes in the Western and Eastern Alps, respectively, the 3D structure suggesting an angle of nearly 100 t` between the strike of the two slabs (Fig. 12); it should be noted that, in three dimensions, the European slab dips to the SE underneath the Western Alps, whereas the Adriatic slab in the Eastern Alps dips to the NE. This complex Alpine subduction system is obviously connected with significant deep-reaching left-lateral strike-slip motion along the Giudicarie line. We tentatively propose that the continental Adriatic lower lithosphere subducted to the NE underneath the Eastern Alps represents a remnant of Eocene orogeny in the Dinarides, characterized by NE-directed subduction according to geological evidence. This piece of lithosphere is interpreted to have been laterally displaced and inserted into the Alps by dextral movements along the Periadriatic line, combined with sinistral transpression along the Giudicarie line during the last 20 Ma (see discussion by Schmid et al. 2004b). The slab detachment at shallow depth (120 kin) under the western transects (ECORS-CROP and NFP20 WEST) was probably caused by the adjacent strong Adriatic lower crust with the Ivrea body as its surface expression. This again shows that crust and lower lithosphere have to be studied in the context of a plate that consists of layers of different mechanical strength, and that possibly exhibits strong lateral variations, when modelling orogenic subduction-collision processes. The new 3D lithospheric picture for the wider Alpine region to 400 km depth demonstrates the tight connection between the deep structure of the lithosphere-asthenosphere system and near-surface tectonic features. It provides new and unexpected evidence for the younger Alpine tectonic evolution, a process that obviously induced significant changes in the architecture of the Alps and their lithospheric roots from west to east.

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Listing the names of all the colleagues in geology and geophysics who contributed over the years with data and valuable research would go almost beyond the length of this paper. We would like to collectively thank all of them. Special thanks go to all colleagues and seismological services in the wider Alpine region who provided us with the teleseismic data. G. Schtnborn is thanked for providing us with the visualization depicted in Figure 6. Careful reviews by U. Achauer and an anonymous reviewer are gratefully acknowledged.

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Tertiary-Quaternary intra-plate magmatism in Europe and its relationship to mantle dynamics MARJORIE WILSON t & HILARY DOWNES 2

llnstitute of Geophysics and Tectonics, School of Earth and Environment, Leeds University, Leeds LS2 9JT, UK (e-mail: M. Wilson @earth. leeds.ac, uk) 2School of Earth Sciences, Birkbeck College, London WC1E 7HX, UK

Abstract: Anorogenic intra-plate magmatism was widespread in Europe from early Tertiary to Recent times, extending west to east from Spain to Bulgaria, and south to north from Sicily to central Germany. Magmatism is spatially and temporally associated with Alpine-Pyrenean collisional tectonics, the development of an extensive lithospheric rift system in the northern foreland of the Alps, and, locally, with uplift of Variscan basement massifs (Massif Central, Rhenish Massif, Bohemian Massif). The volcanic regions vary in volume from large central volcanoes (e.g. Cantal, Massif Central;Vogelsberg, central Germany), to small isolated plugs (e.g. Urach and Hegau provinces in southern Germany). Within the Mediterranean region, the Dinarides, the Pannonian Basin and Bulgaria, anorogenic volcanism locally post-dates an earlier phase of subduction-related magmatism. The major and trace element and Sr-Nd-Pb isotope characteristics of the most primitive mafic magmatic rocks (MgO > 6 wt%) provide important constraints on the nature of the mantle source and the conditions of partial melting. These are predominantly sodic (melilitites, nephelinites, basanites and alkali olivine basalts); however, locally, potassic magma types (olivine leucitites, leucite nephelinites) also occur. In several localities (e.g. Sicily; Vogelsberg and the Rhine Graben, Germany; Calatrava, central Spain) olivine and quartz tholeiites form a significant component of the magmatism. The sodic magmas were derived by variable degrees of partial melting (c. 0.5-5%) within a transitional zone between garnet-peridotite and spinel-peridotite mantle facies, close to the base of the lithosphere; the potassic magma types are interpreted as partial melts of enriched domains within the lithospheric mantle. Mantle partial melting was induced by adiabatic decompression of the asthenosphere, locally in small-scale, plume-like, diapirs, which appear to upwell from c. 400 km depth.

Tertiary and Quaternary volcanic activity within Europe occurs in two principal geotectonic settings, referred to as orogenic and anorogenic by Wilson & Bianchini (1999). Occurrences of calc-alkaline volcanism (orogenic) in the Alpine chain, the Carpathians and the Mediterranean region (Harangi et al. 2006) can be explained geodynamically in terms of contemporaneous subduction, and will not be considered further in this review. Here emphasis is placed on the extensive anorogenic, dominantly alkaline, volcanic province to the north of the Alpine collision zone, including the Massif Central of France, the Rhenish Massif of central Germany, the Rhine Graben, and the Eger Graben within the northern part of the Bohemian Massif in the Czech Republic (Figs 1 and 2). Further to the east mafic alkaline volcanism (anorogenic) post-dates a major phase of subduction-related volcanism in the Pannonian Basin, the Dinarides (Serbia, Slovenia, Croatia, northern Bosnia), Bulgaria and western Turkey. Further south, within the Mediterranean region, anorogenic volcanism occurs in Sicily, Sardinia, Monte Vulture and the Veneto area of Italy, in the Alboran Sea and along the northern coast of Africa, locally post-dating earlier phases of subduction-related magmatism. Anorogenic magmatism also occurs in the Iberian peninsula, mainly in the Calatrava province of south central Spain, and within the southeastern Pyrenees near Olot. Although magmatism initiated locally in the latest CretaceousPalaeocene, the major phase of activity in Western and Central Europe occurred in the Neogene ( 2 0 - 5 Ma), with a subsidiary peak in the Pliocene ( 4 - 2 Ma) (Figs 3 and 4); eruptions continued locally to a few thousand years ~p. Magmatic activity within the European foreland of the Alpine orogen (Fig. 2) is typically mafic and occurs as small-volume monogenetic centres (e.g. Eifel, Urach and Hegau provinces of central Germany), scattered necks and plugs (e.g. North Hessian Depression, Germany) and fissure-controlled plateau basalts (e.g. C~zallier, Aubrac and Coirons in the French Massif Central). Rarer central volcanic complexes (e.g. Cantal and Mont Dote in the Massif Central; Vogelsberg in central Germany) include significant volumes of more differentiated magmas that can be related to processes of magmatic differentiation in sub-volcanic m a g m a chambers (e.g. Wilson et al. 1995a).

The volcanic fields are generally concentrated in lithospheric basement terranes that have experienced tectonothermal events within the last 3 0 0 - 4 0 0 Ma (e.g. the Variscan belt of Europe); these typically have higher heat flow and thinner lithosphere than the surrounding cratons (e.g. Baltic Shield) (Prodehl et al. 1992). A number are located on uplifted basement massifs (e.g. Massif Central; Rhenish Massif; Bohemian Massif) that appear to be dynamically supported by upwelling asthenospheric mantle diapirs (e.g. Granet et al. 1995; Ritter et al. 2001; Wilson & Patterson 2001). Magmatic activity was broadly synchronous with the evolution of an extensive intra-continental rift system in Western and Central Europe, subsequently referred to as the ECRIS (East and Central European Rift System), the origins of which are intimately linked to the collision of Africa with Eurasia (Wilson & Downes 1992; Ziegler 1992; D~zes et al. 2004). The most common primitive mafic m a g m a types are sodic basanites and alkali basalts; highly silica-undersaturated, small melt fraction, nephelinites and melilitites occur much less frequently, although they are the dominant m a g m a type in some areas (e.g. Urach province of Germany). Potassic alkaline mafic magmatism (e.g. leucitites, leucite nephelinites) occurs at scattered localities throughout the province (e.g. Calatrava, Spain; Cantal and the Sillon Houiller in the Massif Central; the West Eifel, Germany; Doupovsk6 Hory and Cesk6 Stredohori in the Bohemian Massif); however, only in the Quaternary East West Eifel do potassic magmas predominate over sodic m a g m a types. More exotic m a g m a types, such as carbonatite, occur very rarely (e.g. Kaiserstuhl in the Rhine Graben, Germany; Monte Vulture, Italy). Felsic magmatic rocks occur in most of the volcanic fields, sometimes in a bimodal association with the basalts; complete differentiation series are, however, rare, and occur only in the central volcanic complexes (e.g. Cantal).

The geodynamic setting of the magmatism The Late Cretaceous-Cenozoic convergence of A f r i c a - A r a b i a with Eurasia resulted in the progressive closure of oceanic

From: GEE, D. G. & STEPHZNSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 147-166. 0435-4052/06/$15.00

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Fig. 1. False coloured topographic map of Europe and the Mediterranean region indicating the locations of the main Tertiary-Quaternary volcanic fields discussed in this review. Also shown is the location of a zone of high-velocitymantle within the mantle Transition Zone (500-600 kin) beneath Europe (from Piromallo et al. 2001), which may be a region of subducted slabs at the base of the upper mantle. The ages of such slabs cannot be constrained; however, they most probably represent remnants of subducted Tethyan oceanic lithosphere. The variable size of the red asterisks, marking the location of individual volcanic fields, indicates, schematically,the relative volume of magmatism. TTZ, Tomquist-Teisseyre Zone.

basins in the Mediterranean region and the collision of the Alpine orogen with the southern passive margin of Europe. Compressional deformation of the lithosphere within Western and Central Europe occurred as a response to the collisional coupling of the Alpine and Pyrenean orogens with their forelands (Ziegler et al. 1995; Dbzes et al. 2004). Throughout the Tertiary there was a gradual shift of compressional tectonic activity away from the foreland of the Carpathians and Eastern Alps to the foreland of the Central and Western Alps, partly as a consequence of dextral translation between the converging blocks during the late Eocene to Pliocene. Stresses related to the collision of Iberia and Europe interfered with stresses transmitted from the Alpine collision front (D~zes et al. 2004); these stresses played an important role in the Eocene reactivation of Permo-Carboniferous fracture systems and the localization of the Cenozoic rifts (e.g. Rhine Graben). Convergence rates between Africa and Europe decreased rapidly during the late Cretaceous and Palaeocene (67-55 Ma) as the African and European plates became mechanically coupled (Rosenbaum et al. 2002). During the late Palaeocene (61-55 Ma) a pulse of intense intra-plate compression affected Western and Central Europe, the East European Craton and North Africa (D~zes et al. 2004; Fig. 3). Compressional stresses

exerted by the evolving Alpine and Pyrenean orogenic belts caused lithospheric buckling and basin inversion up to 1700 km north of the orogenic fronts. This deformation was accompanied by local intrusion of small-degree partial melts (e.g. melilitites and nephelinites) in the Massif Central, Vosges, Black Forest, Rhenish Massif and Bohemian Massif. During the early Eocene (c. 52 Ma) the convergence rate between Africa and Europe gradually increased, followed by a decrease in the early Miocene (c. 19 Ma) (Rosenbaum et al. 2002). Scattered volcanic activity occurred during the early and mid-Eocene in the Massif Central (Michon & Merle 2001), the Rhenish Massif (Lippolt 1982) and the Bohemian Massif (Ulrych & Pivec 1997). During the late Eocene extension initiated along the Massif Central, Bresse and Rhine grabens by transtensional reactivation of older Permo-Carboniferous fracture systems in a northerly directed compressional stress field (Dbzes et al. 2004). At the Eocene-Oligocene boundary, convergence of the West Alpine orogenic wedge with the European foreland changed to a NW direction (Ceriani et al. 2001), coincident with the detachment of a southerly subducted lithospheric slab beneath the Central and Eastern Alps and associated isostatic rebound of the European foreland lithosphere (von Blanckenburg & Davies 1995).

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Fig. 2. Relationship of the main Tertiary-Quaternaryvolcanic fields in Western and Central Europe to zones of uplifted basement, major rift systems and the Variscan basement terranes of Central Europe. AfterWilson & Dowries (1991) and Wilson & Patterson (2001). PBF, Pas de Bray Fault; SHF, Sillon Houiller Fault; V/BF Dome, Vosges-Black Forest dome; RG, Rhine Graben; RH, Rheno-Hercynian terrane; S, Saxo-Thuringian terrane; M, Moldanubian terrane; NHD, North Hessian Depression.

North-directed compressional stresses from the Pyrenees, combined with the forces exerted by collisional tectonics in the Central Alps, induced the main Oligocene extensional stage of the ECRIS. The Pyrenean component of compressive stress relaxed during the late Oligocene (D~zes et al. 2004). During the Oligocene the Rhine Graben propagated northwards, bifurcating into the Ruhr and Leine grabens (Fig. 2); rift propagation was associated with an intensification of volcanic activity in the Rhenish Massif. In the northern parts of the Massif Central rifting was accompanied from the late Oligocene by scattered volcanic activity (Michon & Merle 2001). In contrast, subsidence of the Eger Graben within the northern part of the Bohemian Massif commenced only towards the end of its Oligocene magmatic phase. During the Miocene extension continued in the Rhine, Ruhr and Leine grabens (Schumacher 2002); their triple junction (Fig. 2) was gradually uplifted and became the focus of increased volcanic activity (Sissingh 2003). Uplift of the Vosges-Black Forest dome commenced between 19 and 20 Ma; this has been attributed to lithospheric flexuring in the Alpine foreland (Schureacher 2002; D~zes et al. 2004). Minor volcanic activity within the Upper Rhine Graben, including its rifted flanks, was associated with this phase of uplift from 18 to 7 Ma (Jung 1999). Uplift and northward tilting of the Massif Central also commenced during the early Miocene, followed by a rapid increase in volcanic activity during the mid- and late Miocene (Fig. 3; Michon & Merle 2001). Minor compressional deformation of the European lithosphere also occurred during the late Miocene-early Pliocene and in Pliocene-Quaternary times (Fig. 3; D~zes et al. 2004). Extension continues to the present day along the Rhine and Ruhr grabens, whereas subsidence of the Massif Central grabens ceased during the Miocene. Within the Rhenish Massif, volcanic activity

shifted to the Eifel region during the Pliocene and Quaternary, coinciding with an acceleration of uplift (Garcia-Castellanos et al. 2000). Uplift of the northern Bohemian Massif, which initiated during the early Miocene, continued throughout PlioQuaternary times, accompanied by renewed volcanic activity (Ulrych & Pivec 1997; Michon & Merle 2001). This uplift has been attributed to lithospheric flexuring (Ziegler & D~zes 2006). In the northern Massif Central volcanic activity resumed at the beginning of the Pliocene (peaking between 4 and 1 Ma), whereas in the south a second peak of activity occurred between 3.5 and 0.5Ma (Michon & Merle 2001); volcanism was accompanied by renewed uplift. On a regional scale there appears to be a broad correlation between the timing of magmatic activity within the northern foreland of the Alps and changes in the regional stress field (Fig. 3). A detailed compilation of the available geochronological data for the Massif Central suggests that the main volcanic phases may be associated with periods of compressional stress relaxation in the foreland of the Alpine orogenic belt (Wilson & Patterson 2001). Magmas must rise through the crust and upper part of the lithospheric mantle through fracture systems; consequently, it is possible that the distribution of Cenozoic magmatism within Europe could be related to reactivation of pre-existing lithospheric discontinuities (e.g. Permo-Carboniferous sutures and fault systems) in response to changes in the regional stress field. The maximum horizontal stress direction within Western and Central Europe rotated from NNE-SSW to NNW-SSE during the Late Eocene-Early Oligocene to N W - S E in the Late Oligocene (Schreiber & Rotsch 1998). The orientations of linear chains of volcanic necks and scoria cones commonly reflect the orientation of the contemporary stress field. In most areas Cenozoic rifting initiated earlier than the main phase of magmatic activity and is frequently offset spatially from both

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M. WILSON & H. DOWNES at the boundary between the Saxo-Thuringian and Moldanubian Variscan basement terranes (Fig. 2).

Age and characteristics of the volcanic fields Rhenish Massif

Fig. 3. Chronologyof Tertiary-Quaternary volcanism within Europe and the Mediterranean region in relation to major periods of lithospheric extension, basement uplift and phases of Alpine compression. (For data sources see text.) It should be noted that the time scale from 65 to 25 Ma is compressed. CS, Cesk6 Stredohori; DH, Doupovsk6 Hory; TS, Tyrrhenian Sea.

magmatism and areas of basement uplift. In a number of areas (e.g. Rhine Graben; Massif Central), however, there is evidence for minor early Tertiary magmatic activity, which pre-dates the onset of rifting (Fig. 3). The largest Cenozoic rift within Central Europe, the Rhine Graben, is about 300 km long and 35-40 km wide. It trends N N E - S S W oblique to the NE-trending structural grain of the Variscan crystalline basement of Europe (Moldanubian and Saxo-Thuringian terranes; Fig. 2). The northern end of the rift is located to the SE of the Rhenish Massif at the boundary between the Saxo-Thuringian and Rheno-Hercynian Variscan basement terranes. Extension and subsidence occurred mainly between Oligocene (c. 35 Ma) and Miocene times. Subsidence in the southern part of the rift was interrupted by basement uplift and the magmatic activity of the Kaiserstuhl volcano (Keller et al. 1990). It is notable that although the Rhine Graben is the most highly extended part of the European rift system, it is largely non-magmatic for much of its length, suggesting that lithospheric extension and decompression-induced partial melting of the upper mantle is not necessarily the main cause of magma generation within the European volcanic province. The Miocene volcanic complex of the Vogelsberg is located at the northern end of the Rhine Graben where it splits into two branches

Cenozoic volcanism in central Germany is concentrated in a 350km long, east-west-trending zone extending from the Eifel in the west to the Rh6n-Heldberg area in the east (Fig. 2; Wedepohl & Baumann 1999). Volcanic activity started during the Eocene and Oligocene in both the eastern and western extremities of the belt (e.g. Hocheifel (45-24 Ma); Fekiacova et al. 2006). The climax of volcanic activity occurred between 16 and 18 Ma in the Vogelsberg volcano in the central part of the belt. Volcanism ceased about 5 Ma ago followed by a Quaternary resurgence of activity in the East and West Eifel. Most of the volcanic rocks are relatively primitive alkali olivine basalts, nepheline basanites and olivine nephelinites; quartz tholeiites, however, occur in the Vogelsberg and North Hessian Depression. Locally (e.g. Eifel, Siebengebirge, Westerwald and Rh6n) extreme differentiation of the parental mafic magmas produced phonolites and trachytes. In the Eifel district predominantly potassic magmas, including leucitites, were erupted during the Quaternary. The Miocene Vogelsberg is a shield volcano (Bogaard & W6rner 2003), which erupted basanites, alkali basalts, quartz tholeiites and limited volumes of highly evolved magmas ranging from hawaiite to trachyte. It has an eruptive volume of c. 600 km 3, probably making it the largest volcanic centre within the European volcanic province (Jung & Masberg 1998). The volcano is located to the east of the Rhenish Massif, close to the triple junction of the Rhine, Ruhr and Leine grabens (Fig. 2). Volcanism commenced in the Early Miocene (c. 22-23 Ma); however, the main phase of activity began at c. 18 Ma and peaked between 16 and 17 Ma. The Rh6n and Northern Hessian Depression volcanic fields are closely related, both spatially and temporally, to the Vogelsberg and erupted a similar range of magma types (Wedepohl et al. 1994; Jung & Hoernes 2000). In the East and West Eifel volcanic fields about 300, typically small-volume, eruptions occurred from monogenetic centres between 700 and 10.8 ka By (Schmincke et al. 1983; W6rner et al. 1986; Schmincke 2006), associated with about 250 m of uplift. The volume of magma erupted is small (about 15 kin3), but the actual volume of magma generated at mantle depths must have been significantly greater (70-100 kin3; G. W6rner, pers. comm.). Two geochemically, spatially and temporally distinct groups of sodic-potassic alkaline volcanic rocks were erupted in the East Eifel; in the NW these include nephelinites, leucitites and their differentiates (erupted >400 ka), whereas in the SE basanites and their differentiates predominate (erupted between 400 and 10 ka; Lippolt et al. 1990). The West Eifel volcanic field covers an area ofc. 600 km 2 and comprises about 240 volcanic centres; these erupted predominantly leucitites and nephelinites with subordinate basanites. At c. 12.9 ka there was a major Plinian eruption of the Laacher See volcano, which produced c. 6.3 km 3 of phonolitic tephra, causing a major environmental impact (Litt et al. 2003). Volcanic activity in the Westerwald started in the Oligocene with the eruption of basalts and trachytes (Schreiber & Rotsch 1998); the main phase of activity had ended by 20 Ma, although there were short periods of reactivation, with the eruption of basalts, in the Miocene and Pleistocene (Fuhrmann & Lippolt 1990). Volcanic activity appears to have been synchronous with minor uplift of the Rhenish Massif, which commenced at the end of the Oligocene, strengthened during the Quaternary and continues to the present day (Meyer et al. 1983). On the basis of palaeomagnetic data, Schreiber & Rotsch (1998) proposed that the northeastern part of the Rhenish Massif has rotated clockwise by 10-16 ~ since the late Oligocene, associated

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151

Fig. 4. Summary of the changing distribution and intensity of Tertiary-Quaternary magmatism throughout Europe and the Mediterranean region.

(See text for details). with a system of dextral strike-slip faults. The Quaternary West and East Eifel volcanic fields are located in the non-rotated western Rhenish Massif block. Block rotation is considered to have initiated in the Late Oligocene as a result of small changes in the direction of the maximum horizontal compressive stress from NNW to NW. Southern Germany

Volcanism in southern Germany is confined to a few small regions including the Urach and Hegau provinces to the east of the Rhine Graben, the rift flanks of the Rhine Graben and the Kaiserstuhl volcano, which is axially located within the graben where it bisects the Vosges-Black Forest dome (Keller et al. 1990; Glahn et al. 1992; Fig. 2). On the basis of K - A r dating and stratigraphic constraints it is likely that the main phase of volcanic activity ranges from about 45 to 15 Ma (Keller et al. 2002). However, Keller et al. (2002) have recently dated amphibole phenocrysts from an olivine melilitite dyke (Trois Epis) in the Vosges at 6 0 . 9 _ 0.6 Ma; this suggests that magmatism began some 15 Ma before the onset of graben formation, contemporaneous with the onset of major horizontal crustal shortening in the Western Alps (Gebauer 1999) and a major phase of foreland compression (Ziegler et al. 1995).

Magmatism occurs as a series of dykes, plugs or necks and diatreme pipes, concentrated in two sectors: (1) the Vosges-Black Forest Dome, which is the location of the maximum updoming of the Rhine Graben rift flanks, and the only axially located volcano in the graben (Kaiserstuhl); (2) in the north between Heidelberg and Frankfurt, mostly in the crystalline basement of the Odenwald. Scattered volcanic centres also occur along the flanks of the rift (e.g. Mahlberg). The primary magmas are highly undersaturated mafic alkaline types, predominantly olivine nephelinites and olivine melilitites with high Mg-numbers, Ni and Cr contents. The Miocene Kaiserstuhl complex (15-18 Ma) is an alkaline carbonatite complex that also includes potassic magmas (Schleicher et al. 1990). The Urach province is an olivine melilitite diatreme field with more than 350 individual volcanic necks for which K - A t ages range from 11 to 17 Ma (Lippolt et al. 1973). Most of the diatremes are composed of tufts of olivine melilitite and olivine melilite nephelinite. The main period of activity is in the mid-Miocene, from 16 to 17 Ma. There does not appear to be any correlation between fault tectonics and the location of the diatremes, although the majority are located in a synclinal structure, the 'Urach Trough', in which subsidence has occurred since mid-Triassic times. The Hegau volcanic field, some 100 km further south, has a greater variety of magmatic rock types including olivine

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melilitite, olivine-nepheline melilitite and phonolite, diatremefacies pyroclastic deposits and carbonatites. K - A r ages range from 7 to 15 Ma, with emplacement of olivine melilitites between 8.5 and 12 Ma.

M a s s i f Central

Most of the alkaline magmatic activity in France has occurred within the uplifted Variscan basement of the Massif Central, with subordinate amounts further south in the Languedoc (Wilson & Patterson 2001). The Massif is divided by a NNE-SSW-trending late Variscan strike-slip fault, the Sillon Houiller, which represents a major discontinuity between two distinct lithospheric domains (Fig. 2; Alard et al. 1996). The eastern Vosges-Auvergne domain is distinguished from the western Limousin domain by thinner crust (<29 km) overlying upper mantle with low seismic velocity (Nicolas et al. 1987; Granet et al. 1995; Zeyen et al. 1997). The characteristics of the Vosges-Auvergne domain have been ascribed to lithospheric thinning above an upwelling asthenospheric mantle diapir. Cenozoic volcanic activity is concentrated in two main areas (Wilson & Patterson 2001). A chain of volcanoes extends in an approximately north-south direction along the western edge of the Limagne graben. From north to south these are the Cha~ne des Puys, the large stratovolcano of Mont Dore, the basaltic plateau of Crzallier, the central volcano of Cantal and the Aubrac mountains. Southeast of the Limagne graben, surrounding the Le Puy basin, the volcanic areas of Velay, Dev~s and Vivarais are the locations of some of the youngest activity. Many of the volcanic fields show pronounced NW-SE-trending alignments of eruptive centres (e.g. Aubrac, Coirons, Velay, Dev~s), subparallel to a diffuse system of normal faults and graben segments that may represent the continuation of the north-south-trending Limagne graben to the SE. In contrast, the volcanic fields of the Cha~ne des Puys, in the northern part of the Massif, and Escandorgue, in the extreme south, consist of a large number of volcanic cones aligned in a north-south direction. The changing orientations of these volcanic lineaments reflect the changing orientation of the regional stress field within Europe throughout the Tertiary. Magmatism peaked between the Late Miocene and the Pliocene (10-5 Ma) with widespread eruption of plateau basalts (alkali basalts and basanites) across the province (Dev~s, C~zallier, Coirons, Aubrac and Escandorgue) and the formation of the Cantal central volcano (400km3; Downes 1983). Significant volumes of alkali basalt were associated with this phase of activity and the volume of differentiated lavas increased substantially. A second peak of volcanic activity, between 4 and 2 Ma, resulted in the formation of the Mont Dore volcano (200 km3; Briot et al. 1991), and of the Cha~ne de la Sioule, Devbs and Escandorgue provinces oriented along NE-SW, NW-SE and north-south fractures respectively. Important uplift along the SE border of the Massif coincided with this magmatic peak (Derruau 1971).

Czech R e p u b l i c - S W P o l a n d

Paralleling the Alpine tectonic front, although located much farther north, lies the Eger Graben in the Czech Republic (Fig. 2). Its location, parallel to the Saxo-Thuringian-Moldanubian Variscan suture zone, and indeed utilizing the suture as its southern margin, may be a response to lithospheric doming of the Bohemian Massif. Ulrych & Pivec (1997) have recognized three magmatic phases on the basis of K - A r data: (1) pre-rift melilititenephelinite magmatism (79-49 Ma) on the rift flanks; (2) synrift (42-16 Ma) bimodal basanite-trachyte and nephelinitephonolite magmatism, followed by a second phase of activity from 13 to 4Ma; (3) late-stage melilitite-olivine nephelinite activity (2-0.26 Ma) in the westernmost part of the rift near Cheb.

Cenozoic volcanism extends from the German-Czech border eastwards to Upper Silesia in Poland and northern Moravia (Fig. 2). The magmatic rocks range in composition from melilitites, basanites, alkali basalts and carbonatites to trachytes (Kopecky 1966; Blusztajn & Hart 1989). Local occurrences ofbimodal alkaline magmatism extend for c. 40 km both north and south of the main rift faults; melilite nephelinite magmas are characteristic of the SE border fault. The two main volcanic centres within the graben are the Doupovsk6 Hory and Cesk~ Stredohori complexes. Doupovsk~ Hory, in the west, is a stratovolcano c. 30 km in diameter, containing a mixture of sodic and potassic mafic magmas and their differentiates (Kopecky 1966). The volcano-sedimentary complex of Cesk~ Stredohori, further to the east, is the most significant region of intra-plate alkaline volcanism, characterized by lava flows, tufts, plugs, small sub-volcanic intrusions and dykes of basanite-trachyte with minor nephelinite-phonolite series magmas (Ulrych & Pivec 1997; Ulrych et al. 2001). On the basis of available K - A t data (Shrbeny 1995; Wilson & Rosenbaum, unpubl, data), the main phase of volcanism in the Cesk6 Stredohori was from Late Eocene to Mid-Miocene (42-15Ma); much smaller volumes (<1%) of magmatism occurred from 13 to 9 Ma (Mid-Late Miocene), and there was a resurgence of activity in Pliocene-Quaternary times.

Pannonian Basin

The geodynamic evolution of the Pannonian Basin and its relationship to the Alpine-Carpathian orogenic belt has been the focus of a number of recent studies (e.g. Cloetingh & Lankreijer 2001, and references therein). The development of the Carpathian fold belt and associated back-arc extensional basins is linked to the Mesozoic-Cenozoic collision of Eurasia with a number of continental microplates including the Italo-Dinaride block. Southward subduction of the European plate under the northern margin of the Pannonian Basin during Palaeocene-Eocene times is evidenced by a chain of calc-alkaline volcanoes. Calc-alkaline magmatic activity started in the Eocene and culminated in the Miocene (Downes et al. 1995a). Extension ended in Late Miocene times and was followed by a phase of PlioPleistocene alkali basaltic volcanism. The Neogene alkaline volcanism of the Carpathian-Pannonian region (Fig. 1) has been extensively studied (e.g. Szab6 et al. 1992; Embey-Isztin et al. 1993; Downes et al. 1995b; Embey-Isztin & Dobosi 1995; Vaselli et al. 1995). P~cskay et al. (1995) presented a comprehensive review of the timing of magmatic activity. Alkaline magmatism (including both alkali basalts and rare potassic-ultrapotassic lavas) occurred sporadically from 17 to 0.5 Ma, partly contemporaneously with subduction-related magmatism in some areas (20-0.2Ma). Within the Inner Carpathians calc-alkaline magmatism migrated eastwards in time (Eocene-Miocene in the Northern Carpathians to PlioceneQuaternary in the Eastern Carpathians of Romania; Prcskay et al. 1995). The alkaline volcanism is also oldest in the western part of the Pannonian Basin but youngest in the central regions (Embey-Isztin et al. 1993). In the eastern Transylvanian Basin late Tertiary-Quaternary volcanism associated with both extension and subduction occurred simultaneously (Downes et al. 1995b). The alkaline volcanic rocks exhibit two age groups: an older phase at 17-7 Ma and a younger phase from 6 to 0.5 Ma. Potassic lavas (shoshonites) range in age from 15 to 1 Ma. Episodic Quaternary volcanism has occurred in the West Carpathians, East Carpathians, Persani Mountains and Apuseni Mountains.

Dinarides

Much less attention has been focused on the southern part of the Pannonian Basin and its links with the Dinaride tectonic zone to

TERTIARY-QUATERNARY MAGMATISMIN EUROPE the south (Fig. 1). This province extends from eastern Slovenia, through northern Croatia and northern Bosnia into Serbia. Lateral movements of several hundred kilometres along the Periadriatic Lineament in the Alps, Oligocene and Miocene magmatism, and Neogene block rotations provide evidence for substantial mobility in this transitional region (Sachsenhofer et al. 2001). The northern Dinarides were created by the collision of the northeastern parts of the Apulian plate and the southern margin of the Eurasian plate, which commenced in the Late JurassicEarly Cretaceous (Tari & Pamic 1998); during this stage ophiolites were obducted onto the Apulian plate. Uplift of the Dinarides occurred during the Late Eocene-Early Oligocene associated with andesitic-dacitic volcanism and pyroclastic activity (3222 Ma). Extension started in the Early Miocene but terminated by the end of the Mid-Miocene. High levels of volcanic activity from 16 to 12 Ma resulted in a suite of basalts, andesites, dacites and rhyolites. Eruption of alkali basalts occurred from 10 to 8.5 Ma (Pamic et al. 1995). In Serbia there was a distinct phase of alkaline mafic magmatism (mainly basanites) during the Palaeocene-Eocene (6240 Ma; Jovanovic et al. 2001; Cvetkovic et al. 2004). Magmatism occurred after the termination of Late Cretaceous subduction and cessation of calc-alkaline volcanic activity, and mainly occurs in the former arc and forearc regions. The alkaline mafic magmas, however, display no subduction-related fingerprint in their geochemistry. It is possible that the magmatism was related to slab-break-off of eastward-subducted oceanic lithosphere.

Bulgaria

In Bulgaria there is a 250 krn long, north-south-trending magmatic province, ranging in composition from potassic basanite to alkaline lamprophyre (camptonite), which is closely associated with extensional tectonics. In the north (Moesian platform) the magmatism is of Early Miocene age, whereas in the south (Rhodope Zone) it is Eocene to Oligocene in age. Monogenetic volcanoes and domes dominate the northern part, whereas sills and dyke swarms are more common in the central and southern areas (Vaselli et al. 1997). As in Serbia, the alkaline magmatism post-dates an earlier (Early Oligocene) phase of calc-alkaline magmatism (Marchev et al. 1998).

Italy, Sicily a n d Sardinia

Intra-plate alkaline magmatism, which is geochemically distinct from the orogenic magmatism of the Roman Volcanic Province (Wilson & Bianchini 1999), occurs mainly in Sicily (the Iblean Plateau and Mt. Etna), the Veneto province in the Po Plain region of northern Italy, in central Italy at Monte Vulture, and in Sardinia (Fig. 1). The extension-related, anorogenic, magmatism (5-0.1 Ma) in Sardinia post-dates an earlier phase of subduction-related magmatism (32-13 Ma), and is characterized by the eruption of both subalkaline and alkaline lavas (including primitive and more differentiated magma types; Beccaluva et al. 1977, 1987; Rutter 1987). Although there is no clear temporal trend, the eruption of subalkaline lavas appears to have occurred preferentially during a short period with a climax at about 3.5-3 Ma (Montanini & Villa 1993), followed by more widespread eruptions of alkaline magmas. The Plio-Pleistocene subalkaline basic-acid magmatism of Mt. Arci (Cioni et al. 1982; Dostal et al. 1982) appears to be transitional in chemistry between that of the earlier Oligo-Miocene subductionrelated cycle and the younger alkaline magmas. This may reflect the presence of an inherited 'subduction-related' component in the mantle source of the magmas. Gasperini et al. (2000) have suggested, on the basis of the S r - N d - H f isotope systematics of Pleistocene basalts from Logudoro, that their mantle source may

153

contain a recycled crustal component from a subducted oceanic plateau. Several phases of Cenozoic volcanic activity have been recognized in the Iblean area of southern Sicily. Miocene and PlioPleistocene volcanic rocks occur in the northern part of the Iblean platform, toward the Apennine-Maghrebian compressional front (Beccaluva et al. 1998). The Miocene magmatic phase is predominantly alkaline in composition, characterized by a low melt production rate. After a period of low-level activity from about 6.5 to 4 Ma, a new cycle, with a higher melt production rate, initiated in the Early to Mid-Pliocene, lasting until the Early Pleistocene. This cycle is characterized by an abrupt compositional change to magmas of predominantly tholeiitic affinity. The melt production rate gradually decreased after the climax of volcanic activity and magmas of various geochemical affinities were erupted simultaneously. Sporadic eruptions of highly undersaturated alkaline lavas occurred in the northernmost part of the plateau, towards Mount Etna, during the Early Pleistocene. It is difficult, however, to study the migration of volcanism from the Iblean area to the Etna area, as the transition zone is obscured by recent sediments. Mount Etna is the largest active volcano in Europe with an estimated volume of 500-600 km 3. It is polygenetic, with several distinct stages to its evolution. The oldest volcanic products (c. 600 ka BI~; Gillot et al. 1994) are of tholeiitic affinity and crop out sporadically at considerable distances from the present volcanic focus. Alkaline mafic magmatism commenced around 220 ka Bp (Condomines et al. 1982; Gillot et al. 1994), and for a short period tholeiitic and alkaline magmas erupted simultaneously. More recent magmatism has been entirely alkaline and the erupted lavas have become progressively more differentiated with time, consistent with the development of a high-level magma chamber system beneath the volcano (Clocchiatti et al. 1988; Tanguy et al. 1997; Corsaro & Pompilio 2004). The volcanic activity of Etna has been attributed to differential roll-back of a slab of subducted oceanic lithosphere and the formation of a slab window through which upwelling of deeper upper mantle material has occurred (Armienti et al. 2004). In contrast, Montelli et al. (2004) have proposed the existence of a deep mantle plume beneath Etna based upon a new seismic tomographic model. During the past 10 Ma there has also been extensive magmatism offshore in the Sicily Channel, including the islands of Pantelleria and Linosa. The eruptive products range in composition from nepheline basanite to tholeiitic basalt and their differentiates (Beccaluva et al. 1981; Calanchi et al. 1989). The Late Pleistocene Mt. Vulture stratovolcano has an unusual geodynamic setting at the intersection of NE-SW- and N W SE-trending fault systems at the easternmost border of the Apennine thrust front (Beccaluva et al. 2002). Its eruptive products include pyroclastic deposits and basanitic and melilititic lavas; carbonatites have also been reported. Localized occurrences of Late Palaeocene-Late Oligocene extension-related (anorogenic) alkali basalts, basanites and subordinate transitional basalts occur in the Veneto region of northern Italy (Siena & Coltorti 1989; De Vecchi & Sedea 1995; Milani 1996). The age of the magmatism is based primarily upon stratigraphical constraints and may extend into the Miocene.

Spain

During the Late Miocene-Quaternary extension-related (anorogenic) alkaline magmatism occurred in central Spain in the Calatrava province (Lrpez-Ruiz et al. 1993; Cebrifi & Lrpez-Ruiz 1995), in northeastern Spain near Olot, just south of the Pyrenees (Cebri~i et al. 2000), to the NW of Cartagena (Tallante) within the Betic Cordillera and at Cofrentes and Columbretes Island (Wilson & Bianchini 1999; Fig. 2). Extension post-dated the main phase of Alpine compression, giving rise to a series of basins, some of

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which were magmatically active (Doblas & Oyarzun 1989). In the Olot area Miocene-Quaternary volcanic vents are associated with NW-SE- and NE-SW-trending fault systems. The oldest magmatism follows the trend of the NE-SW-oriented Cenozoic European rift system, whereas the youngest (Garrotxa) magmatism follows N W - S E fault trends related to Alpine convergence in the Pyrenees.

N o r t h A f r i c a n m a r g i n - A l b o r a n Sea

Tertiary-Quaternary magmatic activity also occurs along the whole western Mediterranean margin of North Africa (Fig. 1) from Morocco to Tunisia. This may be divided into two stages: Late Cretaceous-Mid-Eocene (c. 42 Ma) and Early MioceneRecent (Wilson & Guiraud 1998; Wilson & Bianchini 1999). The bulk of the magmatism appears to post-date a major phase of Aquitainian-Burdigalian (c. 20 Ma) compression induced by the Alpine collision, and has many similarities to that of the Betic Cordillera of southern Spain (Hernandez et al. 1987; Lonergan & White 1997). The earliest magmatic rocks have calc-alkaline affinities, becoming progressively more alkaline with time. Given the complex geodynamic setting of the western Mediterranean, and the limited amount of geochemical and geochronological data on the magmatic rocks in both the Betic-Rif and Maghrebide belts, it is difficult to constrain the precise tectonic setting in which the magmas were generated. Some researchers consider that the Miocene calc-alkaline volcanic episode post-dated active subduction (e.g. Hernandez & Lepvrier 1979), whereas others link it directly to subduction (Tricart et al. 1994; Lonergan & White 1997) or to slab detachment in the Tertiary (Moni6 et al. 1992; Zeck et al. 1992). Duggen et al. (2003) have demonstrated that between 6.3 and 4.8 Ma there was a marked change in the geochemistry of the magmatism in the western Mediterranean from subduction-related to intra-plate, which was synchronous with the Messinian salinity crisis. They relate this to westward roll-back of an eastwarddipping slab of Tethyan oceanic lithosphere, delamination of a block of continental lithosphere beneath the Alboran Sea, and associated asthenospheric upwelling. The younger (6.30.65 Ma) volcanic rocks are alkali basalts and basanites plus differentiates. The change from calc-alkaline and shoshonitic magmatism to intra-plate alkaline magmatism was transitional; the youngest shoshonites are dated at 4.8 Ma. The alkaline mafic rocks are similar to ocean island alkali basalts (OIB) from the Canary Islands, and to those from elsewhere in the European volcanic province.

Western Turkey

The Western Anatolian volcanic province of Turkey (Fig. 1) is located at the eastern end of the Aegean arc, which results from the northward subduction of the African plate beneath the Aegean. Calc-alkaline volcanic activity commenced in the Late Oligocene-Early Miocene, followed by alkali basaltic volcanism from Late Miocene to Recent times. This change in the style of the volcanism has been attributed by some workers to a change in the regional stress field from north-south compression to north-south extension (Yilmaz 1990; Gtileq 1991). Seyitoglu & Scott (1992), however, considered that the transition to north-south extensional tectonics actually commenced much earlier in the latest Oligocene-Early Miocene. Seyitoglu et al. (1997) have demonstrated that within the youngest volcanic sequence there is a change from potassic magmatism in the Miocene to more sodic (anorogenic) alkaline magmatism in the Quaternary. The geochemical characteristics of the potassic magmas are inferred to reflect the presence of an inherited subduction-modified component in their mantle source.

Large volumes of trachyandesitic-dacitic lava flows and pyroclastic deposits of Miocene age are associated with small volumes of alkali basalt lava flows in the Galatia volcanic province of NW Central Anatolia (Wilson et al. 1997). The volcanism post-dates continental collision, occurring in a transtensional tectonic setting associated with movement along the North Anatolian Fault zone. Alkali basalts were erupted during two distinct time periods, in the Early Miocene (1719Ma) and Late Miocene (<10Ma); the Early Miocene basalts have geochemical characteristics that suggest the involvement of a subduction-modified mantle source component in their petrogenesis, whereas the late Miocene basalts are identical in their geochemical characteristics to anorogenic alkali basalts erupted throughout Western and Central Europe and the Mediterranean domain.

Temporal distribution of the volcanism From the Late Eocene (c. 40 Ma) to the present day, anorogenic magmatic activity occurred throughout the European and African margins of the Tethyan collision zone and within the Mediterranean region (Wilson & Bianchini 1999). Major phases of volcanism occurred in the Early Miocene and in the Late Miocene-Pliocene, although activity was not widespread until the Late Miocene (Figs 3 and 4). Magmatism locally continued to a few thousand years Be (e.g. Cha~ne des Puys, Massif Central; Laacher See, Eifel province of the Rhenish Massif). This major regional volcanic flare-up may reflect a fundamental reorganization of the convection system within the upper mantle intimately associated with the Alpine collision. In a number of areas there is also evidence for an earlier phase of Palaeocene-Eocene volcanism; for example, in the Massif Central and the Bohemian Massif (Horn et al. 1972; Baranyi et al. 1976; Downes 1987; Malkovsky 1987), eastern Serbia and southern Romania (Downes et al. 1995b; Jovanovic et al. 2001), the Rhine Graben (Keller et al. 2002), the Pannonian Basin (Prcskay et al. 1995), and the Veneto province (De Vecci & Sedea 1995). A number of workers (e.g. Lippolt et al. 1974) have suggested that magmatic activity actually started in the Late Cretaceous; these older ages are, however, suspect as they are based on K - A r age determinations on altered whole-rocks. Keller et al. (2002) have recently demonstrated that supposed c. 85 Ma nephelinites from the Rhine Graben (dated originally by K-Ar) are actually Palaeocene in age (c. 61 Ma) when re-dated by A r - A r geochronology on amphibole phenocryst mineral separates. By the early Eocene, volcanic activity had extended into the northern Rhine Graben region and in the late Eocene commenced in the Rhenish Massif (Lippolt 1982). However, the volume of volcanic rocks at 40 Ma was still small compared with the volume subsequently emplaced in Neogene-Recent times. In the French Massif Central magmatism continued in the Bourgogne region and initiated in the Forez Graben in the NE part of the massif (Patterson 1996). The Veneto district of northern Italy also became magmatically active at this time (De Vecchi & Sedea 1995). Dykes were intruded along the flanks of the Rhine Graben (Lippolt 1982; Wilson & Keller, unpubl, data). The onset of major volcanism occurred in the Oligocene in the Bohemian Massif (Doupovsk6 Hory and Cesk6 Stredohori; Ulrych & Pivec 1997), the Massif Central (Cantal; Patterson 1996), the Rhenish Massif (Siebengebirge, Westerwald and Rhrn subprovinces; Lippolt 1982; Wedepohl et al. 1994; Jung & Hoernes 2000), the Odenwald in the Rhine Graben region (Lippolt 1982), and in southern Bulgaria (Marchev et al. 1998). Magmatic activity was intimately associated with regional rifting (Ziegler 1992). Magmatic activity continued in the Veneto district throughout the Oligocene (De Vecci & Sedea 1995).

TERTIARY-QUATERNARY MAGMATISMIN EUROPE The Neogene (Miocene-Pliocene) sub-period was characterized by a peak in volcanic activity throughout Europe. Magmatism occurred in southern and central Germany (Urach, Hegau, North Hessia, Upper Palatinate, Vogelsberg; Lippolt et al. 1973; Jung & Masberg 1998; Bogaard & W6rner 2003), Spain (Calatrava, Olot; Lrpez-Ruiz et aI. 1993; Cebrifi & L6pez-Ruiz 1995, 2000), the French Massif Central (Patterson 1996; Wilson & Patterson 2001), southern Italy (Sicily Channel, Iblean Plateau in Sicily; Wilson & Bianchini 1999), Burgenland (eastern Austria), central Slovakia and northern Bulgaria. Some of the largest volcanic edifices in the province (e.g. Cantal, Massif Central) were formed at this time. Concurrently, activity in the Siebengebirge, Westerwald and Rh6n regions of northern Germany abated, and the focus of volcanic activity moved eastwards with the eruption of quartz tholeiites in the North Hessian Depression. Soon afterwards, eruptions began to construct the Vogelsberg and Kaiserstuhl volcanic complexes (Lippolt 1982; Schleicher et al. 1990; Jung & Masberg 1998). Magmatic activity continued, and became more widespread, in the Bohemian Massif throughout the Miocene (Kopecky 1966; Ulrych & Pivec 1997). During the Pliocene the Massif Central, the Calatrava province of central Spain and the Bohemian Massif remained volcanically active, whereas in the Rhenish Massif and south German fields magmatic activity diminished considerably (Patterson 1996; L6pez-Ruiz et al. 1993; Wedepohl et al. 1994; Ulrych & Pivec 1997). Alkali basaltic volcanism continued throughout the Pliocene and into the Quaternary in the Massif Central (e.g. Cha~ne des Puys, Vivarais), Italy (Etna), Sardinia, central Germany (Eifel), Spain (Tallente, Olot, Calatrava), Hungary (Balaton Highlands, Little Hungarian Plain), Romania (Banat and Persani Mountains), eastern Austria (Graz Basin), central and southern Slovakia, SW Poland and the Czech Republic (Eger Graben). Quaternary magmatism is not known in Calatrava or the south German volcanic fields (Wedepohl et al. 1994; Crbria & L6pez-Ruiz 1995). Important Pleistocene activity in the Bohemian Massif occurs in west Bohemia and northern Moravia (Wimmenauer 1974). Quaternary eruptions in the Rhenish Massif are concentrated in the Eifel volcanic field on the western margin of the massif, and are characterized by a high proportion of potassic magmatic rocks (Mertes & Schmincke 1985). In the Massif Central, the most northerly volcanic fields (Cha~ne des Puys and Mont Dore) were active alongside those in the extreme east (Vivarais) and extreme south (Escandorgue and Languedoc) (Patterson 1996). In all cases, the volume of magmatism is subordinate to that of the Neogene phase of activity, suggesting a general decline in the amount of melt being produced. Volcanic activity continued to c. 4000 a BP in the Cha~ne des Puys. There does not appear to be any evidence of a geographical progression in the main locus of magmatism with time (Fig. 4). This suggests that magmatic activity is not occurring in response to a single process, such as lithospheric extension, but rather in response to a combination of processes in this complex geodynamic setting.

Relationship between magmatism and basement uplift The geology of Westem and Central Europe is characterized by a series of domal uplifts (Massif Central, Rhenish Massif, Bohemian Massif, Armorican Massif, Vosges-Black Forest dome) of Variscan basement, up to 500-600 km in diameter, surrounded and onlapped by younger sedimentary sequences (Fig. 2). Basement uplift is generally considered to have initiated during the Neogene, extending over a period of some 15-20 Ma (Ziegler 1990, 1992). In detail, however, the timing of uplift is not well constrained; preliminary results from a programme of fission-track studies to characterize the uplift history from east to west across Europe (A. Hurford, pers. comm.) suggest that parts of the Variscan basement were already exhumed by the Early

155

Cretaceous. Additionally, there is stratigraphic evidence to suggest that both the Rhenish Massif and the Massif Central were close to sea level during the Oligocene (D~zes et al. 2004). The Rhenish Massif has probably been a topographic high since the end-Carboniferous Variscan orogeny (c. 300 Ma; GarciaCastellanos et al. 2000). Tertiary uplift started with large-scale tilting during the Palaeocene following a long period of tectonic stability during the Mesozoic. Doming initiated during the Eocene, with the major phase of uplift during the Late Oligocene. Uplift continued during the Neogene and Quaternary, accelerating during the Mid-Pleistocene. The locus of maximum uplift during the last 800 ka (250 m) is broadly coincident with the location of a low-velocity anomaly in the upper 400 km of the mantle, identified by Ritter et al. (2001) as a mantle plume on the basis of a local seismic tomography experiment. A number of the Tertiary-Quaternary volcanic fields within Europe are located on uplifted basement massifs (e.g. Massif Central, Eifel, Bohemian Massif), some of which appear to be dynamically supported by convective upwellings within the upper mantle (e.g. Granet et al. 1995; Wilson & Patterson 2001; Ritter et al. 2001). Elsewhere, however, areas of uplifted basement (e.g. the Ardennes within the Rhenish Massif; Armorican Massif; Fig. 2) are devoid of Tertiary-Quatemary magmatic activity. Equally, not all of the volcanically active areas are located on uplifted basement (e.g. Pannonian Basin). Consequently, it seems clear that there is not a simple correlation between basement uplift, diapiric upwelling of the asthenosphere and magma generation processes. The presence of a regional erosional hiatus at the MesozoicCenozoic boundary throughout Western and Central Europe may mark the initiation of basement uplift in the Massif Central, Vosges-Black Forest region, North Hessian Depression, Rhenish Shield, Bohemian Massif, Armorican Massif and the Ardennes (Ziegler 1990); no European rifting or Alpine collision events are known this early. 'Laramide' basin inversion in the central and west Netherlands Basin (Van Wijhe 1987), Lower Saxony Basin (Betz et al. 1987) and on the west margin of the Bohemian Massif (Schrrder 1987) attest to continued uplift into the Mid- and Late Palaeocene. Further inversion between the late Eocene and the Early Oligocene coincides with the first substantial magmatic episode in the Odenwald, south of the Rhenish Massif. Following a period of Mid-Oligocene subsidence (shown by periodic transgressions onto the Rhenish Massif; Meyer et al. 1983), uplift of the Vosges-Black Forest Massif dome and the Massif Central resumed in the Late Oligocene (Illies 1977). Uplift was initially restricted to rift shoulders, but quickly increased in intensity and extended throughout Western and Central Europe (including the Massif Central, eastem Paris Basin, Vosges and Black Forest mountains, Rhenish Massif, Hessian Depression, Harz Mountains and the Bohemian Massif (Ziegler 1990). Strong uplift throughout the Early Miocene to Early Pliocene coincided with inversion of the Leine, south Ruhr valley and French grabens (Meyer et al. 1983; Ziegler 1992), a sharp rise in volcanic activity (Lippolt 1982) and strong uplift in the Alps and Alpine fore-deep (Lemcke 1974; Malkovsky 1987). Uplift continued throughout the Pliocene and Pleistocene to the present day in the Massif Central, Rhenish Massif and Bohemian Massif (Becker 1993). Rhenish shield uplift is currently 0 . 4 - 0 . 6 m m a -1, with a maximum of l m m a -1 broadly coinciding with the distribution of recent magmatism in the Eifel (Fuchs et al. 1983; Garcia-Castellanos et al. 2000). The regional onset of basement uplift at the MesozoicCenozoic boundary across Europe, associated with early dyke intrusions (e.g. Rhine Graben), occurred at least 20 Ma before the main onset of rifting. This suggests that the main trigger for the Tertiary magmatism within Europe is the diapiric upwelling of mantle beneath the base of the lithosphere. No contemporaneous collisional events are known from the Alps at this time, although

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M. WILSON & H. DOWNES

extension along the Reykjanes, Aegir and Mohns ridges in the North Atlantic may have induced compressional stresses within the European lithosphere at this time (Becker 1993). In a number of areas within the European volcanic province (e.g. Massif Central, Rhenish Massif, northern Bohemian Massif) crustal uplift and surface magmatism during the Cenozoic are strongly linked, both spatially and temporally (Patterson 1996). The most recent volcanic eruptions from the Rhenish Massif, for example, coincide with highest measured uplift rates (Garcia-Castellanos et al. 2000). Uplift in these three areas correlates spatially with the locus of low-velocity anomalies in the upper mantle, which have been interpreted as small-scale mantle plumes (e.g. Granet et al. 1995; Wilson & Patterson 2001). The Vosges-Black Forest dome is, however, distinct from the other uplifted areas. It is somewhat smaller (c. 300 km wide), appears to originate by lithopheric flexuring and is not currently rising.

G e o p h y s i c a l s t u d i e s o f the u p p e r m a n t l e

Babuska & Plomerova (1992) have shown that the thickness of the lithosphere beneath Western and Central Europe is typically 100-120 km; this is a regional characteristic of the Variscan basement. In the northern part of the Massif Central the lithospheric thickness increases to 140 km, whereas thinning to 7 0 80 km occurs beneath the main grabens (e.g. Limagne Graben). The Western Alps have a deep lithospheric root (>170 km). The asthenosphere-lithosphere boundary below the SE part of the Rhenish Massif is elevated to c. 60 km depth (Babuska & Plomerova 1992), and to c. 70 km below the Vogelsberg (Braun & Berchemer 1993). The lower crust beneath the Vogelsberg is characterized by a strongly reflective zone at c. 20 km depth and a Moho depth of c. 28 kin, suggesting the presence of a lower crust underplated and intruded by basaltic magma. Babuska et al. (2002) have produced a more detailed lithospheric model for the Massif Central showing that the lithosphere is thick (100-140km) in the northern and western parts but thinner (70-80 km) in the south beneath the main volcanic regions of Cantal and Mont Dore to the east of the Sillon Houiller fault (a late Variscan sinistral transfer fault; Fig. 2). Additionally, there are significant differences in the orientation of seismic anisotropy within the mantle lithosphere of the western and eastern Massif Central. The eastern Massif Central appears to consist of two domains separated by an east-west-trending boundary. Based on studies of spinel peridotite mantle xenoliths exhumed by the Cenozoic basalts, this boundary corresponds to a terrane boundary between a northern cratonic and a younger southern domain (Lenoir et al. 2000), each characterized by differently oriented tectonic structures (Michon & Merle 2001). Babuska et al. (2002) have suggested that the magmas feeding the Cenozoic magmatism of the Massif Central migrated to the surface mainly along such reactivated basement sutures, which are probably translithospheric fault systems. Other areas in which such basement control on the locus of magma migration may be important include the Eger Graben and the volcanic fields of northern Germany (Eifel, Vogelsberg, North Hessian Depression). Both regional and local seismic tomography studies have imaged anomalously low seismic velocities in the upper mantle beneath Western and Central Europe (e.g. Spakman et al. 1993; Zielhuis & Nolet 1994; Granet et al. 1995; Hoernle et al. 1995; Goes et al. 2000; Ritter et al. 2001), consistent with a model of discrete upper mantle diapirs beneath each of the major volcanic fields. None of these studies resolved structure below the 660 km discontinuity. A recent European tomographic model (Bijwaard & Spakman 1999; Goes et al. 1999), however, has provided some evidence that the upper mantle velocity anomalies may be rooted in the lower mantle. Between 900 and 1200 km depth there appears to be a semi-circular low-velocity structure that links the Iceland plume, a Central European velocity

anomaly and a plume-like structure beneath the Canary Islands, and that may continue to the core-mantle boundary. Low seismic velocities in the upper mantle were considered by Goes et al. (1999) to reflect the presence of mantle some 100-200 ~ hotter than its surroundings. They attributed the geometry of the velocity anomaly to deflection of a lower mantle plume by relatively flat-lying subducted slabs on the 660 km discontinuity at the base of the upper mantle (Piromallo et al. 2001). Ritter et al. (2001) have demonstrated the existence of a 100 km wide, finger-like P-wave velocity anomaly in the upper mantle beneath the Eifel volcanic field that extends to a depth of at least 400 kin; this could be about 150-200 ~ hotter than the surrounding mantle if the entire velocity anomaly is translated into a temperature contrast. This is similar to the structure reported by Granet et al. (1995) beneath the Massif Central. Receiver function analysis indicates a depression of the 410 km seismic discontinuity beneath the Eifel consistent with a thermal anomaly of this magnitude (Grunewald et al. 2001). In the lower mantle beneath Europe, Goes et al. (1999) discovered a 500kin • 500km anomaly at 660-2000 km depth, which might be the lower mantle source of this upper mantle thermal anomaly. However, a connection between this lower mantle structure and the upper mantle diapir has not been demonstrated. Keyser et al. (2002) have demonstrated that there is a 'hole' in the low-velocity channel beneath the Eifel at about 200 km depth consistent with an increase in the shear modulus; this might correspond to the depth of onset of partial melting. Those workers also located a low-velocity region further east beneath the Vogelsberg at c. 170-240 km depth. As there has been no volcanic activity in the Vogelsberg since the mid-Miocene, this anomaly is difficult to interpret; it could reflect a waning Miocene thermal anomaly, or a new diapiric upwelling that has not yet reached the base of the lithosphere. It is important to recognize that variations in seismic velocity within the mantle can represent anomalies in composition, temperature and anisotropy. Keyser et al. (2002) indicated that the - 5% (negative) perturbation in S-wave velocity (Vs) between 31 and 170 km depth below the Eifel is consistent with an excess temperature of 100 K plus c. 1% partial melt. In the lower part of the asthenosphere the Vs anomaly is at least - 1%, consistent with a temperature anomaly of > 70 K combined with the presence of water in the upwelling mantle. Piromallo et al. (2001) have used a high-resolution seismic tomographic model to study the structure of the upper mantle beneath Europe. They showed that between 400 and 600 km depth there is an ellipsoidal region, 2000 km • 4000 km in area and c. 100-150 km thick, of higher seismic velocities, which might locally inhibit the vigour of upper mantle convection (Fig. 1). Wortel & Spakman (2000) also identified the presence of a layer of fast material under the Mediterranean region. If this layer represents a region of cold subducted oceanic lithosphere, its presence would cause a temperature inversion in the upper mantle below 500 km, which might greatly modulate the style of mantle convection, locally slowing down the circulation. Such a scenario might invalidate the commonly used assumption of an adiabatic temperature gradient in the upper mantle, as transient convection might not be sufficiently vigorous to create an adiabatic condition (Matyska & Yuen 2001). The presence of a regional 'cool spot' in the Transition Zone (i.e. the region of high seismic velocities) might explain why the major volcanic regions within Europe (Fig. 1) appear to be located peripherally to the high-velocity anomaly in the Transition Zone. On the basis of a detailed seismic tomographic study of the Massif Central, France, Granet et al. (1995) proposed that all of the major Tertiary-Quaternary volcanic fields of Western and Central Europe (Fig. 1) are underlain by finger-like thermal anomalies in the asthenosphere, sourced from a laterally extensive layer close to the base of the upper mantle. Wilson & Guiraud (1998) subsequently suggested that similar features might underlie other Tertiary-Quaternary volcanic fields in northern (Hoggar and

TERTIARY-QUATERNARY MAGMATISMIN EUROPE Tibesti massifs) and central (Darfur Dome) Africa, and in the Canary Islands. The location of these volcanic fields to the north and south of the Alpine collisional tectonic front suggests that the generation of diapiric instabilities in the upper mantle might be linked to subduction processes, including slab break-off, during continental collision of Africa with Eurasia. The length scale of the diapirs (100-500 km diameter) is much smaller than that typically associated with mantle plumes (10002000 km diameter), which have been inferred to trigger the flood basalt volcanism of the so-called Large Igneous Provinces (LIPs; White & McKenzie 1989). Additionally, unlike LIPs, the characteristic volcanism is of alkali basalts, basanites and nephelinites and their differentiates, although locally eruptions of subalkaline (tholeiitic) basalts do occur. In this respect the volcanism shows strong similarities to that of many Atlantic oceanic islands (e.g. Canary Islands, Azores, Cape Verdes) and to island chains within the southwestern Pacific, which are inferred to be plume-related. It has been suggested that individual volcanic island chains within the Pacific Ocean represent the locus of hotspots rising from the upper boundary of a much larger wavelength upwelling or 'super-swell' within the mantle (McNutt 1998). Such a model may also be applicable to the convective instabilities that have been inferred to underlie the European volcanic province. On the basis of seismic tomographic images, Goes e t al. (1999) have suggested that a low-velocity structure between 660 and 2000 km depth represents a lower mantle plume upwelling beneath Central Europe, which may feed smaller-scale upper mantle plumes beneath each of the volcanic fields.

157

more effective discriminants are trace element ratio plots such as Th/Yb v. Ta/Yb as proposed by Wilson & Bianchini (1999).

The range of magma

types observed

Within the European volcanic province the a n o r o g e n i c suites include both alkaline and sub-alkaline magma series, ranging from basalts to more silica-rich compositions. In general, only the most primitive mafic magmas (basalts s e n s u l a t o ; SIO2<55 wt%, MgO > 6 wt%) can provide information about the nature of their mantle source, and, consequently, we focus on the geochemical characteristics of such rock types in subsequent discussions. Table 1 lists the literature sources of the major and trace element and S r - N d - P b isotope data that we have used to characterize the various European volcanic provinces. A total-alkalis (Na20-k-K20) v. silica (SiO2) variation diagram (Fig. 5) indicates a wide spectrum of magma compositions ranging from silica-poor nephelinites and melilitites (foidites), through alkali basalts and basanites to transitional and sub-alkaline basalts (tholeiites). The geochemical characteristics of the most primitive alkaline mafic magmas erupted throughout the region are, in general, remarkably similar to those of OIB inferred to be related to the activity of mantle plumes, and of other regions of Cenozoic

Table 1.

Sources of major and trace element, S r - N d - P b isotope and geochronological data f o r European Cenozoic anorogenic igneous rocks

Relationship between the location of magmatism and basement structure

Area

Within the Variscan basement of Europe, a series of elongate, east-west-trending terranes (Rheno-Hercynian, Saxo-Thuringian and Moldanubian) have been identified (Franke 1989), separated by deep, laterally persistent fault zones, interpreted as tectonic sutures (Fig. 2). These major structural units of the European lithosphere resulted from the collision, during the Devonian and Carboniferous, of Laurasia with Gondwana and a number of intervening microplates. The terrane boundaries may be regions of anomalously thin, irregular or weak lithosphere and appear to have exerted a significant control on the location of subsequent Cenozoic magmatism within Europe, possibly acting as pathways for magma ascent through the lithosphere. Huismans (1999) has demonstrated that asthenospheric diapirism can be generated by rifting processes. Extension of heterogeneous lithosphere could create differential topography at the base of the lithosphere, which might induce further upwelling of low-density asthenosphere, resulting in magma generation and surface uplift. Thus, it is possible that the Variscan structural fabric of Europe may have pre-conditioned the subsequent locations of Tertiary- Quaternary volcanism.

Rhenish Massif

Southern Germany

Massif Central

Czech Republic-Poland Pannonian Basin

Dinarides

The geochemical characteristics of the magmatic rocks How

do we classify magmatic

rocks as anorogenic

?

Magmas of o r o g e n i c and a n o r o g e n i c affinity can be distinguished based on their major and trace element and S r - N d - P b isotope geochemical characteristics (Wilson & Bianchini 1999). The variation of K 2 0 / N a 2 0 (weight ratio) as a function of SiO2 (wt%) content can sometimes be an effective tool for discrimination; magmas of orogenic affinity frequently have a K 2 0 / N a 2 0 ratio > 1.5, whereas anorogenic magmas always have a KzO/Na20 ratio < 1. Caution must be exercised, however, as there are also examples of low-K subduction-related volcanic suites. Much

Bulgaria Italy, Sicily and Sardinia

Spain

Sources of Data Schmincke et al. (1983); Mertes & Schmincke (1985); Wrrner et al. (1986); Lippolt et al. (1990); Wedepohl et al. (1994); Jung & Masberg (1998); Wedepohl & Baumann (1999); Jung & Hoernes (2000); Wedepohl (2000); Bogaard & W~3rner(2003); Wilson & Rosenbaum (unpubl. data) Lippolt et al. (1973); Keller et al. (1990, 2002); Schleicher et al. (1990, 1991); Hegner et al. (1995); Wilson et al. (1995b) Patterson (1996); Wilson & Rosenbaum(unpubl. data) Downes (1983, 1987); Wrrner et al. (1986); Briot et al. (1991); Wilson et al. (1995a); Patterson (1996); Wilson & Patterson (2001); Wilson & Rosenbaum (unpubl. data) Kopecky (1966); Blusztajn & Hart (1989); Shrbeny (1995); Ulrych & Pivec (1997); Ulrych et al. (2001); Wilson & Rosenbaum (unpubl. data) Salters et al. (1988); Szabo et al. (1992); Embey-lsztinet al. (1993); Downes et al. (1995a,b); Dobosi et al. (1995); Embey-Isztin& Dobosi (1995); Harangi et al. (1995); Pecskay et al. (1995); Wilson& Rosenbaum(unpubl. data) Pamic et al. (1995); Tari & Pamic (1998); Jovanovic et al. (2001) Vaselli et al. (1997); Marchev et al. (1998) Beccaluva et al. (1977, 1981, 1987, 1998, 2002); Carter & Civetta (1977); Cioni et al. (1982); Condomines et al. (1982); Dostal et al. (1982); Rutter (1987); Clochiatti et al. (1988); Calanchi et al. (1989); Montanini& Villa (1993); Gillot et al. (1994); De Vecchi & Sedea (1995); Milani (1996); D'Orazio et al. (1997); Tanguy et al. (1997); Gasperiniet al. (2000); Armienti et al. (2004); Corsaro & Pompilio (2004) L6pez-Ruiz et al. (1993); Cebri~ & L6pez-Ruiz (1995); Cebrifiet al. (2000)

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M. WILSON & H. DOWNES

Fig. 5. Total alkali-silica variation diagrams for the most primitive mafic magmatic rocks. The classification boundaries are from Le Bas et al. (1986). The dashed grey line in (c) marks the subdivision between alkaline and sub-alkaline magma series. Sources of data are given in Table I. BTA, basaltic trachyandesite; TB, trachy-basalt.

continental intra-plate volcanism (e.g. Wilson & Downes 1991). Locally where the alkaline magmatism post-dates a recent episode of subduction-related magmatism (e.g. Sardinia, Pannonian Basin, Dinarides, Bulgaria, western Turkey), the earliest magmatic rocks may preserve the fingerprint of earlier fluxing of the upper mantle by subduction-zone fluids. Localized oceanic spreading centres within the Mediterranean domain preferentially sample a depleted mantle source component, similar to the source of mid-ocean ridge basalts (MORB), although in the Tyrrhenian Sea this is clearly modified by a subduction-related fluid flux (Wilson & Bianchini 1999).

S r - N d - P b isotope and trace element chemistry of the most primitive mafic magmas

In Sr-Nd isotope space (Fig. 6) the European basalts define a linear array trending f r o m 143Nd/a44Nd ratios of around 0.5130 towards Bulk Earth values (0.51264); this array may reflect mixing of partial melts derived from different mantle source components (e.g. Wilson & Downes 1991; Wilson & Bianchini 1999; Wilson & Patterson 2001). An isotopically distinct group of quartz tholeiites from central Germany fall below the array; this may be attributable to crustal contamination of the magmas. A series of broadly linear arrays are also evident in Nd-Pb and P b - P b isotope space (Fig. 7), fanning away from a common focal point, although there is considerable dispersion of the data. These S r - N d - P b isotope signatures have been attributed to the mixing of partial melts derived from a common, possibly plume-related, asthenospheric mantle source component known

as the Low Velocity Component (LVC; Hoernle et al. 1995) or European Asthenospheric Reservoir (EAR; Cebrifi & Wilson 1995) and a number of regionally heterogeneous sub-continental lithospheric mantle components (Cebrifi & Wilson 1995; Granet et al. 1995; Hoernle et al. 1995). It is generally accepted that the increasing degree of undersaturation in S i O 2 within the sequence olivine tholeiite-alkali olivine basalt-basanite-nephelinite-melilitite principally results from decreasing degrees of partial melting at increasing depth in the mantle (e.g. Wilson 1989). Constraints on both the depth and degree of partial melting can be provided by the concentration of highly incompatible trace elements in the most primitive mafic magmas, the enrichment of light rare earth elements (LREE) over heavy rare earth elements (HREE) and the CaO/AI203 ratio. Trace element ratios (e.g. La/Yb or La/Sm) where the numerator has a greater degree of incompatibility than the denominator exhibit a regular decrease from nephelinite and basanite to olivine tholeiite, consistent with increasing degrees of partial melting. High CaO/AI203, La/Yb and Nb/Y ratios suggest that the alkali basaltic (sensu lato) magmas were derived from a garnet-bearing mantle source. Enrichment of highly incompatible trace elements and LREE, and strong fractionation of LREE over HREE, can, however, be explained either by moderate degrees of partial melting of an enriched source or smaller degrees of melting of a more depleted source. Figure 8 illustrates the variation of Nb/Y v. Zr/Nb compared with model melting curves for 0.5-5% partial melting of spineland garnet-peridotite facies mantle from Harangi (2001). This clearly indicates that, in most of the volcanic provinces, the mantle partial melting column spans the transition from garnetto spinel-peridotite facies mantle, probably close to the base of

TERTIARY-QUATERNARY MAGMATISMIN EUROPE

159

Fig. 6. Variation of 143Nd/144Ndv. S7Sr/86Srfor the most primitive mafic volcanic rocks. EAR, isotopic composition of the European Asthenospheric Reservoir from Cebri~i& Wilson (1995). BE, Bulk Earth. Sources of data are given in Table 1.

Fig. 7. (a) Variation of 143Nd/144Ndv. 2~176 for the most primitive mafic volcanic rocks. (b) Variation of 2~176 v. 2~176 for the most primitive mafic volcanic rocks. EAR, isotopic composition of the European Asthenospheric Reservoir from Cebrifi & Wilson (1995). Sources of data are given in Table 1.

the continental lithosphere (see Bogaard & Wrrner 2003). Degrees of partial melting are typically less than 1%. Only in those regions in which sub-alkaline tholeiitic basalts occur (e.g. central Germany, Pannonian Basin and the Iblean Plateau, Sicily) does the degree of partial melting approach 5%; these tholeiitic basalts clearly equilibrated at somewhat shallower depths in spinel-peridotite facies mantle and may include a significant lithospheric mantle source component in their petrogenesis. The insets in Figure 8 show the variation of Nb/Y v. K20/Na20. There is generally a poor correlation between these two parameters. However, it is clear that the most sodic basalts (lowest K 2 0 / Na20) are the sub-alkaline tholeiites. The distinctive melilitites from the Urach and Hegau volcanic fields to the east of the Rhine Graben have the highest Nb/Y ratios, consistent with an extremely low degree of partial melting (<0.5%) within the garnet stability field. A REE ratio plot of La/Yb v. Gd/Yb (Fig. 9) also supports a model of mixing of partial melts from garnet- and spinel-peridotite facies mantle in the petrogenesis of the magmas. Magmas derived in equilibrium with a mantle source in which garnet is a residual phase have high La/Yb and Gd/Yb ratios because Yb is a compatible trace element in garnet (i.e. it preferentially partitions into garnet rather than the coexisting melt), but is incompatible in spinel. The variation of CaO/A1203 shows a strong negative correlation with wt% SiO2 (Fig. 10) consistent with increasing degrees of partial melting from the melilitites (Urach-Hegau) to the sub-alkaline tholeiites of northern Germany and the Iblean Plateau (Sicily). A significant proportion of the more silica-rich alkaline mafic magmas also have distinctly low CaO/A1203 ratios. Many of these samples also have variably elevated K20/Na20 ratios, consistent with their derivation from enriched, probably lithospheric, mantle sources that had been fluxed by subduction-related fluids during either Variscan orogenesis or Alpine collision. In primitive mantle normalized trace element variation diagrams (Fig. 11) the alkali basalts, basanites and nephelinites are typically characterized by marked depletions of K and Rb relative to Ba and Nb, consistent with the presence of residual amphibole or phlogopite in their mantle source. The negative K anomaly disappears in the highest degree partial melts (tholeiitic basalts), consistent with the complete melting of the K-bearing phase in the mantle source. The potassic magma

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Fig. 8. Variation of Nb/Y v. Zr/Nb for the most primitive mafic volcanic rocks. Curves are model melting curves from Harangi (2001) for 0.5-5% partial melting of spinel- and garnet-peridotite facies mantle. Insets show variation of Nb/Y v. K20/Na20. Sources of data are given in Table 1. Model parameters (from Harangi 2001): source composition: Zr 11.3 ppm, Nb 0.72 ppm, Y 2,7 ppm; source mineralogy (%): spinel lherzolite is ol58-opx30-cpxl0-sp2; garnet lherzolite is o159.9-opx25.5-cpx8.8-gt5.8; melting modes (%): oll.2-opx8. I -cpx76.4-sp14.3 and ol 1.2-opx8.1-cpx36.4-gt54.3.

types inferred to be partial melts of enriched lithospheric mantle sources (leucite nephelinites and leucite basanites) have relatively fiat trace element patterns and no K anomaly. The absence of a K anomaly suggests that any K-bearing phase in their mantle source (i.e. phlogopite or amphibole) must have completely melted.

Discussion

Fig. 9. Variation of La/Yb v. Gd/Yb for the most primitive mafic volcanic rocks. Sources of data are given in Table 1.

One of the major problems in petrogenetic studies of intracontinental plate alkaline magmatism is the identification of the source of the primary magmas, in particular whether this is located in the mantle part of the lithosphere, the underlying convecting asthenosphere, or both. It has often been argued that because the basalts have S r - N d - P b isotopic characteristics similar to those of oceanic basalts then they must originate in the asthenosphere (e.g. Wilson & D o w n e s 1991). A number of workers have argued that mantle plumes must be involved in the generation of the Central and Eastern European province volcanic rocks (e.g. Wedepohl et al. 1994; Wedepohl & Baumann 1999; Wedepohl 2000). This remains, however, a subject for debate.

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Fig. 10. Variation of CaO/AI203 v. SiOa for the most primitive mafic volcanic rocks. Sources of data are given in Table 1.

widespread but systematic isotopic heterogeneities within the various structural domains within the European lithosphere, which may become imprinted upon any asthenosphere-derived magmas passing through them. Phlogopite-amphibole-bearing lithospheric mantle, probably metasomatized during or after the Variscan orogeny, appears be the predominant source of potassic mafic magmas within the province (Wilson & Downes 1992). Amphibole is relatively common within the European mantle xenolith suites (Downes 2001), whereas phlogopite typically occurs only in those regions in which potassic magma types are erupted. Detailed studies of mantle xenoliths from the Eifel province of northern Germany (Witt-Eickschen et al. 1998) have revealed the presence of at least two metasomatic enrichment events; one related to modification by subduction-related fluids during the Variscan orogeny and a more recent vein-type infiltration of presumed Tertiary-Quaternary age. In the context of the petrogenesis of the primary magma spectrum within the European volcanic province, an important question is whether we can identify specific magma compositions that appear to be derived directly from the asthenosphere (or the thermal boundary layer at the base of the lithosphere) and that do not appear to have experienced significant degrees of interaction with enriched lithospheric mantle domains en route to the surface. Melilite nephelinites and melilitites are, in general, rather rare highly silica-undersaturated mafic magmas, which occur at scattered localities throughout the European Tertiary-Quaternary volcanic province (Fig. 2). The best known occurrences are Ciudad Real in Spain, Marcoux in the French Massif Central, Essey-la-Cote and Grand Valtin in the Vosges, the Ohre rift of the Czech Republic, Kaiserstuhl in the Rhine Graben, the Eifel province of the Rhenish Massif, and the Urach and Hegau provinces to the east of the Rhine Graben (Wilson et al. 1995b). Locally, however (e.g. Urach and Hegau provinces of Germany, Eger Graben of the Czech Republic), they are a volumetrically significant component of the spectrum of mafic magmas erupted. Melilitites may, therefore, be the most likely candidates for primary partial melts of the thermal boundary layer at the base of the lithosphere, representing the incipient melting of the most fusible parts of the upper mantle (Hegner et al. 1995; Wilson et al. 1995b). In some cases their S r - N d Pb isotope characteristics appear to represent the primary asthenospheric end-member of the European mafic-ultramafic magma spectrum (e.g. Wilson & Downes 1991, 1992); however, in many cases the melilitites appear to have a hybrid lithosphere- asthenosphere signature.

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Bogaard & W6rner (2003) have proposed that high-Ti basanites forming the earliest eruptive rocks in the Vogelsberg area of northern Germany are partial melts of lithospheric mantle that had been metasomatized during the initial stages of uplift of the Rhenish Massif at c. 70 Ma. Subsequent eruptions of tholeiitic basalts are considered to represent higher degrees of partial melting of a much more depleted mantle source component, located in either the lower lithosphere or uppermost asthenosphere. The depleted mantle source component is geochemically similar to MORBsource mantle. With time the influence of an EAR-like source component in the petrogenesis of the Vogelsberg magmas appears to increase as the degree of melting decreases. The sequence is similar to that predicted by Gallagher & Hawkesworth (1992) and Hawkesworth & Gallagher (1993) for the sequential evolution of basaltic volcanism in a region extending above a thermal anomaly in the mantle. Magma generation within the European upper mantle during the Cenozoic was most probably triggered by adiabatic decompression partial melting of the asthenosphere, the solidus of which was locally lowered by the infiltration of slab-derived fluids above contemporaneous subduction zones. Decompression melting may have been triggered locally by lithospheric flexuring in the fore-bulge of the Alps (e.g. Urach province of Germany; Vosges-Black Forest dome), by lithospheric extension (e.g. Limagne Graben; Pannonian Basin; Eger Graben) or by upwelling of asthenospheric mantle diapirs (e.g. Massif Central; Eifel). If such diapirs are hotter than the ambient mantle, the rising mantle will intersect the solidus at greater depths; thus much less lithospheric thinning is required to produce relatively large volumes of melt. The amount of melting, and hence the composition of the magmas, is strongly influenced by the local thickness of the lithosphere (e.g. Ellam 1992). For the generally small amounts of lithospheric extension observed within the European volcanic province (e.g. Merle et al. 1998), it is clear that partial melting requires either an anomalously hot or a volatile-rich mantle source. The convective instabilities (mantle diapirs) that have been inferred to provide the dominant control on magmatism in a number of areas have been imaged by high-resolution seismic tomography (e.g. Granet et al. 1995; Ritter et al. 2001). Their length scale suggests that they most probably originate as instabilities within the upper mantle. Their velocity contrast with the surrounding mantle is consistent with a temperature anomaly of up to 100-150 ~ hotter than the ambient mantle. Convective destabilization of the upper mantle may have been initiated by the Alpine collision and the consequent global reorganization of plate motions. Small-degree, CO2-HzO-enriched partial melts from the ascending mantle diapirs may have frozen at the base of the lithosphere, creating a heterogeneous carbonated phlogopite-lherzolite layer (Wilson et al. 1995b). Partial melting of this layer, during subsequent lithospheric extension or thermal thinning, could produce melilitite-like magmas. Enrichment of the base of the lithosphere by infiltration of smalldegree partial melts could have taken place in several stages, related to Neogene diapiric upwelling of the asthenosphere. It is possible, however, that widespread pollution of the shallow mantle beneath Europe with a plume-like component could have occurred in the Late Cretaceous-Early Tertiary associated with outflow from the Iceland plume system in the North Atlantic, from the Canary Islands plume, or perhaps even earlier during the Mesozoic when there was widespread plume-related magmatism globally (Wilson 1992). Clearly, the simplest model would be to relate the enrichment of the base of the lithosphere to Tertiary mantle upwelling.

Summary The distribution of the major volcanic provinces within Europe is broadly anti-correlated with the location of a zone of high-velocity, presumed subducted slab material, in the base of

the upper mantle (500-600 km depth). Many of the major volcanic fields are located around the periphery of this velocity anomaly (Fig. 1), coincident with the distribution of small-scale convective instabilities (mantle diapirs) in the upper mantle imaged by local seismic tomography experiments (e.g. Granet et al. 1995; Ritter et al. 2001). Triggering of these upwellings may partially be a response to plate subduction and slab detachment during the Alpine collision. Alternatively, it is possible that there is a more complex relationship between extensional tectonics within the Alpine foreland and the presence of discontinuities in lithospheric thickness between adjacent terrane blocks (e.g. Massif Central, Rhenish Massif); extension of heterogeneous lithosphere, consisting of a collage of accreted Variscan teranes, might have induced localized convective instabilities in the upper mantle. Based on the above scenario, melt generation in the asthenosphere and the base of the lithosphere is inferred to be the consequence of decompression partial melting, triggered by mantle upwelling. Locally, melt generation is enhanced by the presence of hydrous fluids from contemporaneous or earlier subduction zones (e.g. Pannonian Basin, Sicily, Western Turkey). Although the mantle part of the lithosphere may locally contribute to the magmatism, the main magma source region is sub-lithospheric. Lithospheric architecture, however, clearly plays an important role in the location of the volcanic fields. Plate-scale stresses have controlled the development of the Cenozoic rift system within Europe, often reactivating older Carboniferous fault systems and terrane boundaries. Most of the major rifts (e.g. Rhine Graben, Limagne Graben) are, however, only weakly magmatic. This demonstrates that passive lithospheric extension alone cannot be the main trigger for magma generation. Changes in the orientation of the regional stress field may play an important role in controlling the migration of mantle-derived magmas through the crust and, thus, the periodicity of the magmatism. During the past decade we have discussed our ideas on the petrogenesis of magmas within the European volcanic province with numerous workers, all of whom are thanked for their insightful comments. J. Rosenbaum contributed a significant volume of S r - N d - P b isotope data, which we report in the figures, some of which has still to be published. R. Patterson provided important insights on the uplift history of Europe during the Cenozoic and its relationship to tiffing and magmatism. J. Ulrych inspired our work on the Eger Graben in the Czech Republic and provided invaluable support in the field. Discussions with P. Ziegler over many years have clarified our ideas about the geodynamic setting of the magmatism. J.-M. Cebri~i helped to constrain the isotopic characteristics of the EAR and focused our ideas about the contribution of enriched lithospheric mantle sources to the petrogenesis of potassic magmas. Collaboration with U. Achauer and M. Granet demonstrated the value of seismic tomography in understanding the pattern of upper mantle convection beneath Europe. M. Lustrino and H.-U. Schminke are acknowledged for helpful discussions on the geodynamic setting of Cenozoic magmatism in Europe.

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Tertiary-Quaternary subduction processes and related magmatism in the Alpine-Mediterranean region S Z A B O L C S H A R A N G I 1, H I L A R Y D O W N E S 2 & I O A N S E G H E D I 3

tDepartment of Petrology and Geochemistry, Ertvrs University, H-1117 Budapest, Pdzmdny P~ter s~tdny 1 / C, Hungary (e-mail: szabolcs.harangi @geology, elte.hu) 2Birkbeck/UCL Research School of Earth Sciences, Birkbeck University of London, Malet Street, London WC1E 7HX, UK 3Institute of Geodynamics, 19-21, str. Jean-Luis Calderon, Bucharest 70201, Romania

Abstract: DuringTertiary to Quaternary times, convergence between Eurasia and Africa resulted in a variety of collisional orogens and

different styles of subduction in the Alpine-Mediterranean region. Characteristic features of this area include arcuate orogenic belts and extensional basins, both of which can be explained by roll-back of subducted slabs and retreating subduction zones. After cessation of active subduction, slab detachment and post-collisional gravitational collapse of the overthickened lithosphere took place. This complex tectonic history was accompanied by the generation of a wide variety of magmas. Most of these magmas (e.g. low-K tholeiitic, calc-alkaline, shoshonitic and ultrapotassic types) have trace element and isotopic fingerprints that are commonly interpreted to reflect enrichment of their source regions by subduction-related fluids. Thus, they can be considered as 'subduction-related' magmas irrespective of their geodynamic relationships. Intraplate alkali basalts are also found in the region and generally postdated the 'subduction-related' volcanism. These mantle-derived magmas have not (or only slightly) been influenced by subduction-related enrichment. This paper summarizes the geodynamic setting of the Tertiary-Quaternary 'subduction-related' magmatism in the various segments of the Alpine-Mediterranean region (Betic-Alboran-Rif province, Central Mediterranean, the Alps, Carpathian-Pannonian region, Dinarides and Hellenides, Aegean and Western Anatolia), and discusses the main characteristics and compositional variation of the magmatic rocks. Radiogenic and stable isotope data indicate the importance of continental crustal material in the genesis of these magmas. Interaction with crustal material probably occurred both in the upper mantle during subduction ('source contamination') and in the continental crust during ascent of mantle-derived magmas (either by mixing with crustal melts or by crustal contamination). The 87Sr/86Sr and 2~176 isotope ratios indicate that an enriched mantle component, akin to the source of intraplate alkali mafic magmas along the Alpine foreland, played a key role in the petrogenesis of the 'subduction-related' magmas of the AlpineMediterranean region. This enriched mantle component could be related to mantle plumes or to long-term pollution (deflection of the central Atlantic plume and recycling of crustal material during subduction) of the shallow mantle beneath Europe since the late Mesozoic. In the first case, subduction processes could have had an influence in generating asthenospheric flow by deflecting nearby mantle plumes as a result of slab roll-back or slab break-off. In the second case, the variation in the chemical composition of the volcanic rocks in the Mediterranean region can be explained by 'statistical sampling' of the strongly inhomogeneousmantle followed by variable degrees of crustal contamination.

The A l p i n e - M e d i t e r r a n e a n region is one of the most complex geodynamic settings on Earth. Subduction of oceanic plates, collision of continents, opening of extensional basins and possible upwelling of mantle plumes have all occurred associated with the formation of a wide variety of igneous rocks during the Tertiary and Quaternary. These processes are still active in some parts of this region. The geodynamic processes and volcanic activity have been the focus of research for a long time. During the last decade a number of papers have been published using the results of new techniques such as seismic tomography and isotope geochemistry (see summary papers of Doglioni et al. 1999; Wilson & Bianchini 1999; Lustrino 2000; Wortel & Spakman 2000). Convergent margins are the sites where subduction of oceanic lithosphere occurs beneath oceanic or continental plates. The style of subduction depends upon various parameters, including the rate of convergence, the rate of subduction, the nature of the subducted lithosphere and the polarity of subduction (Jarrard 1986; Royden & Burchfiel 1989; Doglioni 1993). T e r t i a r y Quaternary subduction in the A l p i n e - M e d i t e r r a n e a n region was governed by the convergence between Eurasia and Africa in an area where continental and oceanic microplates were trapped between the converging continental plates. This resulted in various styles of subduction and collision (Royden & Burchfiel 1989; Royden 1993; Doglioni et al. 1999). Royden & Burchfiel (1989) proposed that orogenic belts with high topographic elevation were formed where the rate of convergence exceeded the rate of subduction (advancing subduction; e.g. Alps). In contrast, low topographic relief and regional extension in the upper plate are considered to characterize subduction boundaries

where the rate of subduction exceeded the rate of overall plate convergence (retreating subduction; e.g. B e t i c s - A l b o r a n - R i f , Apennines, Hellenic and Carpathian thrust belts). Doglioni (1991, 1993) and Doglioni et al. (1999) emphasized the importance of subduction polarity. Westward-directed subduction zones oppose mantle flow and have similar features to retreating subduction boundaries, that is steep angle of subduction, slab roll-back, opening of extensional basins and termination of subduction when the buoyant continental lithosphere enters the trench. Eastward-directed subduction zones are reinforced by mantle flow and show a lower angle of subduction, together with a lack of extension in the overlying plate. Following subduction of oceanic lithosphere, c o n t i n e n t - c o n t i n e n t collision occurs, resulting in thickening of the continental crust and lithosphere. Detachment of the dense oceanic slab (Davies & von Blanckenburg 1995), delamination of the thick lithospheric mantle (Bird 1979), sometimes with the dense mafic lower crust (Lustrino et al. 2000) or convective removal of the lower lithosphere (Houseman et al. 1981; Platt & Vissers 1989; Turner et al. 1999) could be responsible for postcollisional extension and related magmatism. The geochemistry of the magmas is dependent on the rheology of the continental plates involved in the collision, the extent of collision and the velocity of plate convergence (Wang et al. 2004). The complex tectonic evolution of the A l p i n e - M e d i t e r r a n e a n region has been associated with formation of a wide range of Tertiary to Quaternary and even recent igneous rocks (Fig. 1). Wilson & Bianchini (1999) divided the magmatic activity of this area into 'orogenic' and 'anorogenic' types. Alkali basaltic magmas of anorogenic type erupted mainly along the foreland of the Alps (see Wilson & Downes 2006), but can be found also throughout

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 167-190. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Simplified map for the distribution of Tertiary to Quaternary volcanic rocks in Central and Southern Europe. Dark blue areas indicate oceanic crust. ECRIS, European Cenozoic Rift System; BAR, Betic-Alboran-Rif province (Ab, Alboran; Be, Betic; RTG, Gourougou-Trois Furches-Ras Taft (RID; Or, Oranie; Ca, Calatrava; O1, Olot); CM, Central Mediterranean (Sa, Sardinia; Si, Sisco; Tu, Tuscany; Rp, Roman province; Ca, Campania; Vu, Vulture; Va, Vavilov; Ma, Marsili; Us, Ustica; Ai, Aeolian Islands; Et, Etna; Hy, Iyblei; Pa, Pantelleria); PIL, Periadriatic-Insubric Line (Be, Bergell; Ad, Adamello; Po, Pohorje; Ve, Veneto); CPR, Carpathian-Pannonian region (WC, western Carpathians; NEC, northeastern Carpathians; EC, eastern Carpathians; Ap, Apuseni); DR, Dinarides and Rhodope (Di, Dinarides; RT, Rhodope-Thrace); AA, Aegean-Anatolia (Sa, Santorini; WA, Western Anatolia; Ku, Kula; Af, Afyon; Ko, Konya; Ga, Galatia; CA, Central Anatolia).

the Mediterranean region. Orogenic calc-alkaline, potassic to ultrapotassic and silicic magmas erupted in the convergent margins of the Alpine-Mediterranean region. These volcanic rocks show indeed a subduction-related geochemical composition. In this paper, we highlight the results of the most recent publications on the Tertiary to Quaternary volcanism and show that geodynamic and petrogenetic models are still highly controversial in spite of the emerging data. One of the aims of this paper is to present the competing tectonic and petrogenetic models, the volcanic histories of the various areas of this region, and finally to search for the origin of the magmas and the reason for melt generation processes in this complex tectonic setting.

Generation of magmas with 'subduction-related' geochemistry Subduction of oceanic lithosphere often results in a linear chain of volcanoes along an island arc or active continental margin (Fig. 2a; Gill 1981; Thorpe 1982; Wilson 1989). Most of the volcanic rocks, which build up these volcanoes, are calc-alkaline in composition. In addition, low-K tholeiitic and potassic magmas can also be associated with active subduction. The primary magmas are considered to form by melting of hydrous peridotite either in the asthenospheric mantle wedge or in the subcontinental lithospheric mantle (Gill 1981). The complex processes beneath active volcanic arcs seem

to develop fairly quickly (Gill & Williams 1990; Gill et al. 1993; Reagan et al. 1994; Elliott et al. 1997). Dehydration of the descending slab, metasomatism of the mantle wedge by aqueous fluids, partial melting and eruption of magmas can take place in < 3 0 ka (Elliott et al. 1997; Turner et al. 2000). Therefore, the presence of calc-alkaline magmatism is widely considered as evidence for the existence of an active subduction zone (e.g. Pearce & Cann 1973; Wood 1980). Most calc-alkaline magmas are intermediate in composition (andesite to dacite) and true basalts are rare in most suites as a result of the differentiation processes in shallow-level magma chambers. Subduction-related magmas worldwide have trace element and isotopic fingerprints that are usually interpreted as reflecting fluid involvement in their genesis. Such fluids are either aqueous solutions or silicate melts released from the subducted oceanic lithosphere and its overlying sediments. Their effects are seen in high ratios of large ion lithophile elements (LILE; such as Cs, Rb, Ba, K, Sr) and Pb to high field strength elements (HFSE; e.g. Nb, Ta, Zr, Hf and Ti, Fig. 3). The LILE are soluble in aqueous fluids (Tatsumi et al. 1986), therefore they are enriched relative to the immobile HFSE and rare earth elements (REE; Gill 1981; Pearce 1982; Ellam & Hawkesworth 1988; Hawkesworth et al. 1994; Pearce & Peate 1995). This is reflected by negative anomalies in HFSE and the relative enrichment of LILE in trace element patterns in mantle-normalized multielement diagrams (Fig. 3). Addition of pelagic or terrestrial sediment to the mantle has a fairly similar geochemical effect to

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM

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Fig. 2. Geodynamic settings for generation of calc-alkaline magmas: (a) magmatism at active subduction; (b) lithospheric extension-related magmatism following an earlier subduction event; (e) slab break-off magmatism; (d) post-collisional lithospheric delamination-related magmatism.

aqueous fluid metasomatism (increase of LILE); however, it usually also results in a marked enrichment of Th relative to Nb (Elliott et al. 1997) and it strongly influences the radiogenic and stable isotope ratios. Subduction-related magmas often have

Fig. 3. N-MORB (Pearce & Parkinson 1993) normalized trace element pattern of selected 'subduction-related' volcanic rocks. Mariana (volcanic arc): Elliott et al. (1997); Central Andes (active continental margin): Davidson et al. (1990); Columbia River Trough (intracontinental extensional area): Bradshaw et al. (1993). The arrows represent the typical depletion and enrichment features in 'subduction-related' volcanic rocks.

relatively high 87Sr/ 86Sr, low 143Nd/ 144Nd and relatively high 207 Pb/ 204Pb and 208Pb/ 204Pb isotope ratios compared with midocean ridge basalts (MORB), consistent with interaction with a continental crustal component. This component could be introduced to the mantle via subduction of sediment, or could enter the mantle-derived magmas as they pass through the continental crust. Such 'subduction-related' geochemical fingerprints characterize, however, not only the calc-alkaline and low-K tholeiitic magmas in active subduction zones, but also the potassic and ultrapotassic rocks as well as silicic igneous rocks, which are formed in anorogenic or post-collisional settings. In this case, these trace element and isotopic signatures were inherited from the mantle source region modified previously by fluids released from subducted slab (Johnson et al. 1978; Hawkesworth et al. 1995). Reactivation (i.e. partial melting of such metasomatized mantle) could occur as a result of decompression in a thinning lithosphere during extension (Fig. 2b) or owing to the heat flux of upwelling hot asthenospheric mantle material. Magmas with 'subduction-related' geochemical signatures can also be generated in syn- and post-collisional tectonic setting. A particular case is the generation of a slab-free region beneath a continental margin as a result of detachment of subducted oceanic lithosphere ('slab break-off'; Fig. 2c). This can occur when continental lithosphere enters the subduction zone (Davies & v o n Blanckenburg 1995; Wong et al. 1997). It results from the tensional force between the buoyant continental lithosphere and the denser oceanic slab, and can cause the formation of calc-alkaline and ultrapotassic magmas and crustal-derived silicic melts (Davies & von Blanckenburg 1995; von Blanckenburg & Davies 1995). Magmatism related to slab break-off is usually localized and instantaneous following the detachment. Melt generation occurs as a consequence of upwelling of hot

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asthenosphere into the void left between the separating lithospheres. The lithosphere of the overriding plate is conductively heated by the asthenospheric mantle flow and this can result in various degrees of melting within the metasomatized lithospheric mantle. Intrusion of mantle-derived mafic magmas into the thick continental crust could initiate crustal anatexis and formation of silicic magmas ('anatectic rhyolites or granites'). Delamination or convective removal of the lithospheric mantle could also lead to generation of magmas with 'subduction-related' geochemical features (e.g. Turner et al. 1996, 1999; Chalot-Prat & G~rbacea 2000; Fig. 2d). Removal of lower lithosphere results in upwelling of hot asthenosphere, which heats the overlying cooler continental lithosphere. Partial melting could occur first in the volatile-rich portions of the lithospheric mantle, producing LILE-enriched magmas. In summary, magmas with 'subduction-related' geochemical signatures could be formed in various tectonic settings. The common feature of these magmas is explained by the similar nature of their mantle source regions. Fluid-related metasomatism could take place just before the melt generation or be associated with an older subduction event.

Cenozoic subduction zones in the Alpine-Mediterranean region Figure 4 shows the main Tertiary-Quaternary subduction systems in Southern Europe. In general, their development can be explained in the context of convergence of the European and African plates, whereas in the easternmost part of the region, continental collision between Arabia and Eurasia could have played an important role in the evolution of the Anatolian-Aegean system (Jolivet et al. 1999). However, Gueguen et al. (1998) and Doglioni et al. (1999) argued that this relative convergence between Africa and Europe could not be the main mechanism resulting in subduction in the Western-Central Mediterranean, because the Apenninic arc migrated eastward faster than the postulated north-south convergence of Africa relative to Europe. Within Southern Europe, three main orogenic systems can be distinguished that can be divided into further local subduction systems (Fig. 1), as follows. (I) Western-Central Mediterranean: (I. 1) Beric- Alboran- Rif province (Western Mediterranean); (I.2) Apennines-Maghrebides subduction zone (Central Mediterranean);

Fig. 4. Migrationof subduction zones in the Alpine-Mediterranean region during the Tertiaryto Quaternary (after Wortel & Spakman 2000).

(II) Alps- Carpathian-Pannonian (ALCAPA) system; (II. 1) Alpine zone (Periadriatic-Insubric Line); (II.2) Carpathian-Pannonian region; (III) Dinarides-Eastern Mediterranean system: (III. 1) Dinarides-Hellenides; (III.2) Aegean-Anatolian region (Eastern Mediterranean). The most characteristic features of this area are the arcuate orogenic belts and the opening of extensional basins within the overall compressional regime. This region is characterized by mobile subduction zones, where migration of the arcs could be up to 800 km (Gueguen et al. 1998; Fig. 4). At the present day, active subduction occurs only beneath Calabria and the Aegean arc; in the other regions subduction has ceased and in most cases the current geodynamic setting is post-collisional. Magmatic activities occurred at various stages during the evolution of these subduction systems, but much of it appears to belong to the postcollisional stages. Subduction started in the Alpine region during the Early Cretaceous, when closure of the Tethyan oceanic basins took place by eastward to southeastward subduction beneath the AustroalpineApulian plate (Dercourt et al. 1986). In the Dinarides, eastward subduction of part of the Tethyan Vardar ocean occurred during the Late Mesozoic to Early Palaeogene (Karamata & Krsti6 1996; Karamata et al. 2000). Further east, north-dipping subduction formed the Pontide volcanic arc in Western Anatolia from the Late Cretaceous to the Palaeocene (~eng6r & Yilmaz 1981). These subduction processes were followed by continental collision stage during the Eocene in each region. In the eastern part of the Alpine region, roughly south-dipping subduction was still active (Carpathian subduction zone), where an oceanic embayment was present (Csontos et al. 1992; Fodor et al. 1999). This weak lateral boundary could have allowed eastward lateral extrusion of the crustal block from the compressive Alpine regime (Ratschbacher et al. 1991). The advancing subduction style changed to a retreating one during the Early Miocene (Royden 1993). Retreat of the subduction zone beneath the Carpathians enhanced the lateral movement of the North Pannonian block towards the NE, followed by back-arc extension during the Mid-Miocene (Horvfith 1993). Core-complex-type and back-arc extension resulted in the formation of the Pannonian Basin underlain by thin (50-80km) continental lithosphere (Tari et al. 1999). 'Soft' collision (Sperner et al. 2002) of the North Pannonian block with the European continent occurred during the Late Badenian (c. 13 Ma; Jifi6ek 1979), whereas subduction was still active along the Eastern Carpathians. Roll-back of the subducted slab resulted in further east-west extension in the back-arc area (Royden et al. 1982). Subduction ceased beneath the Eastern Carpathians during the Late Miocene (c. 13 Ma). Post-collisional slab detachment is considered to have occurred gradually from west to ESE as a zipper-like process (Tomek & Hall 1993; Mason et al. 1998; Seghedi et al. 1998; Wortel & Spakman 2000; Sperner et al. 2002). The slab break-off is now in the final stage beneath the southern part of the Eastern Carpathians (Vrancea zone), where the detaching near-vertical subducted slab causes intermediate depth seismicity (Oncescu et al. 1984; Oncescu & Bonjer 1997; Sperner et al. 2001). ChalotPrat & G~rbacea (2000) suggested partial delamination of the lithospheric mantle beneath this region. Formation of the Western-Central Mediterranean subduction system started during the Late Oligocene, when the Alpine subduction terminated. The Apennines-Maghrebides west-directed subduction formed along the back-thrust belt of the pre-existing Alps-Betics orogen (Doglioni et al. 1999). At this time, the Betics reached a collisional stage and, as a result, the orogenic crust thickened considerably (>50kin; Vissers et al. 1995). During the Early Miocene, rapid post-collisional extension took place, forming the Alboran basin underlain by thin continental

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM lithosphere. However, the mechanism of this process is highly debated. The competing geodynamic models involve: (1) back-arc extension behind the westward retreat of an east-directed subduction zone along the present Gibraltar arc or the Horseshoe seamounts in the eastern Atlantic (Royden 1993; Lonergan & White 1997; Gelabert et al. 2002; Gill et al. 2004); (2) detachment of the subducted slab (Spakman 1990; Blanco & Spakman 1993; Carminati et al. 1998; Zeck 1999; Calvert et al. 2000); (3) delamination of the lithospheric mantle (Garcia-Duefias et al. 1992; Docherty & Banda 1995); (4) convective removal of the lower part of the overthickened lithosphere (Platt & Vissers 1989; Turner et al. 1999). To the east, the presence of a Tethyan oceanic basin allowed the retreat of the Apennines-Maghrebides subduction zone (Carminati et al. 1998; Gueguen et al. 1998). Gelabert et al. (2002) argued that longitudinal shortening controlled the development of this arcuate subduction belt. They suggested that the subducting slab was split into two main fragments (Apennines and Kabylian slabs) retreating east and SE, respectively. Slab roll-back is explained by the sinking of dense Mesozoic oceanic lithosphere as a result of gravitational instability (Elsasser 1971; Malinverno & Ryan 1986; Royden 1993), by global eastward asthenospheric mantle flow (Doglioni 1992) or by lateral expulsion of asthenospheric material that was shortened and squeezed by plate convergence (Gelabert et al. 2002). Lithospheric extension behind the retreating subduction zone resulted in the formation of the Liguro-Provenqal, Algerian and Tyrrhenian basins underlain by newly formed oceanic crust, whereas the Valencia trough is underlain by thin continental lithosphere. Continental collision is thought to be associated with slab detachment along the north African margin during the Mid-Miocene (c. 16 Ma; Carminati et al. 1998; Coulon et al. 2002). Removal of the southward component of roll-back induced the eastward migration of the Apenninic arc accompanied by eastward migration of extension behind the subduction zone. Wortel & Spakman (1992) and van der Muelen et al. (1998) suggested Late Miocene-Pleistocene slab detachment beneath the Apennines based on seismic tomographic models and the lateral shifts of Apenninic depocentres. In contrast, Doglioni et al. (1994) emphasized the different roll-back rates along the arc, splitting it at least two 'sub-arc' portions. The subducting slab of the Ionian ocean still continues beneath Calabria. Behind it, new oceanic crust has been formed beneath the Vavilov and Marsili basins. Seismic tomographic models show positive seismic anomalies above the 670 km discontinuity beneath the entire Mediterranean region including the Pannonian and Aegean areas (Wortel & Spakman 2000; Piromallo et al. 2001; Piromallo & Morelli 2003). This is interpreted as accumulation of subducted residual material. In contrast to the widely accepted subduction-related models, a sharply different geodynamic scenario (i.e. a relationship with continental extension and/or upwelling mantle plume) has also been suggested to explain the evolution of the Central Mediterranean (Vollmer 1989; Lavecchia & Stoppa 1996; Ayuso et al. 1998; Lavecchia et al. 2003; Bell et al. 2004). Lavecchia et al. (2003) argued that the deformation style of the central Apennine fold-and-thrust belt, the absence of an accretionary wedge above the assumed subduction plane and the occurrence of ultra-alkaline and carbonatitic magmas within the Apennine mountain chain are evidence against the classic subduction-related models. They proposed that plume-induced lithospheric stretching and local-scale rift push-induced crustal shortening form a viable alternative model for the evolution of the Central Mediterranean region (Lavecchia et al. 2003; Bell et al. 2004). Closure of the Tethyan (Vardar) oceanic branches occurred during the Late Cretaceous-Palaeocene in the Dinaride region followed by collisional (Eocene) and post-collisional (Oligocene-Early Miocene) stages (Cvetkovi6 et al. 2004). In the Aegean-Anatolian region, north-dipping subduction terminated

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by the Eocene, when continental collision occurred along the Vardar-Izmir-Ankara suture zone (~eng6r & Yilmaz 1981). The subduction zone has migrated southwestward and its present position is found along the Hellenic-Cyprus arc (Doglioni et al. 2002). A continuous descending lithospheric slab was detected beneath the Hellenic arc by seismic tomography (Wortel & Spakman 1992, 2000). The subducted slab appears to cross the 670 km discontinuity and penetrates the lower mantle. Further east, collision between the Arabian and Anatolian plates took place during the Eocene ($engrr & Kidd 1979; Pearce et al. 1990). This collision also led to the tectonic escape of the Anatolian plate by right-lateral strike-slip movement along the North Anatolian Fault and left-lateral strike-slip along the East Anatolian Fault during the Mid-Miocene (McKenzie 1972; Dewey & Sengrr 1979). It was accompanied by widespread crustal extension and lithospheric thinning in Western Anatolia and the Aegean. The reason for the extension is, however, highly controversial. Dewey (1988) suggested gravitational collapse of the overthickened lithosphere. The orogenic collapse model was also accepted by Seyito~lu & Scott (1996) and Gautier et al. (1999), but they suggested that it occurred earlier (i.e. during the latest Oligocene to Early Miocene (24-20 Ma) time) and therefore it could not be associated with the tectonic escape of the Anatolian block. Other workers have emphasized the back-arc type extension of the Aegean region behind the retreating subduction along the Hellenic arc (McKenzie 1978; Le Pichon & Angelier 1979; Meulenkamp et al. 1988; Pe-Piper & Piper 1989). In Western Anatolia, Aldanmaz et al. (2000) invoked delamination of the lower lithosphere to explain the extension and related magmatism. Doglioni et al. (2002) ruled out the influence of both the westward extrusion of the Anatolian block and the collapse of overthickened lithosphere, and suggested an alternative geodynamic scenario for the Eastern Mediterranean. They interpreted the extension in the Aegean-Western Anatolian region as a result of the differential convergence rates between the northeastward-dipping subduction of Africa relative to the disrupted Eurasian lithospheres. Extension could be attributed to the faster southeastward motion of Greece relative to Anatolia. Thus, Doglioni et al. (2002) argued that the Aegean region cannot be considered as a classic back-arc basin. The next sections will outline the main characteristics of the Tertiary to Quaternary subduction systems in Southern Europe, the geochemical features of the magmatism, and the possible geodynamic relationships between magmatism, subduction and post-collisional processes. Figure 5 summarizes the age distribution of magmatism in this region, separated into orogenic and anorogenic types. Alpine subduction system

Subduction of Tethyan oceanic slabs occurred beneath the Alps during the Late Cretaceous to Palaeogene; however, no prominent subduction-related volcanism appears to have taken place during this period. The only evidence for subduction-related volcanic eruptions comes from Early Eocene andesitic clasts found in flysch sediments (Waibel 1993; Rahn et al. 1995). A characteristic feature of the Alpine collisional orogen is the occurrence of a chain of Oligocene to Early Miocene intrusions and dykes along the Periadriatic and Insubric lines. They continue eastward in the Pannonian Basin along the Balaton line (Downes et al. 1995; Benedek 2002) and southeastwards in the Dinarides (Pamid et al. 2002). These igneous rocks have a bimodal character (granodioritic-tonalitic intrusions and basaltic dykes; Exner 1976; Cortecci et al. 1979; Bellieni et al. 1981; Dupuy et al. 1982; Beccaluva et al. 1983; Ulmer et al. 1983; Kagami et al. 1991; Mtiller et al. 1992; von Blanckenburg & Davies 1995;

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Paleocene

Eocene

O/igocene Miocene

Pliocene Quat Betics Alboran Rif ValenciaTrough Provence Sardinia-Corsica Tyrrhenianbasin Tuscan Region Roman-CampanianProvinces Aeol:ianIslandsand seamounts Sicily Alps (Periadriatic-Insubricline & Veneto) Western Carpathians& PannonianBasin Eastern Carpathians Apuseni Dinarides Rhodope-Thrace NE-Aegean - WesternAnatolia Central Anatolia South Aegean arc

I

I 70

60

50

40

30

20

lO

0

Age (Ma) t

Calc-alkalineto shoshonitic magmatism ('orogenic type')

j

Alkaline(sodic) magmatlsm ('anorogenic type')

Berger e t al. 1996). In addition, calc-alkaline andesites, shoshonites and ultrapotassic rocks (lamproites) also occur in subvolcanic facies (Deutsch 1984; Venturelli e t al. 1984b; Altherr e t al. 1995). All of these igneous rocks are characterized by 'subduction-related' geochemical features. Suggested models for the origin of these igneous rocks include subduction (e.g. Tollmann 1987; Kagami e t al. 1991; Waibel 1993), extension (e.g. Laubscher 1983), and gradual slab detachment (von Blanckenburg & Davies 1995; von Blanckenburg e t al. 1998). The source regions of the primary magmas of the Periadriatic line are inferred to be in the lithospheric mantle (Venturelli e t al. 1984b; Kagami e t al. 1991 ; yon Blanckenburg 1992). Mafic melts could have subsequently mixed with silicic magmas generated in the lower crust. Alkaline mafic rocks ('anorogenic' type) crop out only south of the Eastern Alps, in the Veneto region (De Vecchi & Sedea 1995; M i l a n e t al. 1999; Macera e t al. 2003; Fig. 1). The volcanism occurred in two stages, from the Late Palaeocene to Early Oligocene (30-35 Ma) and during the Early Miocene. It resulted in alkaline and tholeiitic basalts and basanites with subordinate trachytes and rhyolites (De Vecchi & Sedea 1995). The mafic volcanic rocks show an ocean island basalt (OIB)-like composition without any sign of subduction-related component. De Vecchi & Sedea (1995) and Milani e t al. (1999) interpreted this volcanism as related to lithospheric extension in the Southern Alps (Zampieri 1995). In contrast, Macera e t al. (2003) invoked slab detachment and the ensuing rise of a deep mantle plume into the lithospheric gap.

Betic-Alboran-Rif

province

(Western

Mediterranean)

Tertiary to Quaternary volcanic rocks in the Western Mediterranean are found in central Spain (Calatrava province), the Olot

qr Ultrapotassic magmat~sm

W M

C M

A L C A P A

D E

M

Fig. 5. Age distributionof the Tertiary to Quaternary magmatismin the AlpineMediterranean region. Data are from Bellon et al. (1983), Fytikas et al. (1984), Beccaluva et al. (1985, 1987, 1991), Di Battistini et al. (1987), Peccerilloet al. (1987), Aparico et al. (1991), Conticelli & Peccerillo (1992), Martf et al. (1992), Seyito~lu& Scott (1992), Serri et al. (1993), Louni-Haciniet al. (1995), Pamid et al. (1995, 2002), P~cskay et aI. (1995), Christofideset al. (1998), E1 Bakkali et al. (1998), Harkovska et al. (1998), Marchev et al. (1998), von Blanckenburget al. (1998), Wilson & Bianchini(1999), Aldanmaz et al. (2000), Ro~u et al. (2001), Coulon et al. (2002), Cvetkovid et al. (2004), Duggen et al. (2004), and further references therein. WM, Western Mediterranean; CM, Central Mediterranean;ALCAPA, Alps- Carpathians-Pannonian region; DEM, Dinafides and Eastern Mediterranean.

region, the Valencia trough, SE Spain (Betics), the Alboran basin and along the coast of Northern Africa (Morocco to Algeria; Fig. 1). The Calatrava and Olot regions are characterized by Late Miocene to Quaternary alkaline basaltic and leucititic rocks (Cebrifi & Lopez-Ruiz 1995; Cebri~ e t al. 2000), similar to those occurring in the European Rift Zone (Wilson & Downes 1991). In the other areas calc-alkaline, high-K calc-alkaline, shoshonite and lamproites can be found in addition to late-stage alkali basalts. Duggen e t al. (2003) pointed out that the transition of calc-alkaline ('orogenic') to alkaline ('anorogenic') magmatism (6.3-4.8 Ma) was coeval with the Messinian salinity crisis (5.96-5.33 Ma; i.e. the desiccation of the Mediterranean sea as a result of closure of the marine gateway). The 'orogenic' volcanism started around Malaga with intrusion of tholeiitic (basalts to andesite; Fig. 6) dykes into the Alboran block during the early Oligocene. In addition, high-K dacites also occur in this area. It was followed by volcanism in the Valencia trough during the Late Oligocene that continued during the Miocene and to the present. Calc-alkaline volcanism forming mostly dacitic to rhyolitic pyroclastic deposits characterized the first stage of volcanic activity, whereas alkaline basaltic magmas erupted during the later volcanic stage (Mart/et al. 1992). The alkaline basalts have intraplate (OIB) geochemical affinity and often contain ultramafic xenoliths. This scenario could be explained by progressive extension of the continental lithosphere and a change of the source region from lithospheric to asthenospheric. The Alboran basin is also underlain by thin continental crust similar to the Valencia trough. Volcanic rocks on Alboran island and the sea floor have been dated between 11 and 7 Ma (Duggen e t al. 2004). They are mostly low-K tholeiitic rocks with clear 'subduction-related' geochemical features (Duggen e t al. 2004; Gill e t al. 2004; Fig. 7). More widespread calc-alkaline

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM

173

Fig. 6. SiO2 v. K20 diagrams (Gill 1981; Sh, shoshonite series; HKCA, High-K calc-alkaline series; CA, calc-alkaline series; Th, low-K tholeiitic series; B, basalt; BA, basaltic andesite; A, andesite; D, dacite; R, rhyolite) for the 'subduction-related' volcanic rocks from the various segments of the Alpine-Mediterranean region. The wide compositional variations should be noted. Western Mediterranean sources: Nixon et al. (1984), Venturelli et al. (1984a), Zeck et al. (1998), Benito et al. (1999), Turner et al. (1999), Duggen et al. (2004), Gill et al. (2004). Central Mediterranean sources (A1-Fi-Sa-Pa: Alicudi, Filicudi, Salina and Panarea): Rogers et al. (1985), Ellam et aL (1988), Crisci et al. (1991), Conticelli & Peccerillo (1992), Francalanci et al. (1993), Peccerillo et al. (1993), Ayuso et al. (1998), Del Moro et aL (1998), De Astis et al. (2000), Gertisser & Keller (2000), Downes et al. (2001). Carpathian-Pannonian region sources: Downes et al. (1995a), Mason et al. (1996), Harangi et al. (2001, 2005), Seghedi et al. (2001, 2004). Dinarides-Eastern Mediterranean sources (WA, Western Anatolia; CA, Central Anatolia): Mitropoulos et al. (1987), Huijsmans et al. (1988), Pe-Piper & Piper (1989), Seyito~hi & Scott (1992), Wilson et al. (1997), Francalanci et al. (1998), Kiirkqtioglu et al. (1998), Tankut et al. (1998), Temel et al. (1998a,b), Aldanmaz et al. (2000), Cvetkovi6 et al. (2004).

Fig. 7. Normal-MORB (N-MORB; Pearce & Parkinson, 1993) normalized multi-element diagrams for representative samples of the various segments of the Alpine-Mediterranean region. (For data sources see Fig. 6.) Carp., Carpathians; UP, ultrapotassic.

174

S. HARANGIET AL.

to shoshonitic and ultrapotassic volcanism occurred on the SE coast of Spain and in Northern Africa from the Early Miocene to Pliocene (Zeck 1970, 1992, 1998; Nixon et al. 1984; Venturelli et al. 1984a, 1988; Hertogen et al. 1985; Di Battistini et al. 1987; Louni-Hacini et al. 1995; E1 Bakkali et al. 1998; Benito et al. 1999; Turner et al. 1999; Coulon et al. 2002; Duggen et al. 2004; Gill et al. 2004; Fig. 6). Calc-alkaline volcanism was associated with intrusion of granitoid magmas in Northern Africa (Fourcade et al. 2001) and southern Spain (Zeck et al. 1989; Duggen et al. 2004). The calc-alkaline volcanism resulted in andesites and dacites with subordinate rhyolites and shoshonites (Fig. 6). Sporadic cordierite- and garnet-bearing dacites were interpreted as anatectic magmas (Zeck 1970, 1992). Late Miocene ultrapotassic lamproites are found in the central and northern part of the calc-alkaline volcanic belt of SE Spain (Nixon et al. 1984; Venturelli et al. 1984a, 1988; Hertogen et al. 1985). Throughout the region, sporadic eruptions of alkaline mafic magmas followed the calc-alkaline magmatism (El Bakkali et al. 1998; Coulon et al. 2002; Duggen et al. 2004). The geodynamic setting of the Western Mediterranean calc-alkaline volcanic activity is ambiguous. The models can be divided into the following groups: (1) subduction-related; (2) subduction break-off; (3) delamination of lithospheric mantle as a result of gravitational collapse; (4) convective removal of the lower lithosphere. Torres-Roldfin et al. (1986), Royden (1993), Lonergan & White (1997), Duggen et al. (2003, 2004) and Gill et al. (2004) assumed that contemporaneous subduction occurred with the calc-alkaline volcanism. Geophysical data indicate an east-dipping subducted slab (Gutscher et al. 2002) beneath the Alboran region. Duggen et al. (2004) and Gill et al. (2004) emphasized that the 'subduction-related' nature and particularly the strong depletion in the light REE (LREE) and HFSE of the Alboran tholeiites (Fig. 7) could only be explained by formation in a metasomatized mantle wedge above a subducted slab. Furthermore, they assumed that all the other calc-alkaline volcanic rocks in the Betic-Rif province could be generated in the same geodynamic setting. Blanco & Spakman (1993) and Calvert et al. (2000) argued, however, that the seismic tomography models show a detached near-vertical lithospheric slab from about 180200 km down to the 670 km discontinuity beneath the Alboran region. Zeck (1996) considered that slab break-off could have had a major role in melt generation. Influx of hot asthenospheric mantle into the widening gap above the sinking slab induced partial melting in the overlying lithosphere (particularly in the lower crust). The close relationship between the distribution of volcanism in the Alboran volcanic province and the surface projection of the sinking slab was used by Zeck (1996) to support this model. Fourcade et al. (2001) and Coulon et al. (2002) also invoked slab break-off to explain the calc-alkaline to alkaline magmatism in northern Algeria. Other workers (Venturelli et al. 1984a; Platt & Vissers 1989; Zeck 1996; Benito et al. 1999; Turner et al. 1999) argued that the Betic-Alboran volcanism was post-collisional, following Late Cretaceous to Oligocene subduction and Late Oligocene to Early Miocene continental collision. Benito et al. (1999), Turner et al. (1999) and Coulon et al. (2002) suggested that the primary melts were generated in the lithospheric mantle, which had been metasomatized previously by fluids derived from subducted pelagic sediments. These mantle-derived magmas subsequently mixed with crustal melts. Zeck (1970, 1992, 1998) argued for a crustal anatectic origin for the calc-alkaline magmas of southern Spain. Platt & Vissers (1989), Benito et al. (1999) and Turner et al. (1999) emphasized that melt generation occurred by decompression melting caused by extensional collapse of the overthickened orogenic wedge or convective removal of the lithospheric root. In North Africa, E1 Bakkali et al. (1998) also suggested an extension-related origin for the calc-alkaline to potassic magmas of the Eastern Rif (Morocco).

Central Mediterranean

(Italy)

Tertiary-Quaternary volcanism in the Central Mediterranean resulted in extremely variable magmatic rocks including tholeiites (Vavilov basin, Tyrrhenian basin), calc-alkaline to shoshonitic (Sardinia, Aeolian Islands, Roman Province), ultrapotassic (Corsica, Central Italy) and anatectic rhyolites (Tuscany; Serri 1990; Peccerillo 1999, 2003; Figs 1 and 6). Alkali basaltic rocks with OIB chemistry also occur sporadically in Sardinia (Rutter 1987; Lustrino et al. 2000), the southern Tyrrhenian basin (Serri 1990; Trua et al. 2003), eastern Sicily (Etna and Hyblean plateau; Carter & Civetta 1977; Tonarini et al. 1995; D'Orazio et al. 1997; Tanguy et al. 1997; Trua et al. 1998) and in the Pantelleria rift (Esperanca & Crisci 1995; Civetta et al. 1998). In addition, minor occurrences of carbonatites and melilitites have been described in the central Apennines east of the Roman Province (Stoppa & Lavecchia 1992; Stoppa & Cundari 1995; Stoppa & Woolley 1996). The carbonatitic nature of these rocks has been questioned, however, by Peccerillo (1998) who suggested that they could represent a mixture of silicate magmas and carbonate material, and could be classified as ultrapotassic rocks of kamafugitic affinity. The strongly undersaturated haiiyne-bearing alkaline volcanic rocks of Mt. Vulture (De Fino et al. 1986; Serri 1990; Melluso et al. 1996) also have an exotic position (Fig. 1) and distinct magma source region compared with the Roman Province rocks. Volcanism started in the Early to Mid-Miocene in Sardinia with the eruption of tholeiitic and calc-alkaline magmas (Dostal et al. 1982; Morra et al. 1997; Downes et al. 2001; Fig. 5). The 14 Ma Sisco lamproite in northern Corsica represents the oldest ultrapotassic rock in the Central Mediterranean. After a few million years quiescence, the volcanism rejuvenated in the Tyrrhenian basin, the Tuscan region and in southeastern Sicily (Hyblean Mts) at about 7 - 8 Ma. On the west coast of Italy a gradual younging of the volcanism can be observed towards the SE, with still active volcanoes in Campania (Campi Flegrei, Vesuvius; Santacroce et al. 2003). The distinct alkaline volcanic rocks of Mt. Vulture were formed at 0.7-0.l Ma (Melluso et al. 1996). Volcanic activity in Sardinia rejuvenated with eruption of alkaline mafic magmas from 5.5 to 0.1 Ma (Di Battistini et al. 1990; Lustrino et al. 2000). A southeastward shift of volcanism has been pointed out by Argnani & Savelli (1999). Opening of the southern Tyrrhenian basin was accompanied by formation of seamounts consisting of enriched (E)-MORB to OIB type mafic rocks (Vavilov, Marsili; Serri 1990; Trua et al. 2003). Active volcanism is taking place in the Aeolian Islands (Stromboli, Vulcano), in Etna and the Sicily Channel (Pantelleria). Although most workers suggest that subduction has played an important role in the evolution of the Central Mediterranean (e.g. Keller 1982; Doglioni 1991; Serri et al. 1993), calc-alkaline volcanic rocks are volumetrically subordinate within the magmatic suites. Instead, the characteristic rocks are potassic to ultrapotassic (Fig. 6). Furthermore, the Tertiary to Quaternary volcanic rocks of this region show an extremely variable trace element and isotope chemistry (Peccerillo 2003). The close temporal and spatial relationship of this wide range of magmas indicates a heterogeneous mantle source metasomatized during several distinct events (Peccerillo 1985, 1999; Serri et al. 1993). The strongly potassic character of many of the magmas has been explained either by source contamination by subducted continental crustal material (Peccerillo 1985; Ellam et al. 1989; Conticelli & Peccerillo 1992; Serri et al. 1993; De Astis et al. 2000) or by metasomatism of deep mantle-derived melts (Vollmer 1989; Stoppa & Lavecchia 1992; Ayuso et al. 1998). West-dipping subduction of oceanic lithosphere and possibly thinned continental lithosphere is considered to have terminated in the Late Miocene (c. 13 Ma), thus most of the volcanism in the Central Mediterranean can be regarded as post-collisional.

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM

Evidence for subduction includes the introduction of continental crustal material into the mantle sources of the magmas (Peccerillo 1985, 1999) and the detection of high-velocity material either continuously extending from the surface (beneath Calabria) and accumulating in the transition zone (e.g. Spakman e t al. 1993; Piromallo e t al. 2001; Piromallo & Morelli 2003). Serri e t al. (2003) proposed that delamination and subduction of the Adriatic continental lithosphere related to the continuing collision in the northern Apennines could be a viable mechanism to explain the incorporation of crustal material in the mantle source and the eastward migration of magmatism in central Italy. Active subduction in the region occurs in Calabria. Keller (1982) directly related the recent volcanism in the Aeolian Islands to active subduction process. However, the Aeolian Islands lie 2 0 0 - 3 0 0 km above the Benioff zone (Anderson & Jackson 1987) on thinned continental basement (Schutte 1978). Furthermore, they form rather a ringshaped structure considering also the submarine seamounts at the southern margin of the Marsili basin, which is characterized by oceanic crust. The active volcanoes are located along strike-slip tectonic lines (Beccaluva e t al. 1982; Gasparini e t al. 1982). Thus, an alternative hypothesis for the Aeolian volcanism is a relationship with a back-arc environment, where magma generation is attributed to asthenospheric domal uplift developing along a NW-SE-trending extensional tectonic zone (Crisci e t al. 1991; Mazzuoli e t al. 1995). Nevertheless, subduction could have had a major, probably indirect, influence on the genesis of the magmas (release of aqueous fluids from the downgoing slab and metasomatism of the upper mantle; injection of subducted sedimentary component into the upper mantle; Ellam e t al. 1989; Francalanci e t al. 1993). Compositional features (ratios of incompatible trace elements and radiogenic isotope ratios; Fig. 8) of the potassic rocks of Stromboli are similar to those of

175

the alkaline volcanic rocks of Campania (Vesuvius and Phlegrean Fields), indicating common mantle source regions consisting of a mixture of intraplate and subducted slab-derived (continental sediment) components (De Astis e t al. 2000; Peccerillo 2001). The post-collisional volcanism in Italy has been interpreted by Wortel & Spakman (1992, 2000) as being due to gradual slab detachment, based on the absence of high-velocity structure considered to represent subducted slab beneath the Apennines, whereas a continuous slab was identified beneath southern Italy. However, Piromallo & Morelli (2003) argued that their better resolved model showed more vertical continuity of the fast structure in the top 200 km beneath the northern part of the Apennines. In contrast to the most popular subduction-related models, the presence of a mantle plume and related continental rifting was put forward by Lavecchia e t al. (2003) and Bell e t al. (2004). Gasperini e t al. (2002) also invoked upwelling of deep mantle material beneath southern Italy, but they combined it with the subduction scenario, suggesting a broad window in the Adria plate where deep mantle layers are channelled toward the surface. Lavecchia e t al. (2003) and Bell e t al. (2004) proposed that a plume arising from the core-mantle boundary could be trapped within the transition zone beneath the Ligurian-Tyrrhenian region. Asymmetric growth of the plume head within the transition zone as modelled by Brunet & Yuen (2000) could lead to a volume excess within the asthenosphere and an eastward mantle flow. This is thought to result in an eastward-migrating thinning of the overlying lithosphere. The rift-push forces generated on the eastern side of the extending system could be responsible for the fold-and-thrust belt structure beneath the Apennines. In this model, the high-velocity body above the 670 km depth (Piromallo e t al. 2001) was interpreted as reflection of compositional difference rather then abrupt change in the mantle temperature.

Fig. 8. 87sr/arsr v. 143Nd/144Nd diagrams for the Tertiary-Quaternary volcanic and plutonic rocks of the various segments of the Alpine-Mediterranean region. Data sources are as for Figure 6. Additional data sources are as follows. Western Mediterranean: Cebriit et al. (2000), Central Mediterranean: Carter & Civetta (1977), Hawkesworth & Vollmer (1979), Ellam & Harmon (1990), Esperanca & Crisci (1995), Tonarini et al. (1995, 2001), D'Orazio et al. (1997), Trua et al. (1998), Castorina et al. (2000), Lustrino et al. (2000), Alps-Carpathian-Pannonian region: Juteau et al. (1986), Kagami et al. (1991), von Blanckenburg (1992), yon Blanckenburg et al. (1992, 1998), Embey-Isztin et al. (1993), Harangi et al. (1995), Macera et al. (2003), Dinarides-Eastern Mediterranean: Briqueu et al. (1986), Gtilec (1991), Pamid et al. (1995). SV-Po-CF-Cu, melilitite-carbonatite association in San Venanzo, Polino, Colle Fabri, Cupaello; CAV, calc-alkaline volcanic rocks; DMM, depleted MORB mantle; EMI and EMIl, enriched mantle I and II; HIMU, high Ix (238U/2~ mantle components (Zindler & Hart 1986).

176

S. HARANGIETAL.

Lavecchia et al. (2003) suggested that the fast zones in the transition zone could be a highly depleted and dehydrated plume head, whereas the overlying asthenosphere was enriched by HzO-COz-rich fluids. However, Goes et al. (2000) proposed that the velocity variation in the mantle could be attributed mostly to changes in temperature, whereas the effect of mantle composition could be negligible (< 1%). Further integrated geophysical, structural and geophysical studies are needed to test the contrasting models for the evolution of the Central Mediterranean.

Carpathian-Pannonian

region

The Carpathian-Pannonian region (Fig. 1) shows many features that are similar to those of the Mediterranean subduction systems, such as arcuate and retreating subduction zones, formation of back-arc extensional basins and a wide range of magma types (e.g. Horv~ith & Berckhemer 1982; Csontos et al. 1992; Szab6 et al. 1992; Seghedi et al. 1998; Fodor et al. 1999; Tari et al. 1999; Bada & Horv~ith 2001; Harangi 2001a). Volcanic activity in this region started with eruption of Early Miocene rhyolitic magmas followed by contemporaneous calc-alkaline, silicic and sporadic ultrapotassic volcanism in the Mid- and Late Miocene (Fig. 5). Coeval calc-alkaline and alkaline mafic magmas were erupted during the Late Miocene to Quaternary (Szab6 et al. 1992; Prcskay et al. 1995; Seghedi et al. 1998, 2004; Harangi 2001b). The Miocene (21 - 13 Ma) rhyolitic volcanism resulted in extensive ignimbrite sheets. The rhyodacitic to rhyolitic pumices have 'subduction-related' geochemical features consistent with both mantle and crustal origin. The pyroclastic deposits also contain basaltic andesite and andesite lithic clasts, which are considered as cogenetic with the rhyolites. Harangi et al. (2005) interpreted their petrogenesis inferring mantle-derived mafic magmas mixed with variable amount of crustal melts. The silicic volcanism could represent the initiation of back-arc lithosphere extension (Lexa & Konern~ 1998; Harangi 2001a) or delamination of the lowermost lithosphere beneath the Pannonian Basin (Downes 1996; Seghedi et al. 1998). The decreasing age-corrected 87Sr/86Sr ratios of the pumices indicate a gradually decreasing crustal component in their genesis. A major feature of the region is the Carpathian arc, an arcuate belt of calc-alkaline volcanic complexes composed mostly of andesites and dacites (Fig. 6) along the northern and eastern margin of the Pannonian Basin. They were formed from the MidMiocene to the Quaternary and the last volcanic eruption occurred in the southernmost part of the East Carpathians only 10-40 ka ago (Fig. 5; Prcskay et al. 1995). The major and trace element compositions of these rocks show fairly similar character compared with the large variability of the Western and Central Mediterranean volcanic suites (Figs 6 and 7). However, there are major differences in spatial and temporal evolution and underlying lithospheric structure between the western and eastern segments of the Carpathian volcanic arc. These differences also appear in the geochemistry of the volcanic products, leading Harangi & Downes (2000) and Harangi (2001a) to suggest contrasting origins for the calc-alkaline magmas in the different segments. Calc-alkaline volcanism in the western Carpathian arc could be related directly to the main extensional phase of the Pannonian Basin (Lexa & Konern.~ 1998; Harangi 2001a; Harangi et al. 2001), whereas calc-alkaline magmas in the eastern Carpathian arc could have a closer relationship with subduction, particularly with gradual slab break-off (Mason et al. 1998; Seghedi et al. 1998, 2004). Gradual slab detachment was also proposed in the evolution of the Carpathian arc by other workers (Nem~ok et al. 1998; von Blanckenburg et al. 1998; Wortel & Spakman 2000; Sperner et al. 2002). Coexisting eruptions of alkaline basaltic and shoshonitic magmas in the southernmost part of

the east Carpathians led G~rbacea & Frisch (1998) and Chalot-Prat & G~rbacea (2000) to suggest partial delamination of the lower lithosphere beneath this area. In contrast to these models, Balla (1981), Szab6 et al. (1992) and Downes et al. (1995a) considered that melt generation in the whole calc-alkaline suite was a direct consequence of subduction of the European plate and occurred in the metasomatized mantle wedge above the downgoing slab. Calc-alkaline volcanic rocks also occur far from the Carpathian arc, in the inner part of the Pannonian Basin. Mid-Miocene andesites of the Apuseni Mountains are found about 200 km behind the volcanic front. Ro~u et al. (2001) and Seghedi et al. (2004) suggested that the location and compositions of these rocks is inconsistent with a typical subduction model and can be explained rather by decompression melting of the lower crust and/or of the enriched lithospheric mantle in an extensional regime. Seismic tomographic images indicate a low-velocity anomaly beneath the Carpathian-Pannonian region (except at the southeastern margin of the Carpathians) at shallow depth (Spakman 1990; Wortel & Spakman 2000; Piromallo & Morelli 2003). A fast anomaly was detected beneath the Eastern Carpathians in a depth range of 100-300 km, but it cannot be followed beneath the western Carpathian chain (Piromallo & Morelli 2003). Thus, no evidence is present for a detached subducted slab under the latter area, although Tomek & Hall (1993) interpreted the deep seismic reflection data as evidence for subducted European continental crust beneath the western Carpathians. In the southeastern part of the Carpathians, beneath the Vrancea zone, a weak fast anomaly is present, becoming more pronounced with increasing depth. The localized Vrancea slab is considered to represent the final stage of slab break-off (Wenzel et aL 1998; Sperner et al. 2001) beneath the east Carpathians. The oceanic slab is considered as either already detached from the surface (Wortel & Spakman 2000) or still attached to the continental lithosphere (Fan et al. 1998; Sperner et al. 2001). An approximately 150-200 km thick positive anomaly occurs between 400 and 600 km beneath the entire Carpathian-Pannonian region that is interpreted as accumulation of Mesozoic subducted slab material (Wortel & Spakman 2000).

Dinarides and Hellenides

A continuous belt of Tertiary igneous activity is present from the eastern Alps to the north Aegean crossing the southern Pannonian Basin (Slovenia and Croatia), the Dinarides (Serbia, Macedonia) and the Rhodope-Thrace region (Bulgaria and Greece; Fig. 1; Pami6 et al. 1995, 2002; Christofides et al. 1998, 2001; Harkovska et al. 1998; Marchev et al. 1998; Yilmaz & Polat 1998; Jovanovi6 et al. 2001; Prelevi6 et al. 2001; Cvetkovi6 et al. 2004). Subduction of part of the Tethyan Vardar Ocean occurred during the Late Mesozoic to Early Palaeogene, followed by Eocene collision and Oligocene to Pliocene post-collisional collapse (Karamata & Krsti6 1996; Karamata et al. 1999). This igneous belt comprises Eocene to Oligocene granitoid bodies and basanites, Oligocene to Miocene shoshonites, high-K calc-alkaline volcanic rocks and ultrapotassic rocks (lamproites and leucitites). Most of the Palaeogene granitoids have been interpreted as syncollisional magmas that underwent various degrees of crustal contamination (Christofides et al. 1998; Marchev et al. 1998; Pami6 et al. 2002). The Palaeocene-Eocene basanites in eastern Serbia often contain ultramafic xenoliths (Cvetkovi6 et al. 2001) and have major and trace element composition akin to OIB (Jovanovi6 et al. 2001). Similar xenolith-bearing alkaline mafic rocks also occur in the Rhodopes (Marchev et al. 1998). The primary magmas are inferred to originate in an enriched asthenospheric mantle source. Melt generation could have been triggered either by detachment of the subducted slab resulting in a slab window or by a short extensional event during the collisional phase (Jovanovi6 et al. 2001; Cvetkovi6 et al. 2004). The Oligocene to Early Miocene high-K calc-alkaline and shoshonitic series

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM comprises basalts, basaltic andesites and trachyandesites (Fig. 6). They have relatively high mg-values and high concentration of Ni and Cr, indicative of near-primary magmas. The presence of phlogopite-bearing ultramafic xenoliths clearly indicates a mantle origin for these melts, which have typical 'subductionrelated' compositions (Fig. 7). This signature could be inherited from the lithospheric mantle source metasomatized possibly during the post-collisional or collapse stage (Cvetkovid et al. 2004). The scattered ultrapotassic rocks (minettes, lamproites, leucitites, analcimites) show the most extreme enrichment of incompatible elements of all the Tertiary volcanic rocks in this region (Prelevid et al. 2001). They have fairly similar trace element patterns in the mantle-normalized diagrams to the high-K volcanic rocks, but with more pronounced anomalies. Thus, they could also represent magmas derived from metasomatized lithospheric mantle. Melt generation could be related either to slab break-off (Pamid et al. 2002) or to delamination of the lithospheric root (Cvetkovid et al. 2004). Seismic tomographic images indicate a high-velocity anomaly from about 100 km to 600 km beneath the southern Dinarides and Hellenides region, whereas a low-velocity anomaly was detected beneath the northern Dinarides (Spakman et al. 1993; Goes et al. 1999; Wortel & Spakman 2000). This feature has been interpreted as detachment of a subducted slab in the north, whereas it is still unbroken in the south and continues towards the south Aegean area (Wortel & Spakman 2000). Beneath the Dinarides a north- to NE-dipping subduction was proposed with the opposite polarity to that inferred beneath the Alps (Pamid et al. 2002). Stampfli et al. (2001) suggested that the Vardar Ocean (the Tethyan oceanic branch in the Dinaride-Hellenide region) and the Piedmont-Penninic Ocean (Alpine Tethyan oceanic branch) were not connected during the Mesozoic. Therefore, the two linear Palaeogene igneous belts along the Periadriatic line and along the Dinarides could not belong to the same subduction system.

Eastern Mediterranean

(Greece and Turkey)

Tertiary-Quaternary volcanic activity in this region was characterized by eruption of various magmas (alkaline mafic, calc-alkaline and high-K intermediate to silicic volcanic rocks and sporadic ultrapotassic rocks) in the Aegean and Western to Central Anatolia (Fig. 1; Fytikas et al. 1984; Doglioni et al. 2002). Volcanism occurred in two main phases: eruption of Oligocene to Mid-Miocene calc-alkaline to shoshonitic magmas followed by eruption of alkaline and calc-alkaline magmas during Pliocene to Recent times (Pe-Piper & Piper 1989; Seyito~lu & Scott 1992; Pe-Piper et al. 1995; Aldanmaz et al. 2000; Doglioni et al. 2002). In the Aegean-Western Anatolian region, volcanism started in the Oligocene following the Tethyan collision (Yilmaz et al. 2001). The calc-alkaline volcanism resulted mostly in andesitic to dacitic rocks, and was associated with emplacement of granitic plutons in NW Anatolia. Most of the volcanism occurred, however, during the Early to Mid-Miocene (20-14 Ma), when high-K andesitic to rhyolitic effusive and explosive volcanism with cogenetic plutons characterized the northern Aegean and the Western Anatolia (Figs 5 and 6; Fytikas et al. 1984; Pe-Piper & Piper 1989; Seyito~lu & Scott 1992; Wilson et al. 1997; Altunkayak & Yilmaz 1998; Aldanmaz et al. 2000; Yilmaz et al. 2001). In Western Anatolia, these igneous rocks are distributed mostly along the Izmir-Ankara suture zone. A Mid-Miocene (14-15 Ma) lamproite was reported by Savas~in et al. (2000). During the Mid- to Late Miocene, granitic plutonism took place in the Cycladic and Menderes massifs (Altherr et al. 1982; Innocenti et al. 1982; Delaloye & Bing61 2000). Following about 4 Ma quiescence, the volcanism resumed in the Late Miocene (10 Ma), when alkaline mafic magmas with OIB-like composition erupted mostly in the eastern Aegean and the western margin of Anatolia (Gtile~ 1991; Seyito~lu et al.

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1997; Aldanmaz et al. 2000; Alici et al. 2002). The last eruption of basanitic to phonotephritic magmas occurred in the Kula region at 1.1-0.02Ma (Gtileq 1991). In the Aegean region Pe-Piper et al. (1995) found a southward migration of volcanism. Calc-alkaline volcanism has characterized the Aegean volcanic arc from Pliocene to Recent times (Mann 1983; Barton & Huijsmans 1986; Briqueu et al. 1986; Mitropoulos et al. 1987; Huijsmans et al. 1988; Francalanci et al. 1998). Central Anatolia shows roughly the same volcanic history. MidMiocene to Pliocene high-K calc-alkaline andesitic to rhyolitic magmas erupted along major fault systems at Afyon, Konya and Cappadocia (Figs 5 and 6; Innocenti et al. 1975; Aydar et al. 1995; Alici et al. 1998; Temel et al. 1998a,b). In Cappadocia extensive dacitic to rhyolitic ignimbrite sheets were deposited from the Late Miocene to Quaternary, associated with large andesitic stratovolcanoes and alkali basaltic scoria cones and maars (Pasquare et al. 1988; Le Pennec et al. 1994; Aydar & Gourgaud 1998; Ktirkqtioglu et al. 1998; Temel et al. 1998b). Tertiary to Quaternary volcanism in the Aegean and Western to Central Anatolian region occurred mostly in a post-collisional setting and partly behind active subduction zones (Hellenic and Cyprean). The origin of the Miocene plutonic igneous rocks was interpreted as crustal anatexis related to high-T-medium-P metamorphism (Altherr et al. 1982; Innocenti et al. 1982; Br6cker et al. 1993; Delaloye & Bing61 2000). In general, the 'orogenic' volcanic rocks are potassic (high-K calc-alkaline to shoshonitic), whereas those that occur along the Aegean island arc are calc-alkaline (Fig. 6). Their trace element and isotopic compositions are consistent with involvement of a subduction component (Figs 7 and 8; Keller 1982; Briqueu et al. 1986; Mitropoulos et al. 1987; Huijmans et al. 1988; Gtileq 1991; Robert et al. 1992; Pe-Piper et al. 1995; Seyito~lu et al. 1997; Aldanmaz et al. 2000). On the other hand, the younger alkaline mafic volcanic rocks show an intraplate OIB nature (Seyito~lu & Scott 1992; Seyito~lu et al. 1997; Alici et al. 2002). Nevertheless, most workers consider that melt generation of both volcanic suites was mostly due to decompression melting because of extension of continental lithosphere. Initially, magmas with 'subductionrelated' geochemical signature were generated in the lithospheric mantle regions metasomatized by fluids and melts during an earlier subduction event. Perturbation of these metasomatic portions of the lower lithosphere could take place as a result of either lithospheric thinning to delamination of the lower thermal boundary layer allowing direct contact with upwelling hot asthenosphere (Seyito~lu & Scott 1996; Aldanmaz et al. 2000). In contrast, a closer relationship with subduction was invoked by Innocenti et al. (1975), Temel et al. (1998a,b) and Doglioni et al. (2002) to explain the calc-alkaline volcanism especially in Central Anatolia. They considered that the primary magmas originated in the mantle wedge above the subducting Africa plate and subsequently underwent assimilation and fractional crystallization to produce the intermediate to rhyolitic magmas. The Pliocene to Recent volcanic rocks in the Aegean arc seem to be a clearer candidate to have formed as a consequence of active subduction. Geophysical data clearly indicate a northeastward to northward dipping slab beneath the Aegean region (Wortel & Spakman 1992; Spakman et al. 1993). The present volcanic arc is located 130-140 km above the seismic Benioff zone (Makropoulos & Burton 1984), 200-250 km behind the subduction front. Lithospheric extension, however, may have played an important role in melt generation along the present arc, indicated by the underlying thin lithosphere and predominance of asthenosphere-derived uncontaminated mafic volcanic rocks in Santorini (Mitropoulos et al. 1987). In contrast, Briqueu et al. (1986) considered that there is a close relationship between the active volcanic arc and subduction, and assumed a contribution of a small amount of subducted sedimentary component in the genesis of the arc magmas. The late-stage alkaline basaltic rocks have a composition akin to OIB; therefore, they are interpreted as representing

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asthenosphere-derived magmas. These could originate either in places where lithospheric extension progressed further allowing partial melting of the uprising asthenosphere (Seyito~lu et al. 1997) or along strike-slip zones, where localized stretching could result in production of alkaline magmas (Aldanmaz et al. 2000). Doglioni et al. (2002) suggested that the shallow subducted slab beneath Anatolia could be folded by the isostatic rebound of the mantle beneath the extensional area. The stretching between Greece and Anatolia and the differential velocity of convergence with the underlying slab could have generated a sort of window, allowing upward rise and partial melting of the asthenosphere.

Discussion Petrogenetic features

The previous sections showed that a wide variety of magmatic rocks can be found in all of the different subprovinces of the Alpine-Mediterranean region. Most of them show typical 'subduction-related' composition as reflected by the elevated potassium content (Fig. 6) along with the enrichment of LILE and depletion of HFSE (Fig. 7). However, the large geochemical variability as shown in the SiO2 v. K20 diagrams (Fig. 6) implies that complex petrogenetic processes have operated, involving different mantle sources, contamination by different crustal material and different degrees of fractional crystallization. Another important feature of the orogenic magmatic suites is the common occurrences of potassic-ultrapotassic rocks in each subprovince. They are the most characteristic of the Central Mediterranean area. OIB-type alkaline sodic mafic magmas akin to those erupted in the European foreland (Wilson & Downes 1991, 2006) overlap spatially and temporally with the 'orogenic' volcanism, although they are most characteristic of the later magmatic phases. The majority of these alkaline mafic rocks clearly indicate a distinct mantle source regions unaffected or only slightly affected by subduction-related fluids. On the other hand, interpretation of the origin of the 'subduction-related' magmatic rocks in the Mediterranean region is more difficult, because of the lack of mafic undifferentiated rocks in many areas. In multi-element diagrams that are normalized to mid-ocean ridge basalts (N-MORB; Fig. 7), the Mediterranean orogenic rocks all show fairly similar features such as enrichment in LILE and depletion in HFSE (e.g. negative Nb anomaly). As discussed previously, these characters are signs of a subduction component in the genesis of the magmas. Subduction and subduction-related metasomatism of the mantle wedge can be contemporaneous with the magmatism, but could also precede the volcanic activity. However, this geochemical feature can also be interpreted as contamination by crustal material at shallow level. Radiogenic isotope ratios (e.g. 87 Sr/ 86 Sr, 143N d / 144Nd and PbPb isotope ratios) are not changed by closed-system petrogenetic processes such as partial melting and crystal fractionation; therefore they can be used to characterize the source region of the magmas and to detect possible crustal contamination. Recognition of the nature of the pre-metasomatized mantle source region could be important to constrain the geodynamic evolution of the volcanic areas. In the 1980s four main mantle end-member reservoirs were distinguished based on the radiogenic isotope variation of oceanic basalts (White 1985; Zindler & Hart 1986): depleted MORB-mantle (DMM), high tx (238U/2~ mantle (HIMU) and enriched mantle end-members (EMI and EMII). In addition to these mantle components a primitive mantle reservoir, called as PREMA (Zindler & Hart 1986) or FOZO (Hart et al. 1992) was also suggested to be present in the mantle. These mantle components could represent distinct parts of the mantle, although they could also be spatially related (Hart 1988). Among them, the HIMU and FOZO are often interpreted to relate to upwelling mantle plumes coming from the core-mantle boundary

(Hofmann & White 1982; Weaver 1991; Chauvel et al. 1992; Hart et al. 1992; Hofmann 1997). Alternatively, this geochemical feature can reflect derivation of magmas from metasomatized lithospheric mantle (Hart 1988" Sun & McDonough 1989; Halliday et al. 1995; Niu & O'Hara 2003) and in this case no mantle plume is needed. In continental and convergent margin magmas, these isotope ratios are masked by the signature of continental crust and therefore it is difficult to discriminate between mantle and crustal sources. Subcontinental lithosphere can preserve long-lived geochemical heterogeneity (e.g. high Rb/Sr, low Sm/Nd) and therefore can develop high 87sr/g6sr and low 143Nd/144Nd values with time. Certain lithospheric mantle-derived magmas (e.g. lamproites, kimberlites) show radiogenic isotope ratios akin to those of crustal-derived silicic melts (Nelson et al. 1986). As shown in Figure 8, Tertiary to Quaternary orogenic volcanic rock suites from the Alpine-Mediterranean region show similar curvilinear trends in the 87Sr/86Sr v. 143Nd/a~Nd diagram. An exception is the calc-alkaline volcanic rocks from the Betics, which have large scatter in the isotopic ratios. Most of the volcanic series define a continuous trend, suggesting a common origin in terms of two-component mixing between a mantle component and an enriched component. Assimilation of crustal material by mantle-derived magma has been suggested for the Periadriatic magmas (Dupuy et al. 1982; Juteau et al. 1986; Kagami et al. 1991; von Blanckenburg et al. 1998) and for the East Carpathians calc-alkaline magmatism (Mason et al. 1996). However, such trends could imply also derivation of magmas from a strongly heterogeneous upper mantle without significant upper crustal assimilation, as has been proposed for the Central Italian magmas (e.g. Peccerillo 1985, 1999) and Sardinia (Downes et al. 2001). The high 87Sr/86Sr isotope ratios can be explained only by involvement of upper crustal material in the genesis of these magmas, and the most plausible explanation is that upper crustal continental material was subducted into the upper mantle and injected into their mantle source (Peccerillo 1985, 1999). To detect the processes of crustal involvement in magmagenesis, 87 Sr/ 86 Sr isotope ratios are often combined with oxygen isotope data (180/160 or 6180, w h e r e 180/160 is expressed relative to Standard Mean Ocean Water (SMOW); Fig. 9). The upper mantle is inferred to have relatively homogeneous 8180 values (+5.5 + 0.8; Mattey et al. 1994), whereas continental and oceanic crust generally display higher ~180 values as a result of weathering processes and interaction with marine or meteoritic water (Taylor 1968; Cerling et al. 1985). Interaction with crustal material could occur in two end-member processes (Fig. 9). At convergent plate margins, crustal material could be added to the upper mantle either via subduction of continental or pelagic sediments with the oceanic lithosphere or by the entry of crustal lithosphere into the mantle during mature subduction. Dehydration and melting of the subducted crustal material result in metasomatism of the upper mantle, the source region of 'subduction-related' magmas. This process is termed 'source contamination' (James 1981" Tera et al. 1986; Wilson 1989; Ellam & Harmon 1990) and is characterized by elevated 87Sr/86Sr, but relatively low ~180 values in the resulting magmas. Involvement of a crustal component could also occur within the continental crust when the ascending mantle-derived magmas assimilate fusible continental material (Taylor 1980; DePaolo 1981" Hildreth & Moorbath 1988; Davidson & Harmon 1989). In magmas formed by this process ('crustal contamination'), variable 87Sr/86Sr ratios are accompanied by high ~180 values (Fig. 9). The mechanism by which the continental crust is involved in orogenic magmatism can be constrained by combining 87Sr/S6Sr isotope ratios with the ~180 values either of phenocrysts from the volcanic rocks or of bulk rocks. The ~180 values of bulk rocks are usually higher than those of phenocrysts, because post-eruptive alteration and low-temperature weathering can increase the 180 contents (Taylor 1968; Davidson & Harmon

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM

179

Fig. 9. 87Sr/S6Sr v. 8180 diagrams for the Tertiary-Quatemary volcanic rocks of the Carpathian-Pannonian region and the Central Mediterranean. Variation of these data indicates different types of contamination ('source contamination' and 'crustal contamination'). Data sources are as for Figure 6. Additional data are from Taylor et al. (1979) and Holm & Munksgaard (1982).

1989; Ellam & Harmon 1990; Dobosi et al. 1998; Downes et al. 1995, 2001). Therefore, 8180 values of phenocrysts reflect better the isotope composition of the host magma. Unfortunately, only sporadic oxygen isotope data are available for mineral separates from the volcanic rocks of the region. In t h e 87Sr/86Sr v. 8 1 8 0 diagram (Fig. 9), the orogenic volcanic rocks of the AlpineMediterranean region show large variations. Calc-alkaline volcanic rocks from Sardinia (Downes et al. 2001) and most from the Pannonian Basin (Mason et al. 1996; Harangi et al. 2001; Seghedi et al. 2001) show only minor elevation of 8180 with increasing 87Sr/86Sr. T h e s e trends can be explained either by source contamination (Dowries et al. 2001; Seghedi et al. 2001) or by mixing of mantle-derived magmas with lower crustal metasedimentary material (Harangi et al. 2001). Mixing between lithospheric mantle-derived magmas and lower crustal melts has also been proposed by von Blanckenburg et al. (1998), for the genesis of the Alpine Periadriatic igneous rocks. For the remaining volcanic fields the higher 6180 values at a given SVSr/S6Sr could indicate upper crustal contamination (Mason et al. 1996). Contamination of the mantle source by subducted crustal material has been proposed also for Stromboli, Roccamonfina and for the potassic rocks of Vulsini (Taylor et al. 1979; Holm & Munksgaard 1982; Ellam & Harmon 1990). In contrast, contamination by upper crustal material combined with crystal fractionation in mantlederived magmas is envisaged for the calc-alkaline volcanic rocks of SE Spain (Benito et al. 1999), for most of the volcanic rocks of the Aeolian arc (Ellam & Harmon 1990) and the Aegean arc (Briqueu et al. 1986). In summary, continental crustal material has played an important role in the genesis of the 'orogenic' magmas of the Mediterranean region. Variation of 87St/ 86~ Sr and ~is O values sugge sts th at large amounts of crustal material of various types were recycled into the upper mantle during subduction and the following collision and post-collisional events. In the following, we attempt to characterize the pre-metasomatized mantle sources. One of the characteristic features of the Tertiary to Quaternary 'subduction-related' volcanic rocks of the Alpine-Mediterranean region is their close spatial and often temporal association with

alkaline sodic mafic volcanic rocks (Fig. 1). Coeval eruption of alkali basalts and calc-alkaline or shoshonitic magmas occurs at present in the Aeolian archipelago and Sicily. A similar process took place at the southeastern part of the Carpathian chain at c. 0.5-1.5 Ma (Mason et al. 1996, 1998; Seghedi et al. 2004). In the Central Mediterranean, the 'orogenic' volcanic rocks define a curvilinear trend in the 87 Sr/ 86Sr v. 206Pb/ 204Pb diagram (Fig. 10; Peccerillo 2003; Bell et al. 2004). One of the endmembers of this trend has high S7Sr/S6Sr and medium 2~176 values and is related to continental crust. The other end-member has low 87Sr/S6Sr and high 2~176 r a t i o s and could be an enriched mantle component that evolved with high U/Pb ratio over a long period of time. This mantle component shows similarities to the HIMU mantle end-member or to FOZO and is characteristic of OIB magmas. The alkaline sodic mafic rocks from Sicily and the Sicily Channel also show isotopic variation trending towards this mantle component, having a mixing trend between DMM and FOZO or HIMU. Mixing of OIB-like intraplate and subducted slab-derived components was also suggested by Peccerillo (2001) for the genesis of potassic rocks from Campania and Stromboli. Pleistocene basalts from Sardinia deviate from all of these volcanic rocks. They show a transitional geochemical character between the 'anorogenic' alkaline sodic mafic rocks and the 'orogenic' volcanic suites. However, the most peculiar feature of these rocks is the very l o w 2~176 isotope values (Gasperini et al. 2000; Lustrino et al. 2000), which are similar to the EMI-type OIB (Fig. 10). Gasperini et al. (2000) interpreted their origin as derivation from recycled oceanic plateaux material. In contrast, Lustrino (2000) and Lustrino & Dallai (2003) argued that the EMI-type nature of the Sardinian basalts can be explained by post-collisional delamination of the lithospheric mantle and the lower crust during the Hercynian orogeny, melting of the lower crust and the contamination of the lithospheric mantle by this silicic melt. In the Carpathian-Pannonian region, the Sr-Pb isotope plot suggests a more complex petrogenetic scenario (Fig. 10; Harangi 2001a). Calc-alkaline volcanic rocks from the western Carpathians (northem Pannonian Basin) fall in a curvilinear trend between a strongly

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eastern Carpathians trend toward a mantle component with lower 2~176 ratios, which is more characteristic of depleted MORB-type mantle (DMM). This mantle source was modified by addition of subducted flysch sediments and the primary magmas underwent high-level crustal Contamination (Mason et al. 1996). The youngest South Harghita shoshonites deviate from this trend, having significantly lower 2~176 ratios, possibly implying involvement of an EMI mantle component in their genesis. Alkali basalts of the Pannonian Basin (Embey-Isztin et al. 1993) form a continuous trend between the DMM and F O Z O - H I M U mantle end-members. Thus, in the CarpathianPannonian region, a multi-component mixing model can be envisaged (Salters et al. 1988; Rosenbaum et al. 1997; Harangi 2001a), where different mantle sources (DMM and F O Z O - H I M U and possibly also EMI) and lower and upper crustal components could have been involved in the genesis of the volcanic rocks. Petrogenesis of the Tertiary to Quaternary magmas in the Mediterranean region is as complex and controversial as the geodynamic evolution of the area. The mantle source regions are extremely heterogeneous, comprising all the identified mantle end-members found in oceanic basalts. This can be observed both in the undifferentiated alkaline mafic rocks and in the compositional variation of the 'orogenic' volcanic suites. In some places (e.g. Aeolian Islands-Sicily and southernmost eastern Carpathians) these alkaline sodic and 'orogenic' magmas erupted contemporaneously and spatially close to one another, implying that heterogeneity in the upper mantle could exist both horizontally and vertically on at least a 10 km scale. Crustal rocks were subducted into the upper mantle and the fluids and melts released from them thoroughly metasomatized the subcontinental mantle, the source of the 'orogenic' magmas. In addition, crustal material was also incorporated into the ascending magmas at higher crustal levels.

Geodynamic

Fig. 10. 87Sr/86Srv. 2~176 diagrams for the Tertiary-Quaternary volcanic rocks of the Carpathian-Pannonian region and the Western and Central Mediterranean. Data sources are as for Figure 6. Additionaldata are from Vollmer (1976).

radiogenic Sr isotopic component (lower crust?) and an enriched mantle component with low 87Sr//86Sr and high 2~176 ratios similar to that inferred for the Central Mediterranean magmas. In contrast, calc-alkaline volcanic rocks from the

implications

The low-K to high-K calc-alkaline volcanic rocks, the shoshonites and ultrapotassic formations, as well as alkaline volcanic rocks of the Alpine-Mediterranean region were formed in a convergent plate margin setting. The geochemistry of these magmas indicates a strongly heterogeneous mantle beneath this area. Most workers suggest that this can be explained by a lengthy period of subduction and subsequent post-collisional processes. Subduction of remnant Tethyan oceanic plates appears to have played an important role in the evolution of this region and has also had a major influence on the regional upper mantle structure. Seismic tomographic models show the presence of high-velocity anomalies beneath the Gibraltar arc, Calabria and the Hellenic arc, interpreted as recently subducted slabs (Spakman et al. 1988; Wortel & Spakman 1992, 2000; Blanco & Spakman 1993; Faccena et al. 2003; Piromallo & Morelli 2003). In addition, these models revealed an extensive coherent mass between 450 and 650 km depth that is interpreted as remnants of subducted Mesozoic oceanic slabs (Spakman et al. 1993; Piromallo et al. 2001; Piromallo & Morelli 2003). Indeed, most of the Tertiary to Quaternary volcanic rocks in the Mediterranean region show 'subduction-related' geochemical features. Some of them are found along volcanic arcs (Aeolian arc, Aegean arc) associated with the active subduction zones (Keller 1982). However, as shown in previous sections, there are debates on whether they could be considered as classic volcanic arcs. Some features would seem to indicate a back-arc tectonic setting (Mitropoulos et al. 1987; Mazzuoli et al. 1995). In this case, subduction could have only an indirect influence on the magmagenesis (Ellam et al. 1989). The principal reason for magma generation could be passive extension of continental lithosphere resulting in decompression melting of the lithospheric and asthenospheric mantle variably metasomatized by previous subduction processes.

ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM Another candidate for volcanism directly related to active subduction is the Late Miocene low-K tholeiitic to calc-alkaline volcanic products of the Alboran basin (Duggen et al. 2004; Gill et al. 2004). However, other workers (e.g. Zeck 1996; Benito et al. 1999; Turner et al. 1999) have emphasized the post-collisional origin of these rocks. Indeed, formation of most of the 'orogenic' volcanic rocks in the Mediterranean regions post-dates the active subduction process and appear to be related to slab break-off, lithospheric mantle delamination or lithospheric extension (e.g. Mason et al. 1998; Seyito~lu et al. 1999; Turner et al. 1999; Aldanmaz et al. 2000; Chalot-Prat & G~rbacea 2000; Wortel & Spakman 2000; Harangi 2001a; Coulon et al. 2002; Seghedi et al. 2004). The 'subduction-related' geochemical character of the volcanic rocks is inherited from the mantle source regions modified previously (a few million to several tens or even hundreds of million years before) by fluids released from subducted slabs. Post-collisional or back-arc extension of the lithosphere could result in the decompression melting of the hydrous portion of the lithospheric mantle first (Gallagher & Hawkesworth 1992), followed by the melting of the deeper asthenosphere. This could explain the initial 'orogenic' magmatism and the subsequent alkali basaltic volcanism in many areas of the Mediterranean region (e.g. Wilson et al. 1997; Seyito~lu et al. 1999; Harangi 2001 a). Alkaline mafic rocks akin to those occurring in Central Europe occur sporadically in this region, often very close to the 'orogenic' volcanic formations. Furthermore, rare volcanic rocks types such as carbonatites, melilitites and/or kamafugites in the Apennines are more characteristic of intra-plate rift settings (Lavecchia & Stoppa 1996; Lavecchia et al. 2003). This may imply also another mechanism for magmatism of the Mediterranean region; that is, upwelling of hot mantle plume. The role of a mantle plume in the genesis of the volcanic rocks of Central Italy was first suggested by Vollmer (1976, 1989). Recognition of an enriched component (FOZO-HIMU) in both the 'anorogenic' and 'orogenic' volcanic rocks in many subprovinces (Hoernle et al. 1995; Ayuso et al. 1998; Wilson & Bianchini 1999; Harangi 2001a; Gasperini et al. 2002; Peccerillo 2003; Bell et al. 2004) also led some researchers to propose mantle plume activity. This could be supported by the extensive low-velocity anomaly beneath most of this area (Hoernle et al. 1995; Wortel & Spakman 2000; Piromallo & Morelli 2003) that could be interpreted as presence of anomalously hot mantle material. The estimated temperature from the P and S velocity anomalies approaches the dry solidus under the Pannonian Basin, Western Mediterranean, Tyrrhenian Basin and Aegean Basin, and the discrepancy between temperature inferred from P and S waves also indicates the presence of partial melt (Goes et al. 2000). Harangi (2001b) supposed also a relatively hot mantle beneath the Pannonian Basin, based on the composition of the Late MiocenePliocene alkali basalts. As an extreme case, Lavecchia et al. (2003) and Bell et al. (2004) argued that the geodynamic evolution of the Central Mediterranean has been controlled by an upwelling plume and subduction had no role whatsoever. Other workers combined the subduction-related models with the existence of OIB-like mantle (Fig. 11). Gasperini et al. (2002) assumed that the HIMU signature of many of the volcanic rocks in the Central Mediterranean could be related typically to an upwelling plume. They invoked a plate window beneath the central Apennines, where deep mantle plume material could be channelled toward the shallow mantle zones (Fig. l la). Decompression melting of this hot plume material could result in the enriched mantle component of many of the volcanic rocks in Central Italy. In the southern Tyrrhenian area, the coexistence of OIB and 'orogenic' magmas was interpreted as the result of lateral flow of African enriched mantle along a tear at the edge of the Ionian plate (Fig. l lb; Trua et al. 2003). Mantle anisotropy studies in this area also indicate a toroidal mantle flow around the Calabrian slab (CiveUo & Margheriti 2004), which could

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Fig. 11. Proposed models for the involvementof an enriched mantle component (EAR) of the 'orogenic magmas' in the Carpathian-Pannonian and Mediterraneanregion. (a) Deep mantle upwelling could occur via a slab window beneath Central Italy as proposed by Gasperini et al. (2002). (b) Toroidal mantle flow around the Calabrianslab from the African mantle was proposed by Trua et al. (2003) and Civello & Margheriti (2004). (e) Carpathian-Pannonian region: slab detachment could result in the suction of a hot, enriched asthenosphericmantle material possibly from the Bohemian mantle plume finger.

supply enriched mantle material beneath Campania and the southern Tyrrhenian area from behind the Calabrian subduction zone. In the Carpathian-Pannonian region the compositional variation of Miocene to Quaternary calc-alkaline volcanic rocks implies contrasting genesis. Isotopic values of andesitic to dacitic rocks in the western segment of the Carpathian arc show a mixing trend between an enriched (FOZO-HIMU-type) mantle and a crustal component (Fig. 10; Harangi 2001a) similar to other volcanic suites in the Mediterranean region, whereas the

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s. HARANGI ETAL.

calc-alkaline volcanic rocks in the eastern segment of the Carpathian arc show a mixing trend between a depleted mantle and a crustal component. A possible explanation for the contrasting mantle source regions is that enriched mantle could flow from the assumed plume finger beneath the Bohemian Massif to the thinned Pannonian Basin through the gap left behind the detached slab under the western Carpathians (Fig. l lc). In the east, no enriched mantle material could penetrate beneath the thick Ukrainian Shield, therefore slab break-off beneath the eastern Carpathians could initiate upwelling of depleted MORB-type mantle material. Deflection of the assumed mantle plume finger beneath the Massif Central towards the SE was also detected by seismic anisotropy pattern (Barroul & Granet 2002). The southeastward asthenospheric flow was explained by a combined effect of Apenninic slab roll-back and the opening of the extensional basins behind it (Barroul & Granet 2002; Barroul et al. 2004). In summary, the F O Z O - H I M U mantle component recognized in the compositional variation of many Mediterranean volcanic suites led many researchers to propose the influence of localized mantle plume(s) in the genesis of the magmas. Whether upwelling of hot mantle material was the ultimate cause of the magmagenesis and also influenced the tectonic evolution of the areas (Lavecchia et al. 2003) or subduction and post-collision processes (slab rollback, slab break-off, delamination of the lower lithosphere) initiated deflection of nearby mantle plumes, requires further combined geochemical, geophysical and tectonic research. Nevertheless, the HIMU signature of the mantle source could also be interpreted as due to metasomatic processes without assuming mantle plumes (Sun & McDonough 1989; Anderson 1994; Halliday et al. 1995; Niu et al. 1999). It is remarkable that the enriched, HIMU-like mantle component was detected also in earlier, pre-Neogene volcanic rocks of this region (Veneto region, Macera et al. 2003; Dinarides, Cvetkovid et al. 2004; C a r p a t h i a n - P a n n o n i a n region, Harangi 1994; Harangi et al. 2003), which may imply its long-lasting (at least from the Early Cretaceous) presence beneath Europe. Oyarzun et al. (1997) and Wilson (1997) suggested that this enriched mantle component could be derived from the Mesozoic Central Atlantic plume being deflected as a result of either the suction of the European thin-spots or the northeastward motion of the African plate. In any case, portions of the deflected plume material could have polluted the shallow upper mantle beneath Europe since the Early Cretaceous. In addition, subduction of crustal material could also contribute to the inhomogeneity of the shallow mantle. Statistical sampling of this heterogeneous mantle (SUMA model, Meibom & Anderson 2004) could be an alternative model for the wide variation of the Tertiary to Quaternary volcanic rocks of the Mediterranean region. Discussions with the PANCARDI Igneous Team over the last decade have helped us to develop our ideas on the Tertiary to Quaternary magmatism and its geodynamic relationships in the Alpine, Pannonian, Carpathian and Dinarides regions. Part of this study belongs to the research project supported by the Hungarian Science Foundation (OTKA # T 037974; to Sz.H.). The critical reviews of C. Doglioni and F. Chalot-Prat helped to improve the paper and are gratefully acknowledged. We also thank R. Stephenson and D. Gee. for patient editorial help, and R. Lukfics for the careful work in producing the reference list.

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ZECK, H. P., ALBAT, F., HANSEN, B. T., TORRES-ROLDAN, R. L. & GARCIA-CASCO,A. 1989. Alpine tourmaline-bearing muscovite leucogranites, intrusion age and petrogenesis, Betic Cordilleras, SE Spain. Neues Jahrbuch fiir Mineralogie, 11, 513-520. ZECK, H. P., KRISTENSEN,A. B. & WILLIAMS, I. S. 1998. Post-collisional volcanism in a sinking slab setting--crustal anatectic origin of pyroxene-andesite magma, Caldear Volcanic Group, Neogene Alboran volcanic province, southeastern Spain. Lithos, 45, 499-522. ZINDLER, A. & HART, S. R. 1986. Chemical geodynamics. Annual Review of Earth and Planetary Sciences, 14, 493-571.

Formation and deformation of the Pannonian Basin: constraints from observational data F R A N K HORV./~TH 1, GJi.BOR B A D A ~'2, PI~TER SZAFIAN ~, G A B O R TARI 3, A N T A L / ~ D A M 4 & SIERD C L O E T I N G H 2

XDepartment of Geophysics, E6tvOs University, Pdzmdny P. s. I /C, 1117 Budapest, Hungary (e-mail: [email protected]) 2Netherlands Research Centre for Integrated Solid Earth Sciences, Department of Tectonics and Structural Geology, Vrije Universiteit, 1081 HV Amsterdam, The Netherlands 3Vanco Energy Company, Houston, TX 77046, USA 4Geodetic and Geophysical Research Institute, Hungarian Academy of Sciences, Sopron, Hungary

The past decade has witnessed spectacular progress in the collection of observational data and their interpretation in the Pannonian Basin and the surrounding Alpine, Carpathian and Dinaric mountain belts. A major driving force behind this progress was the PANCARDIproject of the EUROPROBE programme. The paper reviews tectonic processes, structural styles, stratigraphic records and geochemical data for volcanic rocks. Structural and seismic sections of different scales, seismic tomography and magnetotelluric, gravity and geothermal data are also used to determine the deformational styles, and to compile new crustal and lithospheric thickness maps of the Pannonian Basin and the surrounding fold-and-thrust belts. The Pannonian Basin is superimposed on former Alpine terranes. Its formation is a result of extensional collapse of the overthickened Alpine orogenic wedge during orogen-parallel extrusion towards a 'free boundary' offered by the roll-back of the subducting Carpathian slab. As a conclusion, continental collision and back-arc basin evolution is discussed as a single, complex dynamic process, with minimization of the potential and deformational energy as the driving principle. Abstract:

The Pannonian Basin is located in eastern Central Europe, as part of the Alpine orogenic system. The Alpine, Carpathian and Dinaric mountain belts surround this extensional basin of NeogeneQuaternary age. Its broader geological environs, the Mediterranean region, is a wide zone of convergence between the Eurasian and African plates. Because of the complex kinematics of the two major plates, the region has had a polyphase deformation history since the opening of the Atlantic Ocean, which has primarily controlled the formation of the entire Alpine orogenic belt (Biju-Duval et al. 1977; Dercourt et al. 1986; Dewey et al. 1989; Seng6r 1993; Yilmaz et al. 1996). A remarkable feature in this overall compressional setting is the abundance of extensional basins superimposed on former orogenic terranes and associated with orogen-parallel displacement of internal blocks and oroclinal bending (Horvfith & Berckhemer 1982). From west to east, these basins are the Alboran, Ligurian, Tyrrhenian, Pannonian and Aegean basins (Fig. la). Although their age, deep structure and tectonics show significant differences, there are several common features in their formation and evolution. For instance, their similar position in proximity to a once or still active subduction zone strongly suggests a causal relationship of primary importance (Royden 1993; Giunchi et al. 1996; Meijer & Wortel 1997). Other important geodynamic processes include the extensional collapse of a gravitationally unstable orogenic wedge (Dewey 1988), and the lateral escape of internal terranes (Seng6r et al. 1985; Royden 1993). Following the suggestion of Horvfith (1988), Ratschbacher et al. (1991) argued that extensional collapse occurs during tectonic escape and suggested the term extrusion to describe these interrelated processes. It is to be emphasized that this was not simply a new term but a novel concept, which postulated that orogenparallel extension and orogen-normal compression in the Mediterranean were taking place simultaneously, as in the Asian segment of the Alpine-Himalayan mountain belt (Tapponnier et al. 1986). This implies that extrusion is not a lateral displacement (translation) of one coherent block but, instead, strong internal deformations occur, which involve brittle faulting and ductile flow depending on the rheological layering of the extruding terrane (Ratschbacher et al. 1991). Based on a summary of 3D seismic tomography data, Wortel & Spakman (2000) have offered new constraints for the Cenozoic evolution of the Mediterranean region and a contribution to the

explanation of back-arc basin formation and deformation. Tomography images clearly depict the elevated asthenospheric material below back-arc basins and the lithospheric slabs underthrusting these lithospheric depressions. The geometry of the individual subducted slabs is particularly important as it shows considerable differences in the Mediterranean region (Faccenna et al. 2003). Along the Hellenic arc a continuous slab penetrates down more than 1000 km, well into the lower mantle below the Aegean Basin, and only the upper 200 km portion of this long slab exhibits seismic activity. Elsewhere, the slab dips from the subduction zone towards the interior of the back-arc basin, and then soles out between the 410 km and 660 km seismic discontinuities. This geometry is compatible with the progressive rollback of a subducted slab and arc retreat. An additional important feature is the observation that in active subduction zones the slab appears to be continuous towards the surface, whereas in inactive subduction zones a detached relict of the subducted slab can be imaged in the upper mantle. Slab detachment occurs in the form of a self-perpetuating break-off under the weight of the slab itself (Dvorkin et al. 1993; Davies & von Blanckenburg 1995). This can happen when thick and buoyant continental lithosphere enters the trench zone. The process of subduction is blocked when the buoyancy of the arriving continental lithosphere becomes equal in magnitude to the downward force exerted by the already subducted oceanic lithosphere. Lateral migration of slab detachment leads to the concentration of tensional forces on a continuously decreasing portion of the arc, which can result in a gradual, often accelerating retreat of the arc-trench system (Dvorkin et al. 1993; Royden 1993). This process is responsible for the migration and bending of the arcs and back-arc basin formation from Oligocene to recent times in the Mediterranean and the build-up of the arcuate orogenic chains of the Betics-Rif, Calabrides-Apennines and Hellenides (Faccenna et al. 2004). Formation of the Carpathian arc and the Pannonian Basin took place in a similar setting from the latest Oligocene. This basin has rapidly reached a mature stage of evolution, as extension has come to an end because of the complete consumption of the subductable lithosphere of the European foreland (Horv~th 1993). As a result, unlike other Mediterranean back-arc basins, the change of stress field from extension to compression and positive structural inversion of the Pannonian Basin system has been in progress since Late

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 191-206. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. (a) Sketch map showing the present-day configuration and the late Cenozoic structural features of the Mediterranean and adjacent areas. The overriding plates above subduction zones have undergone rigid body rotation, translation and extension, which has resulted in the formation of a set of back-arc basins. PB, Pannonian Basin; AdP, Adriatic Promontory; AeS, Aegean Sea; A1S, Alboran Sea; B1S, Black Sea; loS, Ionian Sea; LiS, Ligurian Sea; TyS, Tyrrhenian Sea. Darker shading indicates oceanic crust. (b) Dynamic model for the evolution of the same area during the Cenozoic proposed by Wortel & Spakman (2000). Continuous line indicates area of active subduction; thick grey line indicates area where slab detachment has already occurred at the given time interval. E, Eocene; O1, Oligocene; Mt, Early Miocene; M2, Mid-Miocene; M3, Late Miocene; Pr, present. Pliocene times (Horv~ith & Cloetingh 1996; Bada et al. 1998, 2001). The goal of this paper is to review and interpret recent data regarding the formation and deformation of the Pannonian Basin system. This can assist the understanding of other Mediterranean back-arc basins and lithospheric dynamics in general. We attempt to show that although the level of understanding of the formation and deformation of the Pannonian Basin was very high already in the mid-1980s (e.g. Royden & Horv~th 1988), the past decade has witnessed spectacular progress in the collection of observational data and their interpretation. The EUROPROBE programme, particularly the PANCARDI project, has clearly been a most important driving force of this progress.

Tectonic framework The Cenozoic evolution of the Alpine-Pannonian region is primarily controlled by the northward drift and indentation of the Adriatic promontory, which has produced a net convergence of at least 5 0 0 k m in the Alps (Roeder & Bachmann 1996;

Schmid et al. 1996). Adria has been pushed towards the north by the African plate, although it was not necessarily tightly attached to Africa, at least from the Early Tertiary (M~irton et al. 2003). Another important feature of Adria is that it represents a non-rigid indenter, as is shown by its dramatic deformation history throughout the Cenozoic (Fig. lb). It can be reconstructed fairly well (D'Argenio & Horv~ith 1984) that an originally broad bump at the northern boundary of the African plate gradually became a sharp nose by necking as a result of the development of the Tyrrhenian and Aegean back-arc basins (Fig. lb). The northern front of Adria also experienced strong deformations but of different nature. The formation of the West Alpine arc during the Palaeogene (Giglia et aL 1996; Schmid & Kissling 2000) can be best explained by westward extrusion of a fragment of the Adriatic indenter. On the basis of differential global positioning system (GPS) velocities and seismicity, Oldow et al. (2002) have concluded that contemporary Adria is fragmented. A northwestern block has velocities indicative of little or no motion relative to Europe, whereas a southeastern part is moving together with Africa and exhibits a spatially heterogeneous velocity field with northward displacement rates up to 1 0 m m a - t . In the Eastern Alps, the

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Austroalpine basement and cover nappes overriding the deformed European margin are obviously of Adriatic provenance. Clearly, there is not a single Adriatic-European boundary in the Alps either at upper crustal level or at a depth, as we show later in this paper, when discussing the new results of deep seismic profiling and improved mantle tomography. The most pronounced surface expression of strain partitioning (Ziegler & Roure 1996) has been the Late Oligocene to Early Miocene eastward extrusion of an Alpine orogenic wedge, primarily driven by the northward push of the 'soft' Adriatic indenter. This Alpine wedge is called the ALCAPA terrane (Fig. 2). Its boundaries, again, can only be loosely defined, which reflects the nature of the extrusion process rather than the limits of our knowledge (Frisch et al. 1998). There is a second unit in the substrata of the Pannonian Basin system, called the Tisza-Dacia terrane. Detailed description of the contrasting Mesozoic to Early Tertiary stratigraphy of these two terranes can be found in a set of recent papers (e.g. Kovfics et al. 2000; Csontos & V6rrs 2004). It is generally accepted that the Tisza-Dacia terrane rifted apart from the European margin of the Mesozoic Tethys during the Late Jurassic, and it led to the formation of the marine basins in which the AlpineCarpathian flysch complexes were deposited (Yilmaz et al. 1996; Csontos & V r r r s 2004). Although a number of papers have attempted to reconstruct the Tertiary kinematic history of the Tisza-Dacia terrane (Balla 1986; Csontos et al. 1992; Csontos 1995; Csontos & V r r r s 2004) this issue has remained obscure. This is mainly because of an acute space problem between the Bohemian massif and the assumed fixed Moesian platform (Fig. 2), and the lack of adequate kinematic indicators. Palaeomagnetic data tend to indicate counter-clockwise rotations for the ALCAPA domain and clockwise rotations for the T i s z a -

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Dacia unit (Fig. 2), but with the increase in quantity and quality of palaeo-declinations serious rotation anomalies have been revealed within the two terranes (see a review by Mfirton 2001). This is reasonable, as both units experienced strong internal deformations during the Tertiary. It is generally accepted that the two terranes became juxtaposed at the beginning of the Miocene when formation of the Pannonian Basin started (Csontos et al. 1992). Geodynamically, this formation was a stretching of the two terranes towards the eastern Carpathians facilitated by the subduction roll-back of the lithosphere of the Carpathian flysch basin (Horvfith 1993; Bada & Horvfith 2001).

Tectonic evolution The Pannonian Basin and its surroundings are characterized by a polyphase deformation history with a sequence of distinct structural episodes. There is a good knowledge of the principal kinematic features, that is the location of major fault zones, the timing and the amount of deformation (Figs 3 and 4). A rapid and dramatic change in tectonic style started in the Early Miocene (Eggenburgian to Karpatian; see Fig. 4 for local time scale), which initiated the formation of the Pannonian Basin. This process culminated in the Mid-Miocene (Badenian) and was coeval with a large-scale tectonic transport of the external flysch nappes towards the foreland of the Carpathian arc (Royden et al. 1982). It is widely recognized that in regions of continental collision intense shortening and crustal thickening can eventually lead to gravitational instability of the axial zone of the orogen

Fig. 2. Simplified Late Cenozoic tectonic map of the Alpine-Carpathian-Pannonian-Dinaric system (after Bada & Horv~ith 2001). The Adriatic promontory or microplate has been indenting and pushing the Alpine-Dinaric belt since the Cretaceous. The northern domain beneath the Pannonian Basin (ALCAPA unit) underwent a net counter-clockwise rotation of about 50-70 ~ and was translated to the ENE, whereas the southern unit (Tisza-Dacia unit) rotated c. 100~ in an opposite manner and was moving to the ESE. Green and black arrows indicate the translation and rotation, respectively, of various tectonic units. 1, Foreland (molasse) basins; 2, flysch nappes; 3, Neogene volcanic rocks; 4, Southern Alps, Dinarides and Northern Calcareous Alps; 5, pre-Tertiary units of the East Alpine-Carpathian domain and the Jura Mts; 6, Variscan basement of the European plate, and Dinaric, Vardar and Mures ophiolites; 7, Pieniny Klippen Belt; 8, Oligocene tonalites; 9, Penninic basement; 10, Penninic cover; 11, Helvetic basement; 12, Helvetic cover. T, Tauern window; R, Rechnitz window; PAL, Periadriatic lineament (PAL); G, Giudicarie fault; Za, Zagrab fault; B, Brenner fault; Tr, Trotu~ fault; IM, Intramoesian fault.

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Fig. 3. Depth of pre-Neogene basement in the Pannonian Basin system (Horvfith & Royden 1981) and related late Neogene structural pattern (Horvfith 1993). Thick red lines mark the location of structural profiles shown in Figure 6 (A-A'), Figure 7a (B-B'), Figure 7b (C-C') and Figure 8 (D-D'). B6, B6k6s Basin; D, Danube Basin; HM, H6d-Mak6 Basin; M, Mihfilyi high; PB, Pusztaf61dvfir-Battonyablock; TB, Transylvanian Basin; TR, Transdanubian Range; Z, Zala trough. 1, Foreland (molasse) basins; 2, flysch nappes; 3, Neogene volcanic rocks; 4, pre-Tertiary units on the surface; 5, Variscan basement of the European plate; 6, Dinaric and Vardar ophiolites; 7, tectonic windows in the Eastern Alps; 8, normal and low-angle normal fault; 9, thrust, anticline; 10, strike-slip fault.

(Tapponnier et al. 1986; Molnar & Lyon-Caen 1988; Bird 1991). Moreover, the central belts of orogens are often thermally weakened and, hence, are prone to strain localization. Once the gravitational forces exceed the compression exerted by plate convergence, the process of late orogenic collapse of the weakened crust can be initiated. Orogen-parallel extension in the Eastern Alps is well documented (e.g. Selverstone 1988; Ratschbacher et al. 1989) and the idea of a weakened Alpine lithosphere has been confirmed (e.g. Cloetingh & Banda 1992; Genser et al. 1996; Okaya et al. 1996). The large-scale lateral extrusion of the ALCAPA terrane took place from the Late Oligocene towards an eastern, unconstrained margin of the Carpathian flysch basin (Fig. 5). In a strict sense, lateral extrusion is a ductile flow of the lower crust confined between the brittle upper crust and mantle lithosphere (Ranalli 1995) that leads to the relaxation of the topography of the surface as well as the Moho (Bird 1991). A large amount of material was expelled towards the east from between the soft indenter (Adria) and the rigid foreland buttress (Bohemian Massif). Crustal wedges bounded by conjugate sets of strike-slip faults, that is sinistral and dextral in the north and south, respectively, were extruded and stretched in an orogen-parallel direction to the ENE (Figs 2, 3 and 5). Often, these strike-slips are still

active seismically, and magnetotelluric soundings show that they are associated with highly conducting zones containing lowviscosity material (Ad~im 2001). The extrusion process was coeval with continuing north-south compression in the central zone of the Eastern Alps, and rapid exhumation of metamorphic rocks in the Alps (Dunkl & Frisch 2002), the Pannonian Basin (Tari & B a l l y 1990; Tari et al. 1992, 1999), and the East Slovak Basin (Sotak et al. 1993). Kinematic data (Fodor et al. 1999) and numerical modelling (Bada 1999) suggest the predominant role of Carpathian subduction facilitating large-scale lithospheric extension in the Pannonian Basin from latest Early Miocene to Pliocene times. Continuous roll-back of the subducting plate along the contemporaneous Carpathian arc exerted trench-pull forces on the upper plate. The overriding plate in a subduction zone tends to passively follow the retreating hinge of the downgoing lithosphere. This induced tensional stresses and eastward stretching of the A L C A P A and Tisza-Dacia terranes. The upper plate is extending in the direction of maximum gravitational potential energy difference. Trench suction forces acting normal to the curvature of the Carpathian arc, in combination with collisional forces exerted along the Alpine-Dinaric belt, can reproduce well the reconstructed Mio-Pliocene palaeostress pattern (Bada 1999). Because

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195

Fig. 4. Evolutionaryscheme of the PannonianBasin system with major volcanic horizons and events, generalized stratigraphy, and main tectonic phases. The primary depositional environmentsand facies of the PannonianBasin are described in the bottom panel (after Horv~ith& Tari 1999). Major unconformitiesare labelled U1 (oldest) to U4 (youngest).

of the finite strength of the Pannonian lithosphere, tensional stresses were transmitted far behind the arc region and, as a consequence, nearly the whole Pannonian Basin system extended significantly. It is important to emphasize that there are no remarkable differences in either the style or the amount of extension between the formerly distinct ALCAPA and Tisza-Dacia terranes (compare Figs 6 and 7). There is, however, one obvious exception: the presence of the coeval non-extensional Transylvanian Basin in the eastern part of the Tisza-Dacia unit (Ciulavu et al. 2002). The formation of this basin is poorly understood and the original concept of Royden et al. (1982) still seems to be the most plausible. They explained the Transylvanian Basin as a continental sag caused by the suction force exerted to the upper plate by the downbending Carpathian slab. Tension in the Pannonian region caused about 50-120% crustal lithosphere extension and nearly an order of magnitude higher mantle lithosphere extension (Horvfith et al. 1988; Lenkey 1999). Occasionally, extension was concentrated in discrete zones where pull-apart basins developed (Horvfith & Royden 1981; Horvfith 1993; Csontos 1995; Fodor et aL 1999). Heterogeneous extension is reflected by the variation of pre-Neogene basement depth (Fig. 3). Elevated basement blocks separate deep sub-basins

where the thickness of the Neogene-Quaternary sedimentary rocks can reach 6 - 7 km. Such irregular basement morphology is mainly the result of strain localization along pre-existing crustal weakness zones inherited from Late Cretaceous thrust and nappe tectonics. Quaternary differential vertical movements, related erosion and sedimentation have significantly influenced the observed thickness of the basin fill. These processes are related to the late stage of basin evolution. Contemporary stress data, seismicity pattern, seismic profiles and Quaternary subsidence history indicate that the Pannonian Basin is in the phase of structural inversion (Horvfith 1995; Horvfith & Cloetingh 1996; Bada et al. 1999; Gerner et al. 1999). An increase of horizontal compression as a result of the changes of boundary conditions around the basin system causes buckling of the Pannonian lithosphere (Horvfith & Cloetingh 1996; Cloetingh et al. 2006) manifested in the uplift and subsidence of the basin flanks and centre, respectively. Present-day boundary conditions include active collision along the Alps-Dinarides belt (Adria-push), terminated subduction and continuing continental collision in the SE Carpathians and eastward translation of crustal wedges currently squeezed out from the region of the Alpine orogen (Bada et al. 1998, 2001).

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Fig. 5. Schematicillustrationof the Early Miocene extrusion of the ALCAPAterrane. Continentalconvergencein the Alpine collisionalbelt was able to progress because the buoyant crust decoupled from the dense mantle lithosphere. This crustal flake, the ALCAPA terrane, extruded towards a 'free' eastern boundary. This was free in the sense that the lithosphereof the Carpathian flyschBasin was subductableand its progressiveroll-back offeredthe space to be invaded. It should be noted that the extruding crustal flake was directly superimposedon the hot asthenosphere, which led to crustal melting and a dramatic decrease of its integrated strength. 1, European foreland areas, undifferentiated('stable Europe'); 2, foreland (molasse) foredeep; 3, Alpine-Carpathian nappe system; 4, flyschBasin in the Carpathian embayment;5, normal fault, strike-slipfault and thrust plane, thrust front; 6, orogen-paralleldisplacementof the extruding (collapsing)Pannonianterranes. PA indicates the Periadriaticline. The basin system has become completely landlocked and constrained from all directions, which has led to a gradual inversion in the form of multi-scale folding and fault reactivation (Fig. 8).

Basin-fill stratigraphy and magmatism Stratigraphic data provide important information on the basin evolution in terms of timing and characterization of major tectonic events. Furthermore, the petrological and geochemical analyses of magmatic rocks give further constraints on the composition, thermal state and rheological behaviour of the deforming lithosphere-asthenosphere system. The evolutionary scheme of the basin system, with major volcanic events, stratigraphic pattern, tectonic phases and local chronostratigraphic units, is shown in Figure 4. The onset of rifting in the Pannonian Basin is marked by the regional unconformity between the pre-rift strata and the overlying Lower Miocene (Eggenburgian-Ottnangian) deposits. A rhyolite tuff horizon dated at c. 20 Ma (H~mor et al. 1980) is interbedded in the lowermost part of the synrift sedimentary beds. However, in the main depocentres this tuff horizon is missing, and the thick basal conglomerates and other continental beds cannot be precisely dated. Therefore 20 Ma as a start of rifting is a somewhat uncertain, but reasonably good age estimate. The evolution of the Pannonian Basin is traditionally subdivided into a synrift (Early to Mid-Miocene) and a post-rift (Late Miocene to Quaternary) phase (Royden et al. 1983; Horv~th & Rumpler 1984), which is reflected in the sedimentary architecture. An increase of available stratigraphic data and the re-evaluation of numerous seismic profiles has allowed a slight modification of this subdivision. According to Taft (1994), Horvzith (1995) and Tari et al. (1999), a Mid-Badenian unconformity indicates the termination of the synrift period (Fig. 4). The Sarmatian-Pannonian regional unconformity is due to uplift and erosion related to an early

inversion event (Horv~th 1995; Fodor et al. 1999), which may indicate transient changes in the boundary conditions along the Carpathian subduction belt. This compressional event took place shortly after the termination of the synrift phase, that is at about 11-8 Ma. There is a second, more regional compressional event, which started during the Late Pliocene and has continued until recent time (c. 3 - 0 Ma). Widespread upwarping of basement units from below the Neogene succession has resulted in the characteristic 'inselberg' pattern of present-day Mesozoic and Palaeozoic outcrops inside the Pannonian Basin (Fig. 3). In other places, where the basement was uplifted but did not reach the surface, erosion led to a stratigraphic gap and unconformity between Miocene and Quaternary strata (Fig. 4). At the same time, rapid subsidence has taken place in the deepest sub-basins during the Quaternary. Modelling results (Cloetingh et al. 2006) suggest an increase of compressional intraplate stress magnitudes and related lithospheric folding that can fairly well explain the observed pattern of Quaternary subsidence and uplift. Depositional environments in space and time were strongly influenced by tectonics (Fig. 4). The synrift phase was characterized by continental to marine sedimentation, whereas during the post-rift phase the Pannonian Basin became an isolated brackishwater lake, which has been progressively filled up, and lacustrine sedimentary rocks were replaced by terrestrial deposits (e.g. Nagymarosy & Mtiller 1988; Jfimbor 1989; Magyar et al. 1999; H~mor et al. 2001). The late-stage terrestrial depositional environments have been particularly sensitive to climatic fluctuations driven by Milankovid-type cyclicity (Nfidor et al. 2000). However, the effect of global eustatic sea-level changes on the stratigraphy of an isolated large lake is still a matter of debate (Tari et al. 1992; Pogfics~s et al. 1994; Vakarcs et al. 1994; Juh~sz et al. 1999; Sacchi et al. 1999). Volcanic rocks identified in the Pannonian-Carpathian system show large variation in lithology, geochemical composition, and spatial and temporal distribution (P~cskay et al. 1995; Harangi 2001; Kone6n3) et al. 2002). Silicic volcanism started about

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197

Fig. 6. Seismic reflection section A-A' through the Great Hungarian Plain in the centre of the Pannonian Basin and a line drawing interpretation (after Tari et al. 1999). (For location of profile see Fig. 3.)

20 Ma ago with the deposition of ignimbritic flow deposits and rhyolitic tufts. Repeated activity resulted in the formation of two other rhyolitic tuff horizons (Fig. 4). Based on geochemical and isotope data, these silicic volcanic rocks can be best explained as the product of partial melting of the continental lower crust in the Pannonian lithosphere (Salters et al. 1988; Kone~n2~ et al. 2002). Harangi (2001) have argued for a mantle origin of these magmas, but with strong crustal contamination. The considerable involvement of crustal materials in early synrift magmatic processes strongly suggests that an anomalously high temperature gradient was established, leading to the low strength of the deforming lithosphere from the earliest stages of basin evolution. The second main type of magmatic activity during Miocene to Quaternary times (20-0.15 Ma) produced large bodies of calc-alkaline volcanic rocks in the northern Pannonian Basin and at the inner side of the Eastern Carpathians (Fig. 2). In the first plate tectonic models this volcanism was directly or indirectly related to the subduction of the flysch basin lithosphere (Stegena et al. 1975; Lexa & Konetn2~ 1974). Results of detailed geochemical analysis (Salters et al. 1988; Szab6 et al. 1992; Downes et al. 1995; Mason e t al. 1996; Harangi 2001) confirmed the subductionrelated origin and highlighted important spatial and temporal variations in magma generation and evolution. Volatile-rich fluids from the subducted slab migrated into the mantle wedge above it, which created large volumes of magma of mostly andesitic composition. While rising to the surface, as suggested by Downes et al. (1995), the magma was contaminated by crustal

components, resulting in both spatial and vertical variations of the isotopic composition. The widespread occurrence of garnet xenoliths suggests very rapid ascent of the magmas in an overall tensional stress regime (Harangi 2001). Kone~n3~ et al. (2002) have also argued that in the northern parts of the Pannonian Basin system the generation of calc-alkaline magmas was related to the decompressional melting of the mantle lithosphere as a result of the overall extension in the Pannonian Basin. The youngest region of calc-alkaline magmatic activity is a remarkably linear chain along the inner side of the east Carpathians (Fig. 2). A general progression from 12 Ma to 0.2 Ma in the onset and cessation of magmatism from the NW towards the SE has been documented by K - A r geochronology (Ptcskay et al. 1995). Subduction of the flysch basin lithosphere and progressive break-off of the subducted slab has been suggested as the most probable cause of the east Carpathian magmatism and its migration (e.g. Linzer 1996; Mason et al. 1998). As thick continental crust began to enter the trench zone, detachment of the subducted oceanic slab occurred, and the rupture propagated along the slab. This led to the cessation of volcanism as the slab sank out of the magma generation zone. In recent times, this process is in its final stage at the bend of the eastern and southern Carpathians. Analyses of the magmatism, structural evolution, geometry of the seismically active Vrancea slab, and mantle tomography suggest that in this area detachment is not complete and slab continuity is probably still partly maintained (Girbacea & Frisch 1998; Chalot-Prat & Girbacea 2000; Wortel & Spakman 2000; Gvirtzman 2002).

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Fig. 7. (a) Gravity model of a crustal-scale transect B-B' (after Szafifinet al. 1999) across the ALCAPAterrane. (b) Hydrocarbon exploration seismic profile C-C' across the central part of the Danube Basin (after Tari 1994). (For location of profiles see Fig. 3.)

The third main type of magmatic activity in the PannonianCarpathian region took place during Late Miocene-Pleistocene times (12-0.5 Ma) with a climax at 3 - 5 Ma (Balogh et al. 1986; P6cskay et al. 1995), producing mainly alkali basalts and some other mafic rocks. Volcanic products are located throughout the Pannonian and Transylvanian basins and are of much lower volume than the calc-alkaline rocks. Trace elements and isotope ratios indicate a predominantly asthenospheric source for the related magmas (Downes e t al. 1995; Embey-Isztin & Dobosi 1995). Moreover, the volcanic products exhibit strong similarity to other Late Neogene basalts in Western Europe,

and their origin from a common mantle source, the so-called European Asthenospheric Reservoir, has been suggested (Cebri~ & Wilson 1995).

Style of deformation To illustrate the style of deformation related to basin evolution, we present four representative structural profiles (for location see Fig. 3). The first three sections ( A - A ' , B - B ' and C - C ' ) image the basin architecture and structural pattern of the two deepest

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Fig. 8. Interpreted seismic reflectionprofile D-D ~through the Zala Basin in the southwestern part of the PannonianBasin. (For location of profile see Fig. 3.) TWT, two-way travel time.

sub-basins, the Great Hungarian Plain and Danube Basin in the central and western part of the Pannonian Basin system, respectively (Figs 6 and 7). The fourth profile ( D - D f) is located at the boundary zone of the two terranes in SW Hungary, and is a good example of the anatomy of the late-stage basin inversion (Fig. 8). Deep reflection seismic profile A - A ' extends S W - N E across the deepest depressions in the Pannonian basin, the Hdd-Mak6 and B~k~s basins (Fig. 6). This area of the Tisza terrane has undergone considerable crustal and lithospheric extension, which is clearly reflected in the large thickness of the Neogene Basin infill, high heatflow values (D6v~nyi & Horvfith 1988) and gravity anomalies (Szafi~n et al. 1997). The profile images a largely deformed and extended upper crust and the presence of a shallow Moho surface. In this part of the Pannonian Basin the average thickness of the crust and lithosphere is around 25 and 60 km, respectively. Below the B6k6s basin, however, the Moho is at a depth of about 22 km and the thickness of the lithosphere is about 50 km (/~dfim et al. 1996). The line drawing interpretation of the section (Fig. 6) shows the presence of a set of low-angle normal faults that flatten out at mid-crustal depth. This regional detachment level is located at the transition between the bottom of the brittle upper crust and the top of the ductile lower crust. Low-angle normal faults dip to the NE in the Hdd-Mak6 basin, whereas on the NE side of the B~k~s Basin a switch of fault polarity can be observed. The two deep ( > 7 kin) depressions are separated by a basement high, the Pusztaf61dvgtr-Battonya block (PB in Fig. 3). Fission-track data indicate Miocene cooling ages for the Algy6 block, suggesting a genetic relationship between the onset of extension and exhumation of the footwall block (Taft et al. 1999). The style of deformation was controlled by the collapse of an overthickened, warm and thus weak lithosphere inherited from the Alpine orogeny. In this tectonic setting, Cretaceous nappe boundaries and thrust planes served as weakness zones and were reactivated as the low-angle sole of listric normal faults in an extensional (transtensional) regime. A similar structural pattern is revealed in the density model in section B - B ' (Fig. 7a) and the seismic section C - C ~ (Fig. 7b). Section B - B ' (Szafifin et al. 1999) traverses the AlpineCarpathian-Pannonian region, starting in the molasse zone of

the Eastern Alps, running through the Vienna Basin and terminating at the southern boundary of the ALCAPA terrane. Section C - C ' (Fig. 7b; after Tari 1994) shows the interpretation of an industrial seismic section in the southern part of the Danube Basin. The abundance of low-angle normal faults in the central and SE parts of sections B - B I and C - C ~ suggests a primarily extensional origin of the Danube Basin. These SE-dipping normal faults are again reactivated Cretaceous thrusts of the Austroalpine nappe stack. The faults penetrate down to mid-crustal levels where they sole out in regional detachment levels. They separate sub-basins from basement highs of different original structural positions: from NW to SE these are the Lower, Middle and Upper Austroalpine nappes. The stratigraphically highest units of the former nappe pile are on the right sides of the sections and are exposed further to the SE in the Transdanubian Range (Fig. 2). The stratigraphically lowest units are on the left sides of the sections and are represented by Lower Austroalpine nappes with blueschist-facies metamorphism (~rkai 2001). Crustal thickness values are similar to those in section A - A ' . In section B - B t the contact zone between the underplating Bohemian Massif (European plate) and the overriding Austroalpine unit (Adriatic plate) is interpreted as a gently dipping surface. The Bohemian massif can be traced at depth as far as the Mihfilyi high in the centre of the Danube Basin (marked by well M-27 in Fig. 7a), that is some 150 km behind the front of the Alpine orogen (Taft 1996). Seismic section D - D ' (Fig. 8) highlights the main structural features of late-stage basin inversion in the SW part of the Pannonian Basin. The profile traverses the broad boundary zone between the ALCAPA and Tisza-Dacia terranes (Mid-Hungarian shear zone) and images young, even active folding. The pre-rift basement is composed of a series of Palaeozoic and Mesozoic thrust sheets. The interpretation of the profile suggests a multi-phase tectonic evolution with distinct structural episodes and deformation styles. Similarly to the previous two examples, basin formation was initiated and controlled by the reactivation of suitably oriented Cretaceous thrust planes. The orientation of these faults is relatively constant (NE-SW) although their dip may switch (Budafa anticline). During Miocene times several half-grabens formed with considerable thickness of synrift deposits. The fault pattern

200

F. HORVATHETAL.

at Hahrt high indicates an important strike-slip component of faulting. Such a transtensional regime is consistent with current tectonic reconstructions, in which a mid-Hungarian shear zone played a key role in juxtaposing the two main terranes, that is the ALCAPA and Tisza-Dacia units (Csontos & Nagymarosy 1998; Fodor et al. 1999; Csontos & V6rrs 2004). During MioPliocene times, thermal subsidence resulted in the deposition of a thick sequence of post-rift sediments with an almost complete lack of faulting (e.g. Bajcsa syncline). Some grabens (Budafa) are completely inverted, suggesting differential vertical movements of the order of 500-1000 m during the Late Pliocene to Quaternary. Uplifted anticline cores have been considerably eroded. Surface topography follows the domal architecture of basin infill only in the southern part of the section. It is to be noted that basin inversion in this area is in the most advanced stage compared with other parts of the Pannonian Basin. This is due to the close proximity of the Adriatic indenter, which is considered to be the principal source of recent compressional stresses in the Pannonian lithosphere (Bada et al. 1998, 2001).

Lithospheric structure Formation of back-arc basins is a lithosphere-scale process. Knowledge of the crustal and lithospheric thickness maps of the Pannonian Basin system and the surrounding orogens provides further constraints on the mechanism of their formation and subsequent deformations. We use the two maps presented by Horvfith (1993) as starting models, and upgrade them by taking into account new data and interpretations presented in the past decade. The new map of crustal thickness (i.e. depth to the Moho discontinuity) is shown in Figure 9. There are only minor changes in the Pannonian Basin; the 3 0 k m isoline bounding the attenuated crustal domain remains the same. Inside this isoline a few new deep seismic profiles (Posgay et al. 1995, 1996) corroborated the earlier pattern or led to a better definition of a crustal thickness minimum in the southeastern part (compare with Fig. 6).

Furthermore, the improvement by Lenkey (1999) was also taken into account, as he has made the seismically derived pattern compatible with the gravity anomalies of the basin. Similarly, gravity modelling studies (Szafifin 1999), particularly 3D modelling (Szafifin & Horvfith 2006), suggest smaller thickness of crustal root associated with the southern Carpathians and its bend towards the eastern Carpathians (Vrancea area). This map is a significant improvement in the area of the Eastern Alps, and at the transition zone between the Southern Alps and the Dinarides (45-48~ 13-15~ This is because, in addition to a new synthetic map (Waldhauser et al. 1998), all available refraction and reflection seismic data have been reprocessed and reinterpreted, most recently by Cassinis & Scarascia (2003) as a contribution to an ambitious crustal profiling campaign across the Eastern Alps (TRANSALP Working Group 2002). Although this campaign has not yet finished, the first published results (e.g. Lueschen et al. 2003) are already thought provoking because of the debate on the interpretation of the observed reflectors and refractors. The two preliminary crustal profile alternatives, the 'crocodile model' and the 'lateral extrusion model', presented by the TRANSALP Working Group (2002) differ significantly below the Tauern window and the Periadriatic line to a depth of about 30 km. However, the two models agree perfectly in the geometry of the Moho discontinuity below the colliding European and Adriatic plates, and there is agreement in interpreting the results as showing the decoupling of the light continental crust from the denser mantle lithosphere (Fig. 5). In opposition to this consensus, Schmid et al. (2003) put forward a new interpretation for the complex geometry of the subducted slabs. Their results rely fundamentally on an improved tomography image of the mantle below the Eastern Alps, the Southern Alps and their transition towards the Dinarides (Lippitsch 2002). This improvement was possible because mantle tomography studies carried out in the Alpine-Mediterranean area have generally used a simplifying assumption: they supposed that in the calculation of seismic delay-time anomalies the crustal contribution had been negligible because of its fairly constant value all over the imaged territory.

Fig. 9. Depth to the Moho in the PannonianBasin and the surroundingmountains. Values are given in kilometres. 1, foreland (molasse)foredeep; 2, flyschnappes; 3, pre-Tertiary units on the surface; 4, Penninic windows; 5, Pieniny Klippen Belt; 6, trend of abrupt change in crustal thickness. Modifiedafter Horvfith(1993), Lenkey (1999), Cassinis & Scarascia (2003), and Szafifin& Horv~ith(2006).

PANNONIAN BASIN Obviously, this is not the case in the peri-Adriatic region, which is characterized by strongly contrasting crustal domains. Accordingly, the reliability and resolution in teleseismic tomography critically depends on the precise knowledge of the 3D crustal velocity structure. With an a p r i o r i crustal correction and selection of high-quality teleseismic data, Lippitsch (2002) has been able to decrease the cell size by a factor of 10 relative to previous studies and thus obtain a high-resolution image of the upper mantle structure beneath the Alps and surrounding areas. This confirms the earlier results in the Western and Central Alps depicting a high-velocity slab from the European foreland towards the SE and south to a depth of about 300 km beneath the Po plain (Kissling 1993; Solarino et al. 1996; Wortel & Spakman 2000). This European lithospheric slab underthrusting the Adriatic region appears to be continuous, but detached slabs of earlier subduction are also present in the deeper upper mantle (Davies & v o n Blanckenburg 1995). The real novelty of the high-resolution tomography image is the demonstration that there is a change in the polarity of subduction in the Eastern Alps. To the east of the western edge of the Tauern window, the Adriatic lower lithosphere is subducting beneath the Alpine wedge, probably reaching a depth of about 230 km. This northward dipping slab is made up of the continental part of Adria and probably an oceanic segment in the frontal part related to the former Vardar ocean (Lippitsch 2002). It can be considered as an active part of the subducting Adriatic lithosphere, which is apparently inactive and already fully detached further to the SE along the Dinarides (Wortel & Spakman 2000). Schmid et al. (2003) pointed out that two major orogenperpendicular post-collisional features at the surface coincide with the change of subduction polarity at depth: the Guidicarie strike-slip and the Brenner normal faults. They suggested that the northward dipping Adriatic subduction is a young feature (c. 20 Ma), coeval with the formation of these important orogenperpendicular features. It should be noted that the Brenner fault is the western boundary of the extruding ALCAPA wedge at the

201

Tauern window. Furthermore, they interpreted the TRANSALP profile in terms of northward descending Adriatic Moho under the European crust in analogy to the lithospheric configuration revealed by high-resolution tomography. We favour this interpretation, which is reflected in our crustal and lithospheric thickness maps (Figs 9 and 10). Unfortunately, the high-resolution tomography imaging (Lippitsch 2002) did not extend well into the region of the Dinarides and Pannonian Basin. Therefore the lateral extent of the subducted Adriatic slab towards the Dinarides remains unanswered, and hence, open to speculation. We assume that the surface expression of the boundary of the active Adriatic slab should be again an orogen-perpendicular structural feature, such as the Zagreb line (Fig. 10). The presence of a subducted lithospheric slab in the southern east Carpathians has been known for a longer time because of the remarkable deep seismic activity (down to 200 kin) in the Vrancea zone (Oncescu 1984) and early results of seismic tomography (Spakman et al. 1993). More recent studies (Fan et al. 1998; Wortel & Spakman 2000) have imaged the structure with a higher resolution and have shown a high-velocity slab at the same place but to a depth of 300-350 km. A specially devised regional campaign using about 150 broadband seismometers (Wenzel et al. 1998, 2002) revealed a nearly vertical high-velocity slab to a depth of 230-240 km. These results imply that the deeper part of the slab is aseismic. In addition, a horizontal high-velocity zone has been imaged between the 410 and 660 km seismic discontinuities, which is present beneath the entire Pannonian Basin (Wortel & Spakman 2000). The 410 and 660 km seismic discontinuities are also detected as an electrical conductivity increase by magnetotelluric and magnetovariational soundings (Adfim 1993; Semenov et al. 1997). Between them there is a transitional zone in electrical conductivity which may correspond to the high-velocity zone detected by the tomography. The horizontal high-velocity body may also represent a subducted lithospheric slab. If this is the case, it could have derived from subduction related to the consumption of the Mesozoic Vardar ocean, as

Fig. 10. Lithosphericthicknessmap of the PannonianBasin and the surroundingmountains. Values are given in kilometres. 1, Foreland (molasse)foredeep; 2, flysch nappes; 3, pre-Tertiary units on the surface; 4, Penninic windows; 5, Pieniny Klippen Belt; 6, trend of abrupt change in lithosphericthickness. Modified after Horv~ith(1993) and/~dfim & Wesztergom (2001).

202

F. HORVATHETAL.

well as from detachment of subducted slabs around the rest of the Carpathian arc. Detailed studies of the tectonic and magmatic evolution of the Apuseni Mts., Transylvanian Basin and east Carpathians (Linzer 1996; Girbacea & Frisch 1998; Gvirtzman 2002), together with the results of seismic tomography, suggest that the Vrancea slab is the only existing relict of the subduction roll-back that facilitated the extension of the ALCAPA and Tisza-Dacia terranes. Slab break-off is in progress, but the vertical continuity of the slab is still partly established between two orogen-perpendicular boundaries: the Trotu~ fault and the Intramoesian fault (Fig. 10). The lithospheric thickness map has also been improved remarkably in the Pannonian Basin as a consequence of new magnetotelluric soundings and sophisticated inversion of all available apparent resistivity and phase shift sounding curves (Adfim & Wesztergom 2001). This shows that the depth to the low-resistivity asthenosphere is on average 60-65 km in the basin, and locally, as the B6k& depression, rises to 50 km depth (Adfim et al. 1996). Tomography studies (Weber 2000) revealed the presence of relatively high seismic velocities beneath the B~k~s depression. Weber (2000) explained this phenomenon by local compositional differences in the mantle as a result of a mafic intrusion corresponding to the continental rift model as indicated by interpretation of gravity and magnetic measurements (e.g. Adfim & Bielik 1998).

Extrusion tectonics Extrusion of the ALCAPA terrane into the Carpathian embayment is a well-established and widely accepted concept (Fig. 5). However, the other constituent of the substrata of the Pannonian Basin, that is the Tisza-Dacia terrane, is characterized by a more complex and, hence, less understood structural and kinematic history. Paleotectonic reconstructions, albeit different in detail, agree that this terrane underwent Eo- and Meso-Alpine deformation in the Vardar zone, where continental collision took

place in the Early Tertiary (Csontos & V6r6s 2004). It is plausible to assume that the Tisza-Dacia block was extruded from this collision zone under similar geodynamic conditions to those of the ALCAPA terrane. This primarily implies that strain partitioning also occurred there and a Tisza-Dacia crustal flake was detached from its mantle lithospheric root. This means that, from the Late Oligocene, two compressionally deformed crustal wedges were invading the embayment of the Carpathian flysch basin. Evidence to support this view is as follows. Both terranes are characterized by extensive Early to Mid-Miocene silicic volcanism of identical geochemical composition (Harangi 2001). This implies that lower crustal melting and low crustal strength were common features of the two terranes. The anomalously high temperature gradient from the Early Miocene is well constrained by subsidence, thermal and maturation history analysis carried out for Neogene basins on both terranes (Horvfith et al. 1988). It was Sclater et al. (1980) who first arrived at the conclusion that the subsidence and sedimentation history of the Pannonian Basin system could not be explained in terms of homogeneous stretching of a normal lithosphere. Instead, the data were compatible with an inhomogeneous stretching model, whereby a modest crustal extension (stretching factor of 1.2-2.5) accompanied a dramatic attenuation of the mantle lithosphere (stretching factor of 5-50). More sophisticated modelling techniques, including evolution of the thermal field and organic matter maturation data (Horwith et al. 1986, 1988), corroborated this first conclusion about the apparent disappearance of mantle lithosphere at the beginning of the synrift phase in the Pannonian Basin system. Sclater et al. (1980) disagreed on the geodynamic interpretation of this conclusion, which was indeed debatable at the time. Now, we believe that extrusion of crustal flakes rather than entire lithospheric blocks offers the most plausible explanation (Fig. 5). In other words, extrusion tectonics in the Pannonian Basin is more compatible with the concept of orogenic floating (Oldow et al. 1989), rather than extrusion of lithospheric blocks as in Central Asia (Lav6 et al. 1996).

Fig. 11. Geodynamicmodel of the Quaternaryto recent PannonianBasin and surrounding orogens. 1, West European Platform and East European Craton separated by the Tornquist-Teisseyre (TT) suture zone; 2, foreland (molasse) foredeep; 3, Alpine-Carpathian nappe system; 4, exposed metamorphic core complexes (Penninicunits); 5, upliftingpart of the Pannonianbasin; 6, subsidingpart of the PannonianBasin; 7, extinct and active (duringthe Quaternary) calc-alkalinevolcanism; 8, thrust plane, inactive and active thrust front; 9, orogen-paralleldisplacementof the Pannonianterranes. M and L denote base of the crust (Moho) and the mantle lithosphere, respectively.

PANNONIAN BASIN

203

Conclusions

References

It can be seen from this review that a mass of new data has been collected in the Pannonian Basin and the surrounding orogens in the past decade, largely as a result of co-ordinated research efforts in the framework of the EUROPROBE programme. New data and their interpretations have led to significant progress in understanding the lithospheric dynamics associated with the formation and deformation of the Pannonian Basin and surrounding orogens (Fig. 11). It can be recognized that back-arc basin evolution is a single, although very complex, dynamic process of continental collision zones. This is because the various processes involved are strongly interrelated and their fundamental physical principle is the same: to minimize the potential and deformational energy during collision. The most important processes and their mechanisms are as follows: (1) Plate convergence leads to the complete consumption by subduction of the less buoyant oceanic lithosphere between the two continents, and then continent-continent collision occurs. (2) Continuing convergence results in splitting of the two continental lithospheres, preferably along major density a n d / o r rheological boundaries, and strain partitioning occurs. (3) When the light crust is detached from the continental lithosphere, the remaining mantle lid is heavier than the asthenosphere; hence, subduction can continue. Subduction polarity may remain the same as before; however, double subduction or a change of polarity can also take place. (4) At the upper levels, intricate structural processes produce the orogenic wedge, which consists of an interfingering stack of obducted parts of the consumed oceanic lithosphere, the detached continental crustal blocks and their sedimentary cover. (5) The orogenic wedge grows vertically and also in orogennormal direction, provided that convergence continues. Vertical growth implies thickening of the wedge and an increase in its elevation as a result of isostatic forces. Orogen-normal growth leads to the propagation of thrust faults towards both the undeformed foreland and the hinterland, and their incorporation into the orogenic wedge. (6) Parts of the orogenic wedge may extrude in orogen-parallel direction, provided that it decreases the potential energy and deformation energy dissipation of the system. This requires that a 'free boundary' is available sideways, which can be related to the presence of an oceanic basin. (7) Landlocked oceanic basins are particularly important, as their subductable oceanic lithosphere can give way to lateral extrusion of the orogenic wedge. Extrusion is a sideways extension of the orogenic material accompanied by gravitational collapse of the overthickened wedge. (8) Subduction of the oceanic lithosphere of the landlocked oceanic basin occurs by roll-back of the hinge zone (trench). This process continues until the downward force of the subducted slab is balanced by buoyant forces acting on the transitional to continental lithosphere attached to the subducting slab. (9) Roll-back terminates and the subducted slab breaks off progressively along the arc produced by the retreat of the landlocked oceanic lithosphere. This results in the ending of orogen-parallel extension of the extruding wedge, and the stress field in the back-arc basin changes to compressive. The fate of the orogen critically depends on the interaction of forces driving plate convergence and resisting plate collision. This is, however, a global energetic problem of the plate tectonic system.

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This paper is an outcome of the traditional co-operation between the Ertvrs University, Budapest and the Vrije Universiteit, Amsterdam. Financial support was provided by the Hungarian National Science Fund (OTKA) projects T034928, D34598 and F043715, and the Netherlands Research Centre for Integrated Solid Earth Science (ISES). The reviewers of the paper are thanked for their useful comments and suggestions.

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Modes of basin (de)formation, lithospheric strength and vertical motions in the Pannonian-Carpathian system: inferences from thermo-mechanical modelling S. C L O E T I N G H 1, G. B A D A l, L. M A T E N C O 1, A. L A N K R E I J E R 1, F. HORV/~TH 2 & C. DINU 3 1Netherlands Research Centre f o r Integrated Solid Earth Science, Faculty o f Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 H V Amsterdam, The Netherlands (e-mail: [email protected]) 2Department o f Geophysics, EOtvts L. University, Pdzmdny P. s. l / C , 1117 Budapest, Hungary 3Faculty o f Geology and Geophysics, University o f Bucharest, str. Traian Vuia 6, sect. 1, 70139 Bucuresti, Romania

Abstract: After a rapid multiphase evolution and a transition from passive to active rifting during late Early Miocene to Pliocene times, the Pannonian Basin has been subjected to compressional stresses leading to gradual basin inversion during Quaternary times. Stress modelling demonstrates the significance of the interaction of external plate-boundary forces and the effect of gravitational stresses caused by continental topography and crustal thickness variation. Flexural modelling and fission-track studies have elucidated the complex interplay of flexural downloading during collision, followed by rapid unroofing by unflexing and isostatic rebound of the lithosphere. The stretching and subsidence history of the Pannonian Basin, the temporal and spatial evolution of the flexure of the Carpathian lithosphere, and the lithospheric strength of the region reflect a complex history of this segment of the Eurasia-Africa collision zone. The polyphase evolution of the Pannonian-Carpathian system has resulted in strong lateral and temporal variation in thermomechanical properties in the area. Modelling results suggest that, as a whole, the Pannonian Basin has been an area of pronounced lithospheric weakness since Cretaceous time, shedding light on the high degree of strain localization in this region. This basin, the hottest in continental Europe, has a lithosphere of extremely low rigidity, making it prone to multiple tectonic reactivations. Another feature is the noticeable absence of lithospheric strength in the mantle lithosphere of the Pannonian Basin. Modelling studies suggest pronounced lateral variations in lithospheric strength along the Carpathians and their foreland, which have influenced the thrust load kinematics and post-collisional tectonic history. The inferences and models discussed in this paper are constrained by a large geophysical database, including seismic profiles, gravity and heat-flow data.

The Pannonian-Carpathian system in Central and Eastern Europe has been the focus of considerable research efforts concerning the integration of geophysical and geological data, making it a key area for quantitative basin studies. A vast geophysical and geological database has been established in recent decades as a result of a major international research collaboration in this area, largely carried out in the framework of European programmes such as the EU Integrated Basin Studies project (Cloetingh et al. 1995; Durand et al. 1999), the ILP-ALCAPA (Cloetingh et al. 1993; Neubauer et al. 1997) and EUROPROBEPANCARDI (Decker et al. 1998) programmes, and the Peri-Tethys programme (Ziegler & Horwith 1996; Brunet & Cloetingh 2003), partly funded in the context of petroleum exploration. These studies, building on previous compilations (Royden & Horvfith 1988), marked a major advance in applying basin analysis concepts to the Pannonian-Carpathian system. An important asset of this natural laboratory is the existence of high-quality constraints on basin evolution obtained through the systematic acquisition of seismic, gravity, heat-flow and magnetotelluric data by various research groups (see Radulescu et al. 1976; Royden & Horv~ith 1988; D6vtnyi 1994; Demetrescu & Andreescu 1994; Ionescu 1994; Posgay et al. 1995; Szafi~n et al. 1997; Adfim & Bielik 1998; Tari et al. 1999; Wenzel et al. 1999; Hauser et al. 2001; Lenkey et al. 2002). Extensive well coverage in the context of petroleum exploration and surface studies have allowed the construction of a high-resolution stratigraphic framework for the area (e.g. Vakarcs et al. 1994; Sacchi et al. 1999; Vasiliev et al. 2004). At the same time, the fold-and-thrust belt has been the focus of a concentrated effort, highlighting the connection between lateral variations in structural style, basement characteristics and foreland flexure development in various segments of the Carpathian orogen (Sfindulescu 1988; Roure et al. 1993; Matenco et al. 1997a,b, 2003; Taft et al. 1997; Sanders et al. 1999; Zoetemeijer et al. 1999) and its hinterland, the Transylvanian Basin (e.g. Ciulavu et al. 2002). The Pannonian-Carpathian system, therefore, allows us to test models for basin formation and subsequent deformation, for continuing orogeny and continental collision. This system

comprises some of the best documented sedimentary basins in the world, located within the Alpine orogenic belt, at the transition between the Western European lithosphere and the East European Craton. The Pannonian Basin evolved from its synrift to post-tiff phase during Early to Late Miocene times (c. 2 0 5 Ma), when back-arc extension was coupled with subduction and collision dynamics in the Carpathian orogenic arc system (Royden & Horvfith 1988). The lithosphere of the Pannonian Basin is a particularly sensitive recorder of changes in lithospheric stress induced by near-field intra-plate and far-field plate boundary processes (Bada et al. 2001). High-quality constraints exist on the regional (palaeo)stress (Fodor et al. 1999; Gerner et al. 1999) fields in the lithosphere as a result of earthquake focal mechanism studies, analyses of borehole break-outs and studies of kinematic field indicator data. A close relationship has been demonstrated between the timing and nature of stress changes in the extensional basin and structural episodes in the surrounding thrust belts, pointing to a mechanical coupling between the orogen and its back-arc basin. The Pannonian Basin, the hottest in continental Europe, is thought to have gone through a rapid transition from passive to active tiffing during Late Miocene times, simultaneously with the climax of compression in the Carpathian arc (Huismans et al. 2001). In parallel, significant efforts have been made to reconstruct the spatial and temporal variations in thrusting along the Carpathian orogen (Roure et al. 1993; Schmid et al. 1998; Zweigel et al. 1998; Matenco & Bertotti 2000) and its relationship to foredeep depocentres (Meulenkamp et al. 1996; Matenco et al. 2003; Tfirfipoancfi et al. 2003), changes in foreland basin geometry and lateral variations of flexural rigidity. A general feature of the flexural modelling studies carried out for the Carpathian system (e.g. Matenco et al. 1997b; Zoetemeijer et al. 1999) is the inferred low rigidity of the platform lithosphere downbending under the SE Carpathians. Flexural studies constrained by gravity (e.g. Szafifin et al. 1997) also point to an important role of flexural unroofing of the Carpathian mountain chain and its foredeep. A crucial element of the dynamics of lithospheric deformation is the mechanics of coupling back-arc deformation in the

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 207-221. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Pannonian Basin with continental collision and foreland basin evolution along the Carpathian arc. This has been addressed through a combination of dynamic and kinematic modelling studies constrained by integrated basin analysis in the Pannonian sector and by thermochronology, basin modelling and structural field studies in the Carpathian belt. In particular, the inversion events (i.e. Late Miocene and Late Pliocene-Quaternary) recorded in the Pannonian Basin (Horvfith 1995; Fodor et al. 1999) are coeval with the climax of thrusting in the Carpathians (e.g. S~ndulescu 1988; Hippolite et al. 1999) related to continental collision and late-stage out-of-sequence contraction. Horv~ith & Cloetingh (1996) established the importance of Late Pliocene to Quaternary compression in the Pannonian-Carpathian system, explaining its anomalous Quaternary uplift and subsidence pattern as well as its intraplate seismicity, and thus establishing a novel conceptual model for structural reactivation of back-arc basins in orogenic settings. The basin system has reached an advanced stage of evolution with respect to other Mediterranean back-arc basins and its structural inversion has been taking place for the last few millions years. Basin inversion is related to temporal changes in the regional stress field (for a general discussion, see Ziegler et al. 1995, 2002), from one of tension that controlled basin formation and subsidence, to one of compression resulting in contraction and flexure of the lithosphere associated with differential vertical movements. This paper presents an overview of the tectonic evolution of the Pannonian-Carpathian system in terms of various thermomechanical models. First, the focus is placed on the formation of the system as inferred from subsidence history and stretching models of the Pannonian Basin, and the flexural behaviour of the Carpathian lithosphere. Then a lithospheric strength map of the system is presented, which highlights rheological constraints for basin analysis studies and for the reconstruction of the deformation history of the area. The reactivation or the neotectonics of the region is described in terms of anomalous late-stage vertical movements, that is accelerated subsidence in the centre of the Pannonian Basin and fast uplift of the Carpathians orogen as a result of the rebound of the crust in the aftermath of continental convergence. The paper is concluded with a discussion on the thermo-mechanical aspects of basin inversion, lithospheric folding, and related temporal and spatial variations of continental topography in the Pannonian-Carpathian system.

Formation of the Pannonian-Carpathian system Formation and N e o g e n e evolution o f the Pannonian Basin Dynamic models of basin formation. After a long period of continen-

tal convergence in the Alpine belt during Cretaceous to Palaeogene times, a rapid change in tectonic style in the late Early Miocene led to the formation of the Pannonian Basin in the area of the Carpathian embayment. Consequently, the relatively stable Palaeogene to Early Miocene assembly of continental blocks at the axial zone of Adria-Europe convergence was completely disintegrated and these units experienced a significant amount of stretching, rigid body rotation and translation. This process was coeval with the formation and early evolution of the Pannonian Basin and the large-scale tectonic transport of the flysch nappes in the Carpathian arc (e.g. Balla 1984; S~ndulescu 1988; Csontos et al. 1992; Roure et al. 1993; Kov~i~ et al. 1994; Fodor et al. 1999). Several models have been proposed to explain the dynamics of Neogene tiffing in the Pannonian Basin. An active v. passive mode of rifting has been a matter of continuous debate in the scientific literature (for a review, see Bada & Horwith 2001), resulting in the advent of various dynamic models during the last half-century (Fig. 1). These different theoretical scenarios have been provoked by the most evident features of this basin system, that is thinned and hot v. thickened and colder lithosphere in the central and peripheral sectors, respectively. For instance, Sz~deczky-Kardoss (1967) and, at least in his early works, Stegena (1967) argued for the presence of a mantle diapir beneath the Intra-Carpathian area (Fig. 1a). In this model the ascending mantle flow resulted in thinning and subsidence (active rifting) in the central areas, whereas the nappe structure and the root of the surrounding orogens were formed above the descending branch of a local convection cell. Active tiffing was also proposed by, for example, Horv~th et al. (1975) and Stegena et al. (1975) as an ultimate origin of the Pannonian Basin. However, their model already employed the concept of plate tectonics for the Pannonian region, where basin formation and subsidence, intense Neogene-Quaternary volcanic activity, extremely high heat flow, and the presence of an anomalous upper mantle and thinned crust were considered as closely related phenomena. These features were probably controlled by subcrustal erosion of the underlying lithosphere by an active mantle diapir

Fig. 1. Dynamic models proposed for the evolution of the PannonianBasin system (after Bada & Horvfith2001). (a) Asthenospheric doming results in active rifting of the lithosphere above the central axis of the dome, whereas shorteningis taking place in the peripheral areas. (b) Active rifting may also be caused by a subduction-generated mantle diapir. (c) Hinge retreat of the subductingEuropean margin driven by the negativebuoyancyof the slab induces passive rifting in the overriding plate. (d) The same hinge retreat may be sustainedby an eastward mantle flow pushing against the downgoing slab. (For references see text.)

MODELLING THE PANNONIAN-CARPATHIANS SYSTEM above the European and Apulian (Adriatic) plates subducting under the Pannonian plate fragment (Fig. lb). Other dynamic models also suggest the key importance of subduction as a driving force, and the back-arc position of the Pannonian Basin (e.g. Royden & Horwith 1988; Csontos et al. 1992; Horv~th 1993; Csontos 1995; Linzer 1996). Extension and lithosphetic stretching are caused by the hinge retreat of the subducting European plate along the Carpathians that stretches the overriding Pannonian plate (Fig. lc). The process of slab retreat is caused by the negative buoyancy of a subducting slab (Royden & Karner 1984). Another model (Fig. l d) suggests that an eastward mantle flow may have been pushing the subducting slab and, hence, can be responsible for the same hinge retreat (Doglioni 1993). Nevertheless, both models account for the passive tiffing of the Pannonian Basin, where tension is facilitated by trench suction forces exerted at the contact zone of the overriding and subducting plates. More recently, Huismans et al. (2001) used numerical simulation to explain the temporal changes in rifting style of the Pannonian Basin. The results of their thermo-mechanical finiteelement modelling confirm the two-phase evolution scheme of the system, that is a synrift episode followed by a post-rift period. Initiation of basin formation was mainly driven by passive rifting as a result of the effect of Carpathian subduction and the gravitational collapse of the thickened pre-rift Pannonian lithosphere. This triggered a small-scale convective upwelling of the mantle lithosphere that could have led to the second, active tiffing phase of the Pannonian Basin during Late MiocenePliocene times, coeval with the climax of compression in the Carpathian arc. Stretching models, subsidence analysis. Efforts on the quantification of basin evolution started in the early 1980s by means of classical basin analysis techniques. The Pannonian Basin has been a key area for testing stretching models, because of the availability of excellent geological and geophysical constraints. At the same time, the main characteristics of the basin system, such as the extremely high heat flow, the presence of an anomalous thinned lithosphere and its tectonic position in the Alpine regime of overall convergence, made it particularly suitable and challenging for basin analysis. Interest in this research in the Pannonian Basin was mainly generated through activities in the field of deep crustal and mantle processes, local tectonic and regional correlation studies, and hydrocarbon prospecting. The stretching model of McKenzie (1978) was first applied to the intra-Carpathian basins by Sclater et al. (1980). They found that the formation of the peripheral basins could be fairly well simulated by the concept of uniform extension, using a stretching factor of about two (/3 = 2). In the more central basins, however, the considerable thermal subsidence and high heat flow suggested unrealistically high stretching factors (/3 up to five). Thus, they postulated a differential extension within the Pannonian lithosphere, with moderate crustal thinning being accompanied by a higher level of stretching at subcrustal depth. Building on this and using a wealth of well data, Royden et al. (1983) introduced the concept of modified or non-uniform stretching, in which the amount of extension is depth-dependent in the lithosphere. The subsidence and thermal history of major parts of the Pannonian Basin suggests a large attenuation of the mantle lithosphere with respect to a finite crustal extension. Horv~th et al. (1988) further improved this concept by considering radioactive heat generation in the crust, and the thermal effect of basin-scale sedimentation. By reconstructing the subsidence and thermal history, the thermal maturation of organic matter in the central region of the Pannonian Basin (Great Hungarian Plain) was calculated, resulting in a major step forward in the field of hydrocarbon prospecting by means of basin analysis techniques. These studies highlighted the considerable difficulties in explaining basin subsidence and crustal thinning in terms of uniform extension, pointing to the presence of anomalous subcrustal

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thinning. The issue has been a central theme for subsequent investigations by means of quantitative subsidence analyses (backstripping) of a more extended set of sections and wells in the Pannonian Basin, and by means of forward modelling techniques (Lankreijer et al. 1995; Sachsenhofer et al. 1997; Juhfisz et al. 1999; Lenkey 1999). Kinematic modelling incorporating the concept of necking depth and finite strength of the lithosphere during and after rifting (van Balen et al. 1999), as well as dynamic modelling studies (Huismans et al. 2001) suggested a transition from passive to active tiffing as a mechanism for subcrustal flow and small-scale convection in the lithosphere-asthenosphere system. To quantify lithospheric deformation on a whole-basin scale, Lenkey (1999) carried out forward modelling using the concept of non-uniform stretching and taking into account the effects of lateral heat flow, flexure and necking of the lithosphere. Calculated crustal thinning factors (6) indicate large lateral variation of crustal extension in the Pannonian Basin (Fig. 2). This is consistent with the areal pattern of pre-Neogene basement depth (Horvfith et al. 2006). The range of crustal thinning factors indicates 10100% crustal extension, which is in good agreement with the pre-rift palinspastic reconstruction of the Pannonian Basin, and the amount of cumulative shortening in the Carpathian orogen (e.g. Roure et al. 1993; Fodor et al. 1999). As a major outcome of basin analysis studies, Royden et al. (1983) provided a two-stage subdivision for the evolution of the Pannonian Basin, with a synrift (tectonic) phase during Early to Mid-Miocene times and a post-rift (thermal) phase during the Late Miocene-Pliocene period. Further development of the stratigraphic database, however, demonstrated the need to refine this scenario. According to Tari et al. (1999), a regional Mid-Badenian unconformity indicates the termination of the synrift period, which is followed by a post-rift phase with only minor tectonic activity. In either case, the subsidence history of the Pannonian Basin can be subdivided into three main phases as reflected in the subsidence curves of selected sub-basins (Fig. 3). First, the synrift phase characterized by a rapid tectonic subsidence, started synchronously at about 20 Ma in the entire Pannonian Basin. This phase of pronounced crustal extension is recorded everywhere in the basin system and was mostly limited to relatively small, faultbounded grabens or sub-basins. Second, the post-rift phase of extension affected much broader areas, causing general downwarping of the lithosphere manifested in a thermal phase of subsidence. This is particularly well defined in the central parts, suggesting a gradual transition from thin-skinned to wholelithosphere extension towards the centre of the basin system (e.g. Sclater et al. 1980; Royden & Drvrnyi 1988). The third (final) phase of basin evolution is characterized by the gradual structural inversion of the Pannonian Basin system during Late Pliocene-Quaternary times. As a result, the recent build-up of intraplate compressional stresses has caused basin-scale buckling of the Pannonian lithosphere associated with late-stage subsidence anomalies and differential vertical motions (Horwith & Cloetingh 1996). As seen in the subsidence curves (Fig. 3), an accelerated subsidence has been taking place in the central depressions (Little and Great Hungarian Plain; Fig. 3b and c), whereas most peripheral sub-basins have been uplifted by a few hundred metres after mid-Miocene times (Styrian and East Slovakian basins; Fig. 3d and e) or during the Pliocene-Quaternary period (Zala basin; Fig. 3f). The late-stage tectonic reactivation, as well as other episodic inversion events in the Pannonian Basin (Horvfith 1995; Fodor et al. 1999), highlight the importance of tectonic stresses not only in the stretching phase (tension) but also during subsequent inversion phases (compression). Modelling curves for the Carpathian foreland (Fig. 3h-j) indicate an important phase of basin subsidence related to the collision in the east and south Carpathians. Moreover, this period is coeval with the cessation of synrift subsidence in the Pannonian Basin (see also Horv~ith & Cloetingh 1996). Subsidence curves for the intermediate Transylvanian Basin (Fig. 3g) indicate a dual behaviour;

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Fig. 2. Crustal thinning factors (6) calculated by means of forward modelling for the PannonianBasin employingthe concept of non-uniformstretching complemented by the effects of lateral heat flow and the flexure of the lithosphere (after Lenkey 1999). The pronounced lateral variation of crustal extension, leading to the formation of deep sub-basins connected by areas of lower level of deformation, should be noted. AM, Apuseni Mts; DIN, Dinarides;EA, Eastern Alps; TR, TransdanubianRange; SC, WC, southern and western Carpathians, respectively. Local depressions of the Pannonian Basin system: B~, Brk~s; Da, Danube; De, Derecske; D, Drava; ES, East Slovakian; Jfi, Jfiszsfig;Ma, Mak6; Sa, Sava; St, Styrian; Vi, Vienna; Za, Zala. whereas the Badenian-Pannonian stage reflects a subsidence pattern similar to that in the Carpathian foreland, the major Pliocene-Quaternary inversion in the Pannonian Basin is also recorded in the Transylvanian domain.

F o r m a t i o n a n d N e o g e n e evolution o f the Carpathian system

Over the last few years, research has focused on spatial and temporal variations in thrusting along the Carpathian arc and their relationship to unusual foredeep geometry and lateral variations in flexural behaviour. The reconstruction of orogenic uplift and erosion (e.g. Sanders et al. 1999) coupled with foreland subsidence modelling (Matenco et al. 2003) have elucidated the complex interplay of flexural downloading during collision and its lateral variability, followed by isostatic rebound in the orogenic wedge and increased subsidence in the SE Carpathian corner (Matenco et al. 1997b; Zoetemeijer et al. 1999; Bertotti et al. 2003; TS.rfipoancfi et al. 2003; Cloetingh et al. 2004). The Carpathians represent a highly arcuate orogenic belt formed in response to subduction and continental collision between the European and Apulian plates and related microplates during the Alpine orogeny (Sfindulescu 1988; Csontos & Voros 2004). It consists of a nappe pile of crystalline rocks with Late Palaeozoic to Mesozoic sedimentary cover and, in an external position, an Early Cretaceous to Tertiary thin-skinned belt. The Alpine tectonic evolution of the Carpathians is traditionally subdivided into Triassic to Early Cretaceous extension followed by MidCretaceous to Pliocene shortening (e.g. Sandulescu 1988). Three main Tertiary deformation stages are recognized (e.g. Matenco & Bertotti 2000, and references therein). During PalaeogeneEarly Miocene times, the clockwise rotation of the Rhodopian fragment (Balla 1986; Schmid et al. 1998), part of the

Pannonian-Carpathian system, caused N N E - S S W to E N E WSW shortening in the internal Moldavides nappes (Sandulescu 1988). The ENE-WSW-directed shortening acting in Mid-to Late Miocene (Late Burdigalian to Sarmatian) times was responsible for the major nappe emplacement in the east Carpathians. The oroclinal shape of the thrust belt must have been initiated in the later phases of deformation, as a result of the irregular plate boundaries and the lateral variations in thickness of the sedimentary wedge involved in shortening (Matenco & Bertotti 2000). Widespread left-lateral shearing occurred along numerous eastwest-oriented faults in the north and NW-SE-trending dextral transpressive structures in the south, accommodating movement to the ESE of the intervening central sectors. Late Miocene to Pliocene N W - S E (to north-south) shortening in the east Carpathians led to further deformation mainly concentrated in the external parts of the junction zone between the east and south Carpathians (Bend Zone), and the foreland continued to subside throughout the post-collisional period, accumulating up to 6 km of Pliocene-Quaternary sediments in the Focsani Basin area (Tfir~poancfi et al. 2003). The Carpathians represent the birthplace for popular theories such as slab-retreat and slab roll-back based on flexural modelling results (e.g. Royden & Karner 1984; Royden 1993), which account for a gradual, eastward retreat of the plate contact during the Miocene contractional events. Detailed flexural modelling studies indicate that the European lithosphere primarily controls the deflection in the Carpathian foredeep (Zoetemeijer et al. 1999; Fig. 4b). It appears that small remnants of oceanic slab cause slab-pull forces, increasing in magnitude in an eastward direction. Modelling results indicate the regional importance of post-collisional uplift, which is most pronounced in the vicinity of the Bohemian massif, related to flow of mantle material away from the multiple rifting Pannonian Basin (Huismans

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Fig. 3. Subsidence curves for selected sub-basins of the Pannonian Basin and Carpathian foreland (compiled after Horvfith et al. 1988; D6v6nyi 1994; Lankreijer et al. 1995; Sachsenhofer et al. 1997; Ma~enco et al. 2003). It should be noted that after a rapid phase of general subsidence throughout the Pannonian Basin, the sub-basins show a distinct subsidence history from Mid-Miocene time. Arrows indicate generalized vertical movements. Timing of the synand post-rift phases is after Royden et al. (1983). Timing of Carpathian collision after Matenco et al. (2003). For the time scale the central Paratethys stages of R6gl (1996) are used. O, K, Ottnangian and Karpatian, respectively (Early Miocene); BAD, SA, Badenian and Sarmatian, respectively (Mid-Miocene); PAN, PO: Pannonian and Pontian, respectively (Late Miocene); PL, Pliocene; Q, Quaternary. AM, Apuseni Mts; DIN, Dinarides; EA, Eastern Alps; PanBas, Pannonian Basin; SC, WC, southern and western Carpathians, respectively.

e t al. 2001). The European foredeep configuration is clearly controlled by lateral changes in the European lithosphere strength. These changes are dominated by the amount of internal deformation caused by the curvature of the lower plate during the emplacement of the Carpathians, weakening of the elastic plate

coinciding with the areas of m a x i m u m flexural bending stresses (Zoetemeijer e t al. 1999). Flexural modelling studies also predict a general back-stepping of the lower plate system in the transition area between European and Moesian lithosphere (Matenco e t al. 1997b) along the crustal-scale Trotu~ fault

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(Fig. 4c), accounting for crustal lower plate tearing and disruption of the mechanical continuity between the two types of lithosphere involved in collision. The kinematics of this fault records a large-scale sinistral movement during collision, and continuing subsidence in the post-collisional period has induced the downward throw of this fault to the south. Modelling the deflection of the Moesian lithosphere below the south Carpathian nappe pile indicates a plate contact placed in an advanced, foreland position, demonstrating the small amount of general n o r t h - s o u t h oriented shortening in an overall transpressional regime (Matenco e t al. 1997b). Recent 3D flexural studies (Tfir~poancfi e t al. 2004) indicate that the subsidence of the unusual deep SE Carpathian foredeep is affected by pre-orogenic extension followed by a 3D emplacement of the thrust load on top of large lateral changes in lithospheric strength in the downgoing plate. Under these conditions the m a x i m u m deflection moves out of the orogenic load and accounts for large-scale subsidence in the Bend Zone (Fig. 4d).

The subsidence pattern recorded in the Carpathian foreland units during the Tertiary is characterized by significant vertical motions (e.g. Matenco e t al. 2003; Bertotti e t al. 2003; Fig. 3h-j). Opening of the Early Miocene transtensional basin in the Getic depressionwestern Moesian platform led to the deposition of up to 5 km of sediments, whereas the other platform areas were characterized by non-deposition a n d / o r erosion. Starting from the Mid- to Late Miocene, the entire Carpathian foreland started to subside (e.g. at a rate of 1 5 0 0 - 3 0 0 0 m M a -1 in and around the Focsani depression in the Late Miocene) as a direct response to coeval thrust loading, the major depocentre of the foreland basin being located in the Bend Zone. Following the collisional event at the beginning of the Late Miocene, the subsidence continued particularly in the Focsani depression, where a rate of 2 0 0 - 3 0 0 m Ma -1 is recorded for the entire Pliocene-Pleistocene). Subsidence analysis in the Carpathian foreland demonstrated comparable kinematically related vertical motion episodes simultaneously occurring in the frontal part of both the east and south Carpathians

Fig. 4. Flexural studies across the Carpathians foreland. (a) Location of the flexural studies in the Carpathians orogen. (b) Simplified geological map of the western Carpathians, indicating the predicted flexural bending stresses (magnitudes are in GPa) in the 2D modelled profiles (after Zoetemeijer et al. 1999). (c) Computed contour map of the basement shape along the Romanian Outer Carpathian foreland system, juxtaposed over the structural map of the autochthonous platforms (for details see Fig. 9). BF, Bistri~aFault; TF, Trotu~ Fault; LR, lateral ramp. Inset A shows a model of the late Miocene strike-slip escape above a late Cretaceous-Palaeogene lateral ramp. Inset B shows a model of Mid- to late Miocene deformations in the vicinity of the Trotu~ Fault (after Matenco et al. 1997). (d) A 3D flexural model for the east and south Carpathians foreland. Underlined numbers represent the predicted EET for each flexural block. Shaded zones represents simplified topography of the area. CSmax represents the predicted maximum deflection (after T~poanc~ et al. 2004).

MODELLING THE PANNONIAN-CARPATHIANS SYSTEM (Bertotti et al. 2003; Matenco et al. 2003), with limited to no depocentre migration along the east Carpathian foreland, in contrast to previous inferences (e.g. Meulenkamp et al. 1996). Despite the apparent continuity of sedimentary facies and tectonic units at the surface in the thin-skinned units of the east Carpathians, the analysis of kinematic thrust emplacement has revealed a non-cylindrical Miocene shortening, as a result of lateral variations in the strength of d~collement horizons and mechanical properties of the foreland lower plate (Matenco & Bertotti 2000). This non-cylindrical character refers to the amount of internal shortening (higher to the north) and to the variable character of the associated uplift and erosion. This is particularly important in the SE Carpathians, where a reduced amount of internal shortening during collision has failed to induce significant uplift (in any case below fission-track resolution; see Sanders et al. 1999) to induce significant erosion. Thermo-mechanical modelling (Cloetingh et al. 2004) suggests that collision with the Moesian platform in this sector of the belt locked the system in the early phases and led to thermal re-equilibration in the post-collisional phase, when the subducted slab started to stretch vertically. In contrast, continuation of the oceanic-type subduction in the central-northern domain means that a large part of the former passive margin, comprising thinned to stable continental crust of the lower plate, was underplated in the collision process. This induced significant internal shortening in the thin-skinned nappe pile at the contact and 4 - 5 km of uplift followed by erosion (Sanders et al. 1999). During the post-collisional phase in this central-northern part, the system became welded, the minor uplift observed being related to orogenic rebound. In contrast, the post-collisional phase of the Bend Zone is characterized by large subsidence in the foredeep, whereas a large amount of uplift is recorded towards the neighbouring western nappe pile with no orogenic-related shortening, probably in response to crustal folding (Bertotti et al. 2003). The drag of the vertical stretching slab accounts for a reduction in the orogenic load needed for the observed deflection in the foreland, as inferred also by 3D flexural modelling (Tfir~poancfi et al. 2004).

Lithospherie strength in the Pannonian- Carpathian system The mechanical behaviour of the lithosphere can be formulated in different theological models providing depth-dependent stressstrain relationships (see Ranalli 1995). The vertical strength distribution of the lithosphere is often represented by means of rheological profiles, or yield envelopes, in which the yield limit is plotted against depth. These profiles rely on the results of laboratory experiments extrapolated to different levels in the crust and mantle, and a number of assumptions regarding the thermal structure, strain rate, presence of fluids and several other factors. The depth dependence of these parameters results in a rheological stratification of the lithosphere. Accordingly, strain is mostly concentrated in the weak, mostly ductile levels of the lithosphere. These levels are good candidates for necking during extension or the formation of large-scale shear zones, detachment horizons, etc., and can lead to decoupled deformation between layers of higher strength. The Pannonian-Carpathian system shows a remarkable variation of thermo-mechanical properties of the lithosphere. Lithospheric rigidity varies in space and time, giving rise to important differences in the tectonic behaviour of different parts of the system. As theology controls the response of the lithosphere to stresses, and thus the formation and deformation of basins and orogens, the characterization of rheological properties and their temporal changes has been a major challenge to constrain and quantify tectonic models and scenarios. This is particularly valid for the Pannonian-Carpathian region, where tectonic units of different history and rheological properties are in close contact.

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Figure 5a displays three strength envelopes through the western, central and eastern part of the Pannonian lithosphere, constructed on the basis of extrapolation of rock mechanics data and incorporating constraints on crustal and lithospheric structure, and presentday heat flow along the modelled rheological section. These strength profiles show that the average strength of the Pannonian lithosphere is very low (see also Lankreijer 1998); this is mainly due to high heat flow, which is the consequence of an elevated asthenosphere dome below the basin system. The hottest basin in continental Europe has a lithosphere of extremely low rigidity, making it prone to repeated tectonic reactivation. This is the result of earlier phases of tectogenesis, namely nappe emplacement, crustal accretion and related orogenic thickening and destabilization during Cretaceous-Palaeogene times. At the same time, the strength of the Pannonian lithosphere segments gradually decreased, leading to the collapse of these orogenic terranes, a high level of strain localization and, eventually, the formation of the Pannonian Basin, that is rifting and extension of former compressional domains. Another essential feature is the noticeable absence of present-day lithospheric strength in the mantle lithosphere of the Pannonian Basin. Strength appears to be concentrated in the upper 7 - 1 2 km of the lithosphere. This finding is in very good agreement with the distribution of seismicity at depth. Earthquake hypocentres are restricted to uppermost crustal levels, suggesting that brittle deformation in the lithosphere is taking place not deeper than 5-15 km (T6th et al. 2002). Figure 5b shows estimates of the total integrated strength (TIS) of the Pannonian-Carpathian lithosphere along section A - A ' . Rheology calculations suggest major differences in the mechanical properties of different tectonic units within the system (Lankreijer et al. 1997, 1999). In general, there is a gradual increase of TIS from the centre of the basin towards the basin flanks in the peripheral areas (see also Fig. 5c). The centre of the Pannonian Basin and Carpathian foreland are the weakest and strongest parts of the system, respectively. The presence of a relatively strong lithosphere in the Transylvanian Basin is due to the differences in tectonic history with respect to the Pannonian Basin. The absence of large-scale Tertiary extension left this lithosphere segment relatively strong. The Carpathian arc, particularly its western parts, shows a high level of rigidity, with the exception of the southeastern Bend Zone, where a striking decrease of lithospheric strength is noticeable. Calculations for the seismically active Vrancea area indicate the presence of a very weak crust and mantle lithosphere, indicating mechanical decoupling between the Transylvanian Basin and the Carpathian orogen. The pronounced contrast in TIS between the Pannonian Basin (characterized by TIS < 2.0 x 1012 N m -1) and the Carpathian orogen and its foreland (characterized by TIS > 3.0 x 1012 N m -1) indicates that recent lithospheric deformation is more likely to be concentrated in the hot and hence weak Pannonian lithosphere than in the surrounding Carpathians. The results of rheology calculations are often expressed in terms of the effective elastic thickness (EET) of the continental lithosphere by integrating the thickness of the mechanically strong layers of the lithosphere (Burov & Diament 1995). EET values can also be obtained from forward modelling of extensional basin evolution and from flexural models of the lithosphere, which in turn can be compared with the inferences from rheological studies. This provides an important means to test the validity of different modelling techniques. By the conversion of strength predictions to EET values on a regional scale, Lankreijer (1998) constructed a map of EET distribution for the entire PannonianCarpathian system (Fig. 6). Calculated EET values are largely consistent with the spatial variation of lithospheric strength in the system. Lower values are characteristic for the weak central part of the Pannonian Basin (5-10 km), whereas EET increases towards the neighbouring orogens (15-30 kin) and, particularly, in the foreland areas in the Bohemian Massif and Moesian Platform (25-40 km). This trend is in good agreement with EET

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Fig. 5. (a) Typical strength envelopes from the western (A), central (B) and eastern (C) part of the Pannonian Basin (for location see Fig. 6) constrained by data on lithospheric structure, heat flow, strain rate and stress regime (for details see Lankreijer et al. 1997; Sachsenhofer et al. 1997; Lenkey 2002). The nearly complete absence of mantle strength predicted by the model should be noted. (b) Total integrated lithospheric strength (TIS) along a regional profile through the Pannonian-Carpathian system (Lankreijer 1998). (c) Schematic cross-section showing the non-uniform stretching of the Pannonian lithosphere and its effect on depth-dependent rheology. The thickness of the crust (c) and mantle lithosphere (m) is reduced by the stretching factors 6 and/3, respectively, in the basin centre. Because ascending asthenosphere is heating the system, the isotherms become significantly elevated. As a result, the thinned and hot Pannonian lithosphere becomes extremely weak and, thus, prone to subsequent tectonic reactivation.

estimates obtained from flexural studies and forward modelling of extensional basin formation. Systematic differences, however, can exist. This may be the consequence of significant horizontal intraplate stresses (e.g. Cloetingh & Burov 1996) or the mechanical decoupling of the upper crust and uppermost mantle, which can lead to a considerable reduction of EET values. As lithospheric rheology is primarily controlled by temperature, a general relation is expected between the thermal age (time elapsed since the last thermal event) and the rheology of the lithosphere. This relation is straightforward for oceanic crust but is much more complex for continental areas. In Figure 7 estimated EET values and thermal age for selected locations in the Pannon i a n - C a r p a t h i a n system are plotted. In general, EET increases with thermal age, that is parallel to cooling of the lithosphere. Differences may exist, however, between situations when the crust and the mantle lithosphere behave in a coupled and a

decoupled manner. The presence of a weak lower crust can lead to the mechanical decoupling of the lithosphere, which may substantially change the response of the deforming plate. According to Burov & Diament (1995), EET follows the 3 0 0 - 4 0 0 ~ and 6 0 0 - 7 0 0 ~ isotherm in the case of a decoupled and coupled lithospheric configuration, respectively. In spite of significant variations in EET as a function of thermal age, shown in Figure 7, these estimates do not exceed the depth of the 600 ~ isotherm as predicted for the model of a cooling lithosphere (Cloetingh & Burov 1996). The range of calculated EET values reflects the distinct mechanical habitat and behaviour of different domains in the P a n n o n i a n - C a r p a t h i a n system, which is mainly controlled by the memory of the deforming lithosphere. The considerable differences in the tectonic and thermal history of these domains from Cretaceous-Palaeogene times (Alpine orogeny) and through the Neogene (extension in the Pannonian domain) resulted in a wide

MODELLING THE PANNONIAN-CARPATHIANSSYSTEM

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Fig. 6. Effective elastic thickness (EET, in km) of the lithosphere in and around the Pannonian Basin predicted from rheological calculations (Lankreijer 1998). A, B and C indicate the location of strength envelopes shown in Figure 5a. Regional strength profile A-A' is shown in Figure 5b. BM, Bohemian Massif; MP, Moesian Platform; PB, Pannonian Basin; EC, SC, WC, eastern, southern and western Carpathians, respectively.

Fig. 7. Thermal age of the lithosphere v. calculated effective elastic thickness (EET) for selected sites in the Pannonian-Carpathian system (Lankreijer 1998). DDR-EET and Flex-EET represent EET values based on depth-dependent rheology estimates and flexural models, respectively. Upper limits of DDR-EET boxes represent EET calculated for wet rheology; lower limits represent dry rheology. Isotherms (in ~ are calculated for a cooling half-space model taking radiogenic heat production into account (Cloetingh & Burov 1996).

spectrum of lithospheric strength within the system. This in turn strongly influences its present-day behaviour, leading to a complex pattern of continuing tectonic activity.

Deformation of the Pannonian-Carpathian system The present-day deformation pattern and related topography development in the Pannonian-Carpathian system is characterized by pronounced spatial and temporal variation of the stress and strain fields (Fig. 8). Horv~ith & Cloetingh (1996) established the

importance of Late Pliocene to Quaternary compression in the Pannonian Basin, explaining its anomalous uplift and subsidence as well as intraplate seismicity. Those workers established a novel conceptual model for structural reactivation in back-arc basins within orogens through the case study of the PannonianCarpathian system. At present, the basin has reached an advanced stage of its evolution with respect to other Mediterranean back-arc basins, and its structural inversion has been taking place during the last few millions of years. Basin inversion is related to temporal changes in the regional stress field, from one of tension that controlled Miocene basin formation, extension and subsidence, to one of Pliocene-Quaternary compression resulting in basin deformation, contraction and flexure of the lithosphere associated with differential vertical motions. The spatial distribution of uplifting and subsiding areas inside the Pannonian Basin can be therefore interpreted as a result of an increasing level of intraplate compressional stresses. The general inversion of the basin can be explained in terms of stress-induced lithospheric deflection, that is increasing intraplate stresses cause large-scale bending of the lithosphere at various scales. This includes basin-scale positive reactivation of Miocene normal faults, and large-scale folding of the system leading to differential uplift and subsidence at the anticlinal and synclinal segments of the Pannonian crust and lithosphere. Model calculations are in good agreement with the overall topography of the system (Fig. 8). Several flat-lying, low-altitude areas (e.g. the Great Hungarian Plain, and the Sava and Drava troughs) have been continuously subsiding since the onset of basin formation in the Early Miocene, and are filled with a sequence of Quaternary alluvial deposits of 300-1000 m in thickness. In contrast, the periphery of the basin system, as well as the Transdanubian Range, the Transylvanian Basin and the neighbouring Carpathian orogen, have been uplifting and considerably eroded since Miocene-Pliocene times (see Figs 3 and 9). Quantitative subsidence analysis by van Balen et al. (1999) confirmed that compressive stresses can cause accelerated subsidence in the central parts of the Pannonian Basin. Similar studies at the rim of the Pannonian Basin, including the Styrian Basin (Sachsenhofer et al. 1997), the Vienna and East Slovak basins (Lankreijer et al. 1995) and the Transylvanian Basin (Ciulavu et al. 2002), have demonstrated a major uplift of several hundred metres starting from Mio-Pliocene times. The mode and degree of coupling of the Carpathians with their foreland controls the Pliocene-Quaternary deformation

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Fig. 8. Topographyof the PannonianCarpathian system and present-day maximumhorizontal stress (Srt.....) trajectories (after Bada et al. 2001). Plus and minus signs mark areas of Quaternary uplift and subsidence,respectively.BA, Balkanides;BM, BohemianMassif; D, Drava trough; D, Dinarides;EA, Eastern Alps; EC, eastern Carpathians;F, Focsani depression; MP, Moesian Platform;PB, Pannonian Basin; S, Sava trough; SC, southern Carpathians; TB, Transylvanian Basin; TR, TransdanubianRange; WC, western Carpathians.

patterns in the hinterland, and particularly interesting, in the Transylvanian Basin (Ciulavu et al. 2002). The western-eastern Carpathians is coupled with the strong lid of the European lithosphere, showing comparable coeval deformations in both the upper and the lower plate (Matenco & Bertotti 2000; Krzywiec 2001). The Bend Zone is decoupled in terms of deformation from its Moesian lower plate. It records contraction in the upper plate (e.g. Hippolite et al. 1999) and large-scale extensional collapse in the Moesian unit, as demonstrated by structural studies and crustal earthquake focal mechanisms (e.g. Bala et al. 2003). These findings have been recently corroborated by results from fission-track studies in the Romanian Carpathians, demonstrating up to 5 km of erosion with a systematic migration from the northwestern and southwestern part of the Romanian Carpathians, uplifted and eroding since 12 Ma, towards the bend area where uplift and erosion was initiated around 4 Ma ago (Sanders et al. 1999; Fig. 9). This region coincides with the actively deforming Vrancea zone in Romania, where considerable seismic activity is observed both at crustal levels and in the mantle. These findings can now be connected with seismic tomography results highlighting the development of hot mantle material under the Pannonian Basin and the presence of late-stage detachment of the lithosphere in the Vrancea area (Wortel & Spakman 2000). At the same time, these rapid differential motions at the rim of the Pannonian Basin and the Carpathians have important implications for the sediment supply to the depocentres as well as for the hydrocarbon habitat (Dicea 1996; Taft et al. 1997; Horv~ith & Tari 1999). In summary, the results of forward basin modelling show that an increase in the level of compressive tectonic stress during Pliocene-Quaternary times can explain the first-order features of the observed pattern of accelerated subsidence in the centre of the Pannonian Basin and uplift of the basin flanks in the peripheral areas. Therefore both observations (see Horvfith e t al. 2006) and modelling results lead to the conclusion that compressive stresses can cause considerable differential vertical motions across the back-arc basin-orogen system in the Pannonian-Carpathian area.

The sources of compression in the context of basin inversion were investigated by means of finite-element modelling (Bada et al. 1998, 2001). The results suggest that the state of recent stress in the Pannonian-Carpathian system (Fig. 8), particularly in its western part, is controlled by the interplay of plate boundary and intra-plate forces. The former includes the counter-clockwise rotation and northward indentation of the Adriatic microplate against the Alpine-Dinaric orogen, whereas intra-plate buoyancy forces are associated with the elevated topography and related crustal thickness variation of the Alpine-Carpathian-Dinaric belt. Model predictions indicate that uplifted regions surrounding the basin system can exert compression on the thinned Pannonian lithosphere of about 4 0 60 MPa, which is comparable with values calculated for far-field tectonic stresses (Bada et al. 2001). The combined analysis of stress sources of tectonic and gravitational origin has provided estimates for the magnitude of maximum horizontal compression. Significant compressional stresses (up to 100 MPa) are concentrated in the elastic core of the lithosphere, consistent with the continuing structural inversion of the Pannonian Basin. Such high-level stresses are close to the integrated strength of the system, which may lead to its whole-lithosphere failure in the form of large-scale folding and related differential vertical motions, and intense brittle faulting in the form of seismoactive faulting.

Discussion and conclusions Previous studies of the Pannonian Basin-Carpathian system underlined the importance of crustal stretching and lithospheric flexure as controls on the main features of basin formation and present-day lithospheric structure. These studies have revealed major discrepancies between subcrustal and crustal stretching factors inferred from subsidence analyses and seismic reflection data, and predictions from uniform stretching models. The same is true for the late-stage (Pliocene-Quaternary) anomalous acceleration in subsidence and uplift in the Pannonian Basin and

MODELLING THE PANNONIAN- CARPATHIANS SYSTEM

Fig. 9. Contours of amount of erosion (km) inferred from fission-track analyses of samples from the Romanian Carpathians and Apuseni Mts (Sanders et al. 1999) and thickness of foredeep sediments in the foreland. Numbers in elliptical boxes indicate tinting of onset of erosion (Ma); square boxes indicate the main moment of subsidence. The pronounced lateral differences in uplift ages along the arc should be noted; in contrast, the main subsidence period is coeval along the studied area.

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surrounding Carpathian arc. This pattern of vertical motions deviates from scenarios of decaying thermal subsidence and downflexure of the lithosphere inferred from stretching models and foreland flexural models, respectively. Lower crustal flow and small-scale lithosphere-asthenosphere convection appears to be a viable mechanism to explain the occurrence of a second, Late Miocene rift phase in the Pannonian Basin coeval with the climax of compression in the surrounding Carpathian arc. The Pannonian B a s i n - C a r p a t h i a n system displays pronounced lateral variation in lithospheric structure and sedimentary basin configuration. The tectonic evolution of the system is characterized by a polyphase history with Early to Late Miocene extension in the Pannonian Basin and simultaneous contraction in the Carpathian arc. Pre-existing structures and pre-rift rheology of the lithosphere play a key role in basin formation and subsequent reactivation, explaining anomalous features in subsidence characteristics and inferred thinning factors. The P a n n o n i a n - C a r p a t h i a n system has been subjected to repeated tectonic reactivation and, thus, a high level of strain concentration resulting in an overall weakening of the deforming lithosphere. Cretaceous to Palaeogene convergence in the Alpine belt led to nappe stacking and crustal accretion, and related thickening of the orogen. The overthickened and unstable Pannonian lithosphere, with a laterally unconstrained plate boundary towards the Carpathian embayment in the east, underwent gravitational collapse and extrusion and, eventually, the Miocene formation of the Pannonian Basin by the rifting and extension of former contractional domains. Models for the Pannonian Basin therefore underscored the importance of the pre-rift lithospheric structure for the mode of extension in the basin system. Lithospheric memory is important on different spatial scales. On a lithospheric scale, the presence of a thickened continental Alpine crust was manifested in shallow necking levels during subsequent extension (van Balen e t al. 1999). On an upper crustal scale, reactivation of Alpine thrust-faults in basin extension has been widely documented at

Fig. 10. Example of crustal-scale folding in south Transdanubia, Hungary, based on the interpretation and sequence stratigraphic analysis of regional seismic reflection profiles (Sacchi & Horvfith 2002). During the Quaternary the whole area has been uplifting. The basement units of the Bakony and Mecsek Mts, however, represent regions of higher amounts of uplift and, thus, large-scale anticlines, whereas the folded Miocene strata in the Somogy can be regarded as a syncline between the S.S., sensu stricto; S.L., sensu lato.

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Fig. 11. Example of lithosphere-scale folding in the Pannonian Basin system documented by stratigraphic, geophysical and geodetic data. (a) Crustal profile A-B transects the whole Pannonian Basin from the Eastern Alps to the eastern Carpathians, connecting uplifting and subsiding areas. 1, Areas of Quaternary subsidence; 2, areas without Quaternary subsidence or uplift; 3, areas of Quaternary uplift; 4, NeogeneQuaternary volcanic rocks; 5, internal basement units; 6, flysch nappes; 7, AlpineCarpathian foredeep; 8, Tauern-window; 9, direction of present-day maximumhorizontal stress (SHmax).Adr, Adria; AM, Apuseni Mts; D, Drava trough; Da Danube Basin; DIN, Dinarides; EA, Eastern Alps; EC, eastern Carpathians; EEP, East European Platform; GHP, Great Hungarian Plain; S, Sava trough; SA, Southern Alps; SC, southern Carpathians; TB, Transylvanian Basin; TR, Transdanubian Range. (b) Stratigraphic cross-section along transect AD. (e) Basement deflection profile based on forward modelling of stress-induced subsidence-uplift pattern. Compiled after Jo6 (1992), Horvfith (1993), Horvfith & Cloetingh (1996) and Gerner et al. (1999).

both seismic and outcrop scale in tectonic windows at the margins of the Pannonian Basin (e.g. Horvfith 1993). Recently, independent constraints on the bulk rheology of the pre-rift Alpine lithosphere have been obtained from rheological modelling (e.g. Sachsenhofer et al. 1997; Lankreijer 1998; Willingshofer et al. 1999). A marked contrast in recent rheology between the Pannonian Basin area, the surrounding Carpathian orogen and the foreland lithosphere is directly related to crustal configuration and thermal properties. In general, the Bohemian Massif, the East European Platform and, to a lesser extent, the Moesian Platform form strong and rigid buttresses in the foreland areas, whereas the intra-Carpathian regions are characterized by a much weaker rheology. Lateral and temporal variations in lithospheric strength exert a major control on the tectonic evolution of the study area. The extremely low rigidity of the Pannonian Basin and its tectonic setting locked in the interior of the Carpathians arc has made it a sensitive recorder of stress changes induced by various tectonic processes. Stress studies have underlined the close

relationship between the timing and nature of the stress changes in the extensional basins and the timing inferred from kinematic studies of the surrounding thrust belts. This suggests a mechanical coupling between the orogen and the back-arc domain. Quantitative modelling of recent and palaeostress fields constrained by a vast database of stress indicators has demonstrated that the stress fields in the area are primarily controlled by near-field intraplate and far-field plate boundary forces (Bada et al. 1998) in the context of continuing continental collision in the Alpine-Dinaric belt. The strain and stress pattern are also strongly influenced by the effects of topography and crustal thickness variation (Bada et al. 2001), and by asthenospheric ascent (Huismans et al. 2001). The recent tectonic activity of the region is largely controlled by the counter-clockwise rotation of the Adriatic microplate relative to Europe around an Euler pole in the Alps. As a result of the indentation of this crustal block against the Southern Alpine-Dinaric fold-and-thrust belt, intense oblique shortening (dextral transpression) is taking place in these orogens, as shown

MODELLING THE PANNONIAN-CARPATHIANS SYSTEM

by the general seismicity and crustal deformation pattern. The present-day kinematics of the Pannonian Basin shows that the area is pushed from the SSW by the Adriatic microplate and compressional stresses propagate well into the interior of the basin system. As a result, strike-slip to reverse faulting is observed inside the Pannonian Basin. Furthermore, the nearly complete absence of normal faulting throughout the study area suggests that extension in the Pannonian Basin has been terminated and its positive structural inversion is in progress. The intra-plate seismicity and neotectonic deformation indicate that the Pannonian Basin is affected by strong tectonic coupling with the surrounding parts of the African-European collision system. It appears that the polyphase evolution of the Pannonian Carpathian system has resulted in strong lateral variation in thermo-mechanical properties in the area, with weak lithosphere in the Pannonian Basin making it prone to late-stage basin reactivation. A particular feature of the Pannonian Basin is the build-up of a compressional stress regime resulting in a gradual inversion of the basin system. Tectonic inversion is manifested in a relatively high level of seismicity, abundant fault reactivation, late-stage folding and erosion documented by high-resolution reflection seismic data collected in Hungarian rivers and lakes (e.g. T6th & Horvfith 1998; Sacchi et al. 1999). Crustal deformation in the form of folding can take place at several scales from the inversion of a half-graben (e.g. in the Sava folds in Slovenia and SW Hungary; see Horvfith et al. 2006) to crustal-scale buckling. The area of south Transdanubia in Hungary is a key example for upper crustal folding (Fig. 10). Sacchi & Horvfith (2002) found that although the area has been uplifted as a whole during the Quaternary, differential vertical motions indicate continuing flexure, with maximum uplift taking place in the Bakony and Mecsek Mrs. The wavelength and amplitude of crustal folding can be estimated at about 1 0 0 - 1 5 0 k m and 500-1000 m, respectively (Fig. 10). On the other hand, large-scale bending of the Pannonian lithosphere, manifested in the Quaternary subsidence and uplift history (Horvfith & Cloetingh 1996), is also characteristic. As such, the Pannonian Basin has been interpreted in terms of irregular lithosphere folding (Cloetingh et al. 1999; Fig. 11) with a spectrum of wavelengths from several hundreds of kilometres (lithosphere-scale folding) through a few tens of kilometres (crustal-scale folding) to a few kilometres (basin-scale inversion). Folding and related structural development of the Pannonian lithosphere and related differential vertical movements have several important consequences for the hydrocarbon habitat (e.g. sealing of fault systems, reservoir integrity, maturation history and hydrodynamic regime, overpressure zones) and environmental issues (including landscape and slope stability). Recently, much attention has been paid to the spatial and temporal variations in thrusting along the Carpathian arc and its relationship to migrating depocentres, foreland basin geometry and lateral variations in fiexural rigidity (Matenco et al. 1997b; Zoetemeijer et al. 1999). At the same time, it has become increasingly evident that the second rift phase of the Pannonian Basin occurred simultaneously with the climax of compression in the Carpathian arc (Huismans et al. 1999), suggesting a mechanical link in terms of lower crustal flow induced by the rifting process, directed towards the Carpathian orogen. A general feature of all the flexural modelling studies carried out for the Carpathian system is the relatively low rigidity of the platform lithosphere downbending under the Carpathian belt, with effective elastic thickness (EET) estimates consistently below predictions inferred from rheological models of corresponding thermotectonic ages (Cloetingh & Burov 1996). These low EET values may be partly the result of stress-induced weakening associated with steep bending of the platform lithosphere under the arc, and partly the result of inherited weaknesses related to pre-orogenic Mesozoic extensional faults (Zoetemeijer et al. 1999), and the reactivation of inherited deep-seated

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weakness zones may have also played a key role (Matenco et al. 1997b; Lankreijer 1997). This paper reflects concepts developed during a decade-long scientific cooperation between the Vrije Universiteit Amsterdam, the Netherlands, the E6tv6s University, Budapest, Hungary and the University of Bucharest, Romania. Financial support, received from the Netherlands Research Centre for Integrated Solid Earth Science (ISES), the Hungarian National Science Fund (OTKA) projects T034928, D34598 and F043715, and the Romanian Ministry for Education and Research are acknowledged. The International Lithosphere Programme is thanked for providing the framework for scientific collaboration and discussions.

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MODELLING THE PANNONIAN-CARPATHIANS SYSTEM

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Ranges and basins in the Iberian Peninsula: their contribution to the present topography JAUME VERGt~S & MANEL FERNANDEZ

Group of Dynamics of the Lithosphere (GDL), Institute of Earth Sciences 'Jaume Almera', CSIC, 08028 Barcelona, Spain (e-mail: jverges @ ija. csic. es)

Abstract: The Iberian Peninsula,at the western end of the Alpine-Himalayan Belt, displaysa complexstructure with mountainranges of diverse structural trends and sedimentarybasins between them. The Iberian Peninsulaalso shows an elevated mean topography, the highest in Europe. In this short paper, we investigatethe Alpine evolution of the Iberian Peninsula since Mesozoic times, when Iberia was isolated as an independentplate. This occurred from Albian (formationof the northern plate boundary) to Oligocenetimes (end of the PyreneanOrogeny). Iberia was squeezedbetween Africa and Europe during Tertiary times and all previouslyestablishedMesozoic extensional basins were inverted, as were some of the Hercynian structures. The opening of the Valencia Trough, cutting the eastern margin of the Iberian Peninsula, began in Oligocene times. Concomitant crustal and lithospheric stretching during the Neogene along the eastern margin of Iberia produced limited uplifts, some of which are still active. The modern topography of the Iberian Peninsula was developedmainly as the result of three main tectonothermalmechanismssince late Palaeozoictimes: variationsin crustal densities, and possibly mantle depletion, inherited from the Hercynian Orogeny; crustal and lithospheric thickening during Tertiary compression; and upper mantle thinning during the Neogene-Quaternary.

The Iberian Peninsula constitutes the westernmost segment of the 12 000 km long Alpine-Himalayan Belt formed as a result of the Tertiary closure of the Tethys Ocean during the collision of India, Arabia and Africa with Asia and Europe (e.g. Dercourt et al. 1986). Rifting initiated during Triassic times (c. 250 Ma) and culminated in crustal break-up along the Atlantic margin (e.g. Ziegler, 1988, 1992). Continental break-up of the African Plate occurred during the Late Jurassic (c. 156 Ma). The Atlantic Ocean propagated northwards through the proto-Azores-Gibraltar plate boundary, producing the continental rupture of the Iberian Plate in Early Cretaceous times (c. 118 Ma; e.g. Srivastava et al. 1990). The mid-Cretaceous northern boundary of the Iberian Plate formed along the oceanic lithosphere of the Atlantic Ocean and its eastern continuation along the continental lithosphere of the Pyrenees (Fig. 1). Towards the end of Late Cretaceous (chron 33, 80 Ma) Africa shifted its motion northwards, initiating convergence with Eurasia with the consumption of the Tethys Ocean (e.g. Dercourt et al. 1986). At the westernmost termination of the AlpineHimalayan Belt, the Iberian Plate underwent a protracted deformation phase, resulting in orogenic belts along the plate boundaries (Bay of Biscay-Pyrenees and Azores-Gibraltar) and severe intraplate deformation. Several large Tertiary sedimentary basins developed on the Iberian Plate close to the bounding mountain chains (Friend & Dabrio 1996). Most of these basins began as flexural basins and continued as intermontane basins during the growth of the complex Iberian mountain system (Fig. 1). Iberia initially moved together with the African Plate, from latest Cretaceous to mid-Eocene times (chron 19, 42 Ma), deforming mainly the Bay of Biscay-Pyrenees plate boundary. From midEocene to the end of Oligocene times (chron 6c, 24 Ma), it moved independently and both plate boundaries were active. Subsequently, during the last 24 Ma, most of the deformation was accommodated along the complex and poorly understood plate boundary between Iberia and Africa, leading to the formation of the Betics, the Gibraltar Arc, and the Rif. The end of the Oligocene also coincided with extension along the proto-Western Mediterranean Sea, which affected the entire eastern margin of the Iberian Plate. This extension formed the oceanic lithosphere below the Liguro-Proven~al Basin north of the Paul Fallot Fault. To the south of this fault, thinned lithosphere below the Valencia Trough and Alboran Sea and oceanic lithosphere below the Algeria Basin formed (Fig. 1). The present contact

between Africa and Iberia changes progressively from pure fightlateral strike-slip along the Gloria Fault to a diffuse transpressive boundary from the Gorringe Bank to the Gulf of Cadiz region (e.g. Argus et al. 1989). The present structure of the Iberian Peninsula developed through the interplay of several geodynamic processes related to the Atlantic opening, the formation of two plate boundaries limiting the Iberian Plate, the north- south Africa-Europe convergence, and the concomitant rapid retreat and consumption of the oceanic Tethyan realms. Different geodynamic processes related to these large-scale tectonic events were to some extent coeval over particular morphotectonic regions. Both the diversity of geodynamic processes and their potential conjunction complicate the unravelling of the evolution of the Iberian Plate in general and, in particular, the southern plate boundary between Iberia and Africa (Betic Cordillera, Rif, Alboran Sea and Gulf of Cadiz tectonic units). This paper documents in brief the Alpine evolution of the onshore Iberian Peninsula mountain ranges and sedimentary basins, emphasizing the geodynamic processes that created positive topographic relief. This evolution took involved the following major tectonic events: (1) formation of extensional Mesozoic basins at the intersection of the proto-Atlantic and the Tethys oceans; (2) generation of Late Cretaceous-Tertiary fold-and-thrust belts and basins by the northwards motion of Africa; (3) formation of basins by Neogene extension along the eastern margin of the proto-Western Mediterranean. The paper concludes with the present topographic configuration of the Iberian Peninsula and its heritage from Hercynian times including the relatively recent lithosphefic thinning along the Mediterranean province of Spain. Two recently published books on the geology of Spain give a detailed description of the mountains and basins documented in this brief paper (Gibbons & Moreno 2002; Vera 2004). Andeweg (2002) has also illustrated the evolution of the palaeostress field in the Iberian Peninsula through the Cenozoic.

Mesozoic extensional basins

Preceding the opening of the central Atlantic during mid-Jurassic times (chron BSMA at c. 170 Ma), the Iberian Peninsula (Iberian Plate) was deformed by large-scale stretching that resulted in numerous extensional basins with different orientations. Rift systems developed along the western margin of the Iberian Plate

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 223-234. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Map of Western Europe with location of principal orogenic chains related to the Africa-Europe collision (based on Verg6s & Sfibat 1999). The Pyrenean Range corresponds to the westernmost limit of the about 12 000 km long Alpine-Himalayan Belt. CCR, Catalan Coastal Ranges; CM, Cantabrian Mountains; P.F.F., Paul Fallot Fault.

The separation between Europe and Africa for this period is about 240 km along this transect (e.g. Boccaletti et aL 1977; Olivet 1996; see position B for Africa in Fig. 2). The proposed position of Africa provides very little room for the restored Betic domain, thus creating a significant space problem, which has already been recognized (e.g. Andrieux et al. 1971; Mauffret et aL 1989; Frizon de Lamotte et al. 5995; Lonergan & White 5997; Spakman & Worte12000). Andrieux et al. (1971) proposed a model, still used with modifications by a number of workers, in which the Alboran Block was displaced towards the west by lateral extrusion during the north- south convergence of Africa and Iberia. Crustal and lithospheric thinning was the common process that formed the extensional basins. The thinned regions constituted weaker zones at the end of the Cretaceous just before the onset of Tertiary compression. Most of these extensional basins were tectonically inverted, preserving their original basin orientation. An extensive distribution of Triassic evaporites controls the geometry of the thrust system in both previous tectonic basins and structural highs, as in the case of the southern end of the Iberian Chain along the Altomira thrust system (the western boundary of the Iberian Range; Fig. 3).

Alpine Orogeny: Tertiary compressive belts and sedimentary basins (offshore Galicia and Portugal; e.g. Malod & Mauffret 1990), within the continental plate (Iberian and Catalan rifts; e.g. Salas et al. 2005), and along the two plate boundaries (Pyrenean rift in the north and the rifted south Iberian margin; Fig. 2). The opening of the central Atlantic produced an eastwards motion of Africa along the former Azores-Gibraltar transform fault. The Alpine-Tethys Ocean opened along the eastern margin of the Iberian Plate (e.g. Stampfli et al. 2002). This opening led to the cessation of rifting processes occurring within the Iberian Plate at the end of the Late Jurassic. A renewed phase of extension took place before the opening of the North Atlantic during the early Aptian (chron M0 at about 118 Ma). Pyrenean rift events occurred before the onset of ocean formation in the Bay of Biscay during Aptian-Albian times. The northern Iberian Plate transform boundary propagated eastwards and the northern segment of the Alpine-Tethys formed along the southeastern margin of Western Europe (Stampfli et al. 2002). These processes ended by the end of Coniacian at about 86 Ma. The Iberian rift system formed a linked configuration of extensional basins with different orientations as observed in the NE corner of Spain at the eastern end of the east-west-trending Pyrenees, the NW-SE-trending Figueres-Montgrf branch (F-M in Fig. 2), and the NE-SW-directed Catalan Basin (Fig. 2). The Iberian Peninsula at the end of Cretaceous times (Fig. 2), before the onset of Africa-Europe convergence, shows extended regions within the plate as well as along its margins. It is interesting to note that, during this period, at the end of Mesozoic extension, there are about 125-150km of separation between the central and eastern sides of France and Spain, and at least 35 km of extension in the Iberian Basin (Salas & Casas 1993; Salas et al. 2001). The restoration of the Prebetic and Subbetic units shows that their former, common, southeastern boundary was at least 90 km to the SSE of its present position (restoration according to GarcfaHernfindez et al. 1980). If we add the Internal Betics to the reconstruction by unfolding them to a minimum of double their present width (shortening of 50%), the SE boundary of the Internal Betics restores to about 250 km to the SE of its present position. Adding a counter-clockwise rotation of about 25 ~ to fit reasonably well with the palaeomagnetic rotations observed in the Betics (Platt et al. 2003) before the end of the late Tortonian (Krijgsman & Garc6s 2004), then the SE border of the Internal Betics restores to about 300 km to the SE (Fig. 2). This restoration is approximately in agreement with the proposed restoration by Platt et al. (2003).

In Late Cretaceous times, the northern motion of Africa against Europe significantly deformed the Iberian Plate along the previously extended Mesozoic basins. These basins, of diverse orientations and sizes, developed along both the plate boundaries and within the interior of the plate. Along the northern Iberian Plate margin, the Pyrenees represent a continental collisional orogeny with limited northwards subduction, whereas a more complex region deformed in the Southern Iberian Plate, including the Betics and Rif, the Alboran Sea, and the Gulf of Cadiz. The western Iberian margin along the Portuguese coast represents a slightly inverted margin, especially in its southern segment (e.g. Alves et al. 2003; Zitellini et al. 2004). During the roughly north-south Alpine convergence the interior of the Iberian Peninsula deformed while mostly preserving the original trends of the previously extended basins: N E - S W in the Catalan Coastal Ranges, N W - S E in the Iberian Range, N E - S W in the Central System, and north-south in the Altomira Range (Mufioz Martfn & De Vicente 1998; Fig. 3). All these fold-and-thrust systems are connected and at the intersections of any two of them there is always a linking zone in which the two different trends coexist as well as intermediary trends. The linkages between all the compressive fold-and-thrust belts indicate the synchronicity, at least partially, of several of these deformational events (Fig. 3). These connections between the different thrust systems that shape the present Iberian Peninsula have been described since the early 5980s in several transects crossing the Iberian Peninsula. The linkage of different thrust systems and the partial synchronicity of the deformation are agreed upon by most workers (e.g. Guimerfi 1984; Banks & Warburton 5995; Anad6n & Roca 1996; Casas S ainz & Faccenna 2005). The remarkable repetition of N E - S W trending chains in the central part of the Iberian Peninsula was interpreted as being produced by lithospheric folding during the Neogene (Cloetingh et al. 2002), in contrast to crustal and lithospheric thickening. In the following sections, we describe the principal compressive mountain systems of the Iberian Peninsula and their associated sedimentary basins starting in the Pyrenees (see regional transect of Roca et al. 2004) and ending in the Betics (see regional transect of Frizon de Lamotte et al. 2004). Between these, all the smaller mountain ranges developed in an interior tectonic setting, within the Iberian Plate. Casas Sainz & Faccenna (2001) have published an overview of these compressive mountain chains.

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Fig. 2. Reconstructed map at the end of Late Cretaceous times to show the distribution of principal sedimentary basins in and around the Iberian Peninsula. Ca, Cameros Basin; Ma, Maestrat Basin; Col, Columbretes Basin; F-M, Figueres-Montgrf Basin; NPFZ, North Pyrenean Fault Zone. Position and extent of Columbretes Basin after Roca (1996). Sardinia is shown in its restored position prior to Neogene opening of the Gulf of Lyons (Olivet 1996). The Balearic Islands are shown in their restored position prior to Neogene opening of the Valencia Trough (Vergrs & S~bat, 1999). Position A of Africa corresponds to the present position and position B to the restored position (e.g. Boccaletti et al. 1977; Olivet 1996). Pyrenees

The partial subduction to the north of the Iberian lithosphere underneath the European one along the Iberia-Europe plate boundary shaped the large-scale Pyrenean fold-and-thrust belt (Choukroune & Team, 1989; Roure et al. 1989; Mufioz 1992; Beaumont et al. 2000). The Pyrenean orogen is asymmetrical and double-sided, with the most significant thrust system developed towards the south, on top of the subducted zone as in most orogenic belts (e.g. Capote et al. 2002; Verg6s et al. 2002; Fig. 3). Although Africa compressed the entire Iberian Plate during its northwards shift, the Pyrenees recorded the initial stages of generalized compression around Santonian times (e.g. Puigdef~bregas & Souquet 1986), about 35 million years after the end of major rifting events along the Pyrenean branch of the North Atlantic in Albian times. The geometry at different scales of the northern boundary of the Iberian Plate clearly controlled the structural evolution of

the compressive belt. The northwards subduction was initiated along the extremely thinned lithosphere of the North Pyrenean Fault Zone (Mufioz 1992). The northern and southern thrust systems tectonically inverted the earlier Mesozoic basins together with their extensional fault systems. The location and extent of the early Mesozoic (Late Triassic) evaporites also exerted an important influence on the geometry and extent of the fold-and-thrust belts in the Iberian Peninsula. Most of the important Pyrenean shortening processes lasted for about 40 million years and were partitioned along several thrusts describing an overall foreland-directed propagation of deformation. Maximum shortening occurred across the Central Pyrenees and decreased towards the west (e.g. Vergrs et al. 2002). The northwards subduction of Iberia below Europe continued towards the western Pyrenees (Teixell 1998) and the Cantabrian Mountains (Fern~indez-Viejo et al. 2000; Pedreira et al. 2003). However, the degree of shortening decreased towards the west (e.g. Teixell 1998) and the age of initial deformation was younger in the same direction (e.g. Vergrs et al. 2002).

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Fig. 3. Tectonic map of the Iberian Peninsulabased on Rodrfguez-Fern~ndez (2004) and the detailed tectonic map of the Pyrenees of Vergrs et al. (1995). Tertiary mountain ranges: Pyrenees, Cantabrian Mountains (CM), Catalan Coastal Ranges (CCR), Iberian Range, Central System (CS), and Betics. Foreland basins: Ebro, Duero, Tajo, and Guadalquivir (GB). Iberian Massif: South Portuguese Zone, Ossa Morena Zone, Central Iberian Zone, AsturianLeonese Zone (ALZ), and CantabrianZone (CZ). AP, As Pontes Basin, located in NW Spain; CR, Ciudad Rodrigo Basin, located in the SW side of the Duero Basin. A, Altomira fold system, forming the western margin of the IberianRange.

The South Pyrenean flexural foreland basin was underfilled and marine from 55 to 37 Ma and then became overfilled and continental, starting with the deposition of mid-late Eocene Cardona evaporites until the end of shortening during the Oligocene (Puigdeffibregas & Souquet 1986; Puigdeffibregas et al. 1992; Vergrs et al. 1995, 1998). At m i d - l a t e Eocene times (c. 37 Ma), uplift of the western Pyrenees triggered the end of the foreland basin stage and originated an intermontane basin bounded by the Pyrenees, the Catalan Coastal Ranges and the Iberian Range (e.g. Burbank et al. 1992; Garcfa-Castellanos et al. 2003). A long period of lacustrine deposition that lasted through the Oligocene and most of the Miocene characterized this endorheic period (Riba et al. 1983). An internal fluvial network delivered sediments to the Ebro Basin, which was characterized by a large central lake (e.g. Anad6n et al. 1979; Arenas & Pardo 1999). The end of deformation occurred during late Oligocene times (c. 24.7 Ma; Meigs et al. 1996) although major basement uplift, based on fission-track cooling ages, ended at about 30 Ma (Fitzgerald et al. 1999). The late Oligocene-early Miocene age of younger compression in the western Pyrenees was synchronous with extensional processes affecting the eastern Pyrenees related to the formation of the Western Mediterranean basins. During late Miocene times, the endorheic Ebro fluvial system opened towards the Mediterranean Sea (e.g. Coney et al. 1996; Garcfa-Castellanos et al. 2003).

Roca & Guimerfi 1992). This thick-skinned style of tectonics affected the entire crust (Sfibat et al. 1997; Roca et al. 2004) producing a frontal monocline with relatively high topography (L6pez-Blanco et al. 2001). The oblique position of these basins with respect to the direction of compression also produced a limited sinistral strike-slip component (e.g. Anad6n et al. 1985; Guimerfi & Alvaro 1990).

Iberian R a n g e

The Iberian Range, with a N W - S E trend (nearly orthogonal to the Catalan coastal system), shows a complex structure involving cover and basement units (Alvaro et al. 1979; Casas Sainz & Faccenna, 2001; Fig. 3). The thrust system along the chain shows double vergence corresponding to the inversion of previous Mesozoic extensional faults. The NW termination of the range plunges beneath the Duero Basin and continues at depth below the frontal thrust of the Cantabrian Mountains. The northern tectonic displacement of the Iberian Range in its northwestern terminus is about 30kin (Casas Sainz 1993; Guimerfi et al. 1995). The Sierra de Altomira on the southwestern side of the Iberian Range flanks the Tajo Basin to the east (Fig. 3). This almost north-south-trending fold-and-thrust system is detached above Triassic evaporites (Mufioz Marffn & De Vicente 1998).

Catalan Coastal R a n g e s Central S y s t e m

The Catalan Coastal Ranges have a N E - S W trend (AlpineTethys Ocean trend) and display the effects of multiple tectonic events, including Eocene compression and Oligocene-Miocene extension, the latter related to the opening of the Valencia Trough (e.g. Anad6n et al. 1985; Fig. 3). The Catalan Coastal Ranges represent the inversion of earlier extensional structures, with the inversion affecting both cover and basement units (e.g.

The Central System, with a N E - S W structural trend, comprises an uplifted block, like a large pop-up structure (Vegas et al. 1990), with relatively high topography and associated with crustal thickening (Mezcua et al. 1996; Fig. 3). This block bounds two large sedimentary basins in the central part of the Iberian Peninsula: the Duero Basin to the NW and the Tajo Basin to the SE. The

BASINS & RANGES, IBERIAN PENINSULA Central System thrusts the Duero and Tajo basins to the NW and SE, respectively (e.g. Querol 1989; De Vicente et al. 1996). The Duero Basin is filled by a maximum of 2.5 km of Oligocene and Miocene continental deposits, mostly from the Cantabrian Mountains (e.g. Alonso Gavilfin et al. 2004). The age of the younger sediments in the basin is late Miocene at about 9.6 Ma (Krijgsman et al. 1996). The Ciudad Rodrigo Basin, at the SW termination of the Duero Basin, opened to the Atlantic at the Oligocene-Miocene boundary (Santisteban et al. 1996; its location is shown in Fig. 3). Subsequently, after 9.6 Ma, the Duero River captured the closed Duero Basin and opened it to the Atlantic basin. The Tajo Basin filled with 2 - 3 km of continental deposits ranging in age from the latest Oligocene to the latest Miocene (e.g. Alonso Zarza et al. 2004). Other compressive areas

Towards the NW corner of the Iberian Peninsula, small transpressional basins (such as the As Pontes basin) were filled with alluvial to lacustrine deposits during late Oligocene and earliest Miocene times (e.g. Cabrera et al. 1996; Fig. 3). The western boundary of the Duero Basin also formed during this period, closing it. Betics

The Betic Cordillera, trending generally ENE-WSW, corresponds to part of the former northern Africa-Iberia plate boundary, developed on top of Iberian crust and cropping out now in the southern part of the Iberian Peninsula (Figs 1 and 3). Although some aspects of the evolution of the Betic Cordillera are well known, there is still no agreement about the mechanisms that created it (see discussion of models by Calvert et al. 2000). The Betic Cordillera is divided into the Internal Betics, comprising metamorphic basement, the External Betics, consisting of cover rocks, and the Guadalquivir Foreland Basin (e.g. Azafion et al. 2002). The Alboran Sea, with a complex tectonic history, formed as a result of Neogene extension. Neogene to Quaternary depositional sequences fill curved, elongated and deep basins (e.g. Comas et al. 1999). The Internal Betics comprise a tectonic pile of three different tectonic units separated by thrusts (Nevado-Filfibride at the base, Alpujfirride in the middle, and Malfiguide at the top). Each unit displays a different degree of Alpine metamorphism, decreasing from the bottom to the top. The Nevado-Filfibride unit was affected by H P - L T metamorphism indicating that the rocks were located at depths of 50-70 km (the metamorphic evolution of the Internal Betics has been described by Comas et al. (1992)). The metamorphic units crop out as large antiforms that exhibit east-west trends in contrast to the E N E - W S W regional direction of the External Betics. Towards the western end of the Internal Betics, relatively large massifs of peridotites have been incorporated into the Alpujfirride thrust system (e.g. Ronda Peridotites; Tubfa et aL 1997). The External Betics constitute a system of thrust sheets carrying different Mesozoic SE Iberian passive margin palaeogeographical units towards the foreland. The Subbetic Zone is located to the SE and the Prebetic Zone to the NW (Garcfa-Hernfindez et al. 1980). The stratigraphy of the External Betic units includes Triassic to Miocene successions. The Prebetic Zone is mostly composed of shallow-water deposits whereas the Subbetic Zone units are deeper and pelagic. The Campo de Gibraltar Unit consists of terrigenous deposits forming an Oligocene and early Miocene accretionary prism developed above the Subbetic Zone towards the WSW (e.g. Crespo-Blanc & Campos, 2001; Bonardi et al. 2003) that was actively deformed until late Tortonian times (e.g. Grficia et al. 2003). The external part of the Prebetic Zone is composed mainly of Triassic evaporites, unconformably overlain by an incomplete and discontinuous succession of Mesozoic to Neogene deposits. The most external part of the Prebetic system forms an

227

intricate system of tectonic thrust imbricates and chaotic units, which possibly correspond to an imbricate thrust system emplaced in a marine depocentre filled with numerous olistoliths and olistostromes (e.g. Azafion et al. 2002). Contacts between units in the External Betics are principally foreland-directed thrusts involving different cover palaeogeographical domains mostly detached above Triassic evaporites. The cover-basement contact is a major hinterland-directed thrust in the east and centre of the Betics (Banks & Warburton 1991), which changes to a foreland-directed one to the west, where it over-thrusts the Subbetic Unit as well as the Campo de Gibraltar Unit. This hinterland-directed thrust was a response to tectonic wedging produced by the antiformal stack of basement units flattening on top of the Triassic detachment level (Banks & Warburton 1991). As for the rest of the mountain ranges of the Iberian Peninsula, the Triassic evaporites constitute an excellent detachment level between basement and cover rocks. Several relatively small intermontane basins filled with Neogene deposits are located along the contact between the Internal Betics and the External Betics (e.g. Iribarren et al. 2003). The Neogene to Quaternary Guadalquivir Basin, to the NW and WSW of the External Betics, corresponds to a foreland basin in front of the Betics thrust system (e.g. Berfistegui et al. 1998; Garcfa-Castellanos et al. 2002). Hercynian rocks of the Iberian Massif constitute the NW boundary of the basin.

Neogene formation of the Western Mediterranean basins Beginning in mid-Oligocene times the opening of the Valencia Trough created an extensional fault system paralleling most of the eastern coast of northeastern Spain (e.g. Roca et al. 2004; Fig. 3). This system cut obliquely across the Early Tertiary compressive Catalan Coastal Ranges and cut the SE termination of the Iberian Range almost perpendicularly, forming an extensional arrangement of basins parallel to the Mediterranean coast (e.g. Roca et al. 1999). Concomitant uplift of segments of the Catalan Coastal Ranges as well as of the SE margin of the Ebro Basin initiated the development of the present landscape configuration (e.g. Morgan & Fernfindez 1992; Lewis et al. 2000; GasparEscribano et al. 2004). Fission-track studies indicate that more than 1.5 km of uplift occurred, responsible for the significant dissection of this margin (Juez-Larr6 & Andriessen 2002). To the SE of the Iberian Plate, an early Miocene large-scale system of normal faults connected to the Alboran Sea extensional system cuts the rear flank of the Internal Betics antiform. Most of these large normal faults are subparallel to the Internal Betics antiformal thrusts, reactivating some of them (e.g. Platt & Vissers 1989; Garcfa-Duefias et al. 1992; Crespo-Blanc et al. 1994; Comas et al. 1999). The Alboran Basin is filled by up to 8 km of early MioceneQuaternary sedimentary sequences (e.g. Comas et al. 1999).

Present-day topographic configuration of the Iberian Peninsula The present-day mean elevation of the Iberian Peninsula, slightly over 600 m, is almost certainly the highest in Western Europe. Smith (1996) pointed this out but did not propose a good solution for what is sustaining it. However, the integration of the tectonic structure of the Iberian Peninsula with its topography (Fig. 4), Bouguer anomalies (Fig. 5), and existing 2D and 3D lithospheric models can explain most of this high topography. The greatest negative Bouguer anomalies of the Iberian Peninsula show a very good match with its principal Tertiary mountain chains such as the Pyrenees, the Iberian Range, the Central System and the Betics (e.g. Casas Sainz & Faccenna 2001; Cloetingh et al. 2002). Available geophysical modelling

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indicates that these negative anomalies, corresponding to compressive systems, are primarily explained by crustal thickening, which in most cases is combined with lithospheric thickening such as has been inferred for the Pyrenees (Zeyen & Fern~mdez 1994) and the Central Betics (Tome et al. 2000). Cloetingh et al. (2002) proposed an alternative mechanism, lithospheric folding, as the main cause of alternating mountain ranges and basins of the central part of Iberia, but combined crustal and lithospheric modelling, as for the Pyrenees by Zeyen & Fernhndez (1994), does not support this interpretation.

To the west and SW of the Iberian Range there is a very large segment of the Iberian Peninsula with an elevated topography between 600 and 1000 m (Fig. 4). This region corresponds to the Central Iberian Zone of the Variscan Iberian Massif and includes the Tajo and Duero basins, and also shows a significant negative Bouguer anomaly (Fig. 5), suggesting that the crust of this region is either thickened or less dense than the surrounding crust, or a combination of both (Fernandez et al. 1998). The former interpretation can be applied to the NE domain (Duero and Tajo basins) where thick Tertiary sedimentary successions

Fig. 4. Combined tectonic and topographic map of the Iberian Peninsula.

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229

Fig. 5. Combinedtectonic and Bouguer anomaly map of the Iberian Peninsula (from Mezcua et al. 1996).

are present. However, to account for the regional distribution of the negative Bouguer anomaly and elevated topography the hypothesis of a less dense crust for the whole of the Central Iberian Zone is more realistic. Finally, in the southwestern Iberian Peninsula, the Ossa Morena and South Portuguese zones are dominated by a Bouguer anomaly maximum (Fig. 5), coinciding with an average topographic height

of about 200 m. Field studies demonstrate an increase in crustal density for these two Hercynian domains. A 2D lithospheric model, which integrates elevation, gravity, geoid and heat-flow data, indicates that the present lithospheric structure in these two domains, with relatively high crustal density, must be underlain by a thinned lithosphere or by a depleted lithospheric mantle (FernSndez e t al. 2004). According to the model, the mass deficit

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Fig. 6. Lithosphericthicknessfor the Valencia Trough (Ayala et al. 2003) and Alboran Sea (Tome et al. 2000) superimposed on the tectonic map of the Iberian Peninsula.The volcaniccentres along the Mediterranean coast coincide with the main entrances of the 70 km thick lithosphere line (note that all these fields contain alkaline volcanic rocks). The Campo de Calatrava volcanicprovince crops out in the centre of Iberia. (Abbreviations as in Fig. 2; v.p., volcanic province).

related to mantle thinning or depletion implies a lithospheric buoyancy-driven uplift of about 120-150 m to fit the observed elevation and geoid data (Fernandez et al. 2004). On the Mediterranean margin of Iberia, crustal and lithospheric mantle thinning show dissimilar patterns. Thin lithospheric mantle underlies onshore Iberia at three main localities (marked by the 70 km thick lithosphere contour; Fig. 6): in NE Spain, on the southern flank of the Valencia Trough, and in SE Spain. Positive Bouguer anomalies, crustal and lithospheric thinning, high topography, and asthenospheric volcanism characterize these three regions. In NE Iberia, several studies show that thinning was more intense in the lithospheric mantle than in the crust, producing additional uplift as well as basic volcanism (e.g. Cabal & Fernhndez 1995; Lewis et aL 2000; Ayala et al. 2003; Fig. 7). Radiometric ages for volcanic rocks of the NE volcanic province indicate that there was a migration of volcanism to the SW and west from 14 to 0.011 Ma (e.g. Saula et al. 1994; Lewis et aL 2000; Marti 2004). This indicates that lithospheric thinning and concomitant uplift started during late mid-Miocene times and is still active at present. At the southern end of the Valencia Trough, the SE volcanic province, in Levante (Fig. 6), shows ages ranging from 8 to 1 Ma (Ancochea & Huertas 2004). Towards the SE of Iberia, 3D crustal and lithospheric modelling shows that this area is also supported dynamically (Tome et al. 2000). The SE volcanic province in Almer~a and Murcia shows only a few alkaline volcanic edifices (L6pez-Ruiz et aL 2004; Fig. 6). The lithospheric geometry and the existence of onshore alkaline volcanic provinces along the coast as well as in the interior of Iberia (Campo de Calatrava volcanic province; Fig. 6) has been interpreted as being indicative of a long-lived deep process related to the opening of the Atlantic Ocean (Oyarzun et al. 1997). According to Oyarzun et al. (1997), the thinning along the Western Mediterranean region corresponds to a long,

sublithospheric channelling starting in the Cape Verde and ending in the northern North Sea (Ziegler 1990). Thus, the present-day topography of the Iberian Peninsula was mostly acquired by means of three main tectonothermal mechanisms effective since the Late Paleozoic: variations of crustal densities and possibly mantle depletion inherited from the Hercynian Orogeny, crustal and lithospheric thickening during Tertiary compression, and upper mantle thinning during the Neogene-Quaternary. However, the present landscape of the Iberian Peninsula has also been sculpted by the opening of numerous endorheic basins. The most spectacular of these openings occurred in the Ebro Basin during the early late Miocene (Coney et al. 1996; Lewis et al. 2000; Garcfa-Castellanos et al. 2003). The best preserved basin is the Duero Basin, in which fiver incision is still relatively minor.

Summary Between a widespread Triassic phase of extension and the opening of the central Atlantic in Late Jurassic times, Iberia was affected by extensional processes that created numerous sedimentary basins with different orientations, including the Pyrenees, Catalan and Iberian basins, and the Betic basin on the southern margin of Iberia. The Alpine-Tethys Ocean opened at this time, forming the eastern margin of the Iberian Plate. An additional extensional phase, forming the western margin of Iberia, took place before the opening of the North Atlantic during the early Aptian. This transform boundary propagated eastwards from the Bay of Biscay and constituted the northern plate boundary of the Iberian Plate, resulting in its isolation. These processes ended by the end of the Coniacian. Basins formed during these extensional phases were inverted during Late Cretaceous-Tertiary compression, which started in

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231

Fig. 7. Map of the NE corner of the Iberian Peninsula to show two areas affected by Neogene and Quaternary uplift (from Lewis et al. 2000). The SW area is limited to the footwall of the Oligoceneearly Miocene normal faults that parallel the coastline. The NE area is wider and affects part of the previous area and both the footwall and hanging wall of the late Miocene-Quaternary NW-SE-trening system of normal faults. The volcanism migrated westwards from the Empord~t Basin to the la Selva Basin and finally to La Garrotxa, where it is as young as 0.011 Ma.

Santonian times. The extent and orientation of the basins controlled the size, width and trends of the compressive belts, which normally show a double vergence. In almost all the Iberian fold-and-thrust belts, Triassic evaporites provided an excellent decoupling level between the basement and the cover units. The Iberian mountain chains produced lithospheric flexural bending and thus foreland basins, especially in front of the Pyrenees and the Betics: the Ebro and Aquitaine basins flank the Pyrenees and the Guadalquivir Basin lies to the north of the Betics. The Duero and Tajo basins developed to the south of the Cantabrian Mountains and to the south of the Central System, respectively. The Ebro, Duero and Tajo basins became intermontane basins at different stages of their evolution by the closure of their Atlantic connections. The basins were filled by radial fluvial systems feeding central lakes. Starting in the late Oligocene, the opening of the Valencia Trough initiated the development of a system of normal faults and linked basins aligned with the present Mediterranean coastline of Iberia. The deformation history of the Iberian Peninsula produced a significant mean elevation that is the highest in Europe. Tertiary compressional processes contributed strongly to an increase of mean elevation, but inherited Hercynian lithospheric structures as well as late Cenozoic upper mantle thinning related to the opening of the Western Mediterranean also contributed to the high average elevation of the Iberian Peninsula. Some of the thermo-mechanical

processes affecting the eastern margin of the Iberian Peninsula are still active at present. This is a contribution of the Group of Dynamics of the Lithosphere (GDL), Department of Geophysics and Tectonics, Institute of Earth Sciences 'Jaume Almera', CSIC. Partial support for this paper was provided by MCYT projects REN2001-3868-C03-02/MAR, REN2002-11230-E-MAR, NATO Grant EST-CLG978922. 2001 SGR 00339, and project 2001 SGR 00339 Grup d'Estmctura i Processos Litosf~rics, funded by the Comissionat per Universitats i Recerca of the Generalitat de Catalunya, Grups de Recerca Consolidats, II Pla de Recerca de Catalunya. We finally thank an anonymous reviewer for constructive remarks and suggestions.

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Contrasting modes of ophiolite emplacement in the Eastern Mediterranean region A L A S T A I R H. F. R O B E R T S O N Grant Institute of Geology and Geophysics, School of GeoSciences, University of Edinburgh, Edinburgh, EH9 3JW, UK (e-mail: [email protected])

Abstract: The Eastern Mediterranean region is characterized by one of the largest concentrations of ophiolites anywhere in the world. Many of these ophiolites are fragmentary or highly deformed, such that their initial mode of tectonic emplacement cannot easily be inferred from the local field relations. The emplacement of many of these ophiolites can usefully be compared with the intact Oman ophiolite, one of the largest and best-studied ophiolites in the world. The Oman ophiolite is commonly believed to have been created in Late Cretaceous time (c. 95 Ma) above an oceanward-dipping, intra-oceanic subduction zone. This was followed by collision of the subduction zone with the downflexed Arabian passive margin, facilitating the emplacement of the ophiolite onto the continental margin. A less likely alternative is that the Oman ophiolite formed at a mid-ocean ridge that then collapsed, initiating the emplacement of the ophiolite. An Oman-type model is applicable to many of the Mid-Jurassic and the Late Cretaceous ophiolites of the Eastern Mediterranean region that were thrust over former passive continental margins. These ophiolites are again mainly of suprasubductionzone type. Such ophiolites include many of the Jurassic ophiolites of Greece, Albania and former Yugoslavia, and also the Late Cretaceous ophiolites of Turkey and northern Syria. These ophiolites were emplaced from both more northerly and southerly Neotethyan ocean basins. In contrast, the opposing (northerly) margins of these oceanic basins experienced a history of subduction-accretion, marginal arc volcanism and back-arc basin formation ('Cordilleran-type' ophiolites). Ophiolites that were emplaced associated with active margin settings range from large accreted thrust sheets to small slices within accretionary prisms and back-arc basins. Examples include the Late Cretaceous ophiolites that are related both to the northern margin of the southern Neotethys and to the northern margin of the northern Neotethys in Turkey. Not all ophiolites were emplaced in response to large-scale horizontal tectonic transport (e.g. Jurassic Guevgueli ophiolite, northern Greece), and several ophiolites experienced dominantly strike-slip or transpression (e.g. the Late Cretaceous Antalya ophiolites, SW Turkey). In general, the mode of ophiolite emplacement, especially the direction of emplacement relative to the orientation of the adjacent continental margin was influenced by the regional palaeogeographical setting.

The objective of this paper is to discuss and interpret the tectonic processes related to the emplacement of ophiolites exposed in the Eastern Mediterranean region (Fig. 1). Many but by no means all of these ophiolites were emplaced by processes that were similar to that for the emplacement of the Semail ophiolite in the Oman Mountains, the most complete, widely exposed and bestdocumented Tethyan ophiolite. The main focus here is on information gained from the tectonostratigraphy of the ophiolites in various settings and of different ages throughout the Eastern Mediterranean region. A glance at a tectonic map of the world shows a greater density of ophiolites per unit area in the Eastern Mediterranean region than anywhere else, except possibly Alaska or Indonesia (see Hoeck et al. 2002). With the advent of plate tectonics theory, ophiolites came to be seen as allochthonous slices of oceanic crust and mantle emplaced onto continental margins, as exemplified by the Semail ophiolite, Oman (Glennie et al. 1973, 1974) and the Bay of Islands ophiolite, Newfoundland (Williams & Smythe 1973; Jenner et al. 1991). Alternative tectonic scenarios were initially envisaged for the emplacement of various ophiolites (Dewey & Bird 1970; Dewey 1976; Casey & Dewey 1984), including several from the Eastern Mediterranean region (Woodcock & Robertson 1985). A complete understanding of ophiolite emplacement, however, has remained elusive, as few modern analogues exist. Onland and marine evidence of oceanic crust emplacement (potential future ophiolites) in different modern settings remains sparse (e.g. Dilek et al. 2000). Also, some ophiolites have reached their final locations following re-thrusting associated with regional continental collision, thus concealing the processes of initial emplacement from the ocean onto a continental margin. Different ophiolites were emplaced by several different tectonic processes and no one tectonic model is applicable to all ophiolites. Recent reviews of ophiolite emplacement processes worldwide (Dilek & Newcomb 2003; Dilek & Robinson 2003) re-emphasize the wide range of opinions concerning ophiolite genesis and emplacement, and highlight the need for them to be interpreted in the context of their regional tectonic settings. An attempt will be made here to identify several contrasting modes of ophiolite emplacement that mainly apply to Mid-

Jurassic and Late Cretaceous ophiolites of the Eastern Mediterranean region (Mukasa & Ludden 1987; Liati et al. 2004). The first setting relates to the emplacement of ophiolites onto former passive continental margins (Oman-type model). A second setting relates to ophiolite emplacement along active continental margins ('Cordilleran-type' model) by processes including subduction-accretion or the collapse of back-arc marginal basins. Because the ophiolites of the Eastern Mediterranean are commonly incomplete or dismembered, it is useful to draw comparisons with the largest and best-documented Tethyan ophiolite, in Oman (Fig. 1). The Troodos ophiolite, Cyprus, is undeformed by emplacement-related processes (Gass 1990) but cannot be used as a model for ophiolite emplacement, in view of the absence of any exposed base to the ophiolite, a complex local tectonic setting and its insular position. Several tectonic models for the genesis of the Oman ophiolite exist but there is a general consensus concerning its mode of emplacement over the Arabian passive margin. Should the Omantype emplacement be regarded as a 'one-off" or is it more widely applicable; for example, to the Eastern Mediterranean region? A summary of alternative models for the genesis of the Oman ophiolite and its emplacement is given first, followed by more detailed comparisons with Eastern Mediterranean ophiolites. Ophiolites that have formed in settings that differ from that of Oman are discussed later in the paper. The O m a n trench-collision model A widely accepted model for the tectonic setting of ophiolite genesis and emplacement in Oman (Fig. 2) is that presented by Lippard et al. (1986), with minor modifications. The future Oman ophiolite was generated by spreading above an oceanwarddipping (i.e. NE-dipping) intra-oceanic subduction zone in early Late Cretaceous time (c. 95 Ma; Pearce et al. 1981, 1984; Fig. 3). The exact location of subduction zone initiation is still unclear, whether near or some distance from the former spreading axis. A location removed from the ridge axis is probable in view of the argument that very young oceanic lithosphere may not be

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 235-261. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

235

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subductable (e. g. Cloos 1993; Shervais 2001). The well-developed metamorphic sole of the Oman ophiolite was accreted to the base of the oceanic slab beneath the depleted mantle tectonites of the over riding ophiolite. The high-temperature metamorphic sole was probably created some distance from the former spreading axis, if it is again assumed that young oceanic lithosphere is not normally subductable. This high-temperature metamorphism took place either during steady-state subduction (Searle & TERTIARY

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E. MEDITERRANEAN OPHIOLITE EMPLACEMENT

shallow-level thrust (i.e. related to a collapsed spreading axis). Continental margin-derived sediments were thrust deep beneath the oceanic mantle wedge, giving rise to granulite-facies metamorphism, implying the presence of a subduction zone (Searle & Cox 1999). In addition, granitic rocks were locally formed by melting of sediments at depth, followed by intrusion up to the level of the ophiolitic Moho in the northern Oman Mountains (Cox et al. 1999). A model of subduction trench-passive margin collision in Oman can also explain the occurrence of blueschists, south of the Semail Gap (e.g. E1-Shazly et al. 1990). The passive margin in this area underwent attempted subduction (oceanwards) until underthrusting of a thickening wedge halted subduction. This was rapidly followed by buoyancy-driven exhumation (Searle & Cox 1999, and literature cited therein). A more complex scenario is required in the MOR model, involving genesis of a separate intra-oceanic subduction zone and the transfer of the emplacing ophiolite from one convergence zone (i.e. collapsed ridge) to another (subduction zone) prior to final emplacement over the Arabian margin.

constructed as subduction proceeded (Fig. 2). The subduction zone collided with the Arabian passive margin several million years later, as indicated by combined palaeontological and radiometric dating (Tilton et al. 1981; Hacker & Gnos 1997). As the oceanic slab impinged on the passive margin, this then flexurally subsided beneath the advancing thrust load to form a foredeep (Robertson 1987a). This subsidence accommodated the subaqueous emplacement of the ophiolite over the Arabian continental margin as one, or several, vast thrust sheets (Glennie et al. 1973, 1974; Gealey 1977; Rabu et al. 1990; Robertson & Searle 1990; Le Metour et al. 1995; Searle & Cox 1999). A promontory of the Arabian margin in the south (to the south of the Semail Gap) was subducted and metamorphosed to high-pressure facies, then rapidly exhumed (see Searle et al. 2003, for literature review). In contrast, continental margin units further north (e.g. slope sediments of the Sumeini Group) remained unmetamorphosed. There is an alternative tectonic model for the genesis and emplacement of the Oman ophiolite. In this model the ophiolite formed by spreading at a mid-oceanic ridge (MOR), rather than above a subduction zone (Coleman 1971, 1981; Hopson & Pallister 1980; Coleman & Hopson 1981; Boudier & Nicolas 1988; Nicolas 1989), in a setting perhaps akin to the small East Pacific Juan de Fuca spreading axis. In the mid-ocean ridge interpretation the spreading ridge later collapsed to form an intra-oceanic thrust along which the ophiolite and its underlying metamorphic sole were emplaced towards the adjacent continental margin (Nicolas & Le Pichon 1980; Boudier et al. 1988). The view favoured here is that the Oman ophiolite formed above a subduction zone. However, in both the suprasubduction zone (SSZ) model and the MOR model young, hot oceanic crust and upper mantle were emplaced over a passive continental margin. This general mode of ophiolite emplacement contrasts with some other settings where ophiolites were emplaced along active continental margins ('Cordilleran ophiolites'), as discussed later in the paper. Several features of the emplacement of the Oman ophiolite appear to support the subduction-related hypothesis. There is strong evidence that the emplacement-related thrusting was deep-seated (i.e. involving a subduction zone), rather than a !

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Palaeogeography of the Eastern Mediterranean region A brief summary is given below to facilitate comparisons of the settings of ophiolite emplacement (Fig. 4). There is still considerable discussion about the palaeogeography of the Eastern Mediterranean region, especially during the MesozoicEarly Tertiary time of important ophiolite genesis and emplacement. In outline, the African and European continents were separated by generally older and younger oceans known, respectively, as Palaeotethys (Late Palaeozoic-Early Mesozoic) and Neotethys (Mesozoic-Early Tertiary). Palaeotethyan ophiolites are mentioned only briefly here as they are fragmentary, commonly metamorphosed and still poorly understood. However, pre-Late Jurassic 'Palaeotethyan' ophiolites were emplaced in the Pontides, northern Turkey (e.g. Ydmaz et al. 1997; Fig. 4). According to some workers (~engrr et al. 1980, 1984; Okay et al. 1991; Grnctio~lu et aI. 2000) this emplacement was the result of southward closure of a wide Palaeotethys to the north.

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238

A.H.F. ROBERTSON

However, others have inferred emplacement related to northward subduction of a Palaeotethyan ocean located to the south (rooted towards the present Izmir-Ankara-Erzincan suture zone: Robertson & Dixon 1984; Dercourt et al. 1986, 1992, 2000). Recently, the Palaeotethys of the above workers was reinterpreted as one, or several, marginal basins related to back-arc spreading behind a wide 'Palaeotethys' located further south (Stampfii et al. 2001). However, recent work has discounted the existence of a southerly Palaeotethys as in this tectonic model (Robertson 2006). Also, the related Palaeotethyan units of northern Turkey differ strongly from the nature of modern back-arc basins, as documented from the modern SW Pacific regions (see Robertson 1994, for comparisons). It is instead assumed that Palaeotethys was rooted in the north near the Izmir-Ankara-Erzincan suture zone. This persisted in some form or other as an oceanic area from Late Palaeozoic to Early Tertiary time, evolving into the Northern Neotethys, now represented by the Izmir-Ankara-Erzincan zone in Turkey (Robertson & Dixon 1984; Robertson et al. 2004b). The inferred northward subduction of a Palaeotethys rooted to the south of the Eurasian margin in time led to the construction of a regionally extensive accretionary wedge (Tekeli 1981; Pickett & Robertson 1996, 2004), known as the Karakaya Complex. The type area of the Karakaya Complex is located in the western Pontides (Okay et al. 1991, 1996), with counterparts in the central Pontides and eastern Pontides (Yflmaz et al. 1997; Okay & Sahintiirk 1997; Ustarmer & Robertson 1997, 1999). These inferred accretionary units are dominated by mrlange, including accreted volcanic seamounts (Ntilifer Unit) and inferred continental margin units (e.g. ~al Unit), as exposed in the NW Pontides. Most of the accreted material was derived from the upper levels of the oceanic crust and only exceptionally includes plutonic ophiolitic rocks. Small slices of ophiolitic rocks, including MOR-type basalts covered by radiolarites slices (Pickett & Robertson 1996), are found within the Karakaya Complex, presumably representing accreted fragments of Palaeotethyan ocean floor. There is also the much larger Denizgrren ophiolite, which can be correlated with the Chios ophiolite on the offshore Greek island (Pickett & Robertson 1996). The field relations are compatible with the Denizg6ren ophiolite being a Palaeotethyan ophiolite (Okay et al. 1991; Pickett & Robertson 1996). However, Ar-Ar dating of the Denizgrren metamorphic sole has suggested an Early Cretaceous age (Okay et al. 1996), pointing to an origin more related to the emplacement of the Jurassic ophiolites of the Inner Hellenide Ophiolite Belt in northern Greece (see below). Although the Karakaya Complex is dated as Triassic in NW Turkey (Okay et al. 1991, 1996), it is possible that the Eurasian margin was long lived and that similar but older Karakayatype units may exist elsewhere in the Pontides. Palaeotethyan ophiolites are, indeed, also present in other areas adjacent to the Eurasian margin, including Iran (Strcklin 1974) and the Caucasus (Adamia et al. 1981, 1995), but require additional studies to determine the processes of genesis and emplacement in detail. Northward subduction of Palaeotethys, possibly obliquely (Robertson & Dixon 1984), is considered to have given rise to rifting behind a continental margin arc (~angalda~ unit) in the central Pontides (Ustarmer & Robertson 1997), dated as being active during Late Palaeozoic-Early Tertiary time (Kozur et al. 2000). Back-arc rifting culminated in opening of a small marginal basin floored by Eurasian-derived terrigenous sediments (Ustarmer & Robertson 1994; Robertson et al. 2004; Fig. 5). This basin later closed by southward subduction, accreting small ophiolitic slices, known as the Kfire ophiolite, within mainly terfigenous turbidites, followed by covering by Early Jurassic neritic carbonates. In contrast to the Palaeotethyan ophiolites of pre-Late Jurassic age, ophiolites of Jurassic age are well exposed in the Balkan region and large ophiolites of Late Cretaceous age dominate the region further east, including Turkey, Cyprus, Syria, Iran and Oman. The emplacement of these ophiolites is the main subject of this paper. According to some researchers most, or all, of

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Southerly Mediterranean ophiolites The 'southerly' ophiolites (Figs 6 and 7) include the Jurassic Balkan ophiolites, of which the best known are the Pindos, Vourinos and Othris ophiolites in Greece, the Jurassic Mirdita ophiolites in Albania and the Dinaride ophiolites of former Yugoslavia. The Jurassic Balkan ophiolites are discussed first, below. These ophiolites are all associated with the 'Pindos suture' shown in Figure 4. For the purposes of discussion here these 'southerly' ophiolites also include the Late Cretaceous ophiolites that are exposed along the Arabian margin (Croissant peri-Arabe of Ricou 1971) from Oman, through Iran, to Syria and the Late Cretaceous ophiolites of Cyprus and Antalya (southern Turkey). These ophiolites are associated with the SE Turkish, Antalya and Mamonia sutures shown in Figure 4. All of these ophiolites are believed to have been emplaced from a southern Neotethys ocean that was located outboard of the Arabian continent (Robertson et al. 1991). The' southerly' ophiolites also include a belt of Late Cretaceous ophiolites further north in Turkey, associated with the IzmirAnkara-Erzincan suture (e.g. the Lycian and Bey~ehir-Hoyran ophiolites), which are restored to an origin within a more northerly Neotethyan oceanic basin. J u r a s s i c B a l k a n ophiolite c o m p a r e d with O m a n

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E. MEDITERRANEANOPHIOLITEEMPLACEMENT

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continuation into Albania in the Mirdita zone as the 'Western-type' and 'Eastern-type' ophiolites (Shallo et al. 1990; Robertson & Shallo 2000). This model was proposed for the Pindos ophiolites in NW Greece by Jones & Robertson (1991;

239

see also Jones et al. 1991; Robertson et al. 1991) and has been widely applied to other Tethyan ophiolites since then by many workers (with varying modifications). More recently, a similar model was applied, for example, to the Albanian ophiolites (e.g. Robertson & Shallo 2000; Dilek et al. 2005; Koller et al. 2006) and the Greek ophiolites (Clift & Dixon 1998; Smith & Rassios 2003). Many of the Jurassic Balkan ophiolites (e.g. Pindos, Vourinos, Evia, Eastern-type Albanian ophiolite) show geochemical evidence of formation in a subduction-related setting (e.g. Capedri et al. 1980; Noiret et al. 1981; Jones et al. 1991; Beccaluva et al. 1994; Robertson & Karamata 1994; Cliff & Dixon 1998; B~bien et al. 2000; Rassios & Smith 2000; Pami~ et al. 2002; Bortolotti et al. 2004; Koller et al. 2006; Rassios & Moores 2006). Remnants of MOR-type oceanic crust are also rarely preserved, including some of the ophiolitic units of the Othris area (Smith et al. 1975) and the Western-type Albanian ophiolite (Shallo et al. 1990). Many of these ophiolites are underlain by metamorphic soles and volcanic-sedimentary mrlange. Also, in several cases, the sedimentary cover is preserved, and includes deep-sea sediments and debris-flow deposits containing exotic rocks derived from continental margin units. These associated units, as well as the ophiolites themselves, provide important information about the processes of ophiolite emplacement. One of the most problematic aspects, assuming a SSZ origin of the Jurassic Balkan ophiolites is accepted, is the setting of the initiation of subduction to generate these ophiolites. Most researchers have assumed that this subduction was initiated at, or near, the spreading axis for the Oman ophiolite (Lippard et al. 1986; Searle & Cox 1999) and the Jurassic Balkan ophiolites (Jones & Robertson 1991; Robertson et al. 1991). Recently, Smith & Rassios (2003) have adopted an Oman-type model in which ophiolite emplacement was driven by the collision of a subduction trench with the passive margin (Pelagonian zone). However, they suggested that subduction was initiated, not within the ocean, but instead adjacent to the western margin of the ocean basin. Specifically, the subduction zone was located along the eastern side of the marginal Parnassus carbonate platform and within marginal oceanic crust to the north (and presumably to the south). Thus, after subduction began the preexisting ocean basin was, in effect, replaced by new SSZ oceanic crust, although marginal oceanic crust remained to the north and south of the Parnassus block. This obviates the apparent difficulty of initiating spreading within relatively young oceanic lithosphere. Unfortunately, the eastern margin of the Parnassus platform is a thrust contact with the Pelagonian zone, itself part of the eastern margin of the oceanic basin, and thus the model cannot be directly tested in the field in this area. If applicable, however, the above model could apply regionally throughout the Balkans (Pindos-Mirdita ophiolites) and possibly more widely. In this context several points can be made. (1) If spreading originated near the westerly passive margin, a record of this event might be observed, as margin sediments are well preserved in the Pindos-Olonos Nappes (e.g. in the Peloponnese). If subduction began near this margin, significant tectonic disruption could be expected, combined with formation of a bathymetric trench, which could become a sediment sink with a good chance of ultimate preservation. The Pindos-Olonos thrust sheets record proximal to distal successions that prograded over Triassic oceanic crust (or transitional crust). During Triassic-Jurassic time these successions record passive margin subsidence and a switch to pelagic radiolarian deposition, with the first major siliciclastic input not being until Cretaceous time, after genesis and emplacement of the Pindos-Mirdita ophiolites (Degnan & Robertson 1998, 2006). Similar constraints apply further NW throughout Albania and former Yugoslavia (e.g. Robertson & Shallo 2000). There is thus no stratigraphical evidence for the initiation of a westward-dipping subduction zone along the westerly margin of the ocean basin. (2) ff a subduction zone developed near, or along, this passive margin, a sedimentary record of this event could be expected

240

A.H.F. ROBERTSON

within the sedimentary cover of the ophiolitic extrusive rocks, where well preserved. This cover might be expected to resemble that of the distal passive margin, as preserved in the Pindos-Olonos thr-ust sheets (i.e. mixed muddy, siliceous and calcareous sediments). However, the cover of the ophiolific extrusive rocks, as best preserved in the Albanian ophiolites (Mirdita zone), consists of well-dated ribbon radiolarites without any relatively coarse marginderived sediments lying directly on the ophiolite (e.g. Prela et al. 2000; Danelian & Robertson 2001). (3) If nearly the entire width of the Pindos-Mirdita ocean was 'replaced' by SSZ lithosphere it is surprising that the western-type ophiolites in the extension of the Greek ophiolites into Albania are of MOR-type (e.g. Shallo et al. 1990). According to Dilek et al. (2005), radiometric dating of plagiogranites within both the western- and eastern-type ophiolites indicates that both are of similar age (c. 162-165 Ma) and thus that two contrasting magma types were involved in oceanic crust genesis within a relatively short period of time. This is explicable if a MOR-type setting evolved into an SSZ-type setting relatively rapidly following the initiation of a subducfion zone at, or near, a ridge crest. It is less easy to envisage why MOR-type lavas were erupted in an entirely subduction-controlled setting. (4) The model implies that the original Late Triassic-Early Jurassic spreading ridge was subducted beneath the passive margin, but no obvious thermal effect of this is seen along the emplaced passive margin (e.g. in the Parnassus unit). The available evidence, therefore, provides little obvious support for the hypothesis of subduction initiation at or near the western margin of the oceanic basin, although this option cannot be ruled out. It is interesting to apply the above model of subduction initiation along a passive margin more widely. If applied to the Oman region, the entire ocean floor of the Gulf of Makran and related accreted material within the Makran accretionary prism should be of similar SSZ crust. Applied to the southern Neotethys in Turkey, the model would have difficulty in explaining the presence of arc-type rocks associated with the northern margin, where, by comparison, subduction would have initiated. A similar difficulty applies to the northern Neotethys in Turkey, as again arc rocks are present in a northerly part of the basin (see discussion later in the paper). Applied to the modern oceans, the question arises as to why SSZ spreading apparently ceases, followed by construction of a volcanic arc (Bloomer et al. 1995), rather than the subducting slab continuing to roll-back indefinitely. The main objection to the hypothesis of subduction zone initiation at or near a spreading ridge is the theoretical difficulty of subducting young oceanic lithosphere. This problem may, however, be diminished for the Jurassic Balkan ophiolites, as the MOR-type western ophiolite in Albania is interpreted as a rifted ridge (marked by low magma production), which might be subductable, especially if convergence was initiated along a fracture zone. In general, it is possible that subduction was initiated in different places in different oceans, basically where the crust was weakest. On balance, the writer prefers a model of subduction initiation well within the ocean, at least for the Jurassic ophiolites of the Balkan region. This leaves part of the initial ocean basin intact until much later final closure of the ocean. However, subduction initiation near a continental margin may apply in some cases (e.g. Late Cretaceous, northern Neotethys; see below). For the present discussion, we are more concerned with the mode of emplacement of the ophiolite onto a continental margin itself, rather than the initial setting of subduction, which may remain hypothetical as so little evidence has survived following regional convergence events. One problem is how to emplace a vast slab of oceanic crust and mantle, apparently upwards against gravity onto a continental margin, as depicted by Glennie et al. (1973). In Oman it was found from palaeo-environmental analysis of Late Cretaceous synemplacement sediments (Muti Formation) that the abyssal plain-slope-margin underwent thrust loading and flexural collapse into deep water ahead of the advancing ophiolite

(Robertson 1987a,b). Thus, no major vertical uplift of the emplacing ophiolite was required. Similar flexural processes related to ophiolite emplacement can be inferred for many of the Jurassic Balkan ophiolites and the Late Cretaceous ophiolites further east, as discussed below. In Oman, the adjacent carbonate platform underwent flexural uplift and erosion, locally down to near the base of the Mesozoic carbonate platform succession, creating the regional WasiaArumu break (Glennie et al. 1973; Robertson 1987a; Figs 2 and 3 - l b ) . Alkaline volcanic rocks were locally erupted during this stage (Le Metour et al. 1995). The erosion surface was karstified and in places capped by glauconite- siderite ironstones (Robertson 1987c). The unconformity was attributed to the passage of a flexural bulge ahead of the advancing ophiolite. Rather than an idealized flexural bulge, however, the uplift was probably partially related to the reactivation (i.e. inversion) of older (Triassic) rift faults. After passage of the flexural bulge, the uplifted margin gradually submerged and was covered by deepening-upward neritic, to hemi-pelagic, sediments (Robertson 1987c). The platform margin later collapsed dramatically as a result of flexural loading of the advancing thrust pile. This collapse was associated with large-scale slumping and mass wasting of carbonate debris flows derived from the upper levels of the carbonate platform succession. A switch to non-calcareous muds and quartzose sandstone turbidites ensued. The inferred source of these sediments was rift-related, or basement, rocks exposed to erosion by uplift of the continental margin, rather than transport from the Arabian landmass, then far to the SW. Very similar uplift and collapse features are seen associated with the emplacement of the Jurassic Greek ophiolites (Fig. 7a). The underlying Pelagonian carbonate platform (Fig. 4), where exposed, shows evidence of block faulting prior to the arrival of ophiolite-derived clastic sediments. The carbonate platform subsided and was overlain by non-calcareous radiolarian sediments of Bathonian-Callovian age (e.g. in the Kalidrommon Mountains, central Greece; Danelian & Robertson 1995). This was followed by the arrival of ophiolite-derived sediment in Tithonian time (Thiebault et al. 1994). Evidence of the initial emplacement of oceanic units onto the continental margin is documented from successions within the Pelagonian carbonate platform. On the large island of Evia, eastern Greece (Figs 4 & 8), the Late Triassic to Mid-Jurassic carbonate platform succession culminates in an unconformity, interpreted as a subaqueous erosion surface, overlain by tuffaceous sediments and radiolarian sediments (Robertson 1991). Comparable tuffaceous sediments occur on the adjacent mainland, in the Othris Mountains (Price 1976). The local volcanism, as in Oman, is attributed to extensional processes accompanying flexural upwarping of the carbonate platform (Robertson 1991). Elsewhere on Evia (at Achladi beach), the typical Jurassic platform carbonates pass upwards into mudstones and radiolarites of Late Jurassic (Oxfordian) age. These sediments document the collapse of the carbonate platform in response to loading by the advancing allochthon (B aumgartner & Bernoulli 1976; Robertson 1991). The appearance of chromite grains preceded the arrival of ophiolitederived material generally. There are several possible alternative origins for the chromite grains. They were possibly derived from ultramafic rocks at the base of the overriding ophiolite, from slices of ultramafic rocks within an advancing accretionary wedge, or from submarine erosion of diapiric serpentinite formed within the rifted continent-ocean transition zone (e.g. as in the modern Iberia (Atlantic) rifted margin; e.g. Boillot et al. 1980). The last mentioned possibility was not previously considered. In this case, the ultramafic rock was exposed along the rifted margin during the original continental break-up and later reworked during tectonic deformation of the passive margin, prior to the arrival of exotic ophiolitic material. On Evia (e.g. at Achladi beach), the collapsed platform succession is overlain by debris-flow deposits and then by allochthonous oceanic-derived mrlange (Baumgartner & Bernoulli 1976;

E. MEDITERRANEANOPHIOLITE EMPLACEMENT

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•

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Robertson 1991; Scherreiks 2000). Further south, in the Argolis Peninsula, comparable emplacement-related debris flows contain abundant ophiolitic material, especially serpentinite (Baumgartner 1985). These debris flows are seen as the result of gravity TERTIARY

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Fig. 8. Summaryof the tectonostratigraphy of the Early-Mid-Jurassic Evia ophiolite, eastern Greece. Modified after Robertson (1991).

redeposition of material that was shed off advancing thrust sheets of continental margin sedimentary and volcanic rocks and oceanic crust. In Oman, continental margin-derived material forms a wellorganized thrust stack, including sedimentary units (Hawasina Complex; Glennie et aI. 1973; Searle & Stephens 1984; Cooper 1990; Bernoulli & Weissert 1997; Fig. 2) and largely volcanic units (Haybi Complex; Searle & Malpas 1980; Searle et al. 1980). The Hawasina Complex is interpreted to record an original transition from proximal rifted continental margin units to oceanic sediments. This complex is assumed to have been underlain by transitional-type or oceanic crust that has been entirely subducted (Bernoulli & Weissert 1987; Bgchennec et al. 1988; Robertson & Searle 1990). The Hawasina Complex is interpreted as an accretionary prism related to northeastward subduction of the distal edge of the rifted Arabian margin and adjacent, marginal Tethyan oceanic crust of Permo-Triassic age (Glennie et al. 1973; Lippard et al. 1986). The Haybi Complex is dominated by Permian and Triassic limestone exotic units (Oman Exotics; Glennie et al. 1973; Pillevuit et al. 1997) and Late Triassic alkaline volcanic rocks that variously formed both along the rifted Oman margin and as seamounts within the oceanic crust (Searle & Graham 1982; Lippard et al. 1986). The combined Hawasina and Haybi complexes were emplaced over the Oman margin (beneath the Oman ophiolite) when the subduction trench collided with the Arabian continental margin. Similarly, in Greece the Pelagonian carbonate platform was overthrust by rifted passive margin units, as seen in the Othris

242

A . H . F . ROBERTSON

21~

21~

'

urassic ophiolite, mainly peridotite, ultramafic. fic cumulates, DRAMALA OPHIOLITE - Cretaceous platform carbonates and related ~ts; blocks assoc, with Avdella Melange L. Cretaceous accretionary melanges, _A MELANGE - Cretaceous d e e p - sea sediments, Zone, Dio Dendra Group

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Km

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area (Smith et al. 1975, 1979). The well-organized thrust stack there comprises proximal (marginal), to more distally derived thrust sheets of sedimentary (e.g. turbidites; radiolarites), volcanic (e.g. Agrilia lavas) and ophiolitic lithologies (Sipetorrema lavas) (Smith et al. 1975). Elsewhere, coherent thrust sheets are commonly absent and only a large-scale 'mrlange' is exposed, dominated by broken formation (dismembered thrust sheets) and debris-flow deposits ('olistostromes' of Ferri~re et al. 1988), all exposed beneath the ophiolites (e.g. Evia, Atalanti, Pindos). An excellent example of a mrlange, mainly derived from continental margin units, is exposed in the Pindos Mountains (Figs 9 and 10). The Pindos ophiolites are underlain by a thick unit of mrlange (Avdella Mrlange), including large blocks of Triassic W i PL!OCENE u MIOCENE M U ionian

Fig. 9. Simplified geological map of to show the setting of the Jurassic Dramala (Pindos) ophiolite in the Pindos Mtns, NW Greece (from Robertson 2002).

neritic limestone and locally coherent thrust sheets of Mid-Late Triassic volcanic-sedimentary lithologies, together with debrisflow deposits (Jones & Robertson 1991; Fig. 10). Similar, but volumetrically subordinate, mrlange locally underlies the Albanian ophiolites (Robertson & Shallo 2000). Elsewhere, in northern Greece the Pelagonian carbonate platform was overthrust by ophiolite-related mrlange and ophiolite slices (Sharp & Robertson 2006) All of the above mrlange units can be explained by the accretion of a variety of oceanic to continental margin units of sedimentary and volcanic origin, as shown in Figure 11 (Jones & Robertson 1994). Pre-existing topographic highs located along the rifted margin (e.g. rift fault blocks and carbonate build-ups; Fig. 1 la) were preferentially accreted (Fig. 1 lb and c), culminating in collapse of the

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Fig. 10. Summary of the tectonostratigraphy of the Jurassic Dramala (Pindos) ophiolite, NW Greece. Modified after Jones & Robertson (1991).

Fig. 11. Reconstruction of the rift-drift-emplacement history of the Pindos ophiolites and related units, northern Greece. (a) Late Triassic rifting; (b) Early to Mid-Jurassic initial subduction and accretion (Avdella Mrlange); (c) Bathonian-Oxfordian; approach of the Pelagonian continental margin, leading eventually to emplacement of the Pindos ophiolite and underlying Avdella Mrlange. From Jones & Robertson (1994). WPB, within-plate basalt.

E. MEDITERRANEAN OPHIOLITE EMPLACEMENT

adjacent margin of the Pelagonian platform and overthrusting by the accretionary prism and the ophiolites (Fig. 1 lc). The Oman ophiolite is underlain by a metamorphic sole. The sole is typically up to tens of metres thick, with a lower, greenschist-facies thrust sheet and an upper, amphibolite-facies thrust sheet (Searle & Malpas 1980, 1982; Ghent & Stout 1981; Gnos & Peters 1993). In some areas the metamorphic sole is laterally intact (e.g. Sumeini area), whereas in others it is more dismembered (e.g. Hawasina window), or thickened by structural repetition (e.g. Dibba Zone, Northern Oman Mountains; Searle & Cox 1999). In the Dibba Zone, detailed petrological and mineralogical studies indicate that continentally derived material was entrained with hot peridotite to form granulite-facies rocks. The protoliths were potassium-rich terrigenous and calcareous rocks, which, unusually, are exposed directly beneath the hanging wall of the metamorphic sole. In addition, partial melting of these highgrade metamorphic rock produced peraluminous granitic melts that percolated upwards into the overlying ophiolite (Gnos & Nicolas 1996). Searle & Cox (1999) argued that these processes could only have taken place as a result of the attempted subduction of continental crust beneath a hot mantle wedge. This particular setting of granulite-facies rocks and associated melts was not so far documented in the Eastern Mediterranean region. Very similar metamorphic soles to those of the Oman are present beneath the Greek and Albanian Jurassic ophiolites (Spray et al. 1984; Dimo-Lahitte et al. 2001). These range from intact (as in Evia, Vourinos and the Eastern-type Albanian ophiolites), with well-developed amphibolite- and greenschist-facies units, to blocks of variable lithologies strewn through underlying mrlange, as in the Pindos Mountains. The dismembering of the metamorphic sole of the Pindos ophiolite is largely attributable to re-thrusting of the allochthon associated with continental collision in Early Tertiary time (Jones & Robertson 1991). The details of the petrogenesis of the metamorphic sole are outside the scope of this paper. However, it is notable that the protoliths of the metamorphic soles vary considerably based on chemical evidence; they vary from near-MORB (mid-ocean ridge basalt) composition in Oman (Searle & Malpas 1982), to mainly within-plate-type (alkaline) basalts in Baer-Bassit, northern Syria (A1-Riyami et al. 2002a,b) and Mersin, southern Turkey (Parlak et al. 1995), to subduction-influenced tholeiites, for example, in Evia (Simantov et al. 1990; Danelian & Robertson 2001). Also, the protoliths may include various intrusive as well as extrusive ophiolitic rocks. This evidence shows that oceanic crust from different tectonic settings was incorporated into the metamorphic sole in different areas. The nature of the overlying relatively intact ophiolite provides additional information concerning emplacement processes. The emplaced ophiolites range from a single relatively coherent thrust sheet, as in Oman (although buried thrust sheets may exist offshore), to composite units made up of several different thrust sheets originating in different oceanic settings. Specifically, the Pindos ophiolites include a thin (several hundred metres) lower dismembered thrust sheet of boninitic lavas (high-magnesian andesites) (Aspropotamous ophiolite), overlain by the much thicker main Pindos ophiolite (Dramala unit; Fig. 10). The boninites are seen as forming in a forearc-type setting and provide additional evidence for a subduction-related origin of the Pindos ophiolites as a whole (Jones & Robertson 1991; Jones et al. 1991). The Jurassic Balkan ophiolites also provide evidence for the processes of initial deformation related to tectonic emplacement. In the Oman ophiolite it was found that the preferred orientations of olivine and orthopyroxene crystals, located within hightemperature mylonites (at the base of the ultramafic tectonites) could be used to determine the emplacement direction (Boudier et al. 1985). Similarly, in the Vourinos ophiolite, the primary fabric within the depleted upper mantle tectonites is overprinted by evidence of high-temperature deformation. This can be related to initial tectonic displacement of the ophiolite while it

243

cooled through the brittle-ductile transition (Rassios et al. 1994; Rassios & Smith 2000). Similar observations were made in the Pindos, Vourinous and Othris ophiolites. This evidence incidentally indicates that these ophiolites were initially emplaced to the NE (in present co-ordinates), towards the Pelagonian continental margin. In a few cases the structures in the underlying metamorphic soles may also suggest emplacement directions (e.g. Zlatibor ophiolite; Dimitrijevic & Dimitrijevic 1979). Interpretation is, however, difficult as the early transport lineations that developed at high temperatures commonly lack vergence indicators, and the sole may be disrupted and tectonically rotated during or after ophiolite emplacement. During its emplacement, the Oman ophiolite was segmented into crustal blocks and has undergone clockwise rotation, as shown by palaeomagnetic studies (Perrin et al. 1994; Thomas et al. 1988). These rotations are attributable to oblique subduction, or to diachronous collision with the Arabian margin. Comparable, to more extreme, rotations have affected the Late Cretaceous Troodos, Hatay and Baer-Bassit ophiolites (see later discussion). The nature of the sedimentary cover of the ophiolites can provide additional clues concerning emplacement processes. In Oman, the Semail ophiolite is depositionally overlain by a thin (< 10 m thick), but well-dated cover of metalliferous umbers, radiolarites and pelagic carbonates (Fleet & Robertson 1980; Tippett et al. 1981). In comparison, most of the Jurassic Balkan ophiolites lack intact sedimentary covers (see Danelian & Robertson 2001, for a recent review). Exceptionally, both the Eastern- and the Western-type Albanian ophiolites are depositionally overlain by thin (<10 m) radiolarites of Mid-Jurassic age (Marcucci et al. 1994; Prela et al. 2000). In Oman, the pelagic sediments are overlain depositionally (e.g. in Wadi Jizi, north Oman Mountains) by hemipelagic muds and debris-flow deposits derived from the underlying lava units (Zabyat Formation). These debris flows were explained by submarine faulting and mass wasting, as the ophiolite began to be emplaced over the Arabian continental margin (Robertson & Woodcock 1983). In northern Oman (e.g. in the 'Alley' area, north of Wadi Jizi), a mrlange, known as the Batinah Mdlange, is additionally inferred to have been expelled upward through fault-bounded 'cracks' and then redeposited as debris flows on the upper surface of the ophiolite nearby (Robertson & Woodcock 1982a). These protrusions appear to have been lubricated by serpentinite, which forms part of the debris-flow deposits overlying the ophiolite. The protruded material includes continental margin sediments and volcanic rocks. The protrusion appears to have taken place during the emplacement of the Semail ophiolite over the continental margin. Comparable mrlange, within or above the ophiolites, is occasionally observed in the Eastern

NORTHERN

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Fig. 12. Simplified geological map of the Jurassic ophiolites of northern Albania (from Robertson 2002). The Eastern-type ophiolite is subduction influenced whereas the Western-type ophiolite exhibits a MORB-type composition.

244

A . H . F . ROBERTSON

Frontalw.as~ ENE WNW

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Fig. 13. Tectonic model for the emplacement of debris flows with exotic blocks above the deep-sea sedimentary cover of the Jurassic Albanian ophiolites. From Robertson & Shallo (2000).

Sheeted

1

dykes

Late Cretaceous ophiolites (eastern region)

Evidence of ophiolite emplacement, comparable with that in Oman, applies to many of the Late Cretaceous ophiolites forming the 'southerly ophiolite zone' bordering the northern margin of Arabia. This zone runs through Iran, eastern and southern Turkey, and northern Syria (Gass & Masson-Smith 1963; Ricou 1971; Robertson & Dixon 1984; Dercourt et aL 1986, 1992, 2002; Glennie 2000; Robertson 2000; Stampfli et al. 2001). Where chemically analysed, most of these Cretaceous ophiolites show SSZ chemical characteristics, metamorphic soles, underlying mrlanges and collapsed passive margin successions, similar to Oman and the Jurassic Balkan examples discussed in the previous section. Examples in southern Turkey include the Kocali ophiolite (Akta~, & Robertson 1984), the 'Yiiksekova' ophiolite (Yllmaz 1993), and the Hatay ophiolite (Delaloye & Wagner 1984; Dilek & Thy 1998); also the Baer-Bassit ophiolite in northern Syria (e.g. Parrot 1977; Delaune-May~re 1984; A1-Riyami & Robertson 2002; A1-Riyami et al. 2002b; Figs 14 and 15).

36~

Massive and layered gabbro, plagiogranite Metamorphic

Mediterranean region. The Albanian ophiolites and overlying radiolarites (Fig. 12) are widely overlain by debris-flow units, up to several hundred metres thick. These are dominated by clasts and blocks of sedimentary rocks derived from the underlying continental margin succession (Shallo 1990). This could also be explained by material being expelled through the overriding ophiolite when it was dissected into blocks during its emplacement onto the continental margin (Robertson & Shallo 2000; Fig. 13). Another possibility is that this mrlange in Albania relates to reworking of material that was initially emplaced beneath the ophiolite then exposed, perhaps in response to exhumation processes (i.e. 'backthrusting'). In Oman, the Semail ophiolite is known to have been subaerially exposed, associated with the localized accumulation of fluvial conglomerates, up to several hundred metres thick. The existence of serpentinite-derived conglomerates shows that ultramafic rocks were exposed to erosion soon after emplacement. There is, however, little evidence that the ophiolite was ever greatly topographically elevated or deeply eroded, as it was soon transgressed by shallow-marine carbonates, from Maastrichtian time onwards (Skelton et al. 1990). In contrast, in central Greece the ophiolites (e.g. Evia, Atalanti) were subaerially exposed for a long period after Late Jurassic-Early Cretaceous emplacement, associated with the development of Fe-Ni deposits of locally economic importance (e.g. in Evia; Fig. 8). These ophiolites were not finally transgressed until Late Cretaceous (Cenomanian) time related to rising eustatic sea level. During this long period (Early Cretaceous) deep erosion took place, removing all but the lower crustal and upper mantle units of these ophiolites (Robertson et al. 1991).

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Fig. 17. Summary of the tectonostratigraphy of the Late Cretaceous Mersin ophiolite, southern Turkey. Modified after Parlak et al. (1995).

245

In addition, Late Cretaceous ophiolites are also commonly exposed along the northern margin of the Tauride-Anatolide carbonate platform (an inferred microcontinent). These also show similar SSZ-type affinities and emplacement histories (Fig. 4). They include the Lycian ophiolite (de Graciansky 1972; Collins & Robertson 1997, 1998; ~elik & Delaloye 2003; see Fig. 16), the Bey~ehir ophiolite (Monod 1977; Ozgul 1977, 1984; Elitok 2001; Andrew & Robertson 2002) and the Pozantl-Karsantl ophiolite (Juteau 1980; Lwytyn & Casey 1995; Parlak et al. 2000). The Karsantl ophiolite is also known as the Alihoca ophiolite (Dilek et al. 1999). In addition, the Mersin ophiolite (Fig. 17), located along the southern margin of the Tauride carbonate platform at the longitude of Mersin (Fig. 4), was inferred to have been emplaced northwards from the southern Neotethys (Parlak et al. 1996). However, this ophiolite and its underlying mrlange was recently reassigned to an origin within the northern Neotethys, to the north of the Tauride carbonate platform. The main reason is the similarity in tectonostratigraphy with the Bey~ehir-Hoyran nappes exposed further NW, also to the north of the Tauride carbonate platform (Parlak & Robertson 2004). These more northerly ophiolites were all emplaced generally southwards onto the Tauride platform in latest Cretaceous time (~eng6r & Yllmaz 1981; Robertson & Dixon 1984; Dercourt et al. 1986, 1992, 2000; Dilek et al. 2000). Their emplacement was associated with the collapse of the Tauride-Anatolide carbonate platform and the emplacement of mrlange units, in which the metamorphic grade varies from low grade to high pressure-low temperature. In addition, there are also unmetamorphosed Late Cretaceous ophiolitic rocks, mainly derived from high crustal-level units (extrusive rocks, sheeted dykes and gabbros) that overlie the metamorphic Kir~ehir Massif (Floyd et al. 2000). These ophiolites are believed to have been emplaced southwards from the Ankara-Erzincan suture zone further north (Yahnlz et al. 1996). Most interpretations of these ophiolites favour northward underthrusting and subduction of Neotethyan oceanic crust, followed by onset of intra-oceanic SSZ spreading, comparable with the inferred situation in Oman. Recent studies have not supported earlier suggestions of southward Late Cretaceous subduction in southern Turkey (Dilek & Moores 1990). The processes of genesis and emplacement of the above Late Cretaceous ophiolites, rooted in both the southern and northern Neotethyan basins, are therefore similar to those of the Oman and the Jurassic Balkan ophiolites. The regional Tauride carbonate platform was initially submerged during Cenomanian time. The causes of this subsidence, which long predated ophiolite emplacement, could include tectonic extension (possibly related to onset of intra-oceanic spreading), or eustatic sea-level change. For example, Early Jurassic neritic deposition ended in the Bey Da~lan carbonate platform, SW Turkey, with an unconformity and evidence of erosion, followed by submergence and covering by hemipelagic to pelagic carbonates of TuronianEarly Tertiary age (Poisson 1977; Robertson 1993). In northern Syria, the local Arabian carbonate platform, of Late JurassicLate Cretaceous age (Jebel Agra), was submerged and overlain by pelagic carbonates of Maastrichtian age, prior to overthrusting by allochthonous continental margin and ophiolitic units (DelauneMay~re 1984; A1-Riyami & Robertson 2002). As in the case of the Jurassic of central Greece, the arrival of the allochthonous units was preceded by evidence of rapid collapse of the carbonate platform. This can again be attributed to flexural subsidence and a switch to deeper-water siliciclastic deposition. The situation in southeastern Turkey and northern Syria is particularly relevant, as these areas can be considered as a direct westward continuation of the Arabian margin from Oman and through Iran. The Early Cretaceous carbonate platform succession on the Arabian margin in the Mediterranean region ended with an unconformity, equivalent to the Wasia-Aruma break in Oman (Fig. 2), and was followed by a deepening-upward succession related to collapse of the carbonate platform. This was then overlain by a

246

A.H.F. ROBERTSON

thick unit of debris-flow deposits ('olistostromes'), including detached blocks of Early Mesozoic volcanic rocks and neritic carbonates (Besni olistostrome) that were derived from the rifted margin of the southern Neotethys (Rigo de Righi & Cortesini 1964; Fourcade et al. 1991). In contrast, in northern Syria only a very thin unit of debris flows is seen beneath the overriding allochthonous continental margin-oceanic units (A1-Riyami & Robertson 2002). As in Oman, the collapsed Cretaceous platforms in SE Turkey and northern Syria are tectonically overlain by thrust sheets, broken formation and mrlange in different areas. In SE Turkey, the collapsed platform was overlain by thrust sheets, broken formation and mrlange (Kastel Formation), equivalent to the Hawasina Complex and the Haybi Complex in Oman (Rigo di Righi & Cortesini 1964; Fourcade et al. 1991; Yllmaz 1993). Similar units exposed at the extreme western edge of the Arabian margin, in the Baer-Bassit region of northern Syria (Figs 4 and 15) have been studied in considerable detail. There, the emplaced continental margin units are entirely made up of tectonic mrlange and broken formation (Baer-Bassit Mrlange; A1-Riyami & Robertson 2002). These exotic units can be restored as Late Triassic to Late Cretaceous, deep-water sediments and volcanic rocks (Delaune-May~re 1984) formed in slope to basinplain settings, together with large masses of alkaline lava of Mid-Jurassic to Early Cretaceous age (A1-Riyami et al. 2002a), interpreted as emplaced seamounts (A1-Riyami & Robertson 2002). Minor volumes of (undated) polymict ophiolite-derived debris flows, present locally, were generated by mass wasting during tectonic accretion and emplacement of the m~lange. Similar relationships are encountered in southern Turkey, related to the initial southward emplacement of continental margin and ophiolitic units onto the Tauride carbonate platform. For example, the succession in a local part of the Tauride carbonate platform, known as the Bolkar Da~ (north of Mersin), passes upwards from Early Cretaceous neritic carbonates into a succession (several tens of metres thick) of redeposited limestones interbedded with mudstones. The sediments are then overlain by mudstones, sandstones and turbiditic sandstones (Demirta~h et al. 1984), derived from overriding allochthonous continental margin and ophiolitic units. These sediments include occasional detached blocks eroded from the subjacent carbonate platform (Demlrta~li et al. 1984; Parlak & Robertson 2004). In some other areas such relations between the regional carbonate platform and the overlying allochthonous units were destroyed, either by high-pressure-low-temperature (HP-LT) metamorphism, as along the northern margin of the Tauride platform in western Turkey (Anatolide units), or by re-thrusting related to continental collision in Early Tertiary time. In the central Taurus Mountains, the continental margin-derived Bey~ehir-Hoyran Nappes were emplaced southwards over a slice of Mesozoic carbonate platform rocks, known as the Hadim nappe, in latest Cretaceous time. This allochthonous unit is interpreted as part of the regional Tauride carbonate platform (and its substratum) at the time of the initial ophiolite emplacement onto the Tauride margin. This unit was later detached and thrust southwards, together with the overriding Bey~ehir-Hoyran nappes during Early Tertiary collisional deformation (Andrew & Robertson 2002). The Bey~ehir-Hoyran Nappes include large ophiolite thrust sheets (e.g. Bey~ehir ophiolite) and large volumes of ophiolitic mrlange. The ophiolite appears to have been severely dismembered during its initial southward emplacement onto the leading edge of the Tauride carbonate platform. However, the structure was further complicated by re-thrusting during Eocene time. This thrusting placed the ophiolite and mrlange near the base of a stack of continental margin and ocean-derived thrust sheets (Andrew & Robertson 2002). Further east, where the Hadim Nappe is absent (e.g. Pozantl-Karsantl area), the ophiolite remained at the highest structural level above the emplaced continental margin and accretionary units.

Further west, in NW Turkey (e.g. Kutahya area) the uppermost levels of the Anatolide carbonate platform experienced H P - L T metamorphism and strong deformation, as seen in Oman south of the Semail Gap. The Anatolide platform can be considered as the metamorphosed northerly, leading edge of regional Tauride carbonate platform further south. The Anatolide carbonate platform is inferred to have undergone attempted subduction followed by rapid exhumation during the latest Cretaceous, much as in Oman (Okay 1986; Okay et al. 1998, 2001; Onen & Hall 2000; ~)nen 2003). Again, as in Oman, subduction and exhumation predate continental collision, which did not take place until Palaeocene time at the earliest in NW Turkey (Okay et al. 2001). The northern Neotethyan-derived sedimentary and volcanic units again record a rifted passive margin setting of Triassic to Late Cretaceous age. Facies, structure and the thickness of the allochthonous marginal units were influenced by the pre-existing Mesozoic palaeogeography. In the west, the Lycian Thrust Sheets (de Graciansky 1972; Fig. 16) constitute a vast area of allochthonous rocks, comparable in scale (if not in volume) with the Hawasina Complex (Hamrat Duru Group) in Oman. These thrust sheets can be restored as a proximal to distal rifted continental margin of Late Triassic to Late Cretaceous age (Collins & Robertson 1997, 1998). After its initial emplacement in the latest Cretaceous and covering by shallow-water carbonate sediments in Early Tertiary time the entire allochthon was re-thrust southwards during the Early Tertiary continental collision phase that ensued. In contrast, the counterparts of the Lycian Nappes further east, the Bey~ehir-Hoyran Nappes, document a thinner, more sediment-starved passive margin that was dominated by a marginal Triassic rift basin and adjacent carbonate build-ups (Andrew & Robertson 2002). The structure of these emplaced continental margin units, like that of the associated ophiolite, was strongly modified by large-scale re-thrusting during the Early Tertiary collisional deformation. As a result, the Mesozoic palaeogeographical setting cannot simply be restored by unstacking the present pile of thrust sheets. The inferred counterpart of the Bey~ehir-Hoyran nappes further SE, the Mersin M~lange, is strongly dissected into mrlange and broken formation. This disruption mainly relates to processes of subduction-accretion within the northern Neotethys, followed by further mixing of units during emplacement over the continental margin (Parlak & Robertson 2004). Part of the correlative m~lange is also exposed along the northern margin of the Tauride Carbonate Platform (local Bolkar Da~) where it is known as the Alihoca Mrlange (Dilek et al. 2000). The Mersin and Alihoca mrlanges retain a 'normal' tectonostratigraphy, with the ophiolite at the top of the thrust stack (Fig. 17). The probable reason is that these units escaped the Tertiary re-thrusting that affected the areas further west. Metamorphic soles are again present beneath many of the Late Cretaceous ophiolites (Woodcock & Robertson 1977; Whitechurch et al. 1984). However, few if any of these soles are as intact as in Oman, or the Jurassic Balkan ophiolites mentioned earlier. Metamorphic soles are well developed associated with the Baer-Bassit ophiolite, northern Syria (Whitechurch et al. 1984). However, no complete assemblage of both amphibolite- and greenschist-facies metamorphic rocks can be recognized in any one exposed unit, possibly because of disruption during emplacement onto the Arabian continental margin (A1-Riyami et al. 2002b). Metamorphic sole rocks are also well developed beneath the Lycian ophiolite, but are strongly dismembered associated with Tertiary collision-related deformation (Collins & Robertson 1997; (~elik & Delaloye 2003). The tectonic style is similar to the metamorphic sole lithologies and associated mrlange beneath the Jurassic Pindos ophiolite in northern Greece (see above). Elsewhere, within the Bey~ehir-Hoyran nappes, the metamorphic sole ranges from nearly complete (Bey~ehir Lake area; Elitok 2001), to detached blocks within mrlange (east of Bey~ehir; Monod 1977; Ozgul 1984; Andrew

E. MEDITERRANEANOPHIOLITEEMPLACEMENT & Robertson 2002). The large Hatay ophiolite, southern Turkey, lacks a preserved metamorphic sole, possibly because of removal during tectonic emplacement. In some cases the ophiolite soles are cut by isolated tholeiitic dykes (e.g. Mersin, Lycian and Plnarba~l ophiolites) that are attributed to the initial phases of oceanic arc-type magmatism. This followed the onset of tectonic convergence within the ocean, prior to emplacement onto a continental margin (Parlak et al. 1995, 2000; Collins & Robertson 1997; Parlak & Delaloye 1999; ~elik & Delaloye 2003). In contrast, the Baer-Bassit metamorphic sole lacks such crosscutting dykes (A1-Riyami et al. 2002b). Late Cretaceous ophiolites are mainly preserved as fragments of a once regionally extensive thrust sheet, similar in scale to the Oman ophiolite. Although now separate entities, the Hatay and Baer-Bassit ophiolites are interpreted as preserved parts of a single vast thrust sheet that was emplaced southwards onto the Arabian margin in Campanian-Maastrichtian time (Robertson 2002). The Lycian ophiolite likewise represents one part of a huge thrust sheet that was emplaced southwards onto the Tauride-Anatolide carbonate platform in latest Cretaceous time (de Graciansky 1972; Collins & Robertson 1998; Okay et al. 2001). In contrast to the Balkan ophiolites, few of the Late Cretaceous ophiolites exhibit intact deep-sea sediment covers, the main exceptions being the Troodos (Robertson & Hudson 1974; Urquhart & Banner 1994) and Baer-Bassit (Parrot 1977; A1-Riyami & Robertson 2002b) ophiolites. These ophiolites possibly retained sedimentary covers when they were initially emplaced southwards from the northern Neotethys in latest Cretaceous time. However, they soon underwent deep erosion prior to covering by Early Tertiary shelf sediments (e.g. Lycian ophiolite). Similar deep erosion affected some of the Jurassic ophiolites in central Greece (e.g. Evia, Atalanti). Exceptionally, for an ophiolite rooted in the northern Neotethys the Mersin ophiolite retains a largely intact stratigraphy, but without overlying deep-sea sediments (Parlak 1996; Parlak & Robertson 2004). In addition, Late Cretaceous ophiolites that were thrust southwards over the Kir~ehir Massif (e.g. Sarlkaraman and ~iqekda~ ophiolites; Floyd et al. 2000; Yaliniz et al. 2000) retain well-preserved, little-deformed extrusive units with interbedded or overlying pelagic carbonates and 'olistostromes'. The Troodos ophiolite differs from all of the above Late Cretaceous ophiolites that were emplaced onto the Mesozoic carbonate platforms, as its sedimentary cover has remained intact and mainly unaffected by compressional deformation. In contrast to ophiolites further east along the Arabian margin, from Hatay and Baer-Bassit to Oman, there is no evidence that the Troodos was emplaced onto a continental margin in Late Cretaceous time (Gass & Masson-Smith 1963; Robertson & Hudson 1974; Robertson 1990; Robinson & Malpas 1990). Instead, the Troodos ophiolite is envisaged as remaining within the Neotethys ocean and was strongly uplifted only in Neogene-Recent time related to onset of collisional deformation (Robertson 1977; McCallum & Robertson 1990). During Neogene time renewed northward subduction is believed to have activated (or reactivated) a northward-dipping subduction zone to the south of Cyprus (Robertson 1990; Kempler 1998; Mart & Robertson 1998). Eventually the leading edge of the North African margin, represented by a continental fragment (the Eratosthenes Seamount; Woodside 1977), began to collide with the subduction zone, initiating strong uplift, and southward displacement of the Troodos ophiolite and related units during Late Pliocene-Pleistocene time (Robertson et al. 1995). This emplacement process can be seen as essentially a recent example of the trench-passive margin collision model, as inferred for Oman. However, in this case some oceanic crust still remained to the south of the Late Cretaceous oceanic subduction trench. This was not finally consumed until mid-late Cenozoic time, when the North African continental margin (initially the outlying Eratosthenes microcontinental fragment) began to collide with the Eurasian active margin to the north.

247

'Northerly' ophiolitic units: ophiolite emplacement associated with active margins All of the Jurassic and Late Cretaceous examples discussed above relate to ophiolite emplacement onto the former passive margins of continents or microcontinents. In most of these cases subduction was initiated, facing oceanward from the continental margin, followed by the genesis of SSZ-type ophiolites. Subduction-accretion continued until the subduction trench collided with the passive margin, emplacing the ophiolites. From Oman along the Arabian margin to Syria, it is generally believed that the ophiolites were emplaced southwards from the southern Neotethys in latest Cretaceous time (Glennie et al. 1973; ~engrr & Yllmaz 1981; Robertson & Dixon 1984; Dercourt et al. 1986, 1992, 2000; Stampfli et al. 2001). Similarly, as discussed above, Late Cretaceous ophiolites were also emplaced southwards from the northern Neotethys onto the Tauride-Anatolide continent (~engrr & Yllmaz 1981; Robertson & Dixon 1984; Floyd et al. 2000; ()nen & Hall 2000). However, each of the above oceanic basins was characterized by a continental margin on the opposite side of the oceanic basin. These margins commonly exhibit a contrasting tectonic history, including the emplacement of ophiolites associated with active margin processes, especially tectonic accretion of various oceanic units. For Oman, the opposing margin is located in Makran, which was characterized by continentward subduction and active margin processes during Tertiary to Recent time (e.g. Glennie et al. 1990). Thus, in future, detailed comparisons should be made with areas such as this rather than Oman. Ophiolites associated with active margin settings have been termed Cordilleran ophiolites in the circum-Pacific region, and contrasted with Tethyan-type ophiolites (e.g. Moores 1982; see Wakabayashi & Dilek 2003, for a review). Below, we first briefly discsuss possible active margin-related ophiolites of Jurassic age in the Balkan region, then consider those of Late Cretaceous age further east.

Jurassic active m a r g i n ophiolite e m p l a c e m e n t in the B a l k a n region

Recognition of active margin settings related to ophiolite emplacement in the Balkan region is complicated by the conflicting interpretations of the number and location of ophiolite suture zones and the direction of ophiolite emplacement. These questions were discussed over several decades by various workers (e.g. Mercier 1968; Smith et al. 1979; Robertson & Dixon 1984; Smith & Spray 1984; Mountrakis 1986; Smith 1993; Robertson et al. 1996; Channell & Kozur 1997). One interpretation, which was summarized in an earlier section of the paper, is that the Pindos (sub-Pelagonian) zone (Fig. 4) originated as a small, rifted Mesozoic oceanic basin (Pindos-Mirdita ocean). Ophiolites formed in this basin and were emplaced towards the NE (in present co-ordinates) onto the former passive margin of the Pelagonian carbonate platform during Late Jurassic time (Smith et al. 1975, 1979; Robertson et al. 1991; Doutsos et al. 1993; Rassios et al. 1994; Smith & Rassios 2003). In an alternative hypothesis, all of the Jurassic ophiolites were emplaced from a Vardar ocean within the Vardar zone (to the east of the Pelagonian zone) and the Pindos zone is seen as a rift basin (without ophiolite root zones) (Aubouin et al. 1970; Dercourt et al. 1986, 1992, 2000; Stampfli et al. 2001). In either of the above tectonic interpretations the Vardar zone represents an important oceanic basin and it is generally agreed that the northerly (Eurasian) margin of this ocean basins shows evidence of an active margin setting during Late Jurassic-Early Cretaceous time and probably also during latest Cretaceous-Early Tertiary time. There is good evidence from the Vardar zone that a volcanic arc was constructed on a continentally floored unit, the Paikon

248

A.H.F. ROBERTSON

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Fig. 19. Tectonic model for the setting of the Jurassic Guevgueli ophiolite along the southern margin of the Eurasian continent. This ophiolite was preserved relatively in situ without long-distance horizontal transport, in contrast to many other Tethyan ophiolites. From Robertson (2002). sub-zone. In many interpretations, this arc unit is restored to a position adjacent to the Serbo-Macedonian zone, that was effectively part of the southern margin of Eurasia by Late Jurassic time (Mercier 1968; Brbien et al. 1986, 1987; Baroz et al. 1987; Kostopolous et al. 2001; Brown & Robertson 2003, 2004). The Paikon arc is inferred to have rifted from the Eurasian margin to open a small back-arc basin (Guevgueli) in Late Jurassic time (c. 150 Ma; Figs 18 and 19). The Late Jurassic Guevgueli ophiolites and related ophiolites of the Inner Hellenide Ophiolite Belt, mainly exposed elsewhere in the Chalkidiki Peninsula, are inferred to have formed in this marginal basin (Brbien et aL 1986; Danelian et al. 1996; Tsikouras & Hatzipanagiotou 1998; Robertson 2002). The Guevgueli ophiolite exhibits contact metamorphism against adjacent continental crust that is correlated with the Serbo-Macedonian zone (e.g. Smith 1993). This ophiolite lacks exposed mantle sequences and a metamorphic sole, and is intruded by contemporaneous granitic rocks (146-156 Ma Fanos granite). The Guevgueli ophiolite and related ophiolites of the Inner Hellenide Ophiolite Belt were later deformed and uplifted, but show little evidence of being emplaced as far-travelled thrust sheets relative to the metamorphic basement of the SerboMacedonian zone to which this ophiolite is apparently welded. The Guevgueli ophiolite extends northwestwards into the Macedonia area of former Yugoslavia but it is not known beyond this. The Guevgueli ophiolite, therefore, records back-arc extension along the active continental margin of Eurasia, a mode of genesis and extension that differs strongly from the Oman-type model. Comparable ophiolites that formed in a rifted back-arc basin setting without undergoing later thrust emplacement include the Rocas Verde ophiolites of southern South America (Stern & De Wit 2003). If it is accepted that the ophiolites of the Pindos (SubPelagonian) zone originated within a westerly Pindos-Mirdita ocean basin, separate from the Vardar ocean, then how many

Fig. 18. Simplifiedgeological map of the Jurassic Guevgueli ophiolite, Vardar zone, northern Greece. From Robertson (2002).

truly 'Vardarian' ophiolites actually exist? Such Vardar ophiolites do include large bodies in Serbia (Karamata & JankoviE 2000; Karamata 2006) that are located to the north of the DrinaIvanica continental unit (equivalent to the Pelagonian zone in Greece), and there are also scattered dismembered ophiolitic units in Greece (e.g. in Argolis; Clift & Robertson 1989). Any oceanic crust remaining within the Vardar zone to the SW of the Eurasian margin (in present co-ordinates) after Late JurassicEarly Cretaceous time was mainly subducted during Late Mesozoic-Early Tertiary time, prior to continental collision during latest Cretaceous-Early Tertiary time (Robertson & Dixon 1984; Dercourt et al. 1986, 1992; Karamata & JankoviE 2000; Sharp & Robertson 2006). In general, the Neotethys narrowed northwestwards (from central Bosnia through Croatia) making it difficult to separate ophiolite emplacement onto southerly passive margins from emplacement along the active margin to the north. Late Cretaceous active margin ophiolite e m p l a c e m e n t in the eastern region

Evidence for the processes of ophiolite emplacement associated with active (convergent) margins is preserved along the northern margin of the southern Neotethys ocean in SE Turkey, Cyprus and SW Turkey. In contrast to the emplacement of ophiolites onto passive margins (Oman-type model) as discussed earlier in the paper, each of these settings shows specific features such that a single tectonic model is not applicable to all cases. The southern Neotethys, now preserved as the Antalya, Mamonia and SE Turkish sutures (Fig. 4), rifted in the Triassic and closed related to northward subduction during latest Cretaceous-Early Miocene time (Hall 1976; Akta~ & Robertson 1984; Yllmaz 1993; Yilmaz et al. 1995; Robertson 1998; Robertson et al. 2004b). As discussed earlier in the paper, ophiolites were emplaced onto the southern. Arabian passive margin in latest Cretaceous time. However, large ophiolites were also emplaced along the opposing, northerly active margin of the southern Neotethys. These ophiolites now form part of the Tauride allochthon in SE Turkey (Hall 1976; Yazgan & Chessex 1991; Yllmaz 1993; Fig. 20). The thrust sheets were finally emplaced in their present position related to collision of the Tauride and Arabian continents in Miocene time ($engrr & Yllmaz 1981; Dewey et al. 1986; Robertson et al. 2004). From east to west, these Tauride ophiolites include the Ytiksekova ophiolite (near Lake Van), the Guleman, Krmtirhan, Ispendere and Elazl~ ophiolites, and further west again the Befit (or Grksun) ophiolite (Akta~ & Robertson 1984; Michard et al. 1984; Yazgan & Chessex 1991; Yalmaz 1993; Beyarslan & Bingrl 1996, 2002; Parlak et al. 2001, 2006; Parlak & Rizao~lu 2004; Rizao~lu et al. 2004; Robertson et al. 2006).

E. MEDITERRANEANOPHIOLITEEMPLACEMENT

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All of the above Late Cretaceous ophiolites originated as a complete ophiolite sequence, although most were partially dismembered during emplacement. Geochemical evidence indicates formation in a subduction-related setting (Parlak et al. 2001, 2006). In addition, the extrusive units of the ophiolites comprise basic, intermediate and silicic lavas, together with large volumes of volcaniclastic sediments, as exposed in the G6ksun (Befit) ophiolite (Parlak et al. 2006; Robertson et al. 2006), and also associated with the lspendere and Elazi~ ophiolites (Rizao~lu et al. 2004). Some of these ophiolites (e.g. Kfmfirhan ophiolite) are cut by calc-alkaline plutonic rocks of Late Cretaceous age (Yazgan & Chessex 1991), recently redated to around 7 0 85 Ma (Parlak & Rizao~lu 2004). The ophiolites are tectonically overriden by the Malatya and Keban carbonate platforms, which are correlated with the northern margin of the southern Neotethys ocean. Crucially, large calc-alkaline plutons cut both the ophiolites and the overlying margin units, as seen in both the Keban platform (Yazgan & Chessex 1991) and the Malatya platform (Perincek & Kozlu 1984). Figure 21 shows a tectonic model to explain the origin and emplacement of these ophiolites. They were formed above a Late Cretaceous northward-dipping intra-oceanic subduction zone with the Tauride (MalatyaKeban-Tauride) platform to the north. A second subduction zone-convergence zone (of unknown displacement) then carried the ophiolites northward until they docked with the overriding Tauride active margin (Robertson 1998, 2000). Subduction was then impeded, possibly by the young, hot (buoyant) nature of the oceanic lithosphere. Subduction then continued and both the accreted ophiolite and the overriding continental margin were cut by calc-alkaline, arc-type plutonic rocks. The docking of oceanic crust and the intrusion of calk-alkaline intrusions was complete by latest Cretaceous time, as the thrust contact is locally sealed by Early Tertiary clastic sediments. Some of the accreted Late Cretaceous oceanic crust remained deeply

Fig. 20. Outlinetectonic map of SE Turkey showing the setting of the main Late Cretaceous ophiolites. From Robertson (2002).

submerged along the Tauride margin after Late Cretaceous time. Notably, ophiolitic slices preserved in a structurally low, frontal position of the thrust belt include a deformed Late Cretaceous ophiolite that is depositionally overlain by deep-sea sediments and debris flows ('olistostromes') derived the Tauride margin to the north (i.e. the Killan Imbricate Unit; Akta~ & Robertson 1984). The regionally extensive Late Cretaceous ophiolites were, thus, initially emplaced along the northern active margin of the southern Neotethys, where they remained until final closure of the oceanic basin in the Miocene when they were thrust southward over the Arabian margin (Robertson et al. 2007). A second example of ophiolite emplacement in an active margin setting is provided by the Late Cretaceous ophiolites associated with the Pontides of northern Turkey. These ophiolites formed within the northern Neotethys ocean and were emplaced along the margin of this basin in latest Cretaceous-Early Tertiary time. In addition, small dismembered ophiolite exposures in the Armutlu Peninsula, western Pontides, are believed to have formed in a discrete intra-Pontide oceanic basin (G6nctio~lu et al. 1987; Yllmaz et al. 1997; Robertson & Ustafmer 2004), although these will not be discussed further here. As noted above, the southern margin of the northern Neotethys was characterized by the emplacement of Late Cretaceous ophiolites onto the Tauride margin in latest Cretaceous time (Oman-type model). However, the ophiolites associated with the opposing, Pontide margin show a contrasting setting of emplacement. Large ophiolites are present associated with the western, central and eastern Pontides. These ophiolites apparently originated as a complete ophiolite sequence, but only the lower plutonic parts of these ophiolites are now well preserved (e.g. Eski~ehir and Erzincan areas). Late Cretaceous ages are locally indicated by the presence of Late Cretaceous pelagic sediments intercalated with ophiolitic extrusive rocks (i.e. in the Tosya-Kargl area; central Pontides; Ttiystiz 1990; Ustafmer & Robertson 1997). However, most of

250

A.H.F. ROBERTSON

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Fig. 21. Tectonic model for the accretion of Late Cretaceous ophiolites in SE Turkey including the Berit (Grksun), Ispendere, Krmtirhan, Guleman and Yuksekova ophiolites.

these ophiolites remain poorly dated and metamorphic soles (used to infer ophiolite ages) are typically absent. Geochemical evidence indicates a SSZ setting, as indicated by the composition of chromites within mantle harzburgites in the Erzincan area, eastern Pontides (Rice et al. 2006), and by the chemistry of ophiolitic extrusive rocks in the Tosya-Karg~ area, central Pontides (Ustarmer & Robertson 1997; Rice et al. 2006). A key observation in the eastern Pontides, from the Refahiye ophiolite complex in the Erzincan area, is that the ophiolitic plutonic rocks include numerous screens (i.e. inclusions) of highly deformed metamorphic host rocks, which can be correlated with the Pontide basement to the north (Rice et al. 2006). This implies that this ophiolite formed by rifting along the Eurasian margin in Late Cretaceous time. In addition, the ophiolites are imbricated with large thrust sheets of basic-intermediate-silicic extrusive igneous and volcaniclastic rocks, which are well exposed in the central and eastem Pontides. These lithologies are dated as Late Cretaceous in age in the central Pontides (TtiysLiz 1990). This unit is interpreted as part of an emplaced oceanic magmatic arc (Rice et al. 2006). A several kilometres thick associated unit of mixed terrigenous-volcaniclastic sediment, with minor subduction-influenced volcanic rocks of late Cretaceous age in the Central Pontides (Kosda~ unit) is interpreted as forming in a marginal basin setting along the Eurasian margin (Rice et al. 2006). In addition, a several kilometres thick shallowing upward Late Cretaceous-Early Tertiary succession derived from the arc and accretionary prism material is well exposed in the eastern Pontides (~erpaqin unit) and is interpreted as a

collapsed forearc basin. The thrust sheets of oceanic arcmarginal basin-forearc basin units are intercalated with large thrust slices of m~lange, interpreted as emplaced accretionary prism material (Rice et al. 2006). The Late Cretaceous ophiolitic rocks of the central and eastern Pontides are, therefore, interpreted to have formed in a back-arc basin that was located along the southern margin of Eurasia during Late Cretaceous time. This marginal basin was bordered by an oceanic arc, a forearc basin and an accretionary wedge. Subduction is inferred to have been northwards beneath the Eurasian margin (Okay & ~ahintiirk 1997; Ustarmer & Robertson 1997). The ophiolites and other units were initially deformed by northvergent thrusting; this presumably relates to northward thrusting of the ophiolites and other oceanic units onto the Eurasian margin. The driving mechanism is unclear but may relate to southward underthrusting-subduction of the inferred marginal basin in latest Cretaceous time. All of the above units later underwent pervasive southward thrusting and back thrusting that are attributed to final closure of the northern Neotethys and collision of the Eurasian and Tauride continents in Early Tertiary time (Rice et al. 2006). The above settings of Late Cretaceous ophiolite emplacement relate to the northern margins of the southern and northern Neotethys, respectively. In both cases, similar settings of ophiolite emplacement are applicable for hundreds of kilometres along these active margins. However, contrasting modes of ophiolite emplacement are seen along more specific segments of margins that were influenced by a locally variable palaeogeography. The first of these more local setting of ophiolite emplacement is the Late Cretaceous-Early Tertiary deformation and emplacement of Late Cretaceous ophiolites associated with the Isparta Angle, SW Turkey (Fig. 22). Palaeomagnetic data indicate that the Isparta Angle restores as a broad northward re-entrant of the southern Neotethys during Mesozoic time (Kissel & Poisson 1986; Morris & Robertson 1993; see Robertson et al. 2003, for

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Fig. 22. Outlinetectonic map of the Isparta Angle, SW Turkey, showingthe Late Cretaceous ophiolites (Tekirova and Grdene) and other units of the Antalya Complex. From Robertson (2002).

E. MEDITERRANEANOPHIOLITE EMPLACEMENT

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Fig. 23. Setting of ophiolite emplacement in the Isparta Angle, SW Turkey. (a) The area restores as an embayment of the southern Neotethys; (b) oceanic lithosphere formed in a SSZ setting; (c) ophiolites and marginunits were initially displaced in latest Cretaceous time and finally emplaced over the continental margin in Early Tertiary time; (d) the Isparta Angle began to form, while Neotethys still remained partially open to the south. From Robertson (2003).

a review). The Antalya area formed part of the northern, active margin of the southern Neotethys in Late Cretaceous time. Within this segment, the northern margin was oriented at an oblique angle to the generally east-west trend of the southern Neotethys. As a result of this obliquity, the emplacement of the Late Cretaceous ophiolites was strongly influenced by strike-slip faulting (Woodcock & Robertson 1982, 1985; Fig. 23). In the SW area of the Isparta Angle, the Cretaceous Tekirova ophiolite (Juteau 1975), and more marginal (proximal) units further west (i.e. G6dene zone) were dismembered into small (tens of kilometres or less) tectonic 'terranes' separated by highangle shear zones, interpreted as strike-slip faults (Woodcock & Robertson 1982). Internally, the Late Cretaceous Tekirova ophiolite, exposed along the coast (Reuber 1984), is relatively intact, although the extrusive rocks are missing. Geochemical data suggest a SSZ setting of formation (Ba~cl et al. 2006). The ophiolite is depositionally covered by a chaotic unit of

251

Maastrichtian age, composed of detached ophiolitic rocks and clastic sediments (Robertson & Woodcock 1982b; Lagabrielle et al. 1986). These clastic sediments lack a terrigenous component, implying that they formed in an oceanic setting before the ophiolite was emplaced over the adjacent Bey Da~larl carbonate platform. In addition, large-scale mass-wasting of ophiolite-derived debris flow deposits, commonly rich in serpentinite, took place on the adjacent (more inboard) terrane (Kemer zone) that is underlain by Late Palaeozoic pre-rift basement, Triassic synrift sediments and Jurassic-Cretaceous post-rift, mainly carbonate, sediments (Robertson & Woodcock 1982b). This unit is interpreted as one or several small continental fragments, or large rift blocks that were isolated along the rifted continental margin of the southern Neotethys when spreading began in Late Triassic time (Robertson & Woodcock 1980b). The adjacent small 'terrane' further west (G6dene zone) includes strongly dismembered Late Cretaceous ophiolitic rocks, Triassic rift-related volcanic rocks and sediments, and very rare small exposures of metamorphic sole-type rocks (Robertson & Woodcock 1982b; Ydmaz 1984; (~elik & Delaloye 2003). The G6dene zone also includes distinctive ophiolite-derived breccias and mass-flow deposits related to tectonic amalgamation of this unit in a strikeslip setting (Cmarqlk Breccias; Robertson & Woodcock 1980a). The Late Cretaceous Tekirova ophiolite was initially deformed in latest Cretaceous (Maastrichtian) time associated with strikeslip or transpression, outboard of the margin of the Bey Da~lan carbonate platform. The ophiolite was later emplaced onto the adjacent Tauride continental margin (Bey Da~lan platform) in response to partial closure of Neotethys during Early Tertiary (Late Palaeocene-Early Eocene) time (Robertson & Woodcock 1982b; Poisson 1984; Dilek & Rowland 1993; Robertson 1993). The neighbouring Bey Da~lan carbonate platform to the west remained undeformed in latest Cretaceous time as a submerged platform undergoing pelagic deposition into Early Tertiary time (Poisson 1977; Robertson & Woodcock 1982b). In summary, this is an example of ophiolite emplacement in which strike-slip, rather than orthogonal emplacement played an important role, in contrast to most of the other settings of ophiolite emplacement discussed above. The final setting of ophiolite 'emplacement' discussed here is one in which the ophiolite was not actually tectonically emplaced onto a continental margin at all, but instead was rotated as a microplate, still within an oceanic setting, associated with deformation of its margins. An example of this is the latest CretaceousEarly Tertiary palaeorotation of the Troodos ophiolite (Fig. 24). The Troodos ophiolite, Cyprus, was generated around 90 Ma (Mukasa & Ludden 1987) in a SSZ setting within the southern Neotethys ocean (e.g. Pearce et al. 1984). The Troodos ophiolite, together with the Hatay, Baer-Bassit, Tekirova and other Late Cretaceous southern Neotethyan ophiolites formed above a northdipping subduction zone (Robertson 1998). To the east the BaerBassit, Hatay and other ophiolites underwent southward emplacement onto the Arabian margin (Oman-type model) as discussed earlier, and to the west the Antalya ophiolites (e.g. Tekirova) were strongly affected by oblique (strike-slip) emplacement. However, the Late Cretaceous tectonic setting of the Cyprus region was strongly influenced by 90 ~ anticlockwise palaeorotation of the Troodos ophiolite. The rotation was first discovered by palaeomagnetic study of the Troodos ophiolite (Moores & Vine 1971) and was accurately dated by detailed palaeomagnetic studies of the in situ deep-sea sedimentary over of the ophiolite (Clube et al. 1986; Morris et al. 1990). The rotation began in Campanian-Maastrichtian time and continued at an approximately constant rate until Early Eocene time (Morris 1996). Recent work shows that the Hatay ophiolite also underwent anticlockwise rotation, but to a smaller extent (Morris et al. 2006). Also, tectonic rotations that may be extreme are recorded in the more dismembered Baer-Bassit ophiolite further south (Morris et al. 2002).

252

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One explanation for the rotation of the Troodos ophiolite is that this resulted from oblique northward subduction beneath the northern active margin of the southern Neotethys (Clube et al. 1985). Another possibility is that the palaeorotation relates to

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Fig. 25. Summaryof the main models of ophiolite emplacementin the Eastern Mediterranean, Tethyan region. (a) Collision of SSZ ophiolite with a passive margin (Oman-typemodel); (b) emplacementof ophiolites in active margin settings associated with subduction and accretionary complexes.Model (a) produces very similar geological features regardless of age or location (e.g. Jurassic Greek v. Late Cretaceous Turkish ophiolites), whereas model (b) shows considerablevariation in different examples.

collision of the inferred north-dipping subduction zone with the Arabian promontory to the east, thus triggering pivoting and anticlockwise rotation of a Troodos microplate (Clube & Robertson 1986). It was also proposed that the Late Cretaceous tectonic history of the Kyrenia Range, northern Cyprus, and of the Mamonia Complex, western Cyprus, were strongly influenced by strike-slip around the periphery of the Troodos microplate (Robertson 1990). The probable explanation for the sparsity of evidence for tectonic emplacement of the Troodos ophiolite in Late Cretaceous time, compared with ophiolites further east (e.g. Hatay; Baer-Bassit), is that the Troodos remained protected from deformation within an embayment of the Arabia-North Africa continental margin, today known as the Levant Sea (Clube & Robertson 1986; Robertson 1990). The palaeorotation of the Troodos microplate was spatially and temporally associated with the deformation of adjacent units. To the north, the Troodos ophiolite is inferred to terminate abruptly against the Kyrenia Range, based on geophysical evidence (Aubert & Baroz 1978). The Kyrenia Range is restored as part of the former southern margin of Tauride-related continental units to the north (Robertson & Woodcock 1986). The Troodos is interpreted as having been thrust beneath this margin in latest Cretaceous time (Clube & Robertson 1986) in a similar manner to the eastern Tauride ophiolites discussed earlier. The same explanation for the halting of subduction may apply; that is, the arrival of young buoyant SSZ-type crust at a trench. Furthermore, to the west and SW the Troodos is tectonically juxtaposed with Mesozoic rocks of continental margin-oceanic affinities, represented by the Mamonia Complex (Lapierre 1972). These Mamonia units (Dhiarizos and Ayios Photios Groups; Swarbick & Robertson 1980) can be restored to a position to the NE of the present position of the Troodos ophiolite. This inference is partly based on lithological correlation of deep-sea sediments (e.g. Early Cretaceous quartzose sandstones) exposed in the Mamonia Complex (Akamas Sandstones) with counterparts in the southern part of the Antalya Complex (Kumluca area) on the Turkish mainland (Robertson & Woodcock 1982b). The Mamonia Complex is unlikely simply to represent a continental margin located to the south, as there

E. MEDITERRANEANOPHIOLITE EMPLACEMENT is no evidence of southward emplacement of the Troodos ophiolite onto a continent to the south of Cyprus in latest Cretaceous time (see Dilek & Flower 2003). The first evidence of collisional deformation is, instead, related to collision of the Eratosthenes Seamount with the subduction trench in Plio-Quaternary time (Robertson 1998, and references therein). Deformation of the Mamonia Complex took place in latest Cretaceous time by a combination of thrusting, strike-slip faulting and gravity tectonics (e.g. Robertson & Woodcock 1979; Swarbrick 1980; Malpas et al. 1992; Bailey et al. 2000; see Robertson & Xenophontos 1993, for review). The emplacing Mamonia terrane was sealed by debris-flow deposits (Kathikas Mrlange) shed from the Mamonia continental margin-oceanic units during latest Cretaceous time (Swarbrick & Naylor 1980) while still in an oceanic setting. This setting is confirmed by the cover of deep-sea carbonate sediments of latest CretaceousEarly Tertiary age (Urquhart & Banner 1994; Lord et al. 2000). Gravity emplacement is, additionally, exemplified by the Moni Mrlange, southern Cyprus, in which occur large blocks of, for example, Cretaceous continentally derived sedimentary rocks, including proximal quartzose sandstones (Parekklisha Sandstone) and deeper-water siliceous pelagic limestones (Monagroulli Limestones). These exotic units were emplaced in a matrix of Late Cretaceous deep-sea clays (Moni Clay) that depositionally overlie the Troodos ophiolite (Robertson 1977; Gass et al. 1994; Urquhart & Robertson 2000). The deformation of the continental margin units exposed in Cyprus (Mamonia Complex and Kyrenia Range) is, therefore, closely related to the palaeorotation of the Troodos microplate and differs markedly from that seen in adjacent areas of the southern Neotethys. The question remains as to whether the palaeorotation of the Troodos ophiolite relates more to the tectonic evolution of the active Tauride margin to the north, or to the emplacement of ophiolites along an embayment of the North African-Arabian passive margin to the south (as in the pivoting slab model). Much depends on the width of the southern Neotethys remaining by latest Cretaceous time when the palaeorotation began. According to evidence from palaeomagnetic inclinations, the Troodos ophiolite originated at c. 20~ closer to Gondwana than Eurasia (Morris 2004). The Troodos and other southern Neotethyan ophiolites (e.g. Hatay, Baer-Bassit, Kocali) presumably migrated southwards as northward subduction proceeded and Arabia-North Africa drifted northwards. Indeed, the surviving southern Neotethys was probably not much more than a few hundred kilometres wide by latest Cretaceous time, as there is little evidence of arc volcanism associated with the latest stages of Tertiary destruction of the southern Neotethys (Akta~ & Robertson 1984; Robertson et al. 2006, 2007). Remnants of Late Cretaceous oceanic crust, chemically of subduction-related affinity (Floyd et al. 1992), remained within the remnant southern Neotethys. These units were only finally accreted to the Tauride active margin to the north after Late Eocene time, based on evidence from the Misis-Andlnn Complex in coastal southern Turkey (Robertson et al. 2004). It is, therefore, probable that the present scale of Cyprus represents a large proportion of the width of the southern Neotethys ocean that remained to the east of the north-south Levant embayment by latest Cretaceous-Early Tertiary time. If so, the palaeorotation of the Troodos microplate could have been triggered by collision of the subduction trench with the irregular southerly ArabiaNorth African continental margin, while also involving units associated with the northerly, active margin of the southern Neotethys (i.e. Kyrenia Range and Mamonia Complex). In summary, the well-known palaeorotation of the Troodos microplate is an example of displacement of an ophiolite while still within the ocean, but without involving emplacement onto a continental margin. This, in turn, re-emphasizes the diversity of the processes of ophiolite emplacement and displacement exemplified within the Eastern Mediterranean region.

253

Conclusions The Oman-type trench-margin emplacement model can be applied to a wide range of Mid-Jurassic and Late Cretaceous ophiolites throughout the Eastern Mediterranean region (Fig. 25a). In this interpretation, most of these ophiolites were generated above oceanward-dipping subduction zones. These subduction zones consumed Neotethyan oceanic lithosphere, creating an accretionary prism of deep-sea sediments and volcanic rocks, until the subduction trench collided with a passive margin, either a continent or a microcontinent. This margin was then flexurally loaded and collapsed, facilitating final emplacement of the ophiolites onto submerged former passive margins. The Oman-type emplacement model can be successfully applied to many of the ophiolites of the Eastern Mediterranean region. These include those of Mid-Jurassic age that were emplaced onto collapsed continental fragments, such as the Pelagonian carbonate platform and counterparts further NW along strike (e.g. Othris, Vourinos, Pindos-Mirdita and Dinaride ophiolites of Greece and former Yugoslavia). This model applies regardless of different interpretations of the locations of the root zones of these ophiolites. In addition, the Oman-type model applies well to the emplacement of Late Cretaceous ophiolites of Turkey, Cyprus, Syria and Iran. These ophiolites were emplaced separately, from two Neotethyan basins: a southern Neotethys to the south of the Tauride carbonate platform, and a northern Neotethys to the north of this inferred microcontinental unit. The mode of emplacement was influenced by the regional palaeogeography. In some areas the collided passive margin was subducted and metamorphosed under high-pressure conditions and soon exhumed. However, the ophiolites typically remained attached to the overriding upper plate and remained unmetamorphosed. In addition, ophiolites were also formed and emplaced along active continental margins during Late Triassic-Early Jurassic, Mid-Jurassic and Late Cretaceous time. These include 'Cordillerran-type' ophiolites formed in association with accretionary prisms (Fig. 25b) and collapsed marginal basins. These ophiolites, again mainly of SSZ-type, show considerable variation in emplacement style. Throughout the Balkan region, the northern margin of the Jurassic Vardar ocean experienced active margin processes, including subduction, marginal arc volcanism and back-arc basin opening. Oceanic crust formed in a marginal basin setting by rifting of the Eurasian margin and was later emplaced as the 'Inner Hellenide' ophiolites, but without evidence of large-scale horizontal transport. In addition, the pre-existing Late Palaeozoic-Early Mesozoic Palaeotethys, exposed in the Pontides of northern Turkey, shows an active margin history of subductionaccretion (e.g. Karakaya Complex), arc volcanism and back-arc opening during Late Palaeozoic-Triassic time. When this marginal ocean basin later closed, prior to Early Jurassic time, a dismembered ophiolite (Ktire ophiolite) was emplaced northwards onto the Eurasian margin. Contrasting settings of ophiolite emplacement characterize the northerly margins of the Jurassic and Cretaceous Neotethyan ocean basins. These were characterized by active margin processes including subduction, accretion, arc magmatism and opening of back-arc basins. Jurassic ophiolites in Greece (e.g. Guevgueli ophiolite) formed by rifting of the southern margin of Eurasia (Serbo-Macedonian zone), followed by emplacement without major horizontal transport over adjacent continental units. Late Cretaceous ophiolites also originated in a back-arc basin setting associated with the Eurasian margin in Turkey (Pontides). However, these ophiolites formed in a more oceanic setting, bordered by an intra-oceanic arc and forearc basin, and were later emplaced northwards over the Eurasian margin prior to Early Tertiary time.

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Ophiolites were also emplaced along the northern, active margin of the southern Neotethys by a variety of processes in different geographical areas. The Late Cretaceous ophiolites of the SE Turkish thrust belt were accreted to the hanging walls of a northdipping subduction zone, where they were then intruded by Late Cretaceous calc-alkaline plutons related to continuing northward subduction. Where subduction was oblique or at a high angle to the adjacent margin the ophiolites and related units were emplaced as small 'terranes' by dominantly strike-slip, or transpressional processes. These ophiolites retain relatively intact vertical successions (e.g. Late Cretaceous Antalya ophiolites), rather than forming the lowangle thrust sheets that are typical of ophiolites emplaced by orthogonal overthrusting. The presence of a highly irregular palaeogeography remaining from the previous r i f t - p a s s i v e margin stage influenced the mode of ophiolite emplacement in the southern Neotethys. For example, the well-known 90 ~ anticlockwise palaeorotation of the Troodos microplate was possibly triggered by collision of the subduction trench with the Arabian promontory to the east. The southern Neotethys was by then relatively narrow and it is likely that both southern and northern margin units were affected by the regional palaeorotation. Many other ophiolites, commonly of a fragmentary highly deformed nature, are present in the wider Eastern Mediterranean area, including former Yugoslavia, Romania, Bulgaria, Armenia, Georgia and Iran. The discussion has covered many of the larger ophiolites in the Eastern Mediterranean region but has necessarily excluded many others. In future, it would be useful to determine whether the emplacement of these other ophiolites can be related to the two main settings of ophiolite emplacement discussed here: (1) ophiolites emplaced by t r e n c h - m a r g i n collision onto former passive margins (Oman-type model); (2) ophiolites emplaced along active margins undergoing subduction and arc volcanism. D. Baty assisted with drafting many of the figures, and Y. Cooper helped prepare the illustrations for publication. Helpful comments on this manuscript were received from J. Winchester and L. Jolivet.

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Eastern Mediterranean basin systems ZVI B E N - A V R A H A M 1, J O H N W O O D S I D E 2, E M A N U E L E L O D O L O 3, M I C H A E L G A R D O S H 4, M A R I O G R A S S O 5, A N G E L O C A M E R L E N G H I 3 & G I A N B A T T I S T A V A I 6

1Department of Geophysics and Planetary Sciences, Tel Aviv University, Tel Aviv, Israel (e-mail: [email protected]) 2Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1082 HV Amsterdam, Netherlands 3Istituto Nazionale di Oceanografia e di Geofisica Sperimentale (OGS), Trieste, Italy 4Geophysical Institute of Israel, P.O. Box 182, Lod 71100, Israel 5Dipartimento di Scienze Geologiche, Corso Italia 55, 95129, Catania, Italy 6Dipartimento di Scienze della Terra e Geologico-Ambientali, Via Zamboni 67, 40127, Bologna, Italy

Abstract: The basins in the Eastern Mediterranean can be divided into those that were formed mainly in post-Miocene time and those

that were formed during the rifting episodes that led to the formation of the Neotethys. The younger basins can be further divided into those that were formed mainly in post-Miocenetime and those that were formed in post-Pliocene time. The separation is not only one of convenience but also corresponds to major adjustments in the plate tectonic situation in the Eastem Mediterranean. The late Miocene deposition of thick evaporites throughout the Mediterranean region, or, where evaporites are missing, the creation of an important erosional unconformity during the extreme lowstand of the Mediterranean, makes the Miocene-Pliocene boundary relatively easy to identify, especially on seismic reflection records. At about the same time, following the collision of the Arabian plate with Eurasia, the Anatolian and Aegean microplates came into existence between the convergent African and Eurasian plates to accommodate tectonic escape between them. The general configurationof the Eastern Mediterraneanbasins reflects the tectonic and structural gradients between the collisional domain of southeastem Turkey and Iran, and the continuing but increasingly limited subductionalong the Calabrian and Hellenic arcs, with the Cyprus and Levantine zones between them. Several distinct zones can be identified in the Eastern Mediterranean. The Dead Sea Fault system marks the edge between the collisional and pre-collisional zones to the east and west, respectively. The meridian through the Anaximander Mountains (30~ forms a rough boundary between the zone of incipient collision to the east and the zone of continuing but late-stage subduction to the west. The Malta Escarpment forms the Eastern boundary of the Eastern Mediterranean basins. The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea share this progressive evolution, with those containing Messinian evaporites to the east and those without to the west. The Sicily Channel with its associated basins is an extensional zone between the Eastern and Western Mediterranean. The basins discussed in this paper are divided into two groups, the larger and older basins and the smaller and younger basins. In the first group are the Ionian Basin and the Levantine Basin, and in the second group the Cilicia Basin, Antalya Basin, Finike Basin, Rhodes Basin, Aegean basins, Sicily Channel basins, Latakia Basin and Larnaca Basin. The Eastern Mediterranean represents the last stage in the evolution of an ocean basin. Given the current motion between Africa and Eurasia, the Eastern Mediterranean will cease to exist in about 6-8 Ma from now. As a result, the larger and older basins are shrinking, whereas the younger and smaller basins are growing. Eventually the smaller basins will also disappear.

The Central and Eastern Mediterranean is dominated by the convergence of the African plate with the Eurasian plate. The relative motion between these two plates has produced, after the closure of an oceanic-type domain, a complex system of contractional structures (i.e. the Cyprus, Hellenic, Calabrian and Maghrebian arcs) along which the stress field is being dissipated. There are very few places in the world where the o c e a n - c o n t i n e n t crustal transition at a passive continental margin is approaching a subduction zone. A thick pile of sediments as old as Mesozoic is being deformed at the collisional margin (Mediterranean Ridge), although the irregular shape of the colliding continental margins leaves portions of the former basin still relatively u n d e f o r m e d (Ionian and Herodotus basins). In the Eastern Mediterranean, the transition between the remnants of a thick Mesozoic oceanic crust located in the Ionian Basin (DeVoogd et al. 1992) and in the Levantine Basin (Ben-Avraham et al. 2002) and the African continental margin (Fig. 1) is approaching the Calabrian, Aegean and Cyprus subduction zones (Dewey et al. 1973; McKenzie 1978). The relative plate motion between Africa, Eurasia and the Aegean microplate produces convergence. The African plate moves in an approximately northward direction at a rate of 1 cm a -1 (Rebai et al. 1992). The Aegean region undergoes extension at a rate of 3 . 5 4.3 c m a -1 in a SSW direction relative to Africa (Le Pichon et al. 1995). A large doubly vergent accretionary complex, the eastern Mediterranean Ridge, forms in a complicated setting of altemation of zones of continental collision (the Pelagian and Cyrenaica promontories) and zones of terminal subduction (the

Ionian Basin). The irregular shape of the converging plate boundaries causes diachronous continental collision, strain partitioning and lateral escape of both the Ionian and Levantine inner zones (Finetti 1976; Chaumillon & Mascle 1995; Le Pichon et al. 1995; Polonia et al. 2002). The southern, A f r i c a n - A r a b i a n , continental margins were considerably less affected by the closing of the basins and remained more or less stable in their position near the present-day coastal areas of northern Africa, Sinai, Israel and Lebanon (Bein & Gvirtzman 1977; Garfunkel 1998). The structure of the southern Levant margin and the adjacent deep marine basin is relatively simple, and the Mesozoic and Cenozoic rock record found on land and in the sea is more or less continuous. Both structure and stratigraphy preserve the signature of the main tectonic events that shaped the basin and its margin. These can be separated into four distinct tectonostratigraphic phases. Similar tectonostratigraphic phases are observed in the Hyblean continental margin in eastern Sicily and the adjacent deep Ionian Basin. The young basins of the Eastern Mediterranean can be divided conveniently into those that were formed mainly in post-Miocene time and those that were formed in post-Pliocene time. The separation is not only one of convenience but also corresponds to major adjustments in the plate tectonic situation in the Eastern Mediterranean. The late Miocene deposition of thick evaporites throughout the Mediterranean region (Hsu et al. 1973a,b) or, where evaporites are missing, the creation of an important erosional unconformity during the extreme lowstand of the Mediterranean, makes the M i o c e n e - P l i o c e n e boundary relatively

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 263-276. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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264

Z. BEN-AVRAHAMETAL.

Fig. 1. Map of the Eastern Mediterranean region. Bathymetriccontours are at a 500 m interval; circles with numbers mark the location of basins, and squares with letters mark other structures.

easy to identify. At about the same time, following the collision of the Arabian plate with Eurasia, the Anatolian and Aegean microplates came into existence between the convergent African and Eurasian plates to accommodate tectonic escape between them (McKenzie 1972; Angelier et al. 1982). The general configuration of the Eastern Mediterranean basins reflects the tectonic and structural gradients between the collisional domain of southeastern Turkey and Iran (McClusky et al. 2000), and the continuing but increasingly limited subduction along the Calabrian and Hellenic arcs, with the Cyprus or Levantine zone between. Just as the Dead Sea Fault system marks the edge between the collisional and pre-collisional zones to the east and west, respectively, the meridian through the Anaximander Mountains (30~ forms a rough boundary between the zone of incipient collision to the east and the zone of continuing but late-stage subduction to the west (Ben-Avraham & Nur 1976; Nur & Ben-Avraham 1978). The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea share this progressive evolution, with those containing Messinian evaporites to the east and those without to the west. In the western part of the Eastern Mediterranean within the regional convergent system, NW-SE-trending troughs have developed in the Strait of Sicily. These depressions represent relatively young (late Miocene to Present) extensional structures that cut across the undeformed Pelagian block, a continental crust that represents a promontory of the African plate margin. In this paper we briefly describe the large and small basins in the Eastern Mediterranean. We also describe the main tectonostratigraphic stages and review the evolution of the Ionian and Levantine basins as the result of the large-scale plate motions that took place in the Eastern Mediterranean region during the Mesozoic and Cenozoic.

Older basins Ionian Basin

The Ionian Sea is a deep marine basin separated by the conspicuous Malta escarpment from the shallow, Hyblean-Malta Plateau on the west. Most researchers agree on the oceanic nature of the

Ionian lithosphere, although its uppermost layer is made of up to 10 km thick sedimentary cover. Its age, however, is debatable, with suggestions ranging from Early Permian (Ben-Avraham & Ginzburg 1990; Catalano et al. 1991; Vai 1994, 2003), to Late Permian to Early Triassic (Stampfli et aI. 2001b), Triassic (Finetti 1984), mid-Jurassic or broadly Mesozoic (Robertson & Grasso 1995; Cantarella et al. 1997), Cretaceous (Dercourt et al. 1986) or even Messinian (Fabricius & Hieke 1977). Evidence of deepening and/or accelerated subsidence of the continental margin sequences around the long-lasting Sicilian Sicani Basin, before Tertiary and Quaternary deformation took place, are known during mid-Cretaceous, Late Triassic to early Jurassic, Mid-Triassic, and especially Early to Mid-Permian times (Charier et al. 1988; Catalano et al. 1991, 1992; Vai 1994, 2003; Dercourt et al. 2000; Stampfli et al. 2001a,b). This evidence is consistent with rift pulses producing WNW-ESE-trending basins. The Early Mesozoic rifting events are similar to those in other domains in the Central and Eastern Mediterranean, and mark a major continental break-up followed by rapid tectonic subsidence that initiated the opening of the Neotethyan ocean (Yellin-Dror et al. 1997). This early evolutionary stage of the Hyblean margin was followed by slow thermal subsidence (Early JurassicLate Cretaceous), northward movement and thrusting (Late Cretaceous-Palaeogene), and continued convergence, uplift and subsidence (Neogene-Quaternary; Yellin-Dror et al. 1997). The Ionian Basin, of which the crustal structure is shown in Figure 2 and a seismic section in Figure 3, is located near the orthogonal plate convergence zone. In the western Ionian Basin plate convergence is nearly orthogonal in a N E - S W direction. Because of the curved shape of the Mediterranean Ridge deformation front, however, a variable plate convergence angle has to be expected, especially in the Messina foredeep where the deformation front approaches a north-south direction. Thrusting and folding in a direction roughly parallel to the Ionian rigid crustal backstop are reflected in shortening and uplift on the Mediterranean Ridge accretionary complex, which results in the tectonic addition of post-Messinian material to the outer deformation front. The main regional detachment surface is located at the base of the Messinian evaporites at the outer wedge. The evaporites, of variable thickness and lateral extent, were deposited during

EASTERN MEDITERRANEANBASIN SYSTEMS

Fig. 2. Main structural and morphological elements of the Ionian Basin. 1, Alpheus seamount; 2, Malta escarpment; 3, Medina seamounts; 4, Messina abyssal plain (Messina Foredeep); 5, Victor Hensen seahill; 6, Medina-Victor Hensen structure (Hieke & Dehghani 1999), including the Medina-Cephalonia line; 7, Sirte abyssal plain (Sirte Foredeep); 8, Bannock buried seamount and related tectonic lineament; 9, Cyrene seamount; 10, Hellenic trench system on the continental backstop of the Mediterranean ridge accretionary complex. The doubly vergent structure of the Mediterranean ridge accretionary complex, identified by an outer deformation front (line with filled) and an inner deformation front, should be noted. Asterisk marks position of section shown in Figure 3. the Messinian (late Miocene) desiccation of the Mediterranean Sea (Hsu et al. 1973a,b). The rheology of the deforming materials is expected to have undergone drastic changes in the last 5 Ma (Kastens et al. 1992; Kukowsky et al. 2002; Reston et al. 2002a,b), permitting the identification of pre- and post-Messinian wedges. This has led to the post-Messinian development of an accretionary complex whose features are similar to those of salt-bearing fold-and-thrust belts (Davis & Engelder 1985). Below the inner portions of the wedge the detachment becomes progressively deeper, and cuts into the top of the Mesozoic carbonates belonging to the African foredeep and foreland (Cita e t al. 1981; Camerlenghi et al. 1995; Chaumillon & Mascle 1995, 1997). Thrusting to the NE of accreted sediments above the Ionian rigid backstop (Le Pichon et al. 2002) occurs along the inner deformation front. The Sirte and Messina foredeep areas remain completely undeformed. The structural style is affected by geological structures rooted in the incoming continental crust of the African plate. There are SW-NE-aligned isolated structural highs found either buried in the accretionary wedge or having morphological expression as seamounts in the western

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Fig. 3. Seismic section in the Ionian Basin (modified after Cernobori et al. 1996). The stratigraphic interpretation follows regional correlation with main seismic boundaries tied to known wells. The definition of the oceanic basement is correlated with the results of expanded spread seismic profiles (DeVoogd et al. 1992). The location of the section is indicated by a bold asterisk in Figure 2. Similar seismic images of the Ionian foredeep have been obtained by Polonia et al. (2002). foredeep and foreland (Della Vedova & Pellis 1989; Von Huene et al. 1997; Hieke & Dehghani 1999; see Fig. 3). In the northeastern Ionian basin, plate convergence is controlled by the fast rates of N W - S E extension of the Tyrrhenian Sea and associated southeastward migration of the Calabrian-Peloritan arc. The directions of motion of the Africa plate and the Calabrian-Peloritan arc therefore differ by about 90 ~ Deformation is developed offshore as a series of imbricated thrusts involving the entire sedimentary section (Finetti 1976; Cernobori et al. 1996) with detachment cutting into progressively deeper stratigraphic levels, from the Messinian salts at the outer deformation front to Mesozoic rocks and the crystalline basement in the continental blocks of the Calabrian arc in southern Italy (Tortorici 1982; Tortorici et al. 1995). Similarly to the case of the Mediterranean Ridge accretionary complex, it can be inferred that the introduction of a ductile salt layer in the deforming sequence has modified drastically the rheology of the arc. Strain partitioning induced by obliquity of subduction must exist only in the pre-Messinan deforming domain. The viscous behaviour of the detachment surface has produced thrusting trending normal to the shortening direction (thus the high curvature of the deformation front), at much higher rates of outward propagation (Costa & Vendeville 2002; Costa et al. 2004).

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z. BEN-AVRAHAMETAL.

Levantine Basin Regional setting. The present-day southeastern Mediterranean Sea,

also known as the Levantine Basin (Fig. 4), is a remnant of a larger, Neotethyan oceanic basin that opened between several fragments of the Pangaea supercontinent in Early Mesozoic times (Dewey et al. 1973; Bein & Gvirtzman 1977; Garfunkel & Derin 1984; Robertson & Dixon 1984; Garfunkel 1998). During the M i d - L a t e Cretaceous, as the basin started to close, its northern, Eurasian part underwent a dramatic transition from a passive to an active margin, and the basin was subsequently largely subducted or accreted at the present-day areas of Cyprus, southern Turkey, northern Syria and Iran (Ben-Avraham 1989; Robertson 1998). Closing of the northern Levantine Basin continued through the Cenozoic; the active convergent front is at present located at the area of the Cyprian arc (Kempler & Ben-Avraham 1987; Ben-Avraham et al. 1995; Fig. 4).

Fig. 4. Tectonic map of the Eastern Mediterranean area (modifiedafter Garfunkel (1998) and Robertson (1998), and incorporating results from Ben-Avraham et al. (1995) and Woodside et al. (2002). The Levantine Basin is a remnant of a larger, Early Mesozoic (Neotethyan) oceanic basin that was largely consumed during Late Mesozoic-Cenozoic convergence of the African-Arabian and Eurasian plates. Subduction and wrenchingis taking place at the tectonically active, northern part of the basin (Cyprian arc), whereas only minor thrusting and contraction (Syrian arc folds) is recorded on the southeastern, Levantine margin. The Eratosthenes Seamount is a large continental block that was stranded within the basin and is now approaching its northern margin. The Carmel fault (CF) is an active tectonic lineament that separates two crustal types, and probably originated during the Early Mesozoic rifting phase.

The structure of the southern Levant margin and the adjacent deep marine basin is relatively simple and preserves the signature of the main tectonic events associated with opening and closing of the Neotethyan ocean. These can be separated into several distinct tectonostratigraphic phases: (1) Triassic to Early Jurassic continental break-up and rifting; (2) Mid-Jurassic to MidCretaceous subsidence and formation of passive continental margin; (3) M i d - L a t e Cretaceous to Early Tertiary large-scale convergence and contraction; (4) Late Tertiary to Quaternary minor convergence and subsidence. Triassic - E a r l y Jurassic rifting phase. During the Palaeozoic the area of the Levantine Basin was located south of the Palaeotethys Ocean on the northern edge of the Gondwana continental platform. Most workers agree that a break-up of the northern part of Gondwana into several microplates occurred in the latest Palaeozoic to Early Mesozoic (Bein & Gvirtzman 1977; Hirsch et al. 1995; Garfunkel 1998). Continental break-up processes are indicated by two types of observations: (1) rift-related phenomena identified in wells and multi-channel, seismic reflection profiles; (2) variation in crustal thickness and composition interpreted from deep seismic refraction profiles, gravity and magnetic data. An extensive graben and horst system is recognized in the deeper stratigraphic level of central Israel and the Levant margin. Part of this system is observed in multi-channel, seismic reflection profiles (Fig. 5; Gardosh 2002; Gardosh & Druckman 2006). Early Mesozoic structural highs and lows are further interpreted from thickness variations in the Permian to Lower Jurassic strata, identified in seismic and well data (Garfunkel & Derin 1984; Bruner 1991; Druckrnan et al. 1995; Garfunkel 1998). The system is composed of fault lines several tens of kilometres long that are generally oriented N E - S W , roughly perpendicular to the present-day coastline of Israel (Fig. 4). An early tectonic pulse of this fault system is indicated by the northward thickening trend of the Permian to Lower Triassic section, identified in wells at southern and central Israel (Garfunkel & Derin 1984). The continuation of faulting activity is evident from the occurrence of a 350 m thick, polymictic and angular, Middle Triassic carbonate breccia discovered in wells near the southern coastal plain of Israel (Druckrnan 1984). A conspicuous graben, filled with Carnian dolomite and gypsum, was identified in outcrops and wells in central and southern Israel (Druckman 1974; Garfunkel & Derin 1984). Extensive series of the Lower Liassic Asher volcanic rocks, found in the subsurface of northern Israel, were partly accumulated within a 2.5 km deep trough (Gvirtzman & Steinitz 1983; Garfunkel 1989). Abrupt thickness changes are not recognized in Bathonian and younger strata of the Levant margin, suggesting the cession of activity of the Neotethyan graben and horst system (Garfunkel 1998). The upper age limit of the faulting is further indicated in some of the structures where Triassic and older tilted beds are overlain by horizontal Lower to Middle Jurassic strata (Gardosh 2002; Gardosh & Druckman 2006; Fig. 5). The Early Mesozoic rifting and extension resulted in profound changes in crustal composition and thickness. Long-range seismic refraction profiles, gravity and magnetic data show that a major transition between two types of crust takes place several tens of kilometres NW of the present-day coastline of Israel (Ben-Avraham et al. 2002). A 30-35 km thick crust of continental character underlies the area of southern and central Israel. The SE Mediterranean Sea is underlain by a 10 km thick crust of an oceanic character (Ginzburg et al. 1979; Ginzburg & Folkman 1980; Makris et al. 1983; Ginzburg & Ben-Avraham 1987; Ben-Avraham et al. 2002). A 2 5 k i n thick crust of continental character underlies the Eratosthenes Seamount, located south of Cyprus (Fig. 4; Makris et al. 1983; Ben-Avraham et aL 2002). These observations indicate that a major part of Early Mesozoic rifting and extension took place west of the present-day coastline. This activity resulted

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Fig, 5. Multi-channel, migrated, marine seismic reflection profile across the Levant margin showing the main stratigraphic units and faults. The line was shot in 1983, offshore Israel, using seven Bolt air guns (2180 cubic inches) and was recorded with 120 channels on a 2975 m cable. The interpreted seismic horizons are regional seismic markers, most of which are correlated to offshore wells: I, top of Messinian evaporites; II, base of Messinian evaporites; III, Mid-Tertiary unconformity; IV, Upper Cretaceous; V, Mid-Cretaceous unconformity; VI, near top Lower Jurassic, VII, near top Basement. The four major tectonostratigraphic stages identified on the southeastern margin of the Levantine Basin are: Triassic to Early Jurassic continental break-up and rifting (horizons VII-VI); Mid-Jurassic to Mid-Cretaceous subsidence and formation of passive continental margin (horizons VI-V); Mid-Late Cretaceous to Early Tertiary large-scale convergence and contraction (horizons V-III); Late Tertiary to Quaternary minor convergence and subsidence (horizon III-sea bottom). Location of section is shown in Figure 4. YW1, Yam West-1 well.

in the separation of the Tauride block of southern Turkey from the African continental mass (Smith 1971; Ben-Avraham 1989; Robertson et al. 1991, 1996; Garfunkel 1998). Additional tiffing, extension, creation of oceanic crust and development of deep marine basins took place when the Eratosthenes block separated from the African continent, resulting in the creation of 2 0 0 - 3 0 0 km wide deep marine basin in the southern Levant area (Ben-Avraham 1989; Garfunkel 1998). The maximum size of the palaeo-Levantine Basin, extending south of the Tauride block and north of the African-Arabian continental mass, is hard to estimate, as much of its northern part was later consumed. Garfunkel (1998) estimated that it was at least twice its present width. Another pronounced transition of crustal properties is identified across the Carmel fault on the Levant margin (Fig. 4). The 35 km thick continental crust of central Israel thins to about 20 km in the Galilee area, NW of the fault (Ginzburg & Folkman 1980). The Carmel fault is an active tectonic lineament; its origin is associated with Palaeozoic-Mesozoic plate motion (Ben-Avraham & Ginzburg 1990; Ben-Gai & Ben-Avraham 1995). The relationship between the fault and the rifting processes that led to the formation of the Levantine Basin are not clear. Mid-Jurassic-Mid-Cretaceous continental margin phase. During the later part of the Mesozoic, subsidence and the development of passive continental margins dominated the Levant region (Bein & Gvirtzman 1977). The M i d - U p p e r Jurassic strata of the southern margin show no evidence of large-scale faulting and magmatic activity, indicating that the opening motion in the Levantine Basin was considerably reduced. A shallow marine shelf was developed near the present-day coastline, whereas a deeper marine basin prevailed throughout the SE Mediterranean Sea. The Levant margin evolved through the accumulation of various types of low- and high-order depositional sequences; their geometry and stratal pattern reflect the combined effects of eustatic sealevel cycles, local subsidence and uplift, change of environmental conditions and rate of sediment supply (Flexer et al. 1986; Gardosh 2002). The Lower to Middle Jurassic depositional cycles are characterized by rapid growth and aggradation of carbonate margin and relatively minor bypass into the basin, reflecting fast thermal subsidence coupled with long-term eustatic rise. The uppermost Jurassic to Lower Cretaceous depositional cycles are dominated by intense transport of siliciclastic strata

and accumulation of deep-water turbidites in the basin and margin. These are associated with tectonic uplift and erosion SE of the basin (Gvirtzman et al. 1998) coupled with eustatic sealevel drop. Finally, the Middle Cretaceous depositional cycles are characterized first by the progradation of carbonate slopes into the basin, followed by intense growth of carbonate platforms on the margin and restricted bypass into the basin. These were the results of tectonic quiescence, followed by marked eustatic rise during the Cenomanian-Turonian (Gardosh 2002). Late Cretaceous-Early Tertiary main convergence phase. Closing of the Levantine Basin started during the late Mid-Cretaceous and continued through the Late Cretaceous and Early Cenozoic. The relative motion of the African-Arabian plate towards the Eurasian plate resulted in the development of several subduction and collision zones in the northern part of the basin, in the present-day areas of the Cyprian arc and the Taurus Mountain belt (Fig. 4; Ben-Avraham & Nur 1986; Ben-Avraham 1989; Robertson et al. 1991). The most conspicuous tectonic element associated with this motion is a series of contractional structures found throughout the Levant region, termed the Syrian arc folds (Krenkel 1924; Fig. 4). Well and seismic data from the inland part of Israel show that reverse faults are found in the core of many of the Syrian arc folds. These structures are often superimposed on the older graben and horst system, and their formation is controlled by the reactivation, in a reverse direction, of Early Mesozoic normal faults (Freund et al. 1975; Druckman et al. 1995). Seismic data from the Israeli offshore reveal more details on the relation between the Levant margin and the contractional deformation. A belt of dense Syrian arc type folds and reverse, upthrust faults is observed at the southeastern edge of the basin (Figs 4 and 5). The folded zone terminates some 5 0 - 7 0 km west of the present-day coastline. The entire Precambrian to Mesozoic sequence in the offshore fold belt is uplifted, and the lowered area to the west contains a several kilometres thick sedimentary section of assumed Late Cretaceous to Early Tertiary age (Fig. 5). The western edge of the fold belt coincides with a zone of transition between two types of crust identified by Ben-Avraham et al. (2002). Based on the geophysical evidence it is suggested that the western edge of the fold zone marks an area of convergence and possibly thrusting of an oceanic or transitional type crust under continental crust on the east. Ben-Avraham & Nur (1986) have

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suggested a similar process along the continental margin of the Sinai Peninsula. A deep marine flexural foredeep developed west of the thrust front where large amounts of detrital carbonate strata were accumulated (Figs 4 and 5). Late Tertiary-Pleistocene minor convergence phase. Contractional

deformation in the southern Levant became rather limited during the last tectonostratigraphic phase (Fig. 5), indicating that the rate of convergence was reduced. A possible cause for this reduction is the emergence of a new plate boundary that developed east of the Levant margin contemporaneously with the opening in the Red Sea (Freund et al. 1970). Part of the northward motion of the African plate was probably taken up by this new boundary, the Dead Sea transform zone (Fig. 4). The activity along the Dead Sea plate boundary during the Miocene-Pliocene was also associated with considerable uplift and erosion of the transform shoulders (Garfunkel 1989). This was followed by the development of vast drainage systems that carried large amounts of detrital material westward and northward across the Levant margin. Three distinct depositional cycles are recognized during the Late Tertiary-Pleistocene tectonic stage (Fig. 5). The lower, Oligocene-Miocene cycle is characterized by intense transport of siliciclastic strata into the basin. The middle, uppermost Miocene cycle is associated with the accumulation of thick evaporitic section within the basin. The Messinian salinity event is well recognized throughout the Eastern Mediterranean. The deposition of the Messinian salt within Mediterranean basins is considered to be the result of a dramatic sea-level drop, in the range of 800-1300 m (Ben-Gai et al. 2005), as a result of the isolation of the Mediterranean Sea from the Atlantic Ocean (Hsu et al. 1973a,b; Ryan 1978). Quantitative basin analysis of the Levant margin suggests the existence of a deep basin, similar to the present one, in pre-Messinian time (Tibor et al. 1992). A final Plio-Pleistocene cycle is characterized by the development of conspicuous, prograding delta systems along the western margin as well as the formation of the Nile river delta further to the south (Figs 4 and 5). These systems were associated with anomalously high sedimentation and subsidence rates, influenced by flexural response of the lithosphere to the loading of the Messinian salt and Nile-derived sediments as well as to uplift of the Judea Mountains east of the margin (Tibor et al. 1992).

The present-day active northern edge of the Levantine Basin is located along the Cyprian arc in the NE Mediterranean (Fig. 5). The Cyprian arc is divided into three distinct segments (Kempler & Ben-Avraham 1978; Ben-Avraham et al. 1995). In the western segment subduction of the the African lithosphere under the Turkish plate is assumed to have led to the creation of the small Antalya Basin (Fig. 4). In the central segment subduction was interrupted, as a result of the presence of the Eratosthenes continental block in front of the arc. The eastern segment is a system of wrench faults dominated by shear motion, with no active subduction (Ben-Avraham et al. 1995). The small Latakia and Larnaca basins (see below) were formed in this segment of the arc (Fig. 4).

Younger basins The Strait o f Sicily rift s y s t e m

The morphological and structural evidence of elongated bathymetric lows in the Strait of Sicily (Pantelleria, Malta and Linosa troughs; Figs 1 and 6) were mapped during the early exploration of the Mediterranean Sea (Finetti & Morelli 1972, 1973). The interpretation of these troughs as rift-related structures was proposed by various workers (lilies 1981; Finetti 1984), and some of them have emphasized the role played by transcurrent tectonics in their development (Cello et al. 1984; Jongsma et al. 1985; Reuther & Eisbacher 1985; Boccaletti et al. 1987; Cello 1987). The depressions found within the Pelagian block are viewed in general as pull-apart transtensional basins generated in a dextral wrench system. The proposed interpretations, mainly based on structural analyses carried out in the islands within the Strait of Sicily or in restricted parts of the surrounding areas, present some differences, mostly related to the poorly constrained stretching mechanisms and deformational history. The morphostructural features present in the Strait of Sicily do not have an equivalent in the central and eastern regions of the Mediterranean Sea, where major changes in crustal nature, structural trends and tectonic styles occur within short distances. In the Strait of Sicily, a prevalent extensional regime dominated from late Messinian to early Pliocene time, as determined from the analyses of the seismic sequences found in the depocentral

Fig. 6. Bathymetricmap of the Strait of Sicily obtained by combiningunder-way soundings and depth-converted sea-floor reflectors from a seismic grid. The parts of seismic profiles MS-19 and MS-120 that are presented in Figures 7 and 8, respectively, are indicated by bold continuous lines. The main troughs associated with the rift system (Pantelleria, Malta and Linosa) are indicated by grey shading.

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269

areas of the troughs, and was responsible for the development of the three main depressions within the Pelagian block. Regional structural setting. The Strait of Sicily rift system is located within the Pelagian block (Burollet et al. 1978; Boccaletti et al. 1984), which geologically corresponds to the northern

leading edge of the African plate. The stretched continental crust thins to less than 20 km (Colombi et al. 1973) underneath the Pantelleria trough. The sedimentary cover is made up of MesoCenozoic carbonate sequences that crop out in both the northern (southeastern Sicily) and southem (Tunisia and western Libya) domains of the platform. To the east, the shallow Pelagian block is separated from the deep Ionian Basin by the Malta escarpment, a structural boundary where deep-seated faults have throws exceeding 2000 m. The escarpment extends over a length of about 300 km from the eastern coast of Sicily southwards to the Medina seamounts with a steep slope that descends to more than 3000 m below sea level. It separates the 23 km thick continental crust (Finetti 1984) of the Strait of Sicily (Hyblean-Malta Plateau) from the 13 km thick crust beneath the Ionian Sea (Ferrucci et al. 1991). On land, normal and strike-slip faults formed by left-lateral movements exist on the Ionian side of the Hyblean Plateau, where they represent second-order structures superimposed on the dominant vertical displacement along the Malta escarpment. Sinistral wrenching along NNW-SSE-oriented fault planes is accommodated by extension along the N E - S W oriented faults. The contrasting structural behaviour of the Hyblean Plateau and the Ionian crust to the east could account for the kinematics of the Plio-Pleistocene structures along the onshore extension of the Malta escarpment (Grasso 1993). Three subparallel principal troughs (Fig. 6), roughly trending Nll0~ and deeper than 1000 m, form the Strait of Sicily rift system. The northwestern basin of the rift system (Pantelleria trough) is the widest depression, and is separated from the other two troughs (Malta and Linosa troughs) by localized structural highs at around longitude 12~ However, the geometric relationships between these depressions and their relative boundaries are not yet precisely identified because the lack of detailed bathymetric information. These depressions are filled with turbiditic Plio-Pleistocene deposits (Maldonado & Stanley 1977; Biju Duval et al. 1985), reaching thicknesses of about 1000 m in the Pantelleria trough, 2000 m in the Linosa trough, and 1500 in the Malta trough (Winnock 1981). In the other parts of the Pelagian block, the Plio-Pleistocene sequence has a maximum thickness of about 500 m. Geometry of the rift system. The basic information from which we derive the morphostructural and sedimentary setting of the Strait of Sicily region is the bathymetric map and the grid of seismic reflection profiles collected in the region since the early 1970s, calibrated with boreholes and well data. Seismic profiles provide the most striking evidence for analysis of the structural configuration and geometric disposition of the fault-related structures associated with stretching. After the earlier work of Finetti & Morelli (1972, 1973), a large amount of data have been collected in the Pelagian block, mainly in the course of a series of oil exploration surveys. Here we analyse the reprocessed, migrated version of seismic line MS-19 (Fig. 7), originally presented by Finetti (1982). This profile illustrates in detail the structural architecture of a continent-type crust affected by extensional tectonics. It runs N N E - S S W from the offshore part of the Hyblean Plateau, which constitutes the northern sector of the African foreland in the Pelagian block, to the Lampedusa Plateau, a generally flat-lying carbonate platform that borders to the SE the stretched region of the Strait of Sicily. The seismic profile, located between the two horsts constituting the islands of Malta and Lampedusa, crosses the eastern part of the Malta trough and the central-eastern Linosa trough. The stratigraphy along the seismic profile has been derived from correlation with adjacent well logs and

Fig. 7. Above: part of a reprocessed seismic line MS-19, originallypresented by Finetti (1984). This part of the profile crosses the central-eastern portion of the Linosa trough (see location in Fig. 6), and shows the geometry of the rift-related structures of the Pelagian block. The subvertical faults that progressively deepen the trough symmetrically should be noted. TWT, two-way travel time. Below: line drawing interpretation in which only the principal subvertical faults are indicated. (For the seismostratigraphic control along the profile, see Finetti (1984).) boreholes, and extrapolating the information across the entire grid of data (Finetti 1984; Pedley et al. 1993). Here we utilize the same interpretation in terms of seismo-stratigraphic control and sedimentological character of the sequences. In the Linosa trough a set of mostly subvertical, equi-spaced faults, separating rotated and, in some cases, uplifted blocks, dominate the structural framework along seismic line MS-19 (Fig. 7). Block rotation is particularly evident in the central sector of the Linosa trough. The core of the trough, corresponding to the deepest part of the system, is bounded on both sides by prominent faults with opposite polarities. All the subvertical faults reach the sea floor, indicating recent tectonic activity. Towards the SW, within the flat Lampedusa Plateau, another significant graben (the Lampedusa trough), composed of at least two tilted blocks, marks the southern boundary of the rift zone within the Pelagian block. The Malta trough is imaged by the profile MS-120 presented in Figure 8. The two basin shoulders are symmetrical, and possibly are composed of single normal faults. However, the strong erosion and the possible presence of localized slide structures do not allow clear identification of subsidiary discontinuities on the hanging walls. The total throw along these flanking faults is of the order of 1000 m or more. Secondary normal faults are visible mainly on the southern flank of the trough. Significant uplift occurs on the southern shoulder of the graben. The seismic units filling the Malta trough onlap the basin margins. These units may represent post-Messinian sequences, as identified by Ryan (1978), and are covered by Plio-Quaternary strata that are

270

Z. BEN-AVRAHAMETAL. by many workers, are detectable only in local, very detailed seismic surveys, such as in the western offshore of the Maltese islands (Gardiner et al. 1995). Analogue sandbox experiments in oblique rift models (where there is an angle between the rift axis and the extension direction) show remarkable similarities to the fault architecture and geometric disposition of graben structures found in the Strait of Sicily. In particular, the along-strike offsets in depocentres and the en echelon fault pattern parallel to the zone of rifting are the most striking evidence. En echelon stepping and segmentation of the axial depocentre is interpreted to occur across accommodation zones formed by complex interfingering extensional fault systems, as seen in natural examples such as the Central Graben of the North Sea (Roberts et al. 1990). Conceptual models based upon the analogue experiments show footwall uplift on the individual faults and mantle upwelling below the rift zone. Both these features are seen in the Strait of Sicily region. In SE mainland Sicily the Plio-Pleistocene fault pattern observed along the western margin of the Hyblean Plateau implies rightlateral movements along a broad NNE-oriented fault system that traverses the southern Sicilian foreland oblique to the front of the Maghrebian arc. Grasso et al. (1990) have argued that this foreland strike-slip system played the role of a transform fault linking zones of modern rifting within the Strait of Sicily with zones of recent underthrusting in south-central Sicily.

Small basins along the Cyprian a n d Hellenic arcs

Fig. 8. Above: near-trace monitor of part of seismic profileMS-120 (see location in Fig. 6), where it crosses the central-western part of the Malta trough. The basin is flankedby prominent subverticalfaults, and the basin fill is characterizedby mostly horizontal reflectors onlapping the basin margins. Significantuplift of both of the hanging walls of the depression can be seen on the profile. Below: line, drawing interpretation. affected by structural discontinuities and possibly by block rotation underneath the sedimentary cover. Some evidence of asymmetry within the basin fill is evident from the section in association with recent contractional deformation, as noted by Argnani (1990) on the basis of sparker profiles acquired within both the Malta and Linosa troughs. Rift model. Intense block rotation and tilting generated by differen-

tial motion corresponding to the subvertical faults are the most notable tectonic features characterizing the entire rift system. The polarity of the block rotation varies along the rift system and the most pronounced deformation is found within the axial areas of the troughs, where block subsidence and tilting is particularly impressive. In general, the tectonic style that characterizes the Strait of Sicily rift system is remarkably symmetrical, a mode of extension that has been described kinematically as pure shear (McKenzie 1978), with an upper brittle layer overlying a ductile lower layer, producing a symmetrical lithospheric crosssection. The model predicts the formation of sediment-filled grabens, which causes isostatic disequilibrium and the compensatory rise of the asthenosphere, eventually accompanied by surface volcanism in a late evolutionary stage of the rift system. The origin of possible asymmetries within the rift system is generally ascribed to inherited inhomogeneities in the lithosphere or/and in the uppermost crustal layers. The structural configuration of the Strait of Sicily resembles most, if not all of the tectonic elements predicted by the model. In particular, seismic data have shown that the normal fault pattern dominates, and appears to have controlled the evolution of the trough within the Pelagian block. This suggests a mechanism in which the stress field is largely extensional, and appears to have acted mostly as dip-slip faults. Dextral strike-slip mechanisms along N W - S E - or east-west-trending faults, as proposed

The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea can be divided into two groups, those containing Messinian evaporites to the east and those without to the west. In the following, we describe briefly the various small basins in the same sense, from east to west. A number of small basins are found both on- and offshore (Fig. 1) in the northeastern corner of the Mediterranean. They started to form at roughly the same time, probably at least from the early Miocene, but have since been tectonically separated to some degree. On the basis of the presence at depth of shallow marine and continental deposits overlying Messinian evaporates, the Adana and Iskenderun basins may have subsided at least 3 km in the Pliocene-Quaternary; however, the Mut Basin, to the north of the Cilicia-Adana Basin, contains at least 1500 m of almost horizontal Neogene sediments at a present elevation of 1500-2000 m, a vertical difference between the basins of almost 5 km. This is not unusual in the Eastern Mediterranean. The Cilicia Basin. The Neogene Cilicia Basin (Fig. 4), a roughly 3 km deep N E - S W trough (shallower to the SW) lying between Cyprus and Turkey, is a seaward continuation of the Adana Basin, to the NE in Turkey (but offset to the south by the Ecemis Fault, bounding the western part of the Adana Basin on land). It contains Messinian evaporites, which show diapirism and lateral flowage away from the northern and northeastern depocentres (Evans et al. 1978; Aksu et al. 1992a,b). Asymmetry in the sedimentation has developed as a result of the major sources being to the north and northward tilting of the northern Cyprus margin. The two most important structural boundaries to the basin are the transpressional Kyrenia-Misis Ridge to the SE and the A n a m u r Komakiti Ridge, which separates the Cilicia Basin from the deeper Antalya Basin to the west. The Latakia and Larnaca basins. These are relatively shallow Miocene basins lying to the east of Cyprus on steps formed by the development of several sinistral wrench zones along the boundary of the African and Anatolian plates (Ben-Avraham et al. 1995; Vidal & Alvarez-Marrrn 2000; Vidal et al. 2000). They are formed on northward tilting basement rocks that are continuous from Cyprus to Syria and Turkey. Post-Miocene tectonics

EASTERN MEDITERRANEAN BASIN SYSTEMS

resulting from the extrusion of the Anatolian plate to the west (with Cyprus forming the southernmost part of this, pushing more southward) caused both the tilting and uplift of the basins. Uplift has resulted in the basins forming two steps from the Levantine Basin towards the Cicilia Basin to the north: the Latakia Basin forms the southern step and the Larnaca the second step just to the north (Fig. 4). The fault-bounded basins are separated by the Larnaca Ridge, and limited to the north and south by the Kyrenia-Misis Ridge and the Latakia Ridge, respectively (Ben-Avraham et al. 1995).

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The Antalya Basin. Relatively little is known about this deep basin

lying west of the Cilicia Basin between Turkey and the Florence Rise (Figs 4 and 9). It was shown to contain thick Messinian evaporites and display northward thickening sediments suggestive of active and long-term northward tilting (Woodside 1977; Taviani & Rossi 1989; Sage & Letouzey 1990). A Palaeogene and Mesozoic basement is proposed from the relationship of the basin to Alpine nappe structures inferred at the Florence Rise (Sage & Letouzey 1990; Woodside et al. 2002). The ophiolites that have been shown to continue from SE Turkey (Baer-Bassit) into Cyprus (Delaloye & Wagner 1984; Delaune-Mayere 1984; Robinson & Malpas 1990) and to reappear again in the Antalya Nappes Complex (e.g. Robertson & Woodcock 1982) of southwestern Turkey have been thought also to continue through the Florence Rise or the Antalya Basin (e.g. Monod 1976), but there is little evidence for this in the form of typical magnetic or gravity anomalies (Woodside 1977; Woodside et al. 2002). A major roughly north-south structural discontinuity separates Cyprus and the Cilicia Basin to the north from the Antalya Basin, which has substantially greater depth than the western Cilicia Basin. A gravity discontinuity (low to the west, higher to the east) defines this boundary (Woodside 1976). The western boundary of the Antalya Basin is the fault-bounded Anaximander Mountains (Zitter et al. 2003; Ten Veen et al. 2004), and the fault-bounded western margin of the Gulf of Antalya to the north. Within the Gulf of Antalya to the north, there appears to be a buried, abrupt N W - S E boundary (defined also by a change in the gravity field; Woodside 1976) of the deepest part of the basin, with a shallower part to the north. The shallow part is likely to be the offshore extension of the Manavgat Basin, which is located onshore along the eastern margin of the Gulf of Antalya (e.g. Necker et al. 1998; Flecker & Ellam 1999). The Finike Basin. The Finike Basin is a narrow (about 20 km) depression about 80km long lying between southwestern Turkey to the north and the Anaximander Mountains to the south. It appears at first glance to be a continuation of the Antalya Basin to the west, but this is only a morphological continuity. An absence of Messinian evaporites and a relatively thin post-Miocene sedimentary fill (no more than about 1200m, assuming a seismic velocity of 1700 m s -1) indicate that this basin is geologically relatively young, probably no older than late Pliocene to Present. Northward-tilting sediments indicate that it is still forming, with the appearance of a rift basin being created by listric faulting along its northern boundary. This narrow and very linear (WSW-ENE) section of the margin of southwestern Turkey is inferred to mark the upper part of the listric fault. The fault may have originated as a strike-slip fault connected through the Rhodes Basin with transpressional faulting through the Pliny and Strabo trenches to the west (Ten Veen et al. 2004). Parallel faulting is mapped on land in Turkey (Gutnic et al. 1979). Basement rocks in the western part of the basin are shown to be similar to the Suzug Dag and Bey Daglari section in Turkey (Woodside et al. 1997), and in the east they form a continuation of the Antalya Nappes Complex, which is traced as far south as the Anaximander Mountains at 35~ about 60 km south of the Turkish coast (Woodside et al. 1997). Thus the Finike Basin acts as a rift basin separating the Anaximander Mountains from

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Turkey (Nesteroff et al. 1977; Woodside et al. 1997). To some degree, sediments from the Finike Basin can be followed southward onto the Anaximander Mountains that form the basement below an erosional unconformity. The Rhodes Basin. The Rhodes Basin shows many similarities with the Finike Basin, having no Messinian evaporites, and a relatively thin post-Miocene section (no more than about 1000 m, assuming a seismic velocity of 1700 m s -1) overlying inferred Alpine basement (Woodside et al. 2000). It is also, at about 4500 m depth (Fig. 10), one of the deepest basins of the Mediterranean Sea, which implies a relatively recent and rapid subsidence. It is bounded to the west and NE by faults (Woodside et al. 2000). The western boundary fault shows transpression with reverse faulting (Woodside et al. 2000; Kontogianni et al. 2002) and was thought to form the plate boundary between the Aegean and African plates, in continuity with sinistral transpressive shear along the Pliny Trench. Because of the poorly constrained nature of the Fethiye-Burdur fault zone to the NE, and a number of normal fault-plane solutions for earthquakes along it (rather than strike-slip; McKenzie 1978), the idea that the plate boundary continues in this direction has been disputed by Ten Veen et al. (2004), who have suggested instead that the plate boundary crosses the Rhodes Basin and continues along the northern edge of the Finike Basin. The Rhodes Basin is separated into two sub-basins by a zone of deformation that could be seen as the link between the Finike Basin and the Strabo Trench (Ten Veen et al. 2004). The Aegean basins. Within the Aegean Sea are a number of generally shallow basins that formed by extension of Alpine (Hellenide) basement during the southwestward extrusion of the Aegean and Anatolian plates (McKenzie 1972; Mascle & Martin 1990; Jackson 1994). The two most important of these basins are the Crete Basin (over 2000 m deep in the east), just to the north of Crete and parallel to its west-east long axis (Angelier et al. 1982), and the North Aegean Trough (over 1000 m deep in two connected E N E - W S W linear depressions, the Saros Trough to the east and the Sporades Trough to the west; Papanikolaou et al. 2002) which formed probably as a pull-apart along the principal strand of the North Anatolian Fault zone at its westernmost end (Yaltlrak et al. 1998; Yaltlrak & Alpar 2002). The North Anatolian Fault zone, the key tectonic element in the development not only of the Sea of Marmara but also of the North Aegean

Fig. 10. Six-channelseismic profile (using two GI 75 guns as source; with total air chamber volume of 2.451) from the French PRISMED II expedition (line 24, modifiedafter Woodside et al. 2000) running roughly north-south in the southern Rhodes Basin. It should be noted that there are no Messinianevaporites overlying the pre late-Miocene eroded basement, and that the post-Miocene sediments are relatively thin and undeformed.

Trough, abuts the Greek mainland at the western end of the Sporades Basin; however, the motion is absorbed by a system of extension with rapid clockwise motion, which, within the past 1 2 Ma, has passed motion on to the Gulf of Corinth and the Cephalonia Transform Fault to the west (Le Pichon et al. 2002). Thus an age of about 2 Ma can be given to the tectonic regime now prevailing in the Aegean basins (Le Pichon et al. 2002), although the beginnings of extension date back to as early as 25 Ma (Jackson 1994). Between the Cretan and the North Aegean Troughs lie a number of smaller basins with similar structural trends. Included in these are, from north to south, the Edremit Trough, the North IkariaSamos Basin and the North Mikonos-Andros Basins, and the South Ikarian Basin. These are mainly grouped in the eastern half of the Aegean (e.g. Lykousis et al. 1995). To the west are small Plio-Quaternary basins such as the Saronikos Basin (following roughly the Gulf of Corinth structural trend to the east of the Peloponnesus), the Mirthes Basin, and the Argolide Basin. The differing structural trends between west and east are related to the southward migration of the Hellenic Arc and probably rollback in the back-arc, and have been modelled by Jackson (1994), among others (e.g. Kreemer et al. 2003), as well as imaged by seismic tomography (Spakman et al. 1988).

Discussion The Eastern Mediterranean basin systems were formed in several stages. During the first stage of evolution, continental fragments rifted away from Africa to form the Ionian and the Levantine basins. This stage was followed by other rifting events. In the Levantine Basin, during the first stage, continental fragments, now part of Turkey, rifted away from the Levant and Sinai, and in the second stage, the Erathostenes Seamount and possibly other microcontinental blocks rifted away from Africa and moved northward. This process has caused the formation and destruction of oceanic crusts in the basins. The large-scale process that dominates the evolution of the Eastern Mediterranean basin systems is the approach of two large lithospheric plates, the African plate and the Eurasian plate, toward each other (Ben-Avraham 1989). The sea-floor spreading process in the Ionian and Levantine basins was interrupted occasionally by the collision of the rifted fragments with the southern margin of the Eurasian plate in the north. The rifting of continental fragments away from Africa, while the African plate was moving northward relative to the Eurasian plate, means that subduction along the Calabrian, Hellenic and Cyprian arcs had to be faster than the convergence of the two plates. Le Pichon et al. (1982) suggested that because of the large slab-pull force, the subduction of land-locked deep-sea basins will, in general, occur much faster than the collision rate. As a consequence, subduction of the old deep-sea basins in the Eastern Mediterranean will be compensated by the formation of young deep-sea basins behind the subduction zones. This is, in fact, what is taking place in the Eastern Mediterranean. Large basins, such as the Tyrrhenian and Aegean, as well as small basins, such as Cilicia and Antalya, are opening behind the Calabrian, Hellenic and Cyprian arcs. The mechanism responsible for the origin of the rift system in the Strait of Sicily is a crucial point that needs to be addressed. Very few workers have attempted to unravel this question. The difficulty of understanding the genetic evolution of the area is attributed to the inherent geological complexity of the region, which is surrounded by an assemblage of relatively small lithospheric blocks with a wide variety of rheologies and thicknesses that evolved during the pre-Tertiary tectonic evolution of the Africa passive palaeo-margin. The geological nature of these crustal segments is still poorly known because of the lack of distinctive geological constraints, and, in particular, the poor seismic coverage and stratigraphic information on both the

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pre-Mesozoic substratum and the crystalline basement. Ben-Avraham & Grasso (1990, 1991) stressed that one of the most important elements triggering segmentation along collision zones is probably the crustal structure variation along leading edges of the impinging Africa plate. The most important tectonic event that occurred in the central part of the Mediterranean Sea in the latest Tertiary was the rifting and opening of the Tyrrhenian basin, which started in late Miocene time and continued until the early Pleistocene with the possible formation of oceanic-type crust (Kastens et al. 1988). Because the Strait of Sicily rift system and the Tyrrhenian Sea formed in the same time span, Argnani (1990) suggested that the two geodynamic events could be in some way correlated. Rollback of the subducted slab and lithospheric mantle delamination have been proposed as feasible mechanisms that have produced a limited amount of extension within the Pelagian block, as a consequence of slab-pull forces and secondary mantle convection. On the other hand, Reuther & Eisbacher (1985) suggested that the origin of the dramatic change in the stress pattern during the Messinian might be related to an abnormal tectonic context; for example, a northeasterly subduction of the Ionian lithosphere beneath the Aegean Arc, as argued by Le Pichon et al. (1982). According to this hypothesis, crustal extension affecting the lithosphere underlying the Pelagian block occurs where it pulls away from its African anchor, giving rise to graben development and associated basaltic volcanism. Ben-Avraham et al. (1987) have considered the activity of a 1000 km long transcurrent fault running along the north African passive margin to explain the crustal extension within the Strait of Sicily. The seismic data presented here have shown that most of the faults in the Strait of Sicily affect the sea floor, indicating recent tectonic activity. No significant evidence of reactivation, inner deformation of the fault-rotated blocks and azimuth changes within the throws has been detected. Considering that the Strait of Sicily rift zone is a relatively young tectonic feature (early Pliocene to Present), we may assume that the extensional tectonic regime did not change significantly during this time span. Further support is provided by the fact that the volcanic edifices within the rift system are all Quaternary in age, indicating a progressive evolution of the rift from an immature stage to a more developed configuration in which partial melting is taking place.

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The Mesozoic-Cenozoic tectonic evolution of the Greater Caucasus A L I N E S A I N T O T t'2, M A R I E - F R A N ~ O I S E B R U N E T 3, F E D O R Y A K O V L E V 4, M I C H E L S E B R I E R 3, R A N D E L L S T E P H E N S O N l, A N D R E I E R S H O V 5, F R A N ~ O I S E C H A L O T - P R A T 6 & T O M M Y M C C A N N 7

1Netherlands Centre for Integrated Solid Earth Sciences, Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands 2present address: NGU, Leiv Eirikssons vei 39, N-7491 Trondheim, Norway (e-mail: [email protected]) 3UMR 7072 Tectonique CNRS-UPMC, Case 129, Universitd Pierre et Marie Curie, 4 place Jussieu, 75252 Paris cedex 05, France 4Institute of Physics of the Earth, RAS, Moscow, Russia 5Moscow State University, Geological Faculty, Moscow, Russia 6CRPG, 15 rue Notre Dame des Pauvres, BP 20, 54501 Vandoeuvre-Les-Nancy Cedex, France 7Geologisches Institut, Universitdt Bonn, NuBallee 8, 53115 Bonn, Germany

Abstract: The Greater Caucasus (GC) fold-and-thrust belt lies on the southern deformed edge of the Scythian Platform (SP) and results from the Cenozoic structural inversion of a deep marine Mesozoic basin in response to the northward displacement of the Transcaucasus (lying south of the GC) subsequent to the Arabia-Eurasia collision. A review of existing and newly acquired data has allowed a reconstruction of the GC history through the Mesozoic and Cenozoic eras. In Permo(?)-Triassic times, rifting developed along at least the northern part of the belt. Structural inversion of the basin occurred during the Late Triassic corresponding to the Eo-Cimmerian orogeny, documented SE of the GC and probably linked to the accretion of what are now Iranian terranes along the continental margin. Renewed development of extensional basin formation in the area of the present-day GC began in Sinemurian-Pliensbachian times with rift activity encompassing the Mid-Jurassic. Rifting led to extreme thinning of the underlying continental crust by the Aalenian and concomitant extrusion of mid-ocean ridge basalt lavas. A Bathonian unconformity is observed on both sides of the basin and may either correspond to the end of active rifting and the onset of post-rift basin development or be the record of collision further south along the former Mesozoic active margin. The post-rift phase began with deposition of Late Jurassic platform-type sediments onto the margins and a flysch-like unit in its deeper part, which has transgressed the basin during the Cretaceous and Early Cenozoic. An initial phase of shortening occurred in the Late Eocene under a NE-SW compressional stress regime. A second shortening event that began in the Mid-Miocene (Sarmatian), accompanied by significant uplift of the belt, continues at present. It is related to the final collision of Arabia with Eurasia and led to the development of the present-day south-vergent GC fold-and-thrust belt. Some north-vergent retrothrusts are present in the western GC and a few more in the eastern GC, where a fan-shaped belt can be observed. The mechanisms responsible for the large-scale structure of the belt remain a matter of debate because the deep crustal structure of the GC is not well known. Some (mainly Russian) geoscientists have argued that the GC is an inverted basin squeezed between deep (near)-vertical faults representing the boundaries between the GC and the SP to the north and the GC and the Transcaucasus to the south. Another model, supported in part by the distribution of earthquake hypocentres, proposes the existence of south-vergent thrusts flattening at depth, along which the Transcaucasus plunges beneath the GC and the SP. In this model, a thick-skinned mode of deformation prevailed in the central part of the GC whereas the western and eastern parts display the attributes of thin-skinned fold-and-thrust belts, although, in general, the two styles of deformation coexist along the belt. The present-day high elevation observed only in the central part of the belt would have resulted from the delamination of a lithospheric root.

The Greater Caucasus (GC) belt forms a morphological barrier along the southern margin of the Scythian Platform (SP; contiguous with the southern East European Platform, EEP), running from the northern margin of the eastern Black Sea Basin to the South Caspian Basin (Fig. 1). It developed during several phases of deformation in M e s o z o i c - C e n o z o i c times (Milanovsky & Khain 1963; Adamia et al. 1977, 1981; Rastsvetaev 1977; Khain 1984; Muratov et al. 1984; Gamkrelidze 1986; Dotduyev 1989; Zonenshain et al. 1990; Nikishin et al. 1998a,b, 2001). The geology of the GC has been studied for at least 150 years and a significant volume of published literature deals with its evolution, although much of this is difficult to access for the international scientific community. The GC is located in the Black S e a - C a s p i a n Sea region, which is regarded as a mosaic of terranes of Gondwanan, Tethyan and Eurasian affinity that are sometimes controversial in origin (see discussions by ~eng6r 1984; Zonenshain et al. 1990; Dercourt et al. 1993, 2000). Accretion of these blocks along the SP occurred throughout the Phanerozoic and, accordingly, orogenic events developed in the GC as such: the Palaeozoic Variscan orogeny, the Triassic-Jurassic Cimmerian orogeny, and the Cenozoic Alpine orogeny. Structural styles of the GC belt are not yet unequivocally fixed and different proposed geometries exist in

the literature, even for the major boundary faults separating fundamental tectonic units. It follows that there is still considerable disagreement regarding tectonic mechanisms, simply because there are insufficient diagnostic data. The GC orogenic events are also not well understood in terms of the driving mechanisms. There are major discrepancies concerning the rate of shortening and the nature of the M e s o z o i c - C e n o z o i c basement of the GC. Did oceanic crust and lithosphere form during this time or not? In other words, did a complete orogenic Wilson cycle from opening of an ocean to its consumption by subduction and collision of its margins take place along the GC during the M e s o z o i c - C e n o z o i c ? There is rough agreement regarding the continuing Late Cenozoic pulses of mountain building and uplift, which have resulted from collision-accretion of the Transcaucasus continental block along the southern margin of the SP (Fig. 2). At a regional plate tectonic scale, this corresponds to the final stage of the Alpine orogenic cycle involving the collision between Eurasia and Afro-Arabia 'mega'-continental plates, with the main suture zone of the Tethyan Ocean running through Anatolia and the Lesser Caucasus (Fig. 1). The aim of this paper is first to assess and present the existing data, and then to describe and reinterpret them, as necessary, as well as to present some new data to constrain better the M e s o z o i c - C e n o z o i c orogenesis of the

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 277-289. 0435-4052/06/$15.00

9 The Geological Society of London 2006.

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A. SAINTOT ETAL.

Fig. 1. Digital topographic map of the Greater Caucasus, from the global topography 2 minute database; illumination from N135. GC: from the development of basins, oceanic or not, to their inversion and/or the collision of continental blocks subsequent to the consumption of an oceanic plate by subduction.

Structure of the Greater Caucasus The GC belt comprises a basement-core containing strata as old as Proterozoic (Figs 2 and 3), with Jurassic to Eocene formations lying on its flanks. The deep crustal structure of the GC and Transcaucasus are not known, as the resolution of deep seismic sounding (DSS) lines acquired more than 20 years ago and crossing the belt is insufficient to provide an accurate crustal image. Consequently, discrepancies exist regarding the published deep structure of the belt. The two main competing models of the deep structure of the GC are described below. (1) The first model argues for a subvertical geometry of all the GC main faults, including the border faults. Somin (2000), for example, argued for a subvertical disposition of the Main Caucasian Thrust (MCT; shown in Fig. 3) at great depths, given its steep near-surface dip (65-80 ~ along all of its strike and at imaged depths to 3 - 5 kin. Nevertheless, such a steep geometry could also have resulted from deformation of the fault during the final stage of collision. Both reprocessed old and newly acquired geophysical data (Shempelev et al. 2001, 2005; Grekov et al. 2004; Prutsky et al. 2004) were used to show a similarly deep inclination of the MCT at a depth of 80km. Shempelev et al. (2001) and Rastsvetaev et al. (2004) proposed the same subvertical geometry of regional faults along profiles crossing different parts of the GC. It was concluded that the boundary of the western GC with the Black Sea is a steep (60-80 ~ and deep (80 kin) major fault (Shempelev et al. 2001) linking with the Racha-Lechkhumy Fault Zone (RLFZ; see location in Figs 3 and 4) to the east (Yakovlev 2002, 2005). (2) The alternative model, proposed by Gamkrelidze (1986), Dotduyev (1987), Giorgobiani & Zakaraya (1989), Baranov et al. (1990), Zonenshain et al. (1990) and Gustchin et al. (1996), and referenced by many others, differs strongly and favours instead flat-dipping thrusts at depth. This is a thick-skinned tectonic model for the GC (Fig. 3) with the orogen interpreted as a collage

of two or three northward underthmst slabs. These authors have argued that the SP is thrust upon the Transcaucasus continental block. The MCT sensu stricto is thus considered to be a northdipping flat thrust at depth (Fig. 3) along which the pre-Jurassic basement of the Main Range zone and its overlying Mesozoic cover were presumably displaced southwards some 100 km or more during the Cenozoic. (The MCT sensu lato comprises at least two parallel branches at the surface (see Figs 2 and 4) and the northern branch is the MCT sensu stricto, along which the crystalline basement thrusts onto the sedimentary succession (see Fig. 3). Similarly, units of the GC belt have been thrust southwards over the Transcaucasus along the RLFZ (Fig. 3). Accordingly, the GC is regarded as a large south-vergent fold-and-thrust belt with its northern limb forming a gently north-dipping monocline toward the SP. Some north-vergent thrusting could be locally present along the western and central part of the GC (Milanovsky & Khain 1963). Back-thrusting is more developed onto the Terek-Caspian foreland, where the northward propagation of the Dagestan nappes contributes to the fan-shaped structure of the eastern part of the belt (Fig. 4; see Ershov et al. 2003). Thick-skinned deformation is reported along a north-south profile cutting across the internal part of the GC, with thin-skinned deformation prevailing on the southern front in the Rioni and Kura basins. A north-south profile across the western GC (Fig. 5; Robinson et al. 1996) also shows thick-skinned deformation with imbricate structures involving the basement. It can be seen that the N W - S E and W N W ESE faults parallel to the general grain of the belt are thrust faults flattening at depth whereas the NNW-SSE faults transverse to the belt are steeper (Koronovsky 1984; Giorgobiani & Zakaraya 1989; Philip et al. 1989; Giorgobiani 2004; Fig. 4). The Fore-Caucasus region, which lies on the SP, evolved in conjunction with the GC. From Latest Eocene-Oligocene times, two flexural basins developed, separated by the elevated zone called the Stavropol High (Figs 2 and 3). This comprises a north-south elongated and anomalously thick crustal block (see Kostyuchenko et al. 2004) that from early Mesozoic times never significantly subsided. The Terek-Caspian foreland basin to its east and the Kuban foreland basin to its west developed during the Cenozoic, both showing a high subsidence rate during the Oligo-Miocene ('Maykop' facies). Thus, what is peculiar about the Fore-Caucasus area is that 'foreland type-like basins' developed in front of the more topographically subdued eastern and western parts of the belt but not in front of its topographically highest central part (see Ershov et al. 2003). Several basins also developed south of the GC belt, in the Shatsky Ridge-Transcaucasus area, in Oligo-Miocene times. These are, from west to east, the Tuapse, Rioni and Kura basins (Figs 2 and 5). They are reported to be flexural in type and related to Eocene compression (Milanovsky & Khain 1963; Gamkrelidze 1986; Nikishin et al. 1998b). However, the history of the Kura Basin is more complex, as it is the western prolongation of the South Caspian Basin (Brunet et al. 2003). A review of the Early Mesozoic tectonic evolution of the Greater Caucasus The Triassic and Jurassic history of the area is not well constrained and is still a matter of considerable debate. The extent and age of Cimmerian orogenic phases, in Late Triassic or Early Jurassic, MidJurassic and Late Jurassic times, as well as the successive rifting events, are not confidently known (see Nikishin et al. 1998a,b). E a r l y Triassic basin d e v e l o p m e n t a n d Late Triassic E o - C i m m e r i a n tectonics

Permo(?)-Triassic rifting and volcanism (and, probably, magmatism-related doming) are widespread in the Fore-Cancasus

THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS

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Volcanicplateaus (dacites,andesiticbasalts)

region and in the northern part of the GC (Nazarevich et al. 1986; Lordkipanidze et al. 1989; Tikhomirov et al. 2004). The geodynamic setting of such tectonics is still debated: was this a back-arc setting or not? Throughout the area, including the northern GC, there was also a period of Late Triassic compression (Eo-Cimmerian

Fig. 2. Simplified geological map of the Greater Caucasus and Lesser Caucasus (from Milanovsky & Khain 1963) and locations of cross-sections shown in Figures 3 and 5. PTF, Pshekish-Tyrnauz Fault; MCT, Main Caucasian Thrust; RLFZ, Racha-Lechkhumy Fault Zone. Encircled numbers: 1-6 are zones of the Scythian Platform: 1, Stavropol High; 2, Azov-Berezan High; 3, Manych Basin; 4, Kuban Basin; 5, Terek-Caspian Basin; 6, Kussar-Divitchi Basin; 7-13 are zones of the Greater Caucasus: 7, Peredovoy Zone; 8, Betcha Anticline; 9, Svanetia Anticline; 10, Laba-Malka Monocline; 11, Dagestan Folded Zone; 12, Flysch Zone of southeastern GC; 13, Flysch Zone of north-western GC; 14-24 are zones of the Lesser Caucasus: 14, Somketo-Karabakh Zone; 15, Artvin-Bolnisi Zone; 16, Adzharo-Trialet; 17, Talesh; 18, Sevan-Akera; 19, Kafan; 20, Vedin; 21, Zangezur; 22, Mishkhan-Zangezur Massif; 23, Ararat-Djulfa Massif; 24, Araks Basin; 25-29 are intramontane zones of the Transcaucasus and Black Sea: 25, Rioni Basin; 26, Kura Basin; 27, Dzirula Massif; 28, Tuapse Basin; 29, Shatsky Ridge.

tectonic phase) during which all the Permo(?)-Triassic basins were inverted (Nikishin et al. 1998a,b, 2001; Gaetani et al. 2006). The compressive event is probably related to the coll i s i o n - a c c r e t i o n of Gondwana-derived blocks (which together form the composite Iran plate; S ~ d i 1995; Besse et al. 1998) SE of the GC w h e n the Palaeotethys Ocean closed along the

Fig. 3. Section across the central part of the Greater Caucasus showing the southward vergence of the whole belt and the major thrusting of the belt over the Transcaucasus (Dotduyev 1987). (Section location is shown in Fig. 2.)

280

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Fig. 4. Tectonic map of the Greater Caucasus area from Ruppel & McNutt (1990) (other sources: Milanovsky& Khain 1963; Kotansky 1978; Dotduyev 1987; Philip et al. 1989). Talesh-Alborz-Aghdarband zones, in the area of the present-day South Caspian Basin (Sttcklin 1968; Davies et al. 1972; ~engtr 1984; Alavi 1991; Ruttner 1993; Dercourt et al. 2000, and references therein). This Eo-Cimmerian orogeny has also been clearly identified as a major event in the Turkish Pontides (Okay 2000; Okay et al. 2006). It should also be noted that (1) the Black Sea was certainly not developed by the Late Triassic, and the Pontides were therefore close to the Transcaucasus-GC; (2) the Pontides-Transcaucasus- Talesh-Alborz- Aghdarband-GC zones probably together formed a contiguous part of the widespread Eo-Cimmerian orogenic belt.

E a r l y J u r a s s i c to M i d - J u r a s s i c

A field study carried out in 2003 led to the postulation of a model of the GC in Jurassic times by Saintot et al. (2004). A new rifting phase occurred during the Early Jurassic (Zonenshain et al. 1990; Nikishin et al. 1998a,b, 2001, and references therein) under a transtensional stress regime with a nearly east-west-directed tensional stress axis (Stbrier et al. 1997; Saintot et al. 2004). Thus, this transtensional Early Jurassic rifting shares some similarities

with the model of Banks & Robinson (1997) for the Black Sea region, which surmises that the Early Jurassic GC Basin corresponded to an en echelon set of rhomb-shaped depocentres. Early Jurassic rift activity is also reported in the Eastern Pontides, which were adjacent to the GC at that time (e.g. Okay & Sahintfirk 1997) and in the South Caspian Basin (Early(?) to Mid-Jurassic times; Brunet et al. 2003). GC rifting continued through part of the Mid-Jurassic. Extrusive magmatism (mainly rhyolitic) accompanied the GC rifting phase during Sinemurian-Pliensbachian times (Lordkipanidze et aL 1989). Sediments of this age are represented by deltaic(?) coarse sandstones NW of the belt, by deep marine mudstones-sandstones in the central part, and by shallower mudstones-sandstones to the south. In Toarcian times, from north to south, shelf to deep marine mudstones-sandstones were deposited, with no record of volcanic activity. In AalenianBajocian units, sedimentary facies laterally vary from continental to deep marine. The Aalenian and Bajocian periods are also characterized by bimodal rhyolitic and basaltic extrusive rocks (from mantle and crustal sources) in a subaerial as well as a shallow-marine environment. In the model, the western part of the GC evolved during the Early and Mid-Jurassic as the western margin of the rift with shallow-water sedimentation and subaerial extrusion of lava flows. Deep-water sediments are encountered towards the present-day central part of the belt (crossing the inferred, north-south-oriented normal faults), associated with mid-ocean ridge basalt (MORB)-like tholeiitic basalt extrusion during the Aalenian. Not only partial melting of asthenosphere is implied, but also a high degree of extension, approaching that required for oceanic crust development in the present-day central part of the GC belt. The total thickness of the Lower Jurassic to Aalenian unit in some parts of the GC Basin is more than 5000 m, and it is mainly composed of black shales and deep-water sandstone turbidites (as well as the volcanic rocks and pyroclastic deposits). The Aalenian extensional phase has been well documented in the field with, for example, the presence of a large NW-SE-trending normal fault, east of the Kuban Basin. Toarcian to Aalenian units are tilted along this fault (Fig. 6) and the minimum downthrow should be of several hundreds of metres. The age of fault activity is constrained by overlying, sealing Upper Aalenian units. In Bajocian times, a huge quantity of pyroxene-bearing basalts were extruded, and formed a subaerial to shallow-water volcanic chain on the southern margin of the basin (presumably accompanied by uplift at the rift margin). Synchronously with the formation of this relief, conglomerates (reworking the lavas) were deposited toward the depocentre of the basin to the north.

Fig. 5. Section across the western part of the Greater Caucasus showing the basement involved in south-vergent, flat thrusting. Offshoreis shown an interpretation of the seismic line SU8040 (from Robinson et al. 1996). (Section location is shown in Fig. 2.)

THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS

Fig. 6. Photograph of a progressive unconformitycreated by a major synsedimentarynormal fault within Aalenian deposits dipping to the NE. The normal fault, with a NW-SE strike, is located to the left of the photograph and may be followedfor some 10 km along the upper rims of the Upper Kubanvalley, on the northern side of the central GC. This normal fault testifies to the Mid-Jurassic extensionaltectonics of the GC Basin and contributes to the along-strike variation of sedimentarythickness (photograph at 43~ 42~ by M. Srbrier).

The calc-alkaline nature of the Bajocian lavas has been formerly interpreted as indicative of a subduction-related volcanic arc marking the incipient subduction of a large oceanic plate such as the Palaeotethys along the GC (see, for example, the southward subduction of Palaeotethys along the GC as described by ~eng6r (1984)). However, this hypothesis seems very unlikely because the above-mentioned Bajocian volcanic rocks are also spread over the Transcaucasus and there is no evidence for a subduction zone along the southern edge of the GC (i.e. remnants of an accretionary prism, high-pressure metamorphism, ophiolitic fragments, etc.). The calc-alkaline nature of the Bajocian lavas can also be explained by the GC rift being in a back-arc setting relative to a subduction zone located far to the south in the Lesser Caucasus (see Adamia et al. 1981; Gamkrelidze 1986; Panov 2004). The Artvin-Bolnisi zone lying between the Transcaucasus and the suture zone of the Lesser Caucasus (location shown in Fig. 2) is a good candidate for a subduction-related volcanic arc during the Early and Mid-Jurassic, with shallow-water to continental sediments and major calc-alkaline volcanism (Adamia et al. 1981; Gamkrelidze 1986; Panov 2004). It is not uncommon for lavas extruded in a back-arc rift setting, but close to the volcanic arc, to show such calc-alkalinity. Therefore, it cannot be excluded that the Bajocian lavas were extruded during what could still be considered as a synrift stage of basin evolution, continuing from the Aalenian. However, it is noted that, whereas structural constraints (e.g. synsedimentary normal faulting) clearly exist to define the Aalenian succession as synrift, there are no such structural constraints for the Bajocian units. Indeed, the widespread occurrence of Bajocian calc-alkaline volcanic rocks that can be encountered from the Lesser Caucasus to the MCT may also simply suggest an expansion of the subduction-related volcanic arc from some 5 0 - 1 0 0 km width in Aalenian time, restricted by the Artvin-Bolnisi zone, to nearly 200 km in the Bajocian, thus merging with the southern part of the GC Basin. Shallowing of the subducted slab could explain such an encroachment of the arc into the previously back-arc setting. If this were the case, rift activity in the GC Basin would have stopped (given that a rather flat-dipping slab does not favour the opening of a back-arc basin; eg. Lallemand et al. 2005, and references therein) and, therefore, the Bajocian volcanic rocks should be considered as occurring at the onset of the post-rift stage of GC Basin development. The available observations, relating to only

281

the calc-alkaline character of the Bajocian lavas and their widespread occurrence, cannot discriminate between these two possibilities. The Bathonian unit (where not absent) is composed of a greywacke siltstone unit into the basin and regressive coal-bearing terrigeneous sediments on its southern margin. The Upper Jurassic unit lies transgressively and discordantly on the Middle Jurassic unit. It reportedly lies conformably on the Middle Jurassic unit along the present-day southern slope of the central and eastern part of the GC Basin (Gamkrelidze 1986; Zonenshain et al. 1990; Nikishin et al. 1998a,b), although field observations made in the same central area (by M. Srbrier in 2004) revealed an unconformity between the Callovian deposits and underlying units. It is also worth noting that Cenozoic deformation is so intense in the so-called Flysch Zone (see Fig. 2) that no clear conclusion can be made regarding the detailed relationships between Mesozoic units. A compressional event has been proposed to have occurred in Bathonian times, resulting in the inversion of the margins of the basin (Adamia et al. 1981), although it may be that this unconformity is simply related to the cessation of rifting and the onset of post-rift basin development, such as recorded in many rift basins (see, e.g. Coward et al. 1987; Tankard & Balkwill 1989; Frostick & Steel 1993; Williams & Dobb 1993; Busby & Ingersoll 1995; Stephenson et al. 1996; Cloetingh et al. 1997; McCann & Saintot 2003). (Brunet et al. (2003) also pointed out that the regional Bathonian unconformity around the South Caspian Basin may be a 'break-up unconformity' marking the onset of sea-floor spreading rather than the occurrence of a compressive tectonic event.) Nevertheless, in the southernmost part of the GC (in Georgia), Callovian strata overlie open folds in Middle Jurassic strata, constraining a gentle folding event to the Bathonian. In the central part of the GC, north of the MCT, highly folded Early Jurassic strata are overlain by subhorizontal layers of Upper Jurassic and Cretaceous platform-type deposits. According to Belov et al. (1990) and Somin (2000), they are in place and indicate that the Bathonian folding was significant and involved intense shortening. Other authors (e.g. Korsakov et al. 2001) have considered that the Upper Jurassic and Cretaceous strata are in an allochthonous position and, accordingly, that the thrust sheet and the folding developed together during Alpine orogenesis (implying the occurrence of a folding phase, followed by the development of an erosional surface and then thrusting of nappes along a drcollement level). On the northernmost slope of the belt, the angular discordance between transgressive Callovian and older rocks disappears. Published cross-sections (e.g. Panov 2002, 2004) show south-vergent folds and thrusts affecting strata older than and including Bajocian, and no sealing by younger sediments (which are absent). In the northern part of each of these crosssections lies a gentle monocline composed of Upper Jurassic and Cretaceous units underlain by Lower Jurassic units without any angular unconformity as might be expected to be related to a compressional phase during Bathonian times. (A Bathonian stratigraphic gap indeed exists locally, the Bathonian being a time of worldwide regression.) No important or diagnostic compressive structures (such as folds and thrusts) were observed in Middle Jurassic rocks sealed by the Callovian by the senior author during fieldwork in 2003 in the northern part of the belt (see Saintot et al. 2004). What was observed is a gently tilted unit (like the Aalenian unit) below the Callovian transgressive unit. Going southward across the belt, closely and tightly folded Lower and Middle Jurassic strata can be observed (Fig. 7). The same style of folding is observed some 10 km towards the Black Sea coast in Palaeocene rocks (Fig. 8). SE-vergent minor thrusts are also common in Lower and Middle Jurassic units, similar to the younger strata. The systematic analysis of brittle structures within the GC also strongly suggests that only one set of reverse faults developed in Jurassic and younger strata and that this set is related to the Cenozoic palaeo-stress field (Fig. 9; see discussion

282

A. SAINTOT ETAL. indeed, only isostatic readjustments at the syn- and post-rift transition), affecting units from place to place, rather than the complete inversion of the GC Basin. (In Lower Middle Jurassic rocks there is no evidence of intense folding and thrusting that can be ascribed unequivocally to a Bathonian compressional event, most of the deformation being clearly Cenozoic in age). In summary, the pre-Callovian unconformity remains a matter of debate. It could record either the transition between syn- and post-rift phases in the GC Basin or, alternatively, a weak compressive event related to the accretion of crustal blocks along the active continental margin to the south.

The Callovian-Eocene Greater Caucasus Basin

Fig. 7. Photograph of folded Aalenian-Bajocian unit (PshishFormation) of the western Greater Caucasus (photograph by A. Saintot; S. Korsakov for scale). Fold axes strike NW-SE to WNW-ESE. and analyses of structures related to Cenozoic shortening by Saintot & Angelier (2002)). The localized angular unconformity at the base of Upper Jurassic strata thus probably records not more than a phase of gentle compression of the GC Basin (or,

The GC Basin evolved dominantly as a post-rift (thermally subsiding) basin from the Callovian until the Late Eocene following its Early to Mid-Jurassic episodes of rifting. A thickness of 6 - 8 km of calcareous, mainly Cretaceous, flysch-type sediments was then deposited and most of the Greater Caucasus Mountains corresponds to the so-called Flysch Zone of the southern limb of the GC (Fig. 2; Milanovsky & Khain 1963; Lordkipanidze 1980; Koronovsky 1984; Gamkrelidze 1986; Belousov et al. 1988; Adamia & Lordkipanidze 1989; Zonenshain et al. 1990). The nature of the underlying crust has not been established, although Ershov et al. (2003) estimated a crustal thickness of 15-17 km, suggesting that it corresponds to thinned continental crust. Such an interpretation is in agreement with the absence of oceanic crustal remnants in the belt. It follows that the basin was probably not floored by significant oceanic crust (see also the important discussion by Ershov et al. 2003, p. 102). The Callovian conglomerates and calcareous sandstones clearly belong to the post-rift succession of the GC Basin. They unconformably overlie the oldest units on an erosional surface. Upward, the Callovian unit becomes marly, indicating platform subsidence. In Late Jurassic times, sandstones and clays filled in the sedimentary basin and reef limestones developed towards its margins. Kimmeridgian-Tithonian gypsumbearing and lagoonal sediments were deposited on the northern (Laba-Malka zone) and southern margins (in Georgia). A very thick Cretaceous to Eocene greywacke siltstone flysch-like unit with clastic limestones in the Lower Cretaceous interval conformably overlies the Upper Jurassic sequence. (The Lower Cretaceous succession is 750-1600 m thick, the Upper

Fig. 8. Two photographs of the Lower Palaeocene flysch-likefolded unit of the western Greater Caucasus along the Black Sea coast. Fold axes strike NW-SE to WNW-ESE (photographs by A. Saintot).

THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS

283

Shcherba 1998). In general, according to the Russian literature, there are two distinct orogenic processes recorded in the GC. The first of these is understood to have occurred in Late EoceneEarly Oligocene times, and involved folding with significant crustal shortening but without major uplift of rocks at the surface. The second is considered to have occurred in Miocene times, and is characterized by limited shortening (reported to be 5-10%) and no significant folding but rather by significant uplift, thus explaining the deposition of Sarmatian conglomerates.

Fig. 9. (a) Stereoplot of 60 close and tight fold axes collected in rocks of Early and Mid-Jurassic age of the western Greater Caucasus (Schmidt's projection, lower hemisphere). (b) Attitudes (strikes and dips) of 215 bedding planes collected in rocks of the western Greater Caucasus from Late Jurassic to Eocene in age. The attitude of folds is the same in both stratigraphic intervals: the WNW-ESE close folds in Late Jurassic to Cenozoic rocks are also observed in Early and Mid-Jurassicrocks. Most of the compressivestructures and tectonic contacts (such as thrusts) measuredin Early and Mid-Jurassicrocks are consistent with a NNE-SSW Cenozoic compression(see Saintot & Angelier 2002).

Cretaceous 5 0 0 - 9 0 0 m thick, and the Palaeocene-Eocene, 600-850 m thick). Restorations by Yakovlev (2002, 2005) along several profiles crossing the GC clearly show that the RLFZ (Figs 2 and 3) was the southernmost normal fault bordering the GC Basin, involved in controlling the northward increase of sedimentary thickness from the Mesozoic-Cenozoic Transcaucasus stable platform to its subsiding part, the Gagra-Dzhava zone (see Fig. 3). They also show that the cumulative primary normal displacement along the RLFZ was much larger than the secondary reverse one. The most important normal fault at this time, however, was the Utsera fault, limiting the Gagra-Dzhava zone and the flysch zone, in which the thickness of Mesozoic and Cenozoic deposits reaches 12-15 km. It was also during this time that the eastern Black Sea Basin developed, close to the GC Basin and south of the Shatsky Ridge (a western prolongation of the Transcaucasus; see Fig. 5), although the precise timing is still matter of debate: Late Cretaceous and Palaeocene according to Finetti et al. (1988), Eocene according to Lordkipanidze (1980), and Late PalaeoceneEocene according to Robinson et al. (1996) and Shreider et al. (1997). Analysis of kinematic data by Saintot & Angelier (2002) revealed that a transtensional stress field affected the GC Basin during the Eocene (with an east-west trend of extension), which those workers considered to be a far-field effect of rifting (or, at least, of rift reactivation) in the eastern Black Sea Basin. In any case, during Palaeocene-Eocene times, prior to the main shortening event, the GC Basin was a deep-water basin, with a sedimentary infill of 10 km on average (Borsuk & Sholpo 1983), similar to the eastern Black Sea and South Caspian basins (Zonenshain & Le Pichon 1986; Nikishin et al. 1998b) and linked with the latter but apparently separated from the former by the Shatsky Ridge (see Fig. 2).

Cenozoic to present-day shortening along the Greater Caucasus The main orogenic phase is considered to extend from Late Eocene to Early Oligocene time by Shardanov & Peklo (1959), Beliaevsky et al. (1961), Milanovsky & Khain (1963), Grigor'yants et al. (1967), Khain (1975, 1994), Milanovsky et al. (1984), Muratov et al. (1984), Giorgobiani & Zakaraya (1989), Robinson et al. (1996), Lozar & Polino (1997), Robinson (1997), Nikishin et al. (1998a,b, 2001) and Mikhailov et al. (1999), with pulses of orogeny encompassing the rest of Cenozoic to the present day. According to other workers, the orogeny did not begin prior to the Miocene, in Sarmatian time (e.g. Dotduyev 1987; Shcherba 1987, 1989, 1993; Zonenshain et al. 1990; Kopp 1991; Kopp &

Late Eocene

What follows is a summary of the main arguments used by authors to demonstrate that the inversion of the basin began in Late Eocene times. (1) An angular unconformity of Maykop (Oligo-Miocene) on deformed older units is regionally observed in the field (Milanovsky & Khain 1963; Khain 1975, 1994; Borukaev et al. 1981; Rastsvetaev & Marinin 2001; Banks, pers. comm.) and on seismic lines (Tugolesov et al. 1985; Robinson et al. 1996; Banks, pers. comm.) on both sides of the GC. On the southern slope of the central GC, there is (a) a Late Eocene olistostrome unit (10-400 m thick) coeval with southward thrusting of the GC Basin (Khain 1975, 1994) and (b) a deltaic, southward prograding sandy facies in the Oligocene unit coeval with the uplift-emergence of part of the GC (M. S6brier's field observation). Sharafutdinov (2003) dated the folding event as latest Eocene-Early Oligocene on the northern slope of the GC and in the Fore-Caucasus, and he reported Early Oligocene folds, olistostromes (confirming the existence of back-thrusts), angular unconformities and the tectonic removal of a large part of the section with everything being overlain by flat-lying strata of MidOligocene age. The youngest unit prior to the onset of deformation is Late Eocene in age (Khadum Formation). The reported features, including the folds (especially if they are related to slumping), imply synsedimentary deformation in a foreland developing at the front of a propagating back-thrust. (2) A high rate of tectonic subsidence occurred at the beginning of the Oligocene in the Indolo-Kuban and Terek-Caspian basins as shown by burial history modelling (back-stripping analyses of wells and numerical modelling of lithospheric deformation) by Nikishin et al. (1998a), Ershov et al. (1999) and Mikhailov et al. (1999). Those workers assumed that the Indolo-Kuban, TerekCaspian and Tuapse, Kura and Rioni troughs developed as flexural foreland basins in response to lithospheric compression from the south during Late Eocene times (resulting from the closure of Neotethys and collision south of the Transcaucasus area). Ershov et al. (2003) discussed mechanisms other than foreland flexure for the formation of these basins, including the role of mantle processes occurring at the cessation of shortening, related to the underthrusting of thinned continental crust beneath the basins. (3) Lozar & Polino (1997) carried out a study based on nannofossils occurring in Maykop sediments of the Kuban Basin and of Upper Cretaceous rocks on the northern slope of the western GC. The base of the Maykop group is inferred to be Late Eocene-Early Oligocene in age and its lowermost part contains a reworked assemblage (80% of the total assemblage) of Late Cretaceous and Palaeogene nannofossils. These are very well preserved (with, for example, intact spines), implying that they were not transported over long distances. The sediment source was the area of the present GC where, indeed, Late Cretaceous and Palaeogene sediments were eroded. However, although there is general agreement on the Late Eocene uplift of the central part of the GC, this area was not yet actually above sea level by this time according to Kopp & Shcherba (1985) and Ershov et al. (1999, 2003). The types of Maykop nannofossils found in situ suggest restricted environmental conditions, leading to the interpretation that environmental changes occurred during

284

A. SAINTOT ETAL.

Late Eocene-Oligocene times, either with the onset of a cooler climate or as a result of the isolation of the Paratethys domain by the uplift and emergence of an orogenic belt acting as a barrier along the Pontides-Lesser Caucasus. The Late Eocene compressional palaeostress field responsible for the inversion of the GC Basin has been determined through tectonic analysis by Saintot & Angelier (2002). It was oriented N E - S W to N N E - S S W , leading to the development of N W - S E and W N W - E S E dip-slip thrusts in the GC. The main features and chronology of this tectonic phase have been established as follows: (1) on the northern slope of the GC, where the regional structure is a monocline, the Palaeocene strata are clearly affected by the inferred compressional palaeostress field (see details of measurements and site numbers given by Saintot & Angelier (2002)), whereas no related reverse and strike-slip faults can be observed affecting the overlying Miocene rocks studied by Saintot & Angelier; (2) the palaeostress field has also been recorded in Middle Eocene rocks along the southwestern coast (see details of measurements and site numbers given by Saintot & Angelier (2002)); (3) this palaeostress field is the only one recorded during pre- (e.g. Fig. 10), syn- and post-folding phases (Saintot & Angelier 2002).

M i o c e n e to p r e s e n t day

From Sarmatian times (Mid-Miocene) until the present, pulses of compressional deformation have affected the GC (Belousov 1940; Shardanov & Peklo 1959; Beliaevsky et al. 1961; Milanovsky & Khain 1963; Shcherba 1987, 1989, 1993; Giorgobiani & Zakaraya 1989; Kopp 1989, 1991, 1996; Rastsvetaev 1989; Zonenshain et al. 1990; Milanovsky 1991; Khain 1994; Kopp & Shcherba 1998; Nikishin et al. 1998b). However, it appears as though the present-day structure of the GC is inherited mainly from the Sarmatian compressional pulse. The Sarmatian sedimentary unit surrounding the belt comprises syndeformational conglomerates reflecting the growth of topography at this time (Mikhailov et al. 1999) and, indeed, the emergence of the GC belt as a whole, the central GC having already been uplifted since the latest Eocene (Khain 1994; Lozar & Polino 1997; Ershov et al. 1999, 2003). The present-day displacement of Arabia relative to Eurasia by several centimetres per year is recorded throughout the GC. The indentation of Arabia occurs at Bitlis-Zagros and deformation propagates towards the GC. This indentation has produced large strike-slip faults along which the Anatolian block escapes westward. In the GC, both strike-slip faults and thrusts actively accommodate deformation. Earthquake focal mechanisms reveal that the

Fig. 10. Photograph of reverse faults developed prior to the tilting of beds under a NE-SW compression (Late Cretaceous flysch-likeunit of the western GC) and stereoplots of the related stress tensors (calculatedfor both attitudes of beds: present-day and restored to horizontal stress tensors obtained from inversionof the fault slip data as givenby Saintot & Angelier2002). (Photograph by A. Saintot; J. Angelier for scale.)

whole Caucasian area is under a north-south compressional stress regime (Gushtchenko et al. 1993; Gushtchenko & Rebetsky 1994; Mikhailov et al. 2002), a continuation of the inferred Sarmatian palaeostress regime (Saintot & Angelier 2002). The 'Caucasian' N W - S E and W N W - E S E faults act as oblique reverse faults. The depth distribution of earthquakes is limited to the crust and the overlying sedimentary succession; no deeper earthquakes are observed, nor has a Benioff Zone been imaged. Earthquakes at depths of 10-15 km are related to strike-slip faults, whereas deeper hypocentres are along thrust faults. Also, it is observed that along single focal zones, the depths of hypocentres increase northwards (Gamkrelidze 2005) along gently north-dipping planes. In particular, the identified focal plane of the 29 April 1991 Racha earthquake (Mw = 7) exhibits a dip angle of 2 0 40 ~ north (Triep et al. 1995). The distribution of seismicity also indicates the propagation of the GC front southwards to the offshore Shatsky Ridge (a western prolongation of the Transcaucasus; Fig. 5), and to the Rioni and Kura basins (Figs 2 and 4). On seismic lines crossing the offshore western GC (Finetti et al. 1988), it can be observed that, with the continuing compression, the Tuapse Basin as a whole overthrusts the Shatsky Ridge with a southward propagation of the GC deformation front. Active thrusting of the GC also affects the Rioni and Kura basins. The Oligocene-Early Miocene sedimentary infill of these two basins has been incorporated into the south-vergent fold-and-thrust belt during the Mid-Miocene compressional phase. The faults transverse to the GC belt have been invoked as conduits for Quaternary volcanism (Milanovsky et al. 1984; Giorgobiani & Zakaraya 1989; Lordkipanidze et al. 1989; Koronovsky et al. 1997). These faults, which were very active during the Cenozoic, have segmented the GC and Transcaucasus area as well as the W N W - E S E 1250km long south-vergent frontal thrust of the GC (Giorgobiani 2004). Similarly, a large NE-striking left-lateral fault, with a reported offset of 90 km (Philip et aL 1989), was proposed as the conduit for the Kazbek volcano (see location of volcanoes shown in Fig. 4). However, the evidence for large strike-slip displacements along such structures in the GC belt and Transcaucasus area remains very speculative. S o m e characteristics o f the inversion o f the Greater C a u c a s u s Basin

Using simple area-balancing restoration of cross-sections, Ershov et al. (2003) estimated the amount of shortening along the GC to have been 2 0 0 - 3 0 0 k m (as also reported by Khain 1982; Zonenshain & Le Pichon 1986; Shcherba 1993; Nikishin et al. 1998b). Such an estimate is in agreement with the inferred plate kinematics of the area, which suggests a 400 km displacement of Arabia northwards to (fixed) Eurasia from Oligocene times and takes into account the amount of shortening in the Lesser Caucasus area. However, field observations (M. Stbrier) of the structural relationships between Mesozoic GC formations indicate that the shortening accommodated by the MCT sensu lato is of the order of some tens of kilometres and that each of the few other major thrusts should accommodate some 2 - 5 km (e.g. along the eastern part of one of the MCT branches, the Lower Jurassic units are thrust over themselves). It follows that the shortening across the GC as a whole could be much less, as little as 100 km. The central part of the GC belt, which has the highest elevation and the highest rate of Neogene to present uplift, corresponds to the thinnest part of the Aalenian GC rifted lithosphere. Thus, the anomalously high elevation in this area could be a consequence of the subduction of highly thinned continental lithosphere (if not partly oceanic, as mentioned earlier). The lithospheric root might also be comparatively more important in the central part of the GC because collision and shortening was concentrated there, directly in front of the indenting Arabian plate. The Quaternary and still active uplift of the central part of GC could

THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS

285

Fig. 11. Main subsidence-driving mechanisms for the foreland stage of basin evolution and uplift of the central part of the GC as a result of delaminationof the root (from Ershov et al. 2003).

be the result of delamination of a lithospheric root, as suggested by numerical modelling (Ershov e t al. 2003; Fig. 11) and tomography (Brunet e t al. 2000) results, with the Quaternary volcanism being linked to this deep process. Lithospheric roots would not have been so well developed in the western and eastern GC because there was less shortening there and it was accommodated differently. In the eastern GC, shortening is symmetrically accommodated by the fan-shaped development of foreland structures. To the west of the central GC, no large lithospheric root is expected because the cumulative shortening there is significantly less because of the western escape of Anatolia.

The structural style of the GC belt agrees with the inversion of a deep basin developed on very thin continental crust, perhaps similar to what Gamkrelidze & Giorgobiani (1990) referred to as 'intraplate subduction' in an intraplate setting. As such, the GC can be viewed as a Pyrenees or Atlas Mountains analogue (see, e.g. the overview of the Pyrenees by Grup de Geodin~mica i Anfilisi de Conques 2005). No lateral escape during shortening and consequent development of large nappes (and rootless nappes), such as in the Alps, occurred. The absence of any remnants of an ophiolitic suture supports such a model. Furthermore, there is no obvious record of any subduction zone along the GC

Fig. 12. Summaryof the tectonic evolutionof the Greater Caucasus. Absolute ages are from Gradstein et al. (2004).

286

A. SAINTOT ETAL.

during Mesozoic and Cenozoic times. There is no volcanic arc or blueschist and high-grade metamorphic rocks, and no accretionary complexes are present. (It is noted, however, that older, Palaeozoic, ophiolites and associated high-grade metamorphic rocks do crop out in the central part of the belt.)

Conclusions The crustal structure of the Greater Caucasus remains a matter of debate, and two different models have been postulated. One model considers the GC belt as a former deep marine Mesozoic basin that was subsequently squeezed between steep crustal faults, these faults separating the GC from its adjacent tectonic units, the Transcaucasus and the SP. The alternative model considers the GC as a south-vergent, crustal-scale, imbricated fold-and-thrust belt with the SP thrust over the Transcaucasus massif along north-dipping planes, which flatten at depth. More and better geophysical data are needed to discriminate between these two models. However, the latter appears in general to satisfy better the available data, although some interpretations remain questionable (such as the geometry of the boundary fault zone between the GC and the SP and the amount of shortening in the GC belt). The tectonic evolution of the Greater Caucasus during Mesozoic and Cenozoic times can be summarized as follows (see Fig. 12): (1) Permo(?)-Triassic rifting; (2) Eo-Cimmerian shortening related to collision of the Iranian Block with Europe; (3) development of Early-Mid-Jurassic rift basins, possibly related to north-dipping subduction south of the Transcaucasus (i.e. in the Lesser Caucasus); (4) development of a Bathonian (Mid-Cimmerian) unconformity related either to the syn- to post-rift transition or to a collisional event at the active margin; (5) M i d - L a t e Jurassic to Eocene post-rift subsidence; (6) Late Eocene basin inversion related to the final closure of the Tethys oceanic domain; (7) a second shortening phase from Late Miocene time to the present accompanied by uplift and magmatism and corresponding to the final stages of A r a b i a - E u r a s i a collision. This paper has benefited from many fruitful discussions with Russian colleagues, including S. Korsakov, P. Fokin and P. Tikhomirov. Part of the research was funded by the MEBE programme and, in the past, by the Peri-Tethys Programme (A. S. would especially like to thank A. Ilyin for his help in the field during 'PeriTethys years', as well as J. Angelier for her earlier work). The Netherlands Research Organization (NWO/ALW) funded part of A.S.'s research, and F.Y.'s research was partly supported by NATO 1997 (202025D). The authors also thank M. L. Somin, L. M. Rastsvetaev and A. V. Marinin, who kindly discussed some important scientific aspects of the manuscript, as well as the two reviewers, D. Brown and A. Okay, whose comments led to improvements incorporated in the present manuscript, which is Netherlands Research School of Sedimentary Geology contribution 2005.05.02.

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SAINTOT, A. & ANGELIER, J. 2002. Tectonic paleostress fields and structural evolution of the NW-Caucasus fold-and-thrust belt from Late Cretaceous to Quaternary. Tectonophysics, 357, 1-31. SAINTOT, A., CHALOT-PRAT, F., MCCANN, T., FOKIN, P., KORSAKOV, S. & STEPHENSON, R. 2004. The Early Middle Jurassic basins of the Western Greater Caucasus (Russia). AAPG European Region Conference with GSA, Prague, 10-13 October 2004. SI~BRIER, M., POLINO, R., GALKIN, V. A. & YUNGA, S. 1997. Caucasus stress evolution from Mesozoic to present. LUG 9, Strasbourg, 23-27 March 1997. Terra Nova, 9 (Abstracts Supplement 1), 337. ~ENGOR, A. M. C. 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia. Geological Society of America, Special Papers, 195. SHARAFUTDINOV,V. F. 2003. Geological structure and patterns of development of the Maykopian deposits on the North-Eastern Caucasus in connection with oil-and-gas content. DSc thesis, Lomonosov Moscow State University, Moscow [in Russian]. SIqARDANOV, A. N. & PEKLO, V. P. 1959. Tectonics and history of formation of buried folds on southern flank of western Kuban trough and hydrocarbon potential of Mesozoic deposits. Gostoptekhizdat, Moscow, 3-27 [in Russian]. SHCHERBA, I. G. 1987. Olistostromes and problems of Cenozoic tectonics of the Caucasus. In: MILANOVSKY, E. E. & KORONOVSKY, N. V. (eds) Geology and Mineral Resources of the Great Caucasus. Nauka, Moscow, 191-200 [in Russian]. SHCHERBA, I. G. 1989. Paleogeography and tectonics of Maikop basin, Caucasus. Bulletin of MOIP, Geologia, 306(5), 1196-1200 [in Russian]. SHCHERBA,I. G. 1993. Stages and Phases of Cenozoic Evolution of Alpine Domain. Nauka, Moscow [in Russian]. SHEMPELEV, A. G., PRUTSKY,N. I., FELDMAN, I. S. & KUHMAZOV,C. U. 2001. Geological-geophysical model along cross-section TuapseArmavir. The tectonics of Neogey (Pz + Mz + Cz)---common and regional aspects. GEOS, 2, 316-320 [in Russian]. SHEMPELEV, A. G., PRUTSKY, N. 1., KUHMAZOV, C. U., P'YANKOV, V. YA., LIGIN, V. A. & MOROZOVA, A. G. 2005. The materials of geophysical investigation along Near-Elbrus cross-section (volcano Elbrus-Caucasian mineral waters). The tectonics of Earth crust and mantle. The tectonic regularity of mineral resources localization. GEOS, 2, 361-365 [in Russian]. SHREIDER, A. A., KAZMIN, V. G. & LYGIN, V. S. 1997. Magnetic anomalies and age of the Black Sea deep basins. Geotectonics, 31(1), 54-64. SOMIN, M. L. 2000. About structure of axial zones of Central Caucasus. Reports of the Russian Academy of Sciences, 375(5), 662-665 [in Russian]. STEPHENSON, R. A., WILSON, M., DE BOORDER, H. & STAROSTENKO, V. I. (eds) 1996. EUROPROBE: intraplate tectonics and basin dynamics of the Eastern European Platform. Tectonophysics, 268. STOCKLIN, J. 1968. Structural history and tectonics of Iran: a review. AAPG Bulletin, 52(7), 1229-1258. TANKARD, A. J. & BALKWILL, H. R. (eds) 1989. Extensional Tectonics and Stratigraphy of the North Atlantic Margins. AAPG, Memoirs, 46.

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ZONENSHAIN,L. P. & LE PICHON, X. 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. Tectonophysics, 123, 181 - 211. ZONENSHAIN, L. P., KUZMIN, M. I. & NATAPOV, L. M. 1990. Geology of the USSR: a plate tectonics synthesis. American Geophysical Union, Geodynamics Series, 21.

The Western Accretionary Margin of the East European Craton: an overview T. C. P H A R A O H 1, J. A. W I N C H E S T E R 2, J. V E R N I E R S 3, A. L A S S E N 4 & A. S E G H E D I 5

1British Geological Survey, Kingsley Dunham Centre, Keyworth NG12 5GG, UK (e-mail: [email protected]) 2School of Physical and Geographical Sciences, Keele University ST5 5BG, UK 3Ghent University, Palaontologie, Krijgslaan 281/$8, B 9000, Gent, Belgium 4Copenhagen University, Ostervoldgade 10, DK 1350, Copenhagen, Denmark 5Geological Institute of Romania, 1 Caransebes St, 012271 Bucharest 32, Romania

Abstract: Multidisciplinary investigations of the western margin of the East European Craton (EEC) by EUROPROBE projects since

1992 have confirmed that the Trans-European Suture Zone (TESZ) is the most fundamental lithospheric boundary in Europe, extending 2000 km from the North Sea to the Black Sea-Crimean region. The crust of the EEC is thicker and denser than that of Phanerozoic-accreted Europe, and the base of the lithospheric mantle significantly deeper. These characteristics persist throughout the length of the TESZ, despite the variation in age of the accreted crust along strike. Geological studies of key deep borehole cores and the limited outcrop data confirm that the crust of Phanerozoic-accreted Central Europe comprises a number of terranes, each thought to be derived from Gondwana during several episodes of rifting, ocean formation, ocean destruction and sequential accretion to the EEC throughout Palaeozoic time. There is still much discussion about the identity, provenance and history of these orogenic terranes. The process of accretion led to the formation of terrane-bounding orogenic sutures, which may be marked in outcrop by ophiolitic and eclogitic relics. Recognition of concealed sutures is obviously more difficult,and relies on a variety of geophysical techniques, used in an integrated way by multidisciplinaryteams; the evidence from deep seismic reflection and refraction surveys, teleseismic tomography, magnetotelluric experiments and from geophysical potential-field modelling is crucial for such studies. Since the European Geotraverse, much has been learnt about the geometry of the Thor, Iapetus, Rheic, Saxo-Thuringian and Moldanubian oceanic sutures, through the crust and sometimes into the mantle. This has led to a much better understanding of the 3D crustal structure of the Western Accretionary Margin of the EEC, and the lithospheric processes that have shaped it. From this, the influence of tectonic heterogeneities within the orogenic crust on the development of post-orogenic structures and basins can be much better constrained.

The Western Accretionary Margin of the East European Craton (EEC) is the most fundamental lithospheric boundary in Europe, separating the thick, cold, ancient crust and lithosphere of the exposed Baltic Shield and the partly concealed EEC from the younger, warmer and much thinner crust and lithosphere of Western Europe (Gee & Zeyen 1996), and extending deep into the mantle (Zielhuis & Nolet 1994; Babugka et al. 1998; Plomerova et al. 2002), perhaps as deep as 250 km. The transition from the thick crust and lithosphere of the EEC to the thinner crust of the accreted margin takes place over a broad zone some 400 km wide and 2000 km long, extending from the North Sea to the Black Sea (Fig. 1), for which the name 'Trans-European Suture Zone' was coined by the E U R O P R O B E Programme (Gee & Zeyen 1996). This region, recognized as a particularly significant lithospheric boundary, was identified as a key target for EUROPROBE research (Gee & Beckholmen 1993). For the past 10 years, it has been the focus of multidisciplinary investigations within E U R O P R O B E ' s Trans-European Suture Zone (TESZ) Project. Over most of its length this zone is concealed by deep sedimentary basins of Permian to Cenozoic age; thus geophysical experiments and multidisciplinary studies of samples from deep boreholes are crucial to understanding its history. A number of major WNW-trending crustal lineaments, in particular the Sorgenfrei-Tornquist (STZ), Teisseyre-Tornquist (TTZ) and Elbe Lineaments, are present within the TESZ. The significance of these has long been recognized, as a consequence of the influence they have persistently exhibited during the evolution of overlying late Palaeozoic and Mesozoic sedimentary basins, and as loci for Alpine inversion (Berthelsen 1992a). EUROPROBE seismological research has confirmed that the STZ and TTZ are steep features associated with displacements of up to 5 km at the Moho level. Although these are spectacular crustal features, they do not represent the original, orogenic sutures between the orogenic terranes making up the collage of 'Old' and 'Young' Europe, however. For instance, seismic reflection profiling in the Danish and north German areas indicates that the original

(Ordovician) oceanic suture between Baltica and Avalonia dips at an angle of about 15 ~ through the crust, such that the crust of Baltica extends SW some 140 km beneath the crust of Avalonia (Bayer et al. 2002); the Elbe Lineament corresponds to the SW limit of Baltica at the Moho level. In Central Europe too, seismic reflection experiments suggest that younger Palaeozoic orogenic sutures associated with the accretion of the Variscide terranes of the 'Armorican Terrane Assemblage' (Tait et al. 1997) to the EEC also dip at moderate angles through the crust. The relationship between the inclined orogenic sutures and the steep lineaments is complex, and the latter cannot have developed solely as a consequence of crustal-scale reactivation of the former. Rather more likely is reactivation at the lithospheric scale, with the fundamental differences in the lithospheric properties of 'Old' and 'Young' Europe resulting in a variety of reactivation styles along the various early formed lineaments. The crust of Europe to the SW of the EEC comprises a mosaic of orogenic terranes accreted throughout Phanerozoic time (Ziegler 1982, 1990), 'traditionally' (i.e. in the 20th century) regarded as developing during a series of distinct orogenic cycles, notably the Caledonian, Variscan and Alpine orogenic cycles. The application of modern methods of analysis, in particular high-precision radiometric techniques, has revealed that the traditional interpretation is too simplified: each 'orogeny' comprises a number of distinct deformation phases, corresponding to the terrane evolution described above; specifically to phases of ocean destruction, terrane docking (or 'soft collision'), collision (between larger crustal blocks) and dispersal. There is overlap between some cycles, particularly in late Palaeozoic time, when rifting of Gondwana proceeded virtually without interruption (Stampfli & Kozur 2006). In some places (e.g. the Rheno-Hercynian Basin), the continuous record of synorogenic clastic sedimentation from Frasnian to late Westphalian appears to indicate a continuum of deformation, rather than discrete phases (W. Franke, pers. comm.). The principal evidence for the identity (and duration of existence) of individual terranes comes from the uniqueness (endemicity) of

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 291-311. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Revised basement tectonic sketch map of the TESZ and adjacent areas. (Compare with fig. 4.1 of Gee & Zeyen (1996) and fig. 1 of Pharaoh (1999).) Revisions incorporate information from Matte et al. (1990), Dallmeyer et al. ( 1995, 1999), Franke (1995b), Geluk (1997), Bula et al. (1997), Seghedi (1998), Franke & Zelainiewicz (2002), Verniers et al. (2002) and Winchester et al. (2002). Oceanic sutures, filled ticks; orogenic frontal zones, open ticks. Post-Palaeozoic basins and platforms: ADB, Anglo-Dutch Basin; ADF, Alpine Deformation Front; MNSH, Mid-North Sea High; NDO, North Dobrogea Orogen; NGB, North German Basin; POT, Polish Trough; RFH, Ringk0bing-Fyn High; RG, RCnne Graben; RMFZ, R0mr Fracture Zone; SP, Scythian Platform. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: BST, Bruno-Silesian (Brunovistulian) Terrane; BT, Bohemia Terrane; DSHFZ, Dowsing-South Hewett Fault Zone; EL, Elbe Lineament; EMT, East Moesian Terrane; IMF, Intra-Moesian Fault; KLZ, Krak6w-Lubliniec Zone; LRL, Lower Rhine Lineament; LT, Lysogory Terrane; MT, Matopolska Terrane; MDT, Moldanubian terranes; MST, Moravo-Silesian Terrane; NT, Normannian Terrane; PO, Palazu Overthrust; SNSLT, Southern North SeaLtineberg Terrane; SGF, Sfantu Gheorghe Fault; TT, Tulcea (North and Central Dobrogea) Terrane; WMT, West Moesian Terrane. Proterozoic-Palaeozoic tectonic elements: ABDB, Anglo-Brabant Deformation Belt; AD, Ardennes Massifs; AM, Armorican Massif; BB, Brabant Massif; BM, Bohemian Massif; CBT, Central Brittany Terrane; CDF, front of Caledonian deformation (see text for explanation); CM, Cornubian Massif; CPPR, Central Polish Palaeo-Rift; DR, Drosendorf Unit (of BM); EA, Ebbe Anticline; EEC, East European Craton; EFZ, Elbe Fault Zone; GF, Gf6hl Unit (of BM); HM, Harz Mountains; HCM, Holy Cross Mountains; LF, Loire Fault; L-W, Leszno-Wolsztyn Basement High; MC, Massif Central; MMC, Midlands Microcraton; MGCH, Mid-German Crystalline High; MH, Mazurska High; MN, Mtinchberg Nappe (of BM); MO, Moldavian Platform; NASZ, North Armorican Shear Zone; NBT, North Brittany Terrane; Pom, Pomerania; PP, Pripyat Trough; RM, Rhenish Massif; USM, Upper Silesian Massif (=MST); SH, South Hunsrtick; SNF, Sveconorwegian Front; SASZ, South Armorican Shear Zone; S-TZ, Sorgenfrei-Tornquist Zone; TB, Teplfi-Barrandian Basin (of BM); T-TZ, Teisseyre-Tornquist Zone; UM, Ukrainian Massif; VF, Variscan Front.

their palaeo-faunas and palaeo-floras (e.g. Cocks & Fortey 1982); and the principal evidence for their magnitude and direction of motion and rotation comes from palaeomagnetic constraints (e.g. Torsvik 1998). These topics are discussed in more detail elsewhere in this volume. It is relatively simple to establish faunal endemicity for the platform successions of palaeocontinents, from abundant shelly macrofauna present in little-deformed shelf sedimentary strata. It is much more difficult to do this for the accretionary margins of terranes, or indeed suspect terranes, frequently poorly exposed or known only from boreholes, comprising strongly deformed sequences of deep-water strata with sparse (if any) macrofauna. Palaeomagnetic studies of terrane margins suffer similar geological constraints. There is therefore vigorous debate about the status of many of the terranes reviewed here. Another important matter of debate is the provenance of individual terranes. The crystalline basement of most terranes in the

TESZ has historically been referred as being of 'Cadomian', 'Avalonian' or occasionally 'Pan-African' affinity, usually based on rather sparse and imprecise radiometric data. In the literature, usage of these terms has often been applied inconsistently, to lithofacies rather than to precisely dated rock suites. The increasing availability of large numbers of precise U - P b zircon ages has revolutionized the study of terrane provenance. Thus, although there is an abundance of Neoproterozoic zircon grains dated at about 600 Ma, other peaks in the grain population have been used to attribute sources in Eastern (North Africa) or Western Gondwana (Northern A m a z o n i a - G u y a n a ) and in Baltica. Such attribution depends on the presence of a robust database for the Precambrian shield areas for comparative purposes, and the quality of the palaeogeographical reconstructions used. Combined with other evidence, such as palaeomagnetic evidence for palaeolatitude, sedimentological evidence for palaeoclimate, and

WESTERN ACCRETIONARYMARGINOF THE EEC lithological association, this type of analysis can be a very powerful technique. The Palaeozoic time scale used throughout is that published by McKerrow & Van Staal (2000). Accreted terranes of the western margin of the EEC P a l a e o z o i c terrane motions, accretion and deformation phases

Most of the terranes were derived by phases of rifting along the margins of the Gondwana palaeocontinent, which lay at high southerly latitude for much of Palaeozoic time (Torsvik 1998). A large ocean, Iapetus, opened during late Neoproterozoic time, separating Gondwana from other large relics of the RodiniaPannotia supercontinent (Dalziel 1991, 1997), such as Laurentia and Baltica. That part of Iapetus separating the terrane of Avalonia from Baltica is referred to as the Tornquist Sea (Cocks & Fortey 1982). The subsequent history of the newly rifted terranes varies, but typically involves northward transport as a result of the creation of new oceanic crust between the newborn terranes and Gondwana; the destruction of older oceanic crust lying between the newborn terrane and the EEC, principally by subduction; the episodic accretion of the Gondwana-derived terranes to the EEC margin; and finally, dispersal along the margin of the EEC by strike-slip displacement. The location of the litho-tectonic elements described below is indicated in Figure 1. The Bruno-Silesian Terrane may have been one of the first to leave Gondwana, and was certainly accreted to Baltica by the time of the late Cambrian Sandomierz deformation phase (Znosko 1974; Winchester et al. 2002; Nawrocki et al. 2004). Next, Avalonia left Gondwana in the early Ordovician, migrating from high to low southerly latitudes throughout remaining Ordovician time (Trench & Torsvik 1992), driven by the opening of the Rheic Ocean (Cocks & Fortey 1982) to the south of Avalonia ('ridge-push'), as well as by rapid destruction of the Iapetus Ocean to the north ('slab-pull') in a number of subduction systems. Closure of the Tornquist Sea segment of this ocean involved a significant dextral oblique component (Trench & Torsvik 1992; Oliver et al. 1993). Soft collision ('docking') of Avalonia and Baltica, producing Balonia (Torsvik 1998), occurred during the Shelveian Phase in Ashgill time (Samuelsson et al. 2002b), and is associated with amphibolite-facies metamorphism in the Mid-North Sea region (Frost et al. 1981; Pharaoh et al. 1995). Northward drift of Baltica in late Cambrian-early Ordovician time was accompanied by 55 ~ counter-clockwise rotation (Torsvik & Rehnstr6m 2001). Terranes in central Poland with a basement of supposed 'Cadomian' affinity and Acado-Baltic faunal association in the Cambro-Ordovician (e.g. the inferred M a t o p o l s k a Terrane) are the most controversial. They may have migrated from the vicinity of the southern Urals along the Tornquist margin of Baltica (Pharaoh 1999); or crossed the Iapetus Ocean from Gondwana to Baltica prior to mid-Cambrian time, as envisaged for the Bruno-Silesian Terrane (Betka et al. 2000, 2002); or they may always have been located close to their present position (Cocks 2002). A foredeep developed along the margin of Baltica in Silurian time (Dadlez et al. 1994; Berthelsen 1998) as a result of loading caused by the newly accreted crust. Subduction continued to the NW beneath Laurentia, leading to final closure of the Iapetus Ocean in Wenlock time (Leggett et al. 1979; Kneller et al. 1993), marked by the Scandian event in NW Scotland and the amalgamation of Laurussia (Ziegler 1990). The terranes now comprising the internides of the Variscan orogen and exposed in the Bohemian, Armorican and Iberian massifs, were located along the Gondwana margin at high southerly palaeolatitudes in late Ordovician time (Krs et al. 1986; Tait et al. 1995). Persistent plume-induced magmatism (Floyd et al. 2000) resulted in several phases of tiffing from the margin; rapid dispersal northward was driven by subduction of

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the Rheic Ocean (Franke 1995), and the opening of the Saxo-Thuringian and M a s s i f Central oceans closer to Gondwana. An early accretion of at least part of Saxo-Thuringian Armorica to Laurussia is recorded by emplacement of the Lizard Peridotite and the Acadian deformation phase, in Emsian time (Soper et al. 1987; Pharaoh 1999). Some parts of the 'Armorican Archipelago' (Franke et al. 1999) collided at c. 400 Ma, with significant HP metamorphism, predating their amalgamation into the TESZ. Provenance studies indicate that Iberia remained attached to Gondwana until early Devonian time (Martinez-Catal~in et al. 2004). Palaeomagnetic evidence (Tait et al. 1997, 2000) supports independent motion of distinct M o l d a n u b i a n and Perunica terranes until at least late Devonian time (c. 370 Ma) when collision with Saxo-Thuringian Armorica occurred. In early Carboniferous time, widespread HP metamorphism in the Variscan internides records rapid crustal thickening following closure of the Massif Central Ocean and collision with Iberia (Ziegler 1990). The terminal phase of collision between Laurussia (the Old Red Continent) and Gondwana, to produce Pangaea, is recorded by late Carboniferous (Variscan) and early Permian (Alleghenian) orogenic phases in Europe and America. The crust of the TESZ continued to undergo modification as a result of post-orogenic 'reordering' (Meissner 1989), Permian to Mesozoic basin development, rifting along the pre-Alpine Tethyan margin, Cimmerian inversion and subsequent Alpine-Carpathian thrusting, particularly in Romania.

Terrane analysis

The principles of terrane analysis (e.g. Coney et al. 1980) have been successfully applied to the TESZ (Franke 1990; Po~aryski 1990; Pogaryski et al. 1992). The principal characteristics of the major Palaeozoic terranes depicted in Figure 1 are briefly reviewed here. Geographical extent is described using present geographical coordinates, locations and distances, without palinspastic reconstruction. Provenance studies, using characteristic isotopic, geochemical or biostratigraphic assemblages, aim to identify the source continent of a rifted terrane. Of particular value are studies of detrital zircon suites using the single-crystal or SHRIMP (sensitive high-resolution ion microprobe) methods of U - P b isotopic analysis. The internal structure of a terrane is deduced from outcrop studies and seismic reflection data. Particularly informative for this purpose are images generated from geophysical potential fields (e.g. aeromagnetics and gravity; Banka et al. 2002). Examples of such maps, with a structural overview template, are presented in Figures 2 and 3. Of course, not all the features visible on these images are associated with the basement; for example, Permo-Carboniferous volcanic rocks locally cause magnetic anomalies, particularly close to the mid-North Sea rifts and in the North German Basin (Fig. 2), and the expression of younger sedimentary basins (which may represent extensional reactivations of basement structures) is clear in the Bouguer gravity image (Fig. 3). For further discussion of these topics the reader is referred to Banka et al. (2002), who have also listed and acknowledged the numerous sources of these data. Rifting history is typically determined by sequence stratigraphic studies, subsidence rates or magmatic episodes. The period of terrane isolation is most easily deduced from endemic faunal assemblages (e.g. see Cocks & Torsvik 2006), but also sometimes from isotopic signatures (e.g. Thorogood 1990; Samuelsson et al. 2002a). To determine the direction, rotation and rate of drift requires an excellent palaeomagnetic record, but can be done successfully (e.g. for Baltica in Cambro-Ordovician time; Torsvik & Rehnstr6m 2001). The history of ocean closure is deduced from arc-related magmatic suites. Timing and nature of terrane collision are deduced from the breakdown of faunal endemism, the arrival of orogenic flysch sediment at the foreland (e.g. Franke 2000), structural evidence and isotopic

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T.C. PHARAOHET AL.

Fig. 2. Colour shaded relief map of aeromagnetic potential field of the Western AccretionaryMargin of the EEC, after Banka et al. (2002), who listed the data sources. Colour scale ranges from red (+2000 nT) to green (0 nT) to blue (-700 nT). The localizedcontribution of magnetic sources shallowerthan the basement, (e.g. Permian volcanic rocks near the North Sea graben intersections), should be noted. Key to basin, platform, terrane and other tectonic elements as in Figure 1.

data. The geometry and location of terrane-bounding sutures is deduced by mapping all of the above characteristics, structural evidence, and seismic reflection and other geophysical data. Finally, the post-accretion history is deduced from structural evidence, sedimentology of overstep sequences, reworking of microfossils, isotopic age of stitching plutons, etc. T e r r a n e s a c c r e t e d in e a r l y P a l a e o z o i c t i m e Baltica. This

large palaeocontinent comprised the exposed Fennoscandian and Ukrainian shields and other parts of the concealed EEC, extending eastward from the North Sea to the Urals, and northward to the Arctic Ocean to include the Timanides. Its NW boundary lies within the Scandinavian Caledonides, where it was overthrust by terranes of Laurentian (North American) affinity (Fig. 1) in Silurian time. The long and complex evolution of the Precambrian crust of this palaeocontinent has been described by Bogdanova et al. (2006). The crust of Baltica is characterized by high-frequency aeromagnetic anomalies (Fig. 2), whose zonation bears witness to a complex Proterozoic accretion history (Banka et al. 2002; Williamson et al. 2002). A phase of mafic magmatism ('Older Dykes' and Volhyn Basalts) and tiffing ('Sparagmite Basins') at c. 600 Ma (Andr~asson 1998) reflects rifting from the rest of the RodiniaPannotia supercontinent (Dalziel 1991, 1997) in the Neoproterozoic. The Cambro-Ordovician palaeogeographical history deduced from biostratigraphic and palaeomagnetic data has been described by Cocks & Torsvik (2006). The whole palaeocontinent is thought to have rotated counter-clockwise by at least 100 ~ in late

Neoproterozoic to mid-Ordovician time, with at least 55 ~ of this in late Cambrian to early Ordovician time (Torsvik & Rehnstr6m 2001), a fact of considerable importance for palaeogeographical evaluation (Cocks 2002). Thin Cambro-Ordovician strata are of platformal type with rich shelly faunas. After collision with Avalonia in late Ordovician (Ashgill) time, when faunal isolation ended (Cocks et al. 1997), a rapidly subsiding foredeep developed along the SW margin of the EEC (Dadlez et al. 1994; Poprawa et al. 1999). Boreholes in Denmark (Vejbaek 1997), northern Germany (Katzung et al. 1993) and Poland (Dadlez 1982) prove up to 7 km of basinal Silurian strata. The latter were incorporated in a northward-vergent foreland thrust belt of Scandian (late Silurian) age. The so-called 'Caledonian Deformation Front' (Fig. 1) delimits the western edge of the autochthonous, lightly deformed Baltic platformal sequence. The Teisseyre-Tornquist Zone sharply truncates (Figs 2 and 3) NE-trending belts of granulites, anorthosites and granite-gneiss in eastern Poland and the Ukraine (Bogdanova et al. 1996). In Romania, the basement of the Moldavian Platform lithologically resembles that of the Ukrainian Shield, whereas that of the Scythian Platform comprises Neoproterozoic granitic and dioritic rocks (Neaga & Moroz 1987), atypical of the EEC. Subduction-related magmatism (other than ash-fall bentonites) is absent in Baltica; thus dominantly NE-directed subduction appears unlikely (Pharaoh 1999; Balling 2000); however, an alternative view has been given by Meissner et al. (2002). NE-dipping zones of reflectivity in the subcrustal lithosphere of Baltica may represent the relics of Proterozoic subduction or a late switch in Ordovician subduction polarity, or may post-date subduction entirely (Berthelsen 1998).

WESTERN ACCRETIONARYMARGIN OF THE EEC

295

Fig. 3. Colour shaded relief map of Bouguer gravity potential field of the Western Accretionary Margin of the EEC, after Banka et al. (2002), who listed the data sources. Rainbow colour scale ranges from red (+60 mGal) to yellow (0 mGal) to blue (-235 mGal). The local enhancementof 'basement grain' by basins resulting from post-orogenic extensionalreactivation should be noted. (See Banka et al. (2002) for further discussion). Key to basin, platform, terrane and other tectonic elements as in Figure 1.

The Holy Cross Mountains (Fig. 1) in south-central Poland are the largest exposure of Palaeozoic rocks within the TESZ. Two groups of strata, the Lysogdry unit in the north, and the Kielce unit, part of the largely concealed Matopolska Massif in the south, are separated by the WNW-trending Holy Cross Fault. The relationship of these units to the EEC during Cambrian time has been the subject of much recent debate. Po~aryski (1990) and Franke (1995a) recognized distinct L y s o g d r y and M a t o p o l s k a t e r r a n e s on the basis of stratigraphic and structural contrasts between these units (and the EEC). The crystalline basement of both units is unknown, but the aeromagnetic anomaly map (Fig. 2) indicates that highly magnetic basement (of typical EEC type) is absent. Ediacaran (Vendian) silty and volcaniclastic rocks (Buta et al. 1997) resemble those of the adjacent EEC (Vidal & Moczydtowska 1995). A prominent latest CambrianTremadoc unconformity (Matopolska Massif) records folding and low greenschist-facies metamorphism during the Sandomierz deformation phase (Znosko 1974). The traditional view is that the Cambrian shelly fauna is diagnostic of Baltica (Dzik 1983; Bergstr6m 1984; Ortowski 1992). This was challenged by Belka et al. (2002), who claimed that although the Ordovician faunas are certainly Baltican, the Cambrian faunas of the Lysogdry unit are unknown in Baltica, and those of the Matopolska Massif are dominantly Avalonian in aspect. Cocks (2002) pointed out that as Avalonia existed as a separate entity only in the Ordovician, such affinity cannot be assigned in the Cambrian. From a study of detrital zircon and muscovite ages, Betka and colleagues (Betka et al. 2000, 2002; Valverde-Vaquero et al. 2000) concluded

that the Matopolska Massif (and less certainly, the Lysogdry unit) was detached from Gondwana in early Cambrian time and accreted to Baltica by the end of the Cambrian. However, the recognition of Neoproterozoic basement in the southern Uralides (Glasmacher et al. 1999) and in the Scythian Platform suggests that the presence of 'Cadomian' age detritus is not necessarily diagnostic of a Gondwanan provenance (Winchester et al. 2002). Unrug et al. (1999) have speculated that the Matopolska Massif may have formed the accretionary wedge to the BrunoSilesian Terrane. Detrital mineralogical and biostratigraphic evidence apparently do not support such a linkage, however (Belka et al. 2002; Cocks 2002). In the Holy Cross Mountains (as in the rest of Baltica) a highly condensed Arenig-Lower Silurian carbonate-clastic sequence contains many bentonites. The Upper Silurian sequence comprises up to 1500 m of greywackes deposited on the EEC foredeep (Dadlez et al. 1994; Berthelsen 1998) and strongly affected by Scandian phase deformation (Tomczyk 1980; Dadlez et al. 1994). Palaeomagnetic evidence indicates that the Matopolska Massif may have been displaced dextrally along the TTZ (Lewandowski 1993), but the faunal evidence constrains any displacement to a maximum of a few hundred kilometres (Cocks 2002). Another possibility is that the Polish Trough may have a more ancient antecedence than hitherto realized. It might have been initiated as a rift (referred to as the 'Central Polish Palaeo-Rift' in Figs 1, 2, 3 and 6) controlled by the ancestral TTZ, during the Neoproterozoic break-up of the Rodinia-Pannotia supercontinent, analogous to rift structures seen elsewhere in

296

T.C. PHARAOHETAL.

the EEC (Puchkov 1998). Cambrian strata were deposited on the passive margin of the EEC (see Fig. 6) before undergoing at least partial inversion in the late Cambrian Sandomierz phase, as recognized in the Holy Cross Mountains. Subsequently, the rate of subsidence matched that of the rest of the EEC. Only the NW part of this proposed palaeo-rift then subsided in Carboniferous, Permian and Mesozoic time to form the Polish Trough. Bruno-Silesian Terrane. Also known as the Brunovistulian Terrane (Dudek 1980; Aleksandrowski & Mazur 2002; Nawrocki et al. 2004), this terrane lies at the eastern edge of the Bohemian Massif in SW Poland and the NE Czech Republic (Fig. 1), and is concealed towards the east by the Carpathian Orogen and its foreland basin. It has traditionally been regarded as a component of the Rheno-Hercynian Zone of the Variscides, whose Devonian-Carboniferous evolution it strongly resembles (W. Franke, pers. comm.). It comprises two main sub-regions: the 'Brunovistulian' Block (Dudek 1980), extending south to Brno and the Krems-Vienna Line, and a western part, reworked into parautochthonous Variscan nappes. The concealed Upper Silesian Massif is separated from the Matopolska Massif by the Krakdw-Lubliniec tectonic zone, a narrow (about 0.5 km wide) belt of polyphase ductile deformation and magmatic intrusion, representing a possible terrane boundary (Dadlez et al. 1994; Buta et al. 1997). Deep boreholes prove metasedimentary rocks dated at 610-580Ma and granites emplaced at c. 585Ma (Finger et al. 2000; Belka et al. 2002). Amphibolites from the Rzeszotary Horst have been dated at 2.5 Ga by the U-Pb SHRIMP method (Bylina et al. 2000). These characteristics indicate that the basement of this terrane probably has a Gondwanan provenance (Friedl et al. 2000). Finger et al. (2000) correlated it with Avalonia, but a distinct gap in the range of detrital ages at about 570-590 Ma in the latter terrane (Murphy et al. 2004) suggests that such a correlation is unlikely, as does the presence of Acado-Baltican trilobites (Orlowski 1975; Buta et al. 1997) in gently deformed Cambrian strata. Nawrocki et al. (2004) argued that the presence of the endemic taxon S c h m i d t i e l l i u s p a n o w i in both the Matopolska Massif and the Bruno-Silesian Terrane supports their proximity in early Cambrian time. Acritarchs have closest affinity with the EEC (Jachowicz & Pf-ichystal 1998) and with Iberia (Moczydtowska 1995), but apparently not the rest of Gondwana. Detrital zircon and muscovite ages in Cambrian strata suggested a Gondwanan source to Friedl et al. (2000) and Belka et al. (2002). According to Cocks (2002), neither of the Holy Cross blocks were part of the same terrane as Bruno-Silesia; the faunal evidence on whether the latter terrane was derived from Baltica or Gondwana (or neither) is currently inconclusive. Middle-Upper Ordovician carbonates contain Baltican conodont faunas (Belka et al. 2002). Lower Devonian to Namurian A strata, resting with gentle discordance on lightly deformed early Palaeozoic strata, underwent strong Variscan folding and thrusting. Clear contrasts with the adjacent Matopolska Massif support the view that the Krakdw-Lubliniec tectonic zone is a terrane boundary (Franke 1990; Po2aryski 1990; Dadlez et al. 1994; Buta et al. 1997), along which the Bruno-Silesian Terrane was sutured to Baltica prior to early Devonian time (Nawrocki et al. 2004), although probably not at its current location (M. Lewandowski, pers. comm.). Another provenance could be a Neoproterozoic source in the Scythian Platform or the Uralian margin of Baltica (Puchkov 1998; Glasmacher 1999), which in Ediacaran time faced Gondwana. As a compromise, Winchester et al. (2002) suggested that it might represent a 'bridge' between Baltica and Amazonian Gondwana. Palaeomagnetic data suggest that in the early Cambrian, the Bruno-Silesian Terrane lay near the Equator, far from the Avalonian margin of Gondwana (Nawrocki et al. 2004) at 40-50~ and reached its present location on the Tornquist margin of the EEC before mid-Ordovician time. A further alternative, more 'fixist' model, proposed by Zelazfiiewicz et al. (2001), envisages development

of the terrane at the Tornquist margin of Baltica in Neoproterozoic time, although this is not supported by subsidence modelling (Poprawa et al. 1999) and other regional geological considerations (Nawrocki et al. 2004). Thus, the Bruno-Silesian Terrane is a suspect terrane in early Cambrian time and its provenance is hotly debated. Winchester et al. (2002, 2006) have suggested that, once accreted to Baltica, the Bruno-Silesian Terrane may have acted as an orogenic promontory on the EEC margin predating accretion of Avalonia in late Ordovician time. The 'Moldanubian Thrust' defining the western limit of the terrane (Fig. 1), and associated dextral transpression (see later discussion in the section ' M o r a v i a n S u t u r e ' ) , may have played an important part in the geometrical development of the postulated Variscide orocline in this region (Schulmann et al. 1991, 1995; Franke 1995; Franke & Zelainiewicz 2002). Avalonia. The eastern part of this microcontinent extends from

southern Ireland to the Mid-North Sea High, northern Germany and Poland. Avalonian faunas are recognized in Ordovician strata in Britain and Ireland, south of the Iapetus Suture, in the Rheno-Hercynian nappes north of the Lizard Thrust, in the Brabant and Ardennes massifs of Belgium, and in the Northern Phyllite Belt of Germany (Dallmeyer et al. 1995; Cocks et al. 1997). Only that part of Avalonia lying east of the Atlantic Ocean, referred to as Eastern Avalonia, is discussed here. It includes a heterogeneous Neoproterozoic basement comprising metamorphosed magmatic and sedimentary rocks generated in volcanic arcs and marginal basins (Thorpe et al. 1984; Pharaoh & Gibbons 1994) accreted initially to the Rodinia-Pannotia supercontinent, and following break-up of the latter, to protoGondwana, between 680 and 545 Ma. Xenocrystic zircons with ages of about 1.45 Ga (Tucker & Pharaoh 1991) and Nd isotopic studies (Noble et al. 1993; Nance & Murphy 1996) indicate possible involvement of 'Rondonian' type (Northern Amazonia and Guyana) crust, suggesting affinities with the South American (western) part of Gondwana (Murphy et al. 2000; Winchester et al. 2002). The Midlands Microcraton has a thin cover of lightly deformed lower Palaeozoic strata. Flanking deep-water basinal successions in Wales, northern and eastern England and Belgium obscure the Precambrian basement, and at the extremities of the microcontinent (e.g. in the Lake District, beneath the southern North Sea and northern Germany) the latter may be attenuated or absent. Here, the crust probably comprises juvenile lower Palaeozoicaccreted material. A calc-alkaline magmatic arc is traced from northern England to Belgium and is inferred to result from SW-directed subduction of Iapetus-Tornquist oceanic lithosphere beneath Avalonia (Noble et al. 1993; Pharaoh et al. 1993), possibly with a significant oblique component (Pharaoh 1999). The progressive change in age of volcanic onset from Wales (Tremadoc), to northern England (late Llanvirn), eastern England (Caradoc) and Belgium (Ashgill) and rotation of the Welsh Basin from arc to back-arc position (Kokelaar et al. 1984; Stillman 1988) through Ordovician time (Pharaoh 1999) is compatible with palaeomagnetic evidence for the counter-clockwise rotation of Avalonia (Piper 1997) with respect to this subduction zone (Pharaoh et al. 1995). Deformation is strongest in the Acadian (early Devonian) slaty cleavage arc developed in the basinal areas (Turner 1949; Soper et al. 1987; Van Grootel et al. 1997), contiguous with the Anglo-Brabant Deformation Belt (ABDB; Winchester et al. 2002) of eastern England and Belgium (Fig. 1), where, once again, a strong rotational component has been postulated (Verniers et al. 2002). Granite plutonism in northern England is of early Devonian age. An earlier, Shelveian (Ashgill) phase of deformation recognized in the Welsh Borders (Toghill 1992) is localized along major crustal lineaments (Pharaoh et al. 1995). This may correlate with the inferred late Ordovician phase of deformation affecting the Ardennes massifs in Belgium (Verniers et al. 2002). The parautochthonous Rheno-Hercynian nappes of

WESTERN ACCRETIONARYMARGINOF THE EEC Cornubia, the Ardennes in Belgium and central Germany (Fig. 1) represent the southern margin of Eastern Avalonia, extensively reworked by the Variscan Orogeny (Cocks et al. 1997; Verniers et al. 2002) and containing areally restricted, but significant, exposures of early Palaeozoic rocks (Franke 2000). 'Far Eastern Avalonia'. The existence of a distinct southern North Sea-Ltineberg Terrane has been proposed on the basis of geophysical criteria (Franke 1995a; Pharaoh et al. 1995) and more recently dubbed 'Far Eastern Avalonia' (Winchester et al. 2002). A possible terrane boundary with Avalonia proper may lie in the vicinity of the Dowsing-South Hewett and Lower Rhine Lineaments (Lee et al. 1993; Pharaoh et al. 1995). A small, perhaps marginal, oceanic basin may have been subducted here, giving rise to short-lived (Caradoc-Ashgill) volcanism in the ABDB (Verniers et al. 2002). Unfortunately, the available geophysical data provide little information on the internal structure or composition of this crust. One of the very few basement provings by deep boreholes in this region, the A/17-1 granite in the Netherlands sector, emplaced at 410 + 7 Ma (A. Gerdes, pers. comm.), contained no evidence of older crustal inheritance, compatible with the presence of only juvenile crust in this region. Boreholes on the Mid-North Sea and Ringk~bing-Fyn Highs penetrated metamorphic rocks of uncertain provenance. Whole-rock 4~ plateau ages indicate prograde greenschist-amphibolite metamorphism at 450-425 Ma, with retrogression at 415-400 Ma (Frost et al. 1981). The older age group has been interpreted as the age of docking or 'soft collision' of Avalonia with Baltica (Pharaoh et al. 1995; MONA LISA Working Group 1997b; Torsvik 1998; Pharaoh 1999), during the Shelveian deformation phase, an interpretation supported by biostratigraphic evidence for the reworking of microfossils (Samuelsson et al. 2002b). Gneisses in the Hunsrtick (SE Rhenish Massif) within the Rheno-Hercynian Zone (Fig. 1), yield U - P b ages in the range 560-574 Ma (Baumann et al. 1991) and may represent the only exposed Precambrian basement in this terrane (Winchester et al. 2002). Near Rtigen Island, in the southern Baltic (Fig. 1), deep boreholes encounter thrust and deformed anchizonal graptolitic greywackes of Ordovician age (Katzung et al. 1993) in the presumed hanging wall of the Thor Suture. They yield early Ordovician acritarchs similar to those of the English Lake District (Servais & Molyneux 1997). Many species are common to other periGondwanan areas in early Ordovician time (e.g. Spain, Bohemia and the Taurides of Turkey). Lithologically similar rocks and fossil assemblages are found in the Skibn6 Borehole in Pomerania (Cocks 2002; Samuelsson et al. 2002b). As acritarchs are planktonic, they cannot definitively indicate palaeocontinental affinity (Cocks & Verniers 2000), but they do suggest that terranes of probable Avalonian affinity extend eastward as far as northern Poland. It should be noted that Figure 1 represents a significant modification of earlier mapping (e.g. Pharaoh 1999) in this regard. Ashgill strata in the G14 Borehole lying just to north of the Thor Suture contain reworked Llanvirn microfossils with clear Gondwanan affinities (Samuelsson et al. 2002b), thus constraining the docking event to a period of about 10 Ma. Detrital muscovites from these same strata yield an 4~ plateau age of about 609 Ma, compatible with a Neoproterozoic provenance of Gondwanan affinity (Dallmeyer et al. 1999). There is also a significant detrital contribution from an unknown, immature volcanic arc (Giese et al. 1994; McCann 1998). Geophysical evidence, primarily from aeromagnetic data, supporting the presence of such a 'lost arc' concealed within 'Far Eastern Avalonia', was presented by Williamson et al. (2002). The Rtigen sequence was subsequently overthrust onto Silurian strata of the EEC foredeep during the Scandian (late Silurian) deformation phase. Dallmeyer et al. (1999) inferred that the Loissin-1 Borehole proved a culmination of the EEC (Pharaoh 1999; Fig. 1), emphasizing the low dip angle of the suture (see Fig. 5).

297

T e r r a n e s a c c r e t e d in late P a l a e o z o i c t i m e The Armorican Archipelago. The Rheic Suture represents a fundamental divide within the Variscan Orogen, separating crust reworked into a foreland thrust belt along the southern margin of Avalonia (Rheno-Hercynian or Externide Zone) from crust of the Internide Zones, which separated from Gondwana after the Ordovician (Dallmeyer et al. 1995). The Variscan internides have long been referred to as the Saxo-Thuringian and Moldanubian Zones (Kossmat 1927). These are more complex than the Rheno-Hercynian Zone (RHZ), incorporating ancient crust rifted from Gondwana at high southerly palaeolatitudes in early Silurian (Ziegler 1990) to earliest Devonian time (Paris 1998; MartfnezCatalan et al. 2004). Significant plume-related magmatism from about 500 Ma initiated tiffing at the Gondwana margin (Floyd et al. 2000; Crowley et al. 2002b) and facilitated generation of a progression of terranes referred to as the Armorican Terrane Assemblage (Tait et al. 1997) or Armorican Archipelago (Franke et al. 1999). The first convergence of the internide terranes with Laurussia is believed to have occurred in late Silurian-early Devonian time (Cocks & Fortey 1982), and was a possible cause of the Acadian deformation phase found throughout Eastern Avalonia (Soper et al. 1987). Subsequent collisions gave rise to the various phases of the Variscan Orogeny.

Saxo-Thuringian terranes. The Saxo-Thuringian Zone (STZ) as defined by Kossmat (1927) can be traced from SW Poland (Fig. 1), through central Germany and northern France, possibly extending as far as the Man of War rocks off the Lizard in SW England (Sandeman et al. 1997). The only good exposure is in the northern part of the Bohemian Massif, however, Franke (2000) recognized the following components: Franconia Terrane; Vesser Rift Basin; Saxo-Thuringia Terrane; the Saxo-Thuringian Ocean Basin. The microcontinental terranes comprise Neoproterozoic basement of 'Cadomian' affinity, largely greywackes and granitoid intrusions (Hammer et al. 1998; Linnemann et al. 1998); Cambrian-early Ordovician shallow marine clastic strata; bimodal magmatism at 500-480 Ma (Fumes et al. 1994; Sandeman et al. 1997; Krrner & Hegner 1998; Floyd et al. 2000); mid-late Ordovician hemipelagic shales, turbidites and glacigene strata (Erdtmann 1991); and pelagic shales, cherts and carbonates of Silurian to mid-Devonian age. The MidGerman Crystalline High, associated with a prominent aeromagnetic anomaly in Figure 2, represents the active magmatic margin resulting from southward subduction of ocean crust in Silurian-Devonian time (Franke 1998). Late Devonian-Visran flysch was fed NW from the developing orogenic belt. The apparent lateral continuity of the zone (Fig. 1) favours interpretation as (one or several) forearc or arc terranes accreted to a number of microcontinental terranes now forming the Moldanubian Zone (Ziegler 1990). Palaeomagnetic data (Krs et al. 1986; Tait et al. 1995) support derivation of the STZ terranes from Gondwana at high southerly palaeolatitudes after late Ordovician time. Suspect terranes ( s e n s u Coney et al. 1980) are found at the northern and southern margins of the STZ: the Lizard Peridotite (Clark et al. 1998) and Giessen Ophiolite (Franke 1995) formed at c. 397 Ma (early Devonian), either as marginal basins bordering the Rheic Ocean (Ziegler 1982, 1990) or as relics of Rheic mid-ocean ridge (Franke 2006). In the south, the Cambro-Ordovician Marifinsk~ Lfizn~ Complex (St~drfi 1999; Crowley et al. 2002a) and Silurian Sl~2a-Klodzko Ophiolite (Oliver et al. 1993; Floyd et al. 2002), are interpreted as accreted relics of the Saxo-Thuringian Ocean. A prolonged period of accretional thickening in the Sudetes, associated with HP metamorphism at c. 380-365 Ma (Maluski & Pato~ka 1997; Marheine et al. 2002) was followed by rapid uplift and greenschistfacies retrogression to 340 Ma (Kryza et al. 1990). Late and posttectonic granitoids (e.g. Karkonosze) were emplaced at c. 333 Ma

298

T.C. PHARAOHETAL.

(Kr6ner et al. 1994). The granulites of the Saxonian Dome and Erzgebirge (see Fig. 8), which also experienced high-grade metamorphism at c. 340 Ma (Kr6ner & Hegner 1998), were tectonically emplaced (or intruded?) beneath the floor of the contemporaneous early Carboniferous flysch basin (Franke et al. 1999; Franke & Stein 2000). Bohemia Terrane. This terrane (Franke & Zelainiewicz 2000) corresponds to the northern part of the Perunica Terrane recognized by Havff6ek et al. (1994). The boundary of the Moldanubian Zone with the STZ is poorly exposed, except in the western part of the Bohemian Massif (Fig. 1). The upper levels of the Teplfi-Barrandian Unit comprise slightly deformed Neoproterozoic volcanic rocks overlain by a Cambrian-middle Devonian sedimentary cover (Chlupfi6 1993). Deformation is mainly of Variscan age. The Bohemia Terrane forms a crustal block separating NW- and SE-verging parts of the Variscides (Matte et al. 1990), reflecting the opposing polarity of subduction zones closing the inferred Saxo-Thuringian and Massif Central oceans (Fig. 1). Palaeomagnetic evidence indicates 140 ~ counterclockwise rotation of the terrane before late Devonian time (Tait et al. 1997), apparently independently of the STZ terranes. The high-grade part of the Teplfi Unit (Czech Republic) and Erbendorf-Vohenstrauss Zone (Germany) form the root zone to allochthonous nappe outliers of Moldanubian Zone rocks overlying the STZ (Franke 1989), represented by the Mfinchberg, Wildenfels and Frankenberg massifs in Germany. These comprise paragneiss and orthogneiss with ophiolitic protolith metamorphosed to eclogite facies at c. 400-380Ma (Gebauer & Grtinnenfelder 1979) with Nd model ages c. 100 Ma older. Pressures >25 kbar (Klemd et al. 1994; O'Brien, pers. comm.) indicate subduction of the narrow Saxo-Thuringian Ocean to >75 km depth beneath Bohemia (Franke 2000). The G6ry Sowie of the Polish Sudetes (Cymerman et al. 1997; Kr6ner et al. 1994) lie in a similar structural position and are metamorphosed to granulite facies. Further collisions with STZ arc terranes caused amphibolite-facies retrogression at 370 Ma (Timmermann et al. 2000). Further HP granulite metamorphism at c. 340 Ma affecting both intemide zones in central Germany may be a consequence of lithospheric delamination or crustal thickening (Franke et al. 1999), probably during collision with Laurussia. Nappe emplacement occurred in latest Visgan time c. 330-325 Ma (Franke 1998). Moldanubian Terrane. The Teplfi Unit is separated from the high-

grade Gf6hl and Drosendorf units of the Bohemian Massif by steep ductile shear zones on its western and southern sides (Rajlich 1987; Zulauf 1994). The structurally higher Gf6hl Unit comprises anatectic ortho- and paragneisses and felsic granulites. HP metamorphism at c. 400Ma (Pin & Vielzeuf 1983) was followed by widespread amphibolite-facies metamorphism at c. 340 Ma (as described above) and emplacement of late tectonic granites. An accretionary complex of imbricated early Palaeozoic (c. 480 Ma) oceanic crust and passive margin components, it is comparable to the Massif Central Terrane (Matte et al. 1990). The Drosendorf Unit, overthrust by the Gf6hl unit, comprises >6 km of pelitic metasediments with a Neoproterozoic-lower Palaeozoic protolith interpreted as a passive margin sequence, similar to that of the C6vennes-Vend6e Terrane (Matte et al. 1990). The presence of ophiolitic fragments within this ductile shear zone again indicates the likely presence of an oceanic suture here. Romanian Terranes. The Trans-European Suture Zone reappears

from beneath the Carpathian Orogen and its foreland basin in Romania (Fig. 1) NW of the Black Sea. In the Dobrogea region at least three and possibly four, distinct fault-bounded terranes, described below, are recognized within the Carpathian Foreland, all suspect with respect to the Scythian Platform and Moldavian Platform of the EEC. Cimmerian inversion structures control the

disposition of the pre-Mesozoic basement blocks, but do not represent simple reactivations of the original terrane boundaries. (1) Tulcea (North a n d Central Dobrogea) Terrane. In North and Central Dobrogea, north of the (pre-Jurassic) Palazu Overthrust (Visarion et al. 1979), the crystalline basement comprises dismembered ophiolitic and metasedimentary rocks interpreted as a Neoproterozoic accretionary complex (Seghedi et al. 1999); upper Ordovician-Devonian anoxic distal turbidites and radiolarian cherts occupy a younger accretionary prism near Tulcea (Seghedi 1998); Silurian distal shelf strata pass up conformably into Lower Devonian shelf clastic and carbonate strata. The northern part (North Dobrogea Orogen) of this terrane was strongly affected by the Variscan Orogeny, with NE-directed thrusting and granite intrusion (S~ndulescu 1984; Seghedi & Oaie 1995; Liszkowski et al. 1998; Seghedi 1998) although the orogenic front is poorly located (Banks 1997). Monazite ages suggest a late Carboniferous-early Permian age for amphibolite-facies metamorphism (Seghedi et al. 2003). The basement in Central Dobrogea comprises Neoproterozoic metaturbidites affected by 'Cadomian'-age folding (Kr~iutner et al. 1988, and references therein; Seghedi & Oaie 1995), and is regarded by many researchers as a terrane separate from North Dobrogea. The evolution of the pre-Variscan basement of the Tulcea Terrane shows closest parallels to that of the SaxoThuringian (southern margin) of the Rheic Ocean, described in an earlier section. Unconformably overlying upper Palaeozoic strata (Carapelit Fro) have significant volcaniclastic input from a calc-alkaline magmatic arc. In Permian-early Mesozoic time the terrane was dislocated, by rifting along the peri-Tethyan margin (aided by strike-slip along possible correlatives of the TeisseyreTornquist Zone, e.g. the Peceneaga-Camena and Sfante Georghe Faults, crust-penetrating structures associated with offset of the Moho), from the remainder of Variscan Europe (and the TESZ) during break-up of the Pangaea supercontinent (Ziegler 1990). (2) East M o e s i a n Terrane. The Moesian Platform (MesozoicCenozoic) extends SW from the Capidava-Ovidiu Fault in South Dobrogea towards the Carpathian Foreland in Bulgaria (Visarion et al. 1988), but the pre-Mesozoic terrane boundary is the concealed Palazu Overthrust. The Intra-Moesian Fault is the boundary with the West Moesian Terrane (see below). To the north of the latter, in what is referred to here as the East Moesian Terrane, a higher-grade basement of Archaean gneisses and Palaeoproterozoic banded iron formation, similar to that of the Ukrainian shield, is overlain by a low-grade Neoproterozoic volcano-sedimentary succession (Seghedi 1998), comparable with the Volhyn volcanic units. These poorly dated units are overlain by Cambro-Ordovician siliciclastic strata (Iordan & Spassov 1989). The lithostratigraphic similarities to the EEC of Poland and the Ukraine are therefore strong. Claims that mid-Cambrian trilobites show affinities with England, Bohemia and the EEC (Iordan 1999), or support a Baltican affinity (Rushton & McKerrow 2000) cannot at present confirm the faunal provinciality of East Moesia. These rocks are unconformably overlain (following a Llandovery hiatus) by largely pelitic upper Silurian-Lower Devonian strata of North Gondwanan (i.e. Armorican Terrane Assemblage; ATA) affinity (Vaida et al. 2005). Midddle Devonian coarse clastic strata are overlain by upper Devonian-lower Carboniferous carbonate platform strata thickening into a foredeep north of the Craiova High. Westphalian-Stephanian coal measures are unconformably overlain by Permian strata (Banks 1997). Thus it is possible that the East Moesian Terrane may have been displaced from the EEC in early Ordovician time and, following transcurrent displacement and intercalation within the ATA, reincorporated to the EEC in post-Variscan time. (3) West M o e s i a n Terrane. That part of the Moesian Platform west of the Intra-Moesian Fault is referred to here as the West Moesian Terrane (Fig. 1). The Precambrian basement here is overlain by a Palaeozoic sequence up to 6.5 km thick, and is poorly

WESTERNACCRETIONARYMARGINOF THE EEC known and undated. The basement reworked into the internal massifs of the adjacent southern Carpathian Mountains may provide clues to its affinity. The structurally lowest, supposedly parautochthonous, Danubian nappe complexes contain a metaplutonic and metasedimentary basement of Neoproterozoic age and 'Cadomian' affinity, considered by some to represent the exhumed basement (Sandulescu 1994) of the West Moesian Terrane. Arguments have been presented for both Armorican (the traditional view) and Avalonian provenance (Winchester et al. 2006) for this crust. In contrast, the higher, allochthonous Getic Nappes apparently show more certain ATA affinity (Iancu et al. 2006). It is apparent that the Moesian Terrane (as defined by Haydutov & Yanev 1997) had a complex Palaeozoic history, and is probably composite. The larger part of the Carpathian Foreland in Romania (Tulcea and West Moesian terranes) has a basement showing affinities with Gondwana and the ATA, and was probably accreted to the TESZ during the Variscan Orogeny. Tectonically interleaved is a rather narrow wedge of crust about 100 km wide, the East Moesian Terrane, apparently originating in the EEC (Cambrian), but showing increasingly Gondwanide affinity through Ordovician and Silurian time, which was probably reamalgamated with the EEC during or soon after the Variscan Orogeny. The affinities of West Moesia are even more controversial. Subsequent opening of the proto-Pannonian marginal basin in late Triassicearly Jurassic time (Banks & Robinson 1997) may have caused further dispersal of the newly amalgamated Moesian composite terrane along the EEC margin, as did mid-Cretaceous opening of the Black Sea Basin. Correlation of the terranes of the Moesian Platform with the Moravo-Silesian Terrane, which is in a similar structural position with regard to the EEC margin, has been proposed by Burchfiel (1975) and Matte et al. (1990), but is here considered unlikely because of the lithostratigraphic contrasts described above.

Peri-Tethyan dispersal of TESZ terranes The Variscide basement of the Western and Central Pontides in Turkey (Okay et al. 1994), was rifted away from the Moesian Platform in mid-Cretaceous time, following oblique slip along the TTZ, to form the Western Black Sea Basin. It is therefore clear that the TTZ crustal discontinuity also affected post-Variscan tectonics very significantly (C. Tomek, pers. comm.), during Mesozoic extension and Cimmerian and Alpine inversion. Further consideration of the evolution of terranes first accreted to, and then lost from, the Western Accretionary Margin of the EEC (e.g. the Zonguldak Terrane (?Avalonian affinity) and Sakarya and Eastern Pontide Blocks (?Armorican)), has been given by Winchester et al. (2006). The Alpine and Carpathian orogens (Fig. 1) contain numerous internal massifs (e.g. the Tatra Mountains) comprising crust reworked from the TransEuropean Suture Zone following the end of the Variscan Orogeny. Some were displaced from the margin of the Palaeotethys and Neotethys oceans (Stampfli et al. 2001), prior to Alpine thrust displacement. The peri-Gondwanan and Gondwanan affinities of such massifs have been demonstrated by detailed U-Pb zircon dating studies (e.g. Schaltegger et al. 1997; Von Raumer et al. 1999), but a detailed description of them is largely beyond the scope of this review. Further details may be found in the sections of this volume dealing with the younger orogens.

Orogenic sutures within the Western Accretionary Margin of the EEC The most significant terrane boundaries (Fig. 1) are sutures associated with destruction of oceanic lithosphere: the Iapetus Suture, separating Avalonia from the Laurentian terranes (beyond the

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scope of this review); the Thor Suture (Berthelsen 1998) separating Baltica and Avalonia, marking closure of the Ordovician Tornquist Sea; the R h e i c Suture (Cocks & Fortey 1982), separating the early Palaeozoic accreted terranes of Laurussia from an 'archipelago' of Gondwana-derived terranes accreted in late Palaeozoic time. The latter terranes were separated by seaways and larger ocean basins, such as the Saxo-Thuringian and Massif Central oceans (Matte et al. 1990), the closure of each being marked by the suture zones that are now summarized.

Iapetus Suture

The Iapetus Ocean (Harland & Gayer 1972) was initiated during the break-up of the Rodinia-Pannotia Supercontinent (Bond et al. 1984; Dalziel 1997) in late Neoproterozoic time. This event was preceded by the intrusion of tholeiitic dykes throughout the Scandinavian margin of Baltica (Andrrasson 1998) and the conjugate Laurentian margin in Newfoundland, Scotland (Tayvallich) and Ireland at c. 600 Ma. The 'Sparagmite Group' in Scandinavia (Kumpulainen & Nystuen 1985) and the volcanic-rich Volhyn strata of Ukraine and Poland (Moczydlowska 1997) are manifestations of this same phase of rifting. The Tornquist Sea (Cocks & Fortey 1982) is generally regarded as a segment of Iapetus separating Avalonia from Baltica only in Ordovician time (see below). The final closure of the Iapetus Ocean occurred in Silurian time (McKerrow et al. 1990). The location of the Iapetus Suture is well constrained within Britain and Ireland by deep seismic data (Freeman et al. 1988; Soper et al. 1992) and extends into the western part of the North Sea (Fig. 1). However, the thermo-mechanical effects of extension in the North Sea graben have so modified lower crustal reflectivity that it is difficult to identify the location of the suture farther east. Fichler & Hospers (1990) inferred a position on the East Shetland Platform west of the graben, but a location closer to the Norwegian coast (Fig. 1) is equally feasible (Pharaoh 1999).

T h o r Suture

This suture marks the closure of the Tornquist Sea separating Baltica and Avalonia in Ordovician time, and was originally recognized using faunal provinciality criteria (Cocks & Fortey 1982). It closed slightly earlier than the rest of the Iapetus Ocean, probably in late Ordovician time (Picketing 1989; Pharaoh et al. 1995). The apparent absence of any subductionrelated magmatism on the Baltica margin (other than bentonites representing ash-fall deposits) argues for subduction towards the SW. The near-surface location of the suture is constrained by deep boreholes on the Ringkcbing-Fyn High of Denmark and NE Germany, as described earlier (Fig. 1). A set of rather weak, SW-dipping, mid-crustal reflections observed on seismic profiles in the southern Baltic Sea has been correlated with the suture (BABEL Working Group 1993; Tanner & Meissner 1996; DEKORP-BASIN Research Group 1998, 1999). Similar features have also been identified west of Denmark (MONA LISA Working Group 1997a,b) where they maintain a rather constant dip (c. 10-12 ~ into the lower crust. The concept of the suture as a steep 'Trans-European Fault' (Berthelsen 1992b, 1993) is contradicted by evidence from normal-incidence and wide-angle seismic profiles on the North German Plain (EUGENO-S Working Group 1988; Krawczyk et al. 2002), which indicate the presence of a wedge of high-velocity, EEC-type basement in the lower crust as far south as Hamburg (Thybo et al. 1990; Aichroth et al. 1992; Rabbel et al. 1995). The suture is presumed to continue from the vicinity of Rtigen Island into northern Poland (Winchester et al. 2002) towards the Moravian Line marking the western limit of the Matopolska and Moravo-Silesian (=Brunovistulian) Terranes, which had Baltican affinity by mid-Cambrian time at the

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T.C. PHARAOHETAL.

latest. The Sorgenfrei-Tornquist Zone (STZ) has a long history of reactivation (Berthelsen et al. 1992a; Thybo 1997) and is associated with a Moho offset (EUGENO-S Working Group 1988; Kind et al. 1997; Cotte & Pedersen 2002) but does not appear to have a simple genetic link to the Thor Suture. Instead, it may represent a younger 'orogenic back-stop' or 'boundary detachment' defining the northern limit of Variscan crustal deformation (Berthelsen 1998), subsequently reactivated in the Alpine Orogeny. The Teisseyre-Tornquist Zone (TTZ) in Poland (Guterch et al. 1986; Krdlikowki & Petecki 1997) may have originated as part of the passive margin architecture of the EEC as early as the Neoproterozoic, and acted as a transcurrent fault zone for much of Palaeozoic time. It is unlikely to represent an oceanic suture, however (Pharaoh 1999).

2000) at the northern edge of the Bohemia Terrane (=Tepl~Barrandian Unit). Gf6hl Suture

To the south of the Tepl~-Barrandian Unit, which separates NWand SE-verging parts of the Variscides (Matte et al. 1990), a suture within the Moldanubian Zone dips to the north (see Fig. 8), with an opposite sense of vergence (and, presumably, subduction polarity) to that of the Saxo-Thuringian Suture. It is associated with accreted oceanic protoliths (e.g. pods of mantle peridotite and eclogite) in the Gf6hl Unit of the Bohemian Massif. It may represent a continuation of the inferred M a s s i f Central Suture (Matte et al. 1990).

R h e i c Suture

M o r a v i a n Suture

The Rheic Suture is a complex structure, and is almost certainly compound in nature (W. Franke, pers. comm.). The early history of the ocean is poorly known, but it may have begun to open in early Ordovician time with the rifting of Avalonia and the earliest terranes of the Armorican archipelago from Gondwana (Cocks & Fortey 1982; Pharaoh 1999). The Lizard Peridotite and Giessen Ophiolite appear to represent the only likely relics of Rheic ocean floor. Plagiogranite in the former crystallized at 397 _+ 2 Ma (Clark et al. 1998) in early Devonian time. This age is identical within error to a white mica age (Ar-Ar laser microprobe on Acadian cleavage pressure fringes) of 396.1 __ 1.4 Ma obtained by Sherlock et al. (2003). The ophiolites may represent relics of original marginal-basin crust associated with northward subduction at the northern margin of the ocean (Ziegler 1990), apparently contemporaneous with Acadian deformation farther north in Avalonia. If this subduction was at a low angle, it might explain the transfer of compressional stress deep into the interior of Avalonia to produce Acadian deformation, as well as the apparent absence of subduction-related magmatism. Obduction of the Lizard complex, associated with overthrusting by the Normannian Complex, began at about 380 Ma (Clark et al. 1998). Subsequently, in late Devonian to early Carboniferous time, a smaller Rheno-Hercynian Ocean Basin may have opened not quite coincident with the original Rheic Suture (Franke 2000). The Mid-German Crystalline High (Dallmeyer et al. 1995; Franke 2000) is a magmatic arc produced by later southward subduction of this ocean, and it is this later phase of suture development that is imaged by deep seismic reflection profiling. Profiles in the English Channel (Leveridge et al. 1984) show that the suture maintains a constant southward dip of c. 20 ~ into the lower crust and suggest that the Variscan orogen is distinctly thick-skinned in aspect. The DEKORP-2 profile (see Fig. 8) provides a complete transect across the Variscides in Germany. The Rheic Suture Zone is imaged as several SE-dipping reflector zones in the mid-crust, interpreted as thrusts (Meissner & Bortfeld 1990). In North Dobrogea in Romania, terranes showing affinity with the Armorican Terrane Assemblage, including Variscan-age amphibolite-facies metamorphism (Seghedi et al. 2003) directly abut the EEC, and the Rheic Suture probably lies close to the Galati-St. George Fault. Its original geometry has, however, been severely modified by post-Variscan events in this region.

On the eastern flank of the Bohemian Massif, the Gf6hl Unit is in tectonic contact with the Bruno-Silesian Terrane, along a ductile shear zone containing ophiolitic fragments (Schulmann et al. 1991) long referred to as the 'Moldanubian Thrust'. Structures along this line show highly oblique (dextral sense of shear) overthrusting to the east in early Carboniferous times (Schulmann & Gayer 2000). This geometry is inferred to reflect the northward convergence of the Armorican Terrane Assemblage with Laurussia, obliquely against the orogenic promontory on the EEC formed by the former Bruno-Silesian Terrane (Banka et al. 2002; Winchester & PACE TMR Network 2002). However, the ophiolitic fragments are of uncertain age and might be derived from older Giessen, Mfinchberg or Gf6hl-type protoliths. An alternative view (Franke 2006) is therefore that the inclusion of the ophiolitic fragments is fortuitous, and the shear zone does not represent a suture. Furthermore, the Moldanubian Thrust truncates three different terranes along its western edge, and studies of the flysch provenance suggest that it is of late tectonic age. The extension of the 'Moravian Line' toward the NE has also been invoked as the eastern boundary of 'Far Eastern Avalonia' (Winchester & PACE TMR Network 2002). All of the sutures within the ATA, described above, reflect closure of (perhaps small) ocean basins that originally separated the Gondwana-derived elements of the Armorican Archipelago (Franke 2000). Faunal evidence (McKerrow et al. 2000) indicates that no large (i.e. > 1 0 0 0 k m wide) oceans existed in late Palaeozoic Europe.

S a x o - T h u r i n g i a n Suture

This SE-dipping suture forms the root zone of the Mfinchberg Nappe, and is well imaged by seismic reflection data, both by the DEKORP-4 profile at the western margin of the Bohemian Massif (Fig. 1) and the DEKORP-2 profile, which is incorporated in Crustal Transect 5 (see Fig. 8). It is marked by early Palaeozoic mid-ocean ridge basalts (Dallmeyer et al. 1995; Linnemann et al.

Crustal transects through the Western Accretionary Margin of the EEC The crustal structure of the Western Accretionary Margin of the EEC is illustrated by a number of transects, shown in Figures 4 - 8 . These attempt to build on the cross-sections presented in the European Geotraverse (Blundell et al. 1992) by incorporating information from geophysical experiments carried out in the past 15 years. Transect 1 (Fig. 4) extends some 1200 km from the English Midlands to the Stockholm archipelago. The crustal structure is entirely concealed in this region by Carboniferous and younger strata of the Southern North Sea Basin, thickest in the Central Graben, Horn Graben and Norwegian-Danish Basin. The core of the transect is based on interpretations of deep seismic reflection profile SNST 83-07 and refraction profile EUGENO-S2, tested by modelling of the gravity and magnetic potential fields by Williamson et al. (2002). The most significant crustal boundary in this transect is the Thor Suture between crust of Avalonian affinity (in the SW) and crust of Baltican affinity to the NE. A SW-inclined, poorly reflective zone recognized on the MONA

WESTERN ACCRETIONARYMARGIN OF THE EEC

301

Fig. 4. Transect 1, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC in the central North Sea-southern Scandinaviaregion; 5 x vertical exaggeration. Based on an interpretation of deep seismic reflectionprofile SNST 83-07 (Klemperer& Hobbs 1991) and refraction profile EUGENO-S2 (EUGENO-S Working Group 1988). MR, mantle reflectors shown schematicallyfollowingBlundell et al. (1991). Slightly modifiedafter Williamsonet al. (2002). Post-Palaeozoic basins and platforms: CG, Central Graben; HG, Horn Graben; NDB, Norwegian-Danish Basin; RFH, Ringkcbing-Fyn High; RG, R0nne Graben. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: DSHFZ, Dowsing-South Hewett Fault Zone; EEC, East European Craton; ?SNSLT, Inferred Southern North Sea-Luneberg Terrane; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: ABDB, Anglo-BrabantDeformation Belt (=Eastern EnglishCaledonides); 'CDF', front of Caledoniandeformation (see text for explanation);EEC, East European Craton; MMC, MidlandsMicrocraton; SFD, Svecofennian Domain; SND, SveconorwegianDomain; SNF, SveconorwegianFront; S-TZ, Sorgenfrei-TornquistZone; TSB, Trans-ScandinaviazlBatholith;VA, concealed volcanic arc inferredfrom magnetic signaturein Southern North Sea (Williamsonet al. 2002). For mantle reflectors: key boreholes (in red): G1, Glinton;Gr, Grinsted;NC, North Creake; No, Novling; Ro, R0nne; WF, WithycombeFarm.

LISA deep seismic profiles across the RingkCbing-Fyn High has been attributed to the suture (MONA LISA Working Group 1997a). This hypothesis was supported by the modelling work of Williamson et al. (2002), which used the fundamental contrast in magnetic susceptibility of these two types of crust to map the location and geometry of the suture. This analysis indicated that the suture dipped SW at an average angle of about 14 ~ being steeper in the upper crust (up to 40 ~) than in the lower crust (average 7~ This interpretation also invoked tectonic imbrication of platform cover and crust in the footwall of the suture, with strong deformation dying out eastwards towards a 'Caledonian Deformation Front' some 100 km east of the suture. Within the crust of the EEC, strong magnetic contrasts were not found across the Sorgenfrei-Tornquist Zone, suggesting that it does not bound crust of fundamentally different type. MONA LISA Working Group (1997a) also reported inclined zones of reflectivity in the mantle beneath the crustal suture, with both SW and NE dips. Pharaoh (1999) interpreted the SW-dipping set as being related to Avalonia-Baltica collision, invoking lithospheric delamination at or close to the Moho to explain the observed offset of about 200 km (Fig. 4). The crust of Avalonia is laterally heterogeneous. In the SW, crust of the Midlands Microcraton, known from exposures and deep boreholes, virtually unaffected by Caledonian deformation, passes towards the NE into the AngloBrabant Deformation Belt (----Eastern English Caledonides), which experienced much stronger Caledonian deformation. In the west, SW-vergent thrusting in the shallow crust involves the Precambrian basement, and may be of Acadian age. In the mid-crust, SW-dipping zones of inclined reflectivity described by Reston (1990) and Blundell (1993) may have developed during the earlier, Shelveian, 'soft' collision between Avalonia and Baltica (Pharaoh et al. 1995). The Dowsing-South Hewett Fault Zone and Lower Rhine Lineament (Fig. 1) may have been initiated at this time, as a suspect suture between Eastern Avalonia s e n s u s t r i c t o and the poorly known crust of the Southern North

Sea region (Far Eastern Avalonia), for which separate terrane status has been invoked (Franke 1995; Pharaoh et al. 1995; Pharaoh 1999). Williamson et al. (2002) invoked the presence of a concealed volcanic arc within this inferred terrane to explain the linear magnetic anomaly lying along the southern flank of the Mid-North Sea High and R y n k c b i n g - F y n High, clearly seen in Figure 2, possibly the 'lost arc' generated by subduction of the Tornquist Sea. Transect 2 (Fig. 5) runs SSW from the Harz Mountains in Germany, across the North German Plain towards the Baltic Sea in the NNE. The interpretation presented is slightly modified from that presented for the deep seismic reflection profiles DEKORP-BASIN 9601 (onshore) and DSB-9 (offshore extension) by Lassen et al. (2001) and Lassen (2005). In this region, the crystalline basement is almost completely concealed by the thick Permian-Cenozoic fill of the North German Basin. The most important crustal structure is the Thor Suture between Avalonia in the SW and Baltica in the NE. The DEKORP profile clearly images the high-velocity crystalline basement of the EEC extending for 165km as a SW-tapering wedge beneath the North German Basin (Bayer et al. 2002; Lassen 2005) to about the location of Hamburg. The average dip of the suture is about 10 ~ The tip-line of this wedge (at the Moho level) is apparently juxtaposed with the postulated Elbe Lineament, inferred to separate crustal regions with distinct basement tectonic grains and geophysical attributes, referred to as the Pompeckj and Holstein blocks (Aichroth et al. 1992; Rabbel et al. 1995), and a possible suspect terrane boundary (e.g. Franke 1995b; Rabbel et al. 1995; Tanner & Meissner 1996). In the interpretation presented here, the Elbe Lineament does not represent the SW boundary of the EEC, which lies at the base of the crust here. The lineament may represent a postcollisional structure in the upper crust focused at the boundary between contrasting lithospheric types; or it may represent a boundary between contrasting types of upper crust (e.g. Variscan

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T.C. PHARAOHET AL.

Key to Figs 4-8.

and Caledonian deformed crust), in which case the geographical coincidence with the lower crustal wedge of Baltica is fortuitous. Various 'interwedging' structures in the lithosphere of this region are attributed to Caledonian compressional tectonic phases (Meissner et al. 2002). Anorthosite xenoliths of shield type entrained in Permian lavas encountered by the Schwerin-1 Borehole (K/impf et al. 1994; Breitkreuz & Kennedy 1999), and isotopic data from the Loissin Borehole (Dallmeyer et al. 1999), support this geophysical interpretation. No attempt is made to depict the crustal structure of the EEC in this profile, which is poorly constrained by seismic refraction data. Other features of the EEC shown here are the presence of inferred half-grabens infilled by Ediacaran strata (Lassen 2005), probably generated during the rifting and break-up of Rodinia-Pannotia; and Cambro-Ordovician strata of the Baltica rifted passive margin

(Scheck et al. 2002) thinning eastward. The latter are known from deep boreholes to the north of Rtigen Island (Katzung et al. 1993; Vejbaek 1997). The carbonaceous Alum Shale is a likely tectonic detachment horizon, forming, as in the Scandinavian Caledonides, an important geophysical marker: a structural boundary ('O-Horizon') on seismic reflection profiles; and a zone of high conductivity on resistivity profiles (ERCEUGT Group 1992). On Rfigen Island, several boreholes penetrate deep-water strata of Ordovician age, which have been assigned an Avalonian affinity (Verniers et al. 2002). These, together with boreholes further east in Pomerania, and to the west in Heligoland (Frost et al. 1981), represent rare provings of the crust of the suspect terrane of 'Far Eastern Avalonia'. Little detail is shown within the largely concealed Caledoniandeformed crust of Avalonia presumed to underlie much of the

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303

Fig. 5. Transect 2, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC beneath the North German Basin; 5 x vertical exaggeration. Modified after interpretations of seismic reflectionprofiles DEKORP BASIN 9601 (onshore) and DSB-9 (offshore) by Lassen et al. (2001) and Lassen (2005). Post-Palaeozoic basins, platforms and geographical features: ADF, AlpineDeformationFront; GT, GardelegenThrust; NGB, North German Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: EL, Elbe Lineament;TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: CDF, front of Caledoniandeformation(see text for explanation);EEC, East European Craton; HM, Harz Mountains; SHT, Sub-HercynianTrough; VF, VariscanFront. Key boreholes (in red): G14, G14; Lo, Loissin; Pz, Pritzwalk; Ru, Rtigen 5, Other geographical locations mentioned in text: BS, Baltic Sea; ER, Elbe River.

North German Plain. Towards the south, several south-dipping reflector sets have been correlated with Variscan thrusts within the Rheno-Hercynian Zone. Some of these structures (e.g. the Gardelegen Thrust) were reactivated during Alpine compression. Transect 3 (Fig. 6) is an 800 km long profile extending from the Elbe Fault Zone (not to be confused with the Elbe Lineament, defined above) in the SW across the central part of the Polish Trough to the Mazurska High of the EEC in the NE. This interpretation is based on the results from the POLONAISE P4 deep seismic refraction profile published by Grad et al. (2002a,b). The crust of the EEC exhibits characteristic three-layered structure, with average P-wave velocities of about 7.1 km s-1 (lower crust), 6.5 km s -1 (middle crust) and 6.2 km s -t (upper crust). The Suwalki anorthosite complex forms a localized high-velocity (6.4 km s - i ) anomaly within the upper crust. At up to 50 km thick, the crust of the EEC is much thicker than that of young, accreted Europe. The velocity model presented by Grad et al. (2002a) suggests that at the Teisseyre-Tornquist Zone, the high-velocity upper and middle crust terminates rapidly against the Polish Trough. This abrupt termination is also clear in the magnetic potential field image (Fig. 2). The lower crust appears to extend as an attenuated wedge for perhaps 250 km to the SW of the Teisseyre-Tornquist Zone, however (Fig. 6), beyond the Polish Trough, to underlie a less heterogeneous crust with much lower average velocity assigned to Avalonia (Winchester & PAGE TMR Network 2002). It is clear that in this region, the geometry of the Thor Suture has been considerably modified by the effects of Permian-Mesozoic extension in the Polish Trough. The midcrust beneath the Polish Trough has a P-wave velocity of about 5.8 km s- 1. This is too low for crystalline basement, and therefore may represent deeply buried Neoproterozoic-Palaeozoic strata. These could be of Devonian and Carboniferous age, as in other rifts within the EEC; but more probably comprise thick Ediacaran (Vendian) and thinner lower Palaeozoic strata comparable with those of the Lysogory and Matopolska Blocks, exposed in the Holy Cross Mountains some 250 km to the SE of the transect.

They could be filling a 'Central Polish Palaeo-Rift' (see above, for discussion) antecedent to the (Permian-Cenozoic age) Polish Trough, and initiated during the rifting of Rodinia-Pannotia. The eastern bounding fault of the Polish Trough may represent a simple reactivation of an early Palaeozoic syndepositional fault, perhaps associated with the passive margin development of Baltica; or, if the interpretation of these blocks as suspect terranes is correct (see earlier discussion), a reactivated terrane boundary (between the Lysogory Terrane and Baltica). It was a locus of Alpine inversion. The western boundary of the Polish Trough also appears sharp, but not quite as steep as the eastern boundary. It coincides with the northward extension of the Moravian Line, which Winchester et al. (2002) have identified as the boundary at the eastern extremity of (Far Eastern) Avalonia. The crust here is comparatively homogeneous and unlayered, exhibiting a P-wave velocity of about 6.3 km s- 1 to 25 klTIdepth, characteristics that support an Avalonian affinity. At shallow crustal level, deep boreholes on the Leszno-Wolsztyn High sample metamorphic rocks of the Rheno-Hercynian Zone, derived from (Franke & Zelainiewicz 2002), and resting upon, the Avalonian crust. Transect 4 (Fig. 7) extends for 240 km across the Mid-North Sea High off the NE coast of Britain (Fig. 1). It crosses the various accreted terranes bordering the Iapetus Suture between Avalonia and Laurentia. The interpretation presented is adapted from those presented by Freeman et al. (1988) and Chadwick & Holliday (1991). The location of the Iapetus Suture is well constrained in the western part of the North Sea by deep seismic data (Freeman et al. 1988; Soper et al. 1992). The northward dip of reflector packages (IN, IS) in the lower crust and in the mantle (P) at and just below the Moho (Fig. 7) indicates a northward-dipping suture, with a wedge of Avalonia (probably comprising juvenile accreted material, such as the Skiddaw Group) extending some 7 0 k i n beneath the largely Silurian age Southern Upland Acccretionary Complex. The lower crust of the latter is distinguished not by sonic velocity, but by its reflection character. In the north, the Southern Upland Fault is

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T.C. PHARAOHETAL.

Fig. 6. Transect 3, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC in central Poland; 5 x verticalexaggeration. Interpretation based on crustal model of POLONAISE P4 deep seismic refractionprofilepublished by Grad et al. (2002a,b) and Sroda et al. (2002). Post-Palaeozoic basins and platforms: CPT, Central Polish Trough; L-W, Leszno-Wolsztyn Basement High. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: DF, Dolsk Fault; EEC, East European Craton; KLZ, Krakdw-Lubliniec Zone; LT, LysogoryTerrane; ML, MoravianLine; OF, Odra Fault; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: CPPR, Central Polish Palaeo-Rift;MH, Mazurska High; ML, MoravianLine; SU, Suwalki AnorthositeMassif; T-TZ, Teisseyre-TornquistZone; VF, Variscan Front. the terrane boundary with the Midland Valley Terrane, another component of the Laurentian terrane collage. The development of Carboniferous sedimentary basins overlying the accretionary complex (e.g. the Northumberland Trough and Tweed Basin)

Fig. 7. Transect 4, to illustrate inferred crustal structure across the 'Caledonide' accreted crust in the central North Sea region; 2.5 x vertical exaggeration. Modifiedfrom interpretationsof NEC deep seismicreflectionline by Freeman et al. (1988) and Chadwick & Holliday (1991). Post-Palaeozoic basins and platforms: MNSH, Mid-NorthSea High; NFF, Ninety Fathom Fault; NT, NorthumberlandTrough; TB, Tweed Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: IN, IS; Iapetus Suture Zone; MVT, Midland Valley Terane; SNSLT, SouthernNorth Sea Terrane. Mantle reflectors: P.

was facilitated by extensional reactivation of the early Palaeozoic structures (Chadwick & Holliday 1991). Transect 5 (Fig. 8) extends from Kiel, southward across the North German Basin (Variscan Foreland), crossing the various internal zones of the German Variscides towards the Alpine Molasse Basin near Zurich. The interpretation of the crustal structure is based on that published for the central segment of the European Geotraverse (EGT) by Aichroth & Prodehl (1990) and Prodehl & Aichroth (1992). The detail of shallow crustal structure in the central part of the transect is derived from interpretations of the DEKORP-2N and -2S deep seismic reflection profiles, which are slightly divergent from the path of the EGT (Fig. 1), published by Giese (1995) and Oncken et al. (2000). As on Transect 2, a tapering wedge of high-velocity EEC lower crust is depicted extending southward beneath (Far Eastern) Avalonian crust underlying the North German Plain, to the vicinity of the Elbe River. Crossing the Variscan Front near Celle, the DEKORP-2N profile (Meissner & Bortfeld 1990; Meissner et al. 1994) provides an excellent view of the internal structure of the Rheno-Hercynian nappe pile. The profile images a ramp-flat geometry for the basal detachment, unreflective crust of the Brabant Massif forming a south-tapering wedge in the footwall, and a Moho that brightens southward. Despite the considerable shortening and displacement implied, the early Palaeozoic faunal affinities of the RhenoHercynian Zone lie with Avalonia (Cocks et al. 1997) so that the Variscan Front and basal detachment represent an entirely intra-Avalonian boundary. The section balancing reconstruction of Oncken et al. (2000) indicates at least 200 km of horizontal shortening in the upper crust in this zone of the Variscides. The lower crust has a much higher P-wave velocity of about 6.8 km s -1. In Figure 8, following the known inherited ages of many Neoproterozoic Avalonian granitoids, this is speculatively interpreted as an extensive block of Rondonian (Mid-Proterozoic)

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Fig. 8. Transect 5, to illustrate inferred crustal structure of the 'Variscide' accreted crust in the central German region; 5 • vertical exaggeration. Interpretation of full crustal structure from the EGT deep seismic refraction experiment (Aichroth & Prodehl 1990; Prodehl & Aichroth 1992). Detail of shallow crustal structure in central part of transect is after interpretations of DEKORP deep seismic reflection profiles by Giese (1995) and Oncken et al. (2000). Post-Palaeozoic basins and plaOeorms: HD, Hessen Depression; MB, Molasse Basin; NGB, North German Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: MGCH, Mid-German Crystalline High; NPZ, Northern Phyllite Zone; PT, Perunica (Bohemia) Terrane; MGS, Moldanubian-Gf6hl Suture; MT, Moldanubia Terrane; RS, Rheic Suture; STS, Saxo-Thuringian Suture; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: BM, Bohemian Massif; DR, Drosendorf Unit (of BM); ET, Ebbe Thrust; GF, Gf6hl Unit (of BM); GN, Giessen Nappe; MN, Mtinchberg Nappe (of BM); SG, Saxonian Granulites; TBZ, Tepl~-Barrandian Basin (of BM); TT, Taunus Thrust. Other geographical locations mentioned in text: Ki, Kiel; ER, Elbe River; SJ, Swabian Jura; Zu, Zurich.

crust lying at the heart of Avalonia. The transect then crosses the Saxo-Thuringian Zone, passing just east of the Odenwald. The Rheic Suture, defining the southern edge of pre-Variscan Avalonia, appears to be a rather steeply dipping feature in the upper crust (perhaps 70~ but is less steeply dipping in the lower crust, where a wedge of Rheno-Hercynian (=Avalonian protolith) apparently extends beneath the Saxo-Thuringian domain (Giese 1995; Krawczyk et al. 2002). To the south of the Mid-German Crystalline High, the Saxonian Granulite Dome forms a key, but enigmatic, element of the crust (Krawczyk et al. 2000). The Mtinchberg Nappe, resting on the Saxo-Thuringian Suture, is interpreted as an outlier of the Moldanubian Zone, which forms the southern part of the transect. The Moldanubian Gfrhl Suture separates correlatives of the Drosendorf and Gfrhl units of the Bohemian Massif, amongst the latest of the TESZ terranes to leave Gondwana before the opening of the Palaeo-Tethys Ocean.

Conclusions Multidisciplinary studies by the EUROPROBE Programme, including reinterpretation of older geophysical datasets, have led to considerable improvements in knowledge of the structure and evolution of the Trans-European Suture Zone (TESZ), the most fundamental lithospheric boundary in Europe. The TESZ represents a tectonically complex zone of crustal terranes accreted, throughout Palaeozoic time, to the passive margin of the East European Craton (EEC). Various geophysical techniques have been applied to define the geometry of the sutures, representing destroyed ocean basins, which define the boundaries of these various terranes, and most are shown to be non-vertical. Other well-known structures and lineaments (e.g. the SorgenfreiTornquist Zone (STZ) and Teisseyre-Tornquist Zone (TTZ), are steep structures unrelated to suturing, but may have originated during late Neoproterozoic rifting of the Rodinia-Pannotia supercontinent. They have subsequently been reactivated many

times, most recently during Alpine inversion. Standard methods of terrane analysis have been applied to identify and characterize individual terranes within the TESZ, but the status of some remains controversial. Fortunately, the application of more sophisticated studies of isotopic composition (e.g. of detrital grain provenance), palaeontology (e.g. to recognize timing of ocean closure) and palaeomagnetism (e.g. to identify terrane rotation) is helping to resolve some of these controversies, and add detail to the history of accretion and dispersal. These studies have demonstrated that numerous crustal terranes were rifted away from various margins of Gondwana at low southerly palaeolatitude more or less continuously, for much of early Palaeozoic time. After a northward passage, driven by the opening and closure of numerous (perhaps not very large) ocean basins, at least some of these ended up being accreted to the passive margin of the EEC. As studies proceed, the evidence for the rates of rotation and closure amidst this archipelago of Gondwanaderived terranes is starting to contribute to a dynamic model for the evolution of the TESZ. Subsequent to initial accretion, dispersal of some terranes to other locations on the EEC margin was facilitated by major, crust-penetrating steep faults such as the TTZ. This process is most advanced in the SE part of the TESZ in Romania and Turkey, where terranes originally accreted during the Variscan Orogeny have been displaced and re-accreted to the EEC margin during the development of the peri-Tethyan margin, opening of the Black Sea and the subsequent AlpineCarpathian Orogeny. Some of the studies reported here were carried out in the EU-funded PACE (Palaeozoic Amalgamation of Central Europe) TMR Network, no. ERBFMRXCT97-0136. The contribution of T.P. is published with the permission of the Executive Director, British Geological Survey (NERC). Numerous fruitful discussions with EUROPROBE TESZ and PACE project participants are gratefully acknowledged. PACE colleagues P. Williamson and D. Banka contributed significantlyto the quality of the diagrams. J. Carney and W. Franke are thanked for their helpful reviews, and D. Gee for editorial comments. D. Gee is sincerely thanked for his significant contribution to the TESZ project, and for the unstoppable motivation he provided during the 10 EUROPROBE years, 1992-2002.

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Physical differences in the deep lithosphere of Northern and Central Europe SOREN G R E G E R S E N l, PETER VOSS 1'2, Z. H O S S E I N S H O M A L I 3'4, M A R E K G R A D 5, R O L A N D G. ROBERTS 3 & TOR W O R K I N G GROUP

IGEUS, Ostervoldgade 10, DK-1350 Copenhagen K, Denmark (e-mail: [email protected]) 2Department of Geophysics, Niels Bohr Institute, University of Copenhagen, DK-2100 Copenhagen O, Denmark 3Department of Earth Sciences, Uppsala University, SE-752 36 Uppsala, Sweden 4present address: Geological Survey of Sweden, SE-751 28 Uppsala, Sweden 5Institute of Geophysics, University of Warsaw, Pasteura 7, PL-02-094 Warsaw, Poland

A number of large-scale integrated studies, including the TOR and POLONAISE'97 projects, with an emphasis on seismic methods, have been used to elucidate the southwestern boundary (suture zone) between the East European Craton and the Phanerozoic terranes of Western Europe. Results indicate that a thick slab of mantle lithosphere below the craton thins southwestwards beneath the Trans-European Suture Zone and is not seen south of the Variscan front. The thinning is not gradual, but is interrupted by at least two abrupt deep boundaries, the most significant of which corresponds to the surface position of the Tornquist Zone, a major fault. The present geometry of the lithosphere is the result of modification of the margin of the Neoproterozoic continent Baltica by Phanerozoic processes, including the development of the Tornquist Zone and the stretching of the lithosphere in a broad central block SW of this zone. Seismic results and their interpretations from the TOR tomographic project are presented and compared with results from the POLONAISE'97 controlled source project to the SE. Both investigations have shown high-angle, non-symmetrical features extending deep into the mantle. Abstract:

The lithosphere of Eastern Europe is dominated by the East European Craton (EEC), a crystalline complex largely composed of Archaean and Early Proterozoic rocks, assembled in the Late Palaeoproterozoic (Bogdanova et al. 2005). It is flanked to the SW by younger terranes of Neoproterozoic and Phanerozoic age that were accreted to the EEC in the Palaeozoic and subsequently overlain by Mesozoic and Cenozoic successions (Pharaoh et al. 2006). The boundary between the EEC and these western accretionary complexes (Fig. 1) is a wide belt of thrust-emplaced terranes, called the Trans-European Suture Zone (TESZ) (see Fig. 1 and Gee & Zeyen 1996). Here, we present the results of some more recent studies designed to provide information on the lithospheric and asthenospheric mantle across the TESZ. The area investigated covers the southern part of Scandinavia from the exposed Proterozoic crystalline basement rocks of the Baltic (Fennoscandian) Shield, in southern Sweden, to the Palaeozoic, Mesozoic and Cenozoic cover successions of Denmark and northern Germany (Fig. 1). It is an area that has been subject to a variety of very different stress regimes since its establishment as the passive continental margin of the EEC in the Late Neoproterozoic. This thinned southwestern edge of the EEC, composing, in the Early Palaeozoic, the margin of the continent Baltica (Cocks & Torsvik 2006) was subject to mid-Palaeozoic (Caledonian) N E - S W compression, Late Palaeozoic transcurrent faulting, Early Mesozoic extension with lithospheric stretching, and Late Cretaceous to Early Cenozoic Alpine inversion. Thus, the configuration of the lithosphere today is a result of a complex interplay of lithospheric processes that have influenced the edge of the craton over the last 600 Ma. Geophysical investigations of this area were an important component of the European Geotraverse Project (Blundell et al. 1992) and the crustal structure of the area has been investigated by a number of projects (e.g. EUGENO-S Working Group 1988; BABEL Working Group 1993; Rabbel et al. 1995; Abramovitz & Thybo 2000; Thybo 2000; Gregesen et al. 1992; Thybo et al. 1998). These geophysical investigations, supported locally by drilling (e.g. to the Precambrian basement of the Ringkcbing-Fyn High (RFH), across central Denmark, just north of the Thor Suture of Fig. 1), have been interpreted to show that the old crystalline complexes of the Baltic Shield can be followed southwestwards across Denmark beneath the Phanerozoic and partly

Neoproterozoic sedimentary successions at least as far as the Thor Suture (see Fig. 1 and Pharaoh et al. 2006). The crystalline basement of the EEC has been interpreted on crustal seismic data to taper southwestwards from the southern exposures of the Baltic Shield, across the NW-trending fault system of the Sorgenfrei-Tornquist Zone in southernmost Sweden and northern Denmark, and beneath the North German Basin (e.g. Abramovitz & Thybo 2000; Thybo 2000), interrupted only by the thicker crust of the RFH. It is inferred to reach to a line that trends from southernmost Denmark southeastwards across northern Germany into northern Poland. This c. 200 km wide zone of tapering cratonrelated rocks is overthrust by the Neoproterozoic and Early Palaeozoic complexes of Eastern Avalonia and Late Palaeozoic Variscan nappes (Franke 2006). Studies providing information about the deeper structure have also been published, including a number of analyses of fundamental-mode and higher-order Rayleigh waves. The highermode data of Nolet (1977) have been interpreted to distinguish Scandinavia from the Western European structure (see also Zielhuis & Nolet 1994). Concerning the deep lithospheric structure of the Baltic Shield, conflicting results have been published by Cara et al. (1980) and by Dost (1990). Dost (1990) found a low-velocity zone of up to 2% in the depth interval 150-220 km whereas Cara et al.'s (1980) interpretation of the higher modes found no need for a low-velocity zone in the mantle. Some information on very deep structure is also available from earlier P-wave studies. From explosion studies Guggisberg (1986) interpreted several deep low-velocity channels (of about 5% in velocity) in the lithosphere, with the bottom of the lithosphere at about 200 km depth marked by a velocity change of a couple of per cent. Later, Ryaboy (1990) and Thybo & Perchuc (1997) have interpreted low-velocity zones of several per cent beneath the craton at depths of 105-135 km and somewhere between 100 and 280 km, respectively. Because of the limited resolution of earlier studies, the deeper structure of Northern and Central Europe has recently been investigated by a number of projects including TOR, POLONAISE'97, CELEBRATION 2000, SVEKALAPKO and Eifel Plume. The TOR results are discussed below. To the NE of TOR, in Scandinavia, the subcrustal lithosphere of the shield area has been investigated by the SVEKALAPKO project (Sandoval 2002), and to the SW

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 313-322. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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lithosphere. High-quality record sections were obtained for the longest offsets of about 600 km from the shot points, with clear first arrivals and later phases of waves reflected or refracted in the subcrustal lithosphere. The CELEBRATION 2000 project (Guterch et al. 2003) slightly farther to the SE has also provided some information about the upper mantle. Below, results from the POLONAISE'97 project are discussed and compared with results from TOR. With new fieldwork commencing in 2006, this area in Poland will soon be further investigated in a large-scale tomographic project, similar to TOR.

TOR field experiment, data and crustal structure

Fig. 1. The location of seismographs during the field-workof the TOR project, 1996-1997. The red dots indicate short-period seismographs and the blue dots indicate broadband seismographs. The 2D interpretation profile of Figure 5 is perpendicular to the Teisseyre-Tornquist Zone (TTZ) in Poland, whereas the 2D interpretation profile of Figure 4 follows the middle of the cloud of seismographs from 59~ 16~ to 51~ 9~ and is perpendicular to the Sorgenfrei-Tornquist Zone (STZ). The geology (a) and crustal interpretation (b) are from Pharaoh et al. (2006). The Ringk0bing-Fyn High (R-F H), a basement high separating sedimentary basins to the north and south, is located just north of the interpreted Thor Suture. The Elbe Line is a geophysically recognized lineament, locatedjust south of the interpreted Thor Suture.

of TOR, in Germany, by the Eifel Plume project (Ritter et al. 2000). In Poland the POLONAISE'97 project (Guterch et al. 1999) investigated that part of the TESZ known as the Teisseyre-Tornquist Zone (the continuation southeastwards of the the Sorgenfrei-Tornquist Zone in Sweden and Denmark) using a wide-angle controlled source method, revealing seismic velocity structures and reflective interfaces in the subcrustal

In contrast to, for example, POLONAISE'97, where the seismic signals were produced using explosions, TOR was primarily a passive project, recording signals from distant earthquakes and using the characteristics of these recordings to deduce structures below the recording array. Figure 1 shows the position of the recording array, with a strike roughly N E - S W (perpendicular to the Tornquist Zone) from southern Sweden, through Denmark into northern Germany (Gregersen et al. 2002). The position of the array was chosen partly because the crustal structure here had been well investigated by a number of earlier seismic projects, referred to above. This was considered important because the size of the TOR area to be investigated, together with the number of recording instruments available, implied that the distance between recording stations was too large to resolve the crustal structure in any detail. This is further discussed below. The TOR seismic antenna was designed to be relatively long and narrow, as geological evidence and previous geophysical studies have indicated that the large-scale subcrustal lithospheric inhomogeneities in the area are predominantly 2D and oriented N W - S E . However, an array, rather than a single profile, was used to allow some control of the effects of a possibly more complex subsurface geometry. A pilot project was undertaken in 1995 (Kind et al. 1997) consisting of 26 broadband seismographs on a 120kin line from NE Denmark to SW Sweden. It confirmed a rapid change in crustal thickness and seismic velocity across the Tornquist Zone and demonstrated the feasiblity of the proposed main TOR project. The TOR seismic antenna was then deployed in a 900 km long by 100 km wide strip around Profile 1 of the EUGENO-S study and along the German DEKORP reflection seismic line (DEKORP-BASIN Research Group 1999) as depicted in Figure 1. Short-period seismographs (more than 100) were operated from October 1996 to April 1997 (close to half a year) to record distant earthquakes for teleseismic body-wave P- and S-wave travel-time tomography. By chance, they also recorded an earthquake inside the array (Schmidt 1998). The TOR antenna also contained 31 broadband seismographs operated from summer 1996 to summer 1997. These broadband seismographs were mainly used for studies of surface waves (Cotte et al. 2002), anisotropy (Wylegalla et al. 1999; Plomerova et al. 2002), and receiver functions (Gossler et al. 1999; Wilde-Pi6rko et al. 2002). The average station spacing (all stations) is 20 km in the centre of the array in Denmark and southern Sweden, and 2 5 - 3 5 km in northern Germany and central Sweden. The broadband seismographs were placed in triangles of side lengths 4 0 - 6 0 krn. The locations of the triangles were selected so that the data could also be used together with those from permanent broadband seismographs in the area and a few strategically placed temporary stations located off the TOR array, as shown in Figure 1. All the seismographs were placed in isolated areas, as far away as possible from traffic and other man-made disturbances. In the shallow low-velocity lithosphere, rays from distant earthquakes propagate with steep incidence angles. Together with the relatively large station spacing, this severely limits TOR's potential to resolve structures in the sedimentary cover and underlying crystalline crust. It was therefore decided to establish a crustal

DEEP LITHOSPHERE DIFERENCES

model for the area based on other available geophysical data. Two versions of the 3D crustal model were derived, based on different choices of interpolation method between areas with well-resolved crustal velocities, which lie along previous seismic profiles. One of these was derived by Arlitt et al. (1999) following the procedure of Waldhauser et al. (1998), where only interfaces that have confinned seismic reflections are accepted, and smooth mathematical interpolation is done between these. The sediments are included in an a d hoc manner, with little geological reference (Arlitt et al. 1999; Gregersen et al. 1999). In contrast, the model derived by Pedersen (1999) and Pedersen et al. (1999) considers also geological information and other geophysical data, such as gravity. This model was produced by interpolation between structures in published crustal models (EUGENO-S Working Group 1988; Green et al. 1988; Thybo et al. 1989; Stangl 1990; Thybo 1990; Thybo & Sch6nharting 1991; Aichroth et al. 1992; BABEL Working Group 1993; Guterch et al. 1994; Rabbel et al. 1995). Other available information was used to constrain the interpolation. The sediments are included in the modelling through realistic average velocities in the upper layer. Thus, the model of Pedersen et al. (1999) includes more data, which should improve reliability, but also includes more choices in the interpolation procedures, giving greater scope for preconception and bias. The crustal influence on the TOR travel-time anomalies to be interpreted is shown in Figure 2. The Arlitt model version implies crustal corrections of the order of half a second, whereas the Pedersen model involves crustal corrections twice as large. The effects of using these 3D crustal models on the analysis of the mantle structures based on the TOR tomography has been investigated and discussed by Shomali et al. (2002). They showed that, in the specific case of the TOR experiment, because of the character and geometry of the large-scale lithospheric structures, the dominant features of these are well constrained, largely independent of differences between the two crustal models. The largest part of the TOR project was the bodywave tomographic study. Almost 300 earthquakes were well recorded by the array, but not all of these were analysed in detail, partly because earthquakes with sources relatively close to each other do not provide independent information in the tomography, and partly because the inclusion of data with lower signal-to-noise ratios can degrade the results. After the definition of a suitable suite of events, P-wave first arrival times were independently read by workers in the several groups involved in TOR, before being combined into a single dataset for interpretation. An important aspect of this work was a series of tests and comparisons to ensure that students and scientists of the many groups would pick arrival times in a consistent manner. Travel time residuals relative to a standard 1D global model were then computed for each event and station. Examples of P-wave travel time residuals for two events are given in Figure 2 (Pedersen et al. 1999). The lowermost row of Figure 2 presents those residuals caused by the deeper lithosphere-asthenosphere structure after the effect of the crustal model has been removed. Even from each single event, it is clear that major structures exist at depth below the array, and the size and character of these structures is such that they are not masked by the overlying crustal structure.

Interpretation of the TOR data Several different analytical procedures have been applied to the TOR data, including P- and S-wave tomographic, surface-wave, anisotropy, scattering and receiver function analyses. We now discuss some of the main results. Using slightly different methods and datasets, several interpretations, including those by Arlitt (1999), Pedersen et al. (1999) and Shomali et al. (2002), have been made of the P-wave travel time residuals. All of these have found the largest

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subcrustal lateral velocity contrast directly below the Tornquist Zone (see Figs 1 and 4). Although the existence of this feature is very clear, some of the details, including the direction of dip of the boundary (northwards or southwards), are less well resolved. Further independent analyses include that of Horn (2001) using a Monte Carlo inversion and that of Busche (2001) using eigenfunction analysis. A recurring theme in these different analyses is that of the resolution of specific features. Clearly, we would like to extract as much information as possible from the available data, and a natural question concerns which features in the derived models are reliable. Some questions remain, and will be further analysed. However, all analyses confirm the important division of the lithosphere into three blocks, with thick (>200 km) high-velocity lithosphere NE of the Tomquist Zone, a thinner lithosphere (about 100 km) in a central block beneath Denmark, and an even thinner lithosphere in a southwestern block. This model is consistent with analyses of crustal and even some sub-crustal data, but the TOR data now provide resolution down to several hundred kilometres depth. In a separate investigation, Shomali et al. (2006) have extended the tomographic study to include the S waves, in addition to the P waves. Fewer data exist because of lower signal-to-noise ratio for the S phases, so the S-wave velocity model has larger uncertainty than that for the P waves. One outcome of this study has been a confirmation of the most prominent features of the P-wave models. From the P- and S-wave velocity models, the ratio between the two, Vp/V~, has been derived. A difference has been distinguished between the velocity ratios in the three blocks, northern, central and southern. The difference is a couple of per cent, with a ratio of < 1.8 in the two outer blocks, and > 1.8 in the central block. In addition to direct station-to-station analyses, the triangular broadband sub-arrays allow array processing methods to be applied to the surface-wave data, providing phase velocity dispersion information to periods as long as 90 s. The fundamental mode Rayleigh-wave dispersion characteristics of the various paths have been grouped and classified, and the results have been compared with the tomography results. The dispersion characteristics can be grouped into three classes, covering essentially the same areas as those found in the P-wave travel time anomaly study: the North German Basin area, the Danish area almost up to the Tornquist Zone, and the Baltic Shield, NE of the Tomquist Zone. The Rayleigh-wave dispersion curves of each of the three regions have been inverted to estimate the deep S-wave velocity structure. In Figure 3 the S-wave velocities in each area are shown as a function of depth, with a line for each model that is acceptable within the standard deviation of the dispersion curves (i.e. the spread of the lines indicates the resolution). The southern areas (1 and 2) are both interpreted to contain low-velocity channels for the S waves (interpreted as asthenosphere), but at different depths. No low-velocity channel is observed in the shield below Sweden. Exactly where the boundaries between these three areas occur is discussed in a separate paper on the surface-wave results by Cotte et al. (2002). Those researchers have claimed that the northern broadband instrument triangle in Denmark (Fig. 1) shows the same trend of Rayleigh wave dispersion, for periods of 3 0 - 9 0 s, as that in the shield. Therefore they have proposed that the separation between the shield and the Danish lithospheric block dips southwestwards from the northern rim of the Tornquist Zone. However, a consideration of surface-wave theory in laterally inhomogeneous media (e.g. Gregersen 1976) suggests that this dip may not be correctly resolved. Cotte et al. (2002) argued that the dispersion similarity in northeastern Denmark and in the shield to the north indicates similar vertical velocity profiles. The counter-argument, based on surface-wave propagation in laterally inhomogeneous media (e.g. Vaccari & Gregersen 1998), is that this is not a valid argument because

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Fig. 2. P-wave travel time residuals from two earthquakes, in Japan (19 October 1996; 32~ 132~ wave arrival to TOR area from NE) and in Mexico (11 January 1997; 18~ 103~ wave arrival to TOR area from NW), each in one column. For each of the earthquakes, the upper diagram shows total observed residuals (observed arrival times minus expected arrival times according to the global average tables IASP91). The middle diagram shows computed crustal residuals to a depth of 50 km (from Pedersen et al. 1999) which, when subtracted from the observed ones, give the lower lithosphere residuals below 50 km depth; these are shown in the lower diagram.

these long-wavelength surface waves are sensitive to structures at a lateral distance from the site comparable with the depth of the feature. Two studies of shear-wave splitting in the TOR data have been carried out (Wylegalla et al. 1999; Plomerova et al. 2002). Good SKS, SKKS or diffracted and refracted S signals have been collected from the many seismographs, and the polarization of the fast S waves has been determined. One of the studies on horizontal anisotropy (Wylegalla et al. 1999) found the fast S-wave azimuth to be almost e a s t - w e s t in the shield area and beneath the German Basin. Close to the Tornquist Zone, the fast S-wave azimuth is

interpreted to be along the trend of the Tornquist Zone, even where this bends to the south of Sweden (Fig. lb). The other study (Plomerova et al. 2002) retrieved dipping high-velocity structures in three dimensions. Plomerova et al. identified three geographical zones consistent with those found in the isotropic tomography studies. In the NE the fast S-wave structures dip NE, in the central zone NW, and in the southern area the dip is SW. The methods used in the two shear-wave splitting investigations are different, so the two investigations do not necessarily disagree. Using P-wave coda, H o c k et al. (2000) investigated scattering. A difference between the shield and the area to the south of the

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Discussion

Fig. 3. LithosphericS-wave velocity models from the surface-wave investigations of Cotte et al. (2002) in the three blocks of different Rayleigh-wavedispersion.

Tornquist Zone was observed, but no smaller-scale differences were resolvable. Hock et al. deduced that the scattering in the shield is mostly confined to the crust, where the correlation lengths are as short as c. 1 km, with P-wave velocity fluctuations of the order of 4%. Farther south, correlation lengths in the crust are 5 - 1 0 k m , with P-wave velocity fluctuations of c. 8%, and the subcrustal lithosphere also produces some scattering, with correlation lengths 10-20 km and P-wave velocity fluctuations of 6 - 8 % . The scattering attenuation was found to be more important than the anelastic attenuation, which beneath the craton is negligible and further to the SW is small.

All analyses of TOR data are consistent with a sharp and steep boundary that penetrates the whole lithosphere below the Tornquist Zone, and a more shallow near-vertical boundary below the Elbe Line, just south of the Thor Suture of Figure 1, which in recent crustal investigations has been interpreted to define the southwestern boundary of the EEC. There are also some indications of a significant boundary about 100 km NE of the Tornquist Zone within the shield, which could be connected to the known crustal differences across the Protogine Zone (see, e.g. Plomerova et al. 2002) between southwestern and southeastern Sweden. Given the low heat flow of the craton, temperature considerations alone would suggest that there should be significant velocity differences between the northeastern and southwestern parts of the TOR area. However, it is clear that the transition zone is not gradual between the craton in Sweden and the thin lithosphere in Germany, but contains at least two rapid lateral changes in velocity in the upper mantle. In terms of seismic velocity contrasts, the most pronounced lateral variations at sedimentary, deeper crustal and subcrustal depths occur in different geographical locations along the profile. This is illustrated in Figure 4 using the teleseismic tomography model of Shomali et al. (2002). Superimposed on the velocity image are boxes showing where the major structural transitions are deduced to be, based on all available data, not just the P-wave tomography. The teleseismic tomography has poor resolution in the uppermost 50 km or so, and this range is left uncoloured in the figure. However, as referred to above, there is information about shallower structure based on other data (Arlitt et al. 1999; Pedersen et al. 1999). The deduced crustal thickness is shown by the light dotted line. The sloping crustal transition from Baltica to Avalonia in box A, and the wedging out southwestwards of the mantle lithosphere are noteworthy features. The blue and red P-wave velocity anomalies are computed with reference to a 1D global travel-time model (Kennett & Engdahl 1991). The boxes B and C in Figure 4 emphasize two sharp and steep velocity changes. These steps are seen as red-blue steps. The third step is less obvious, between light blue and dark blue, 100 km into Sweden from box C. That the change from red to blue occurs at these points depends, of course, on our choice of presentation parameters. However, a closer examination of the actual velocities involved, or a presentation of lateral velocity gradients (as opposed to estimated velocities) clearly shows that Figure 4 is not misleading because of the choice of colours. The existing uncertainties in the exact locations of the transitions and their slopes are illustrated by broad boxes (B and C) in the transition regions. Further studies are in progress to, as far as possible, enhance and define the resolution of these features. Although the various seismological investigations, such as the tomographic inversions of Arlitt (1999) and Shomali et al. (2002), agree regarding the major features of the derived models (zones B, C, etc.), some important details in the images differ, notably the slope of the deep boundary beneath the Tornquist Zone. Box D has been introduced as the bottom of the low-velocity layer, deduced from higher-mode surface-wave studies (Nolet 1977; Dost 1990). Box D does not extend to the NE, consistent with Cotte et al.'s (2002) interpretation from the fundamental mode Rayleigh waves that there is no low-velocity layer below the shield. This conclusion is not, however, consistent with some previously published studies of higher-mode surface-wave data and some P-wave controlled source data, which deduced lowvelocity layers (see above). Both fundamental-mode and highermode surface waves are mainly sensitive to the S-wave velocity structure, but because of their shorter wavelengths, higher modes are more sensitive to rather thin low-velocity layers. However, the TOR measurements provide a considerable quantity

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Fig. 4. Tomographic image of the lithosphere-asthenosphere system in a profile along the TOR array (from 59~ 16~ to 51~ 9~ STZ, Sorgenfrei-Tornquist Zone; RFH, RingkCbing-Fyn High; EL, Elbe Line. Blue areas show high P-wave velocities; red areas show low P-wave velocities with respect to the laterally homogeneous IASP91 model (Kennett & Engdahl 1991). Box A is the crustal transition region (e.g. Abramovitz & Thybo 2000). The generalization boxes B and C are drafted taking into account the TOR results of Arlitt (1999), Pedersen et al. (1999), (2001), Horn Busche (2001), Cotte et al. (2002), Shomali et al. (2006) and Voss et al. (2006). Moho depths are from Thybo et al. (1998) and Gossler et al. (1999). Box D in the southern part delimits the material below the asthenosphere at depth 230 km, from Nolet (1977) and Dost (1990). of new very high quality data from the area, suggesting that the previously deduced low-velocity zones may be incorrect. In this paper only the most significant well-resolved features have been discussed. There is, of course, more information in the data than this. Thus, for example, several other features in Figure 4 could be geologically meaningful. However, resolution generally decreases with the spatial size and velocity contrast of a feature. It can be difficult to assess just which features in the models should be regarded as well defined and suitable to interpret in geological terms. It follows from this that issues regarding spatial resolution, accuracy and uniqueness of the models are important. Within TOR, considerable effort has therefore been dedicated to examining these issues in, for example, tomographic inversions. In such an inversion it is straightforward to calculate, for example, the variance of model parameters, which quantifies the reliability of each parameter. The resolution is limited by station spacing, ray geometry between earthquakes and seismic stations, and the frequency content of data, and it is different for different parts of the model. The calculated resolution is dependent on choices in the mathematical inversion procedure. Furthermore, the variance of a model parameter describes only a part of the problem, partly because it makes a simplifying mathematical assumption of linearity, and partly because it ignores the often

complicated interactions between model parameters (covariance). The complexity of these problems means that several different approaches can be considered to quantify reliability. Some information on this can be gained by solving the mathematical inverse problem stepwise linearly, as was done by Arlitt (1999), Shomali (2001) and Shomali e t al. (2002). Other approaches used have been relative model testing (Shomali e t al. 2002), resolution kernel and synthetic tests (Shomali 2001), sigmoid function and Fourier component computations (Busche 2001), and Monte Carlo model evaluations (Horn 2001; Voss e t a l . 2006). All the results indicate that the major transitions and blocks that have been discussed here are resolved well by the observed data. It is interesting to compare the TOR results with those from POLONAISE'97, which also crossed the TESZ, but farther to the SE. In POLONAISE'97, 2D interpretation was carried out using a ray tracing approach for the reversed system of P- and S-wave travel times (Grad e t al. 2002). The model of the seismic P-wave velocity structure beneath profile P4, perpendicular to the Tornquist Zone in Poland, includes several reflectors in the lower lithosphere (Fig. 5). The corresponding reflected waves are interpreted as originating from a sub-Moho reflector (PI), the top of a low-velocity zone (PH, usually poor) and from discontinuities at depths of c. 80 and c. 90 km (Pro and Piv, respectively; Fig. 5). The velocity

Fig. 5. Subcrustal reflectors on profile P4 crossing the Tornquist Zone from the POLONAISE'97 experiment (Grad et al. 2002). The displacement of sharp lateral transitions in various depth ranges should be noted.

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beneath the Moho is found to be rather high (c. 8.25 km s -1) in the Palaeozoic terranes in the SW and normal (c. 8.1 km s -1) in the East European Craton to the NE. The thickness of the crust is 30 km in the SW and 4 0 - 5 0 km in the EEC. The subcrustal, almost vertical inhomogeneities between the various regions (Fig. 5) are displaced into the EEC, with respect to the crustal inhomogeneities, similar to the sloping derived by Pedersen et al. (1999), Busche (2001), Shomali et al. (2002) and Voss et al. (2006). The variations in P- to S-wave velocity ratio reported by Shomali et al. (2005) from TOR P- and S-wave tomographic studies, with an anomalous central block, are consistent with receiver function results for the uppermost mantle in Poland (Wilde-Pirrko et al. 2002), where the central block is the Tornquist Zone. Thus, the TOR and POLONAISE'97 data both reveal deep structures of the TESZ, and several significant components in the derived models show similar properties, despite the considerable geographical separation of the projects. In the TOR area, the extreme southeastern boundary of the EEC crust deduced from geological and geophysical data lies below the Thor Suture (just north of EL in Fig. 4), but there is a clear subcrustal transition zone as far north as the Tornquist Zone, which can be characterized as the edge of the undisturbed (non-stretched) craton lithosphere. A different and complementary way to describe the edge of the craton is by its rheological behaviour (i.e. its reaction to regional stress). In studies on the small, local earthquakes of the region Gregersen et al. (1996, 1998) suggested that the rheological edge of the craton is located between the STZ and RFH. Furthermore, analyses of recent crustal movement data (Lykke-Andersen & Borre 2000; Gregersen & Schmidt 2001) have shown that the STZ is a separation between different geodetic movement patterns. As shown in the papers by Cotte et al. (2002) and Gregersen et al. (2002), we interpret the surface-wave results of Figure 3

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and the tomography results of Figure 4, as well as the other geophysical data, to show a lithosphere of thickness a little under 100 km in the southwestern part of the profile, a little over 100 km in the central part, and more than 200 km in the northeastern part of the profile. The Baltic (or Fennoscandian) Shield (i.e. the area of the craton where old, Precambrian, mostly Palaeoproterozoic-Archaean crystalline rocks are exposed), terminates to the SW at the Tornquist Zone. From deep drilling data and seismic crustal studies, it is deduced that material originating from the craton extends southwards from the Tornquist Zone, gradually thinning below the overlying sediments, and terminating at the Thor Suture, just south of the RFH. A seismic transition is seen in this area, consistent with craton material extending to a depth of the order of 100 km. In the latest Precambrian to early Palaeozoic (c. 600 Ma), the EEC composed the core of the palaeo-continent Baltica, and the TOR area was a passive continental margin. The tomographic images reveal the structure originating from this episode, reworked by a number of later events, as shown in Figure 6. The large-scale development of the transition zone between the Proterozoic craton to the NE (right side of Fig. 6) and the Phanerozoic lithosphere to the SW (left side of Fig. 6) can be described in the block diagrams of Figure 6. Stage 1 is the situation of two separate lithospheric plates colliding in mid-late Palaeozoic times. In stage 2, the area was deformed by lithospheric stretching and transcurrent faulting, the R i n g k c b i n g - F y n High (RFH) with its thick crust detached from the Baltic Craton, and the various blocks of the RFH rotated slightly, separately. In the ensuing stage 3, the lithospheric stretching was perpendicular to the trend of the plate transition and the RFH. The compressional stage 4 is very different, involving inversion of the SorgenfreiTornquist Zone.

Fig. 6. Generalizedsummarydiagram of the broad-scale geologicaldevelopmentof the TOR area (Fig. 1). Blue shows sediments, orange shows crystallinecrust, and red is the uppermost mantle lithosphere. The Toruquist Zone (STZ) and the Ringkoebing-Fyn High (RFH) act through time as compression, spreading, shearing and compressionzones. Very thick arrows show regional stress field. (For further explanation, see text).

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Conclusions The TOR project successfully collected an extensive high-quality seismic dataset, suitable for many seismological analysis methods and allowing the production of high resolution P- and S-wave velocity models for the transition across the Trans-European Suture Zone, which separates the East European Craton from the Phanerozoic lithosphere of Western Europe. The models produced are consistent with previous models based primarily on geological modelling of controlled source crustal seismic and potential field data. They relate surface geological observations and crustal structure with the deeper lithosphere. A major, very deep, near-vertical boundary is observed below the Tornquist Zone. The velocity contrast occurs over too short a lateral distance to be explained by increasing temperature farther from the Shield. Although anisotropic velocity structures in the mantle could contribute to the observations, the feature suggests a compositionally significant boundary. If, as indicated by crustal seismic data and drilling, material from the proto-continent Baltica exists as far south as the Thor Suture, this suggests that this deep feature below the Tornquist Zone cannot have recent origins. Presumably, it originated in some form at the time of development of the Baltica passive margin during Caledonian collision, and has been altered as a result of later phases of regional deformation. The boundary in the neighbourhood of the Thor Suture appears to be much shallower, but still extends to a depth of over 100 km. We infer this to mark the boundary of stretched material from Baltica. No 'seismic asthenosphere' in the sense of a low-velocity layer (relative to the material both above and below) appears to be present below the craton, but is apparent to the SW. Striking parallels in the models from the TOR and POLONAISE projects are seen, suggesting that significant parts of the derived models cannot be related to some local component of the geological evolution, but rather reflect a more fundamental aspect of the large-scale evolution of the area. Clearly, the geological evolution of the crust is dependent upon the evolution of the entire lithosphere in the area, and, to fully understand the crustal evolution, we must understand that of the underlying lithosphere. TOR was a large-scale project, but nevertheless (as we can see from e.g. Fig. 4) there are lithospheric-asthenospheric features on a scale comparable with that of the TOR array. It seems therefore that, to fully understand the TOR data, it will be necessary to place the data into a larger spatial perspective. A profile for the TOR area extracted from the global tomographic results of Bijwaard et al. (1998) shows some similar features to the TOR teleseismic travel-time tomographic results (W. Spakman, pers. comm.), but the locations of the transitions and the regional differences are much better delineated in Figure 4 than in the large-scale global model. This supports the reliability of both models, and it follows that a combination of the two datasets will be very valuable. Similarly, the TOR model fits very well with later, as yet unpublished, tomographic studies to the north in Sweden (R. Roberts, pers. comm.). The data from the SVEKALAPKO (Sandoval 2002) and Eifel Plume (Ritter et al. 2000) projects may also be incorporated. Clearly, one future direction will be to integrate the data from these different projects to create a geophysical and geological model including length scales of thousands of kilometres, depths to 1000 km, and resolution of a few tens of kilometres, even at great depth. The TOR work has been carried out within the framework of EUROPROBE's international TOR Working Group with the following members, besides the five main authors: L. B. Pedersen, A. Berthelsen, H. Thybo, K. Mosegaard, T. Pedersen, R. Kind, G. Bock, J. Gossler, K. Wylegala, W. Rabbel, I. Woelbern, M. Budweg, H. Busche, M. Korn, S. Hock, A. Guterch, M. Wilde-Pi6rko, M. Zuchniak, J. Plomerova, J. Ansorge, E. Kissling, R. Arlitt, F. Waldhauser, P. Ziegler, U. Achauer, H. Pedersen, N. Cotte, H. Paulssen and E. R. Engdahl. Many scientists have supported the project as members of the original TOR planning group or through the field-work with

advice. The following scientists are thanked for their participation: D. Gee, R. Gorbatschev, A. Tryggvason, N. Juhojuntti, H. Wagner, N. Balling, B. H. Jacobsen, P. H. Nielsen, W. Hanka, P. and E. Bankwitz, M. Weber, H.-P. Harjes, A. Biegling, J. Skamletz, E. Perchuc, W. Spakman, J. Zednik, T. Hyvonen, S.-E. Hjelt and L. N. Solodilov. The TOR project has been supported in Germany by the GeoForschungsZentrum Potsdam through personnel and seismographs from the German seismograph pool, and by the Deutsche Forschungsgemeinschaft; in Switzerland by the Swiss National Science Foundation under contract 21-43444.95; in Denmark by the Danish Natural Science Research Council, grant 9401105; and in Sweden by the Swedish Natural Science Research Council, contract G-AA/GU 04990-336. The field-work of the Polish groups was partly supported by the Institute of Geophysics, Polish Academy of Sciences. The research of the Czech group was partly supported by a grant from the Czech Academy of Sciences.

References ABRAMOVITZ, T. & THYBO, H. 2000. Seismic images of Caledonian.

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Palaeozoic accretion of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories J. A. WINCHESTER a, T. C. PHARAOH 2, J. VERNIERS 3, D. IOANE 4 & A. S E G H E D I 5 1School of Physical and Geographical Sciences, Keele University ST5 5BG, UK (e-mail j. a. winchester @esci. keele, ac. uk) 2British Geological Survey, Kingsley Dunham Centre, Keyworth NG12 5GG, UK 3Ghent University, Palaontologie, Krijgslaan 281/$8, B 9000, Gent, Belgium 4Faculty of Geology and Geophysics, University of Bucharest, Bucharest, Romania 5Geological Institute of Romania, 1 Caransebes St, 012271 Bucharest 32, Romania

Abstract: Recent work in Central Europe, combined with emerging information about basement massifs in SE Europe and NW Turkey, permits a new look at the relationships between crustal blocks abutting the East European Craton (EEC) along the Trans-European Suture Zone (TESZ). The simplest model indicates that the end-Cambrian establishment of the Bruno-Silesian, Lysogory and Matopolska terranes close to their present location on the SW margin of the EEC formed a major promontory on this margin of the continent. Moesia may also have formed part of this block. Both late Ordovician accretion of Avalonia and early Carboniferous accretion of the Armorican Terrane Assemblage (ATA) attached new continental material around the Bruno-Silesian Promontory (BSP). Palaeozoic faunal affinities and inherited isotopic signatures similar to those of Avalonia seen in the Istanbul block of NW Turkey, and in minor thrust slices in Moravia and Romania, suggest that easternmost Avalonia was severed, on collision with the BSP, and migrated east along the southern margin of the EEC. Likewise, the similarities to the ATA of the Balkan, Istranca, Sakarya and eastern Pontides blocks suggests that more easterly components of the ATA were detached at the BSP and migrated east. All the newly accreted blocks contain similar Neoproterozoic basement indicating a peri-Gondwanan origin; Palaeozoic plume-influenced metabasite geochemistry in the Bohemian Massif may explain their progressive separation from Gondwana before their accretion to the EEC. Inherited ages from Avalonia contain a 1.5 Ga 'Rondonian' component arguing for proximity to the Amazonian Craton at the end of the Neoproterozoic; Armorican terranes lack such a component, suggesting that they have closer affinities with the West African Craton. Models showing the former locations of these terranes and the larger continents from which they rifted, or later became attached to, must conform to both these constraints and those provided by palaeomagnetic data. In the late Neoproterozoic and Palaeozoic, these smaller terranes, some containing Neoproterozoic ophiolitic marginal basin and magmatic arc remnants, probably fringed the end-Proterozoic supercontinent as part of a 'Pacific-type' margin. When this margin fragmented, most resulting fragments accreted to the EEC.

The S W margin of the East European Craton (EEC) is marked by the Trans-European Suture Zone (TESZ), traceable from the Black Sea coast of Romania to North Germany, the Baltic Sea, Denmark and the North Sea (Fig. 1), despite being everywhere concealed beneath thick sedimentary cover (Gee & Zeyen 1996; Pharaoh 1999). A description of the nature, age and geometry of this fundamental feature of European geology has been provided by Pharaoh et al. 2006). On the SW side of this zone a collage of blocks accreted to the EEC margin during the Palaeozoic following the end of the Cambrian. SW of the TESZ and north of the A l p i n e - C a r p a t h i a n Front, the basement structure of Central Europe has long been known in outline. Evidence from geophysical compilations, geological information provided by deep boreholes, and outcrops of Palaeozoic and older rocks across central Germany and in the Bohemian Massif reveals a mosaic of microcontinental blocks, derived from different sources and shown by isotopic dating, and biostratigraphic and palaeomagnetic evidence, to have become attached to the EEC in their present locations during the Palaeozoic. An early phase of terrane emplacement was largely complete by the end of the Cambrian. These terranes, including the BrunoSilesian Terrane, with the Lysogory and Ma~opolska blocks of the Holy Cross Mountains in Poland appear to have been situated in approximately their present position since that time. They may extend southwards to the Danube, approximately as far as the K r e m s - V i e n n a Line in Austria (Dudek 1980) and also be linked to the SE beyond the Carpathians with the central and southern Dobrogea and the Moesian Platform in Romania. Whether these terranes comprise displaced portions of the EEC (Cocks 2002) or an early accreted fragment derived from Gondwana (e.g.

Belka et al. 2000, 2002) has been hotly debated, and the arguments have been set out more fully by Pharaoh et al. (2006). However, the end-Cambrian attachment of these blocks to the EEC also precludes any pre-Ordovician association with terranes accreted later, particularly Avalonia, which was still attached to Gondwana in the early Ordovician (Winchester et al. 2002). Whatever their derivation, if it is accepted that the subsequent Devonian displacement of these terranes, suggested from palaeomagnetic and structural evidence (Lewandowski 1993; Mizerski 1995), was restricted in extent (Cocks 2002), they must have formed a major promontory extending from the SW margin of the EEC during most of the Palaeozoic. This is referred to below as the Bruno-Silesian Promontory (BSP) and its geometry is crucial in explaining the mechanisms of attachment of the microcontinents that subsequently accreted to the EEC. Excluding the small portions of Laurentian crust forming Scotland and N W Ireland, the main microcontinental blocks that subsequently accreted to the SW margin of Europe during the Palaeozoic are known as Avalonia and the Armorican Terrane Assemblage (ATA; Franke 2000; Tait et al. 2000). Both of the latter were derived from Gondwana, but rifted from it at different times. They therefore possess characteristic Proterozoic basement, affected by end-Proterozoic Panafrican (locally termed Cadomian) magmatism and deformation, which therefore does not distinguish between them. Factors distinguishing these microcontinental blocks are: (1) the timing of accretion to the EEC; (2) the presence or absence of an inherited c. 1.5 Ga 'Rondonian' event, which seems, in particular, also to be a characteristic feature of the 'Ganderian' portion of Avalonia; (3) the occurrence of either distinctive

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 323-332. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. A map showing the distribution of crustal blocks and Palaeozoic deformation belts in Central and SE Europe. ABDB, Anglo-Brabant Deformation Belt; AD, Ardennes; ADF, Alpine Deformation Front; AM, Armorican Massif; BB, Brabant Massif; BK, Balkan Terrane; BM, Bohemian Massif; BSM, Bruno-Silesian Massif; CACC, Central Anatolian Crystalline Complex; CAU, Caucasus; CD, Central Dobrogea; CDF, Caledonian Deformation Front; CDO, Central Dobrogea; CM, Cornubian Massif; COF, Capidava-Ovidiu Fault; DR, Dronsendorf Unit; EA, Ebbe Anticline; EL, Elbe Lineament; EP, Eastern Pontides; GF, Gf6hl Unit; HCM, Holy Cross Mountains; HPDB, Heligoland-Pomerania Deformation Belt; IB, Istanbul Block; IMF, Intra-Moesian Fault; Istr, Istranca Terrane; KLZ, Krakow-Lubliniec Zone; LU, Lysogory Unit; L-W, Leszno-Wolsztyn High; MC, Midlands Microcraton; MM, Matopolska Massif; MN, Mfinchberg Nappe; MNSH, Mid-North Sea High; MP, Moesian Platform; MST, Moravo-Silesian Terrane; NASZ, North Armorican Shear Zone; NBT, North Brittany Terrane; NDO, North Dobrogea; NGB, North German Basin; PCF, Peceneaga-Camena Fault; Pom, Pomerania; POT, Polish Trough; R, Rfigen Island; RFH, Rynkcbing-Fyn High; RG, RCnne Graben; Rh, Rhodope; SASZ, South Armorican Shear Zone; SBT, South Brittany Terrane; SDO, South Dobrogea; SGF, Sfantu Gheorghe Fault; SNSLT, South North Sea-Luneberg Terrane; SP, Scythian Platform; S-TZ, Sorgenfrei-Tornquist Zone; Su, Sudetes; TB, Teplfi-Barrandia; T-TZ, TeisseyreTornquist Line; VF, Variscan Front; ZZ, Zonguldak Zone. 'Celtic' (e.g. Avalonian) faunas that were unique to Avalonia during the mid- and late Ordovician, or of distinctive mixed Siluro-Devonian faunas characteristic of the ATA; (4) the presence of late Ordovician glaciogenic sediments, a Gondwana feature shared by the ATA, but not by Avalonia, which had by that time already migrated into lower latitudes (Cocks et al. 1997).

Avalonia Precambrian and early Palaeozoic basement exposed in central England, Belgium and western Germany is widely accepted as part of Avalonia, this Ordovician microcontinent extending west as far as New England, and being best exposed in the Avalon Peninsula of Newfoundland, after which it is named. Avalonian basement in central England typically consists of late Proterozoic intrusive, volcanic and sedimentary rocks (e.g. Thorpe et al. 1984; Pharaoh & Gibbons 1994; Strachan et al. 1996), affected by end-Proterozoic or pre-Early Cambrian deformation. In the English Midlands it was little affected by later movements, and is overlain by a thin Lower Palaeozoic shallow marine sedimentary sequence, succeeded conformably by Devonian terrestrial deposits: the 'Old Red Sandstone'. For this reason it has sometimes been called the 'Midlands Microcraton' (e.g. Turner

1949; Pharaoh et al. 1987). A fuller description has been given by Pharaoh et al. (2006). The Midlands Microcraton is flanked to the NW by much thicker Lower Palaeozoic successions, strongly deformed in an early Devonian 'Acadian' event (Soper et al. 1987), deposited on Avalonian basement and exposed in Wales, the English Lake District and SE Ireland. Boreholes in eastern England reveal that similarly deformed rocks (Pharaoh et al. 1987) containing Upper Ordovician calc-alkaline volcanic rocks, as an apparent continuation of the Lake District Arc, also extend from eastern England to Belgium, where they are exposed in the Brabant Massif (Andr6 et al. 1986; Pharaoh et al. 1991). These rocks have been termed (Winchester et al. 2002) the Anglo-Brabant Deformation Belt (ABDB), in which the deformation is inferred to have developed in the early Devonian (Acadian). This deformation belt is thought to mark a zone of crustal suturing inherited from the late Ordovician soft collision of Avalonia and Baltica (Verniers et al. 2002). Unusually thick lower Cambrian deposits in Brabant suggest the presence of rifting, which may have thinned and weakened the Proterozoic basement, thereby controlling the location of this deformation belt (Winchester et al. 2002). Because the ABDB contains no known ophiolitic material, it is not thought to mark a zone of microcontinent collision and seaway destruction, even though,

DETACHED TERRANE FRAGMENTS IN EEC as suggested by Pharaoh et al. (1993), it may separate crusts with somewhat differing structures. An area of stable basement, indicated by seismic traverses and termed the Southern North Sea-Luneberg Terrane (SNSLT) by Pharaoh et aL (1995), lies east of the ABDB, NE of the DowsingSouth Hewett Fault Zone-Lower Rhine Lineament, themselves younger reactivations of earlier major fault-lines (Pharaoh 1999). It has been more recently dubbed 'Far Eastern Avalonia' (Winchester et al. 2002), because of its geological similarities to Avalonia. Although crystalline basement is mostly concealed, one outcrop area, the 574 _+ 3 Ma Wartenstein Gneiss (Molzahn et al. 1998), is exposed far to the south in the South Hunsriick at the SE margin of the Rhenish Massif; lying to the south of the Variscan Front, this typically calc-alkaline granitoid gneiss of late Neoproterozoic age is broadly comparable with granitoid rocks in the Avalonian basement in central England. In addition, Samuelsson et al. (2002a) noted that the eNd(t) trends of Ordovician sedimentary rocks from the Ebbe Anticline of NW Germany (Fig. 1), situated NE of the Lower Rhine Lineament and therefore on SNSLT basement, match those from the Welsh Basin and the Brabant Massif, but are different from those in Brittany and Iberia. They also concluded that these rocks formed part of Avalonia. Hence the ABDB is interpreted as an intra-Avalonian zone of local subduction, initiated above a failed Cambrian rift, where the basement had been thinned and weakened, in response to the distortion and anticlockwise rotation of part of Avalonia as it moulded itself on to the margins of the EEC and Laurentia (Verniers et al. 2002). The late Ordovician timing of the accretion of the SNSLT to the EEC (Vecoli & Samuelsson 2001; Samuelsson et al. 2000b), which only slightly predates Avalonian convergence with Laurentia, based on evidence from Atlantic Canada (e.g. Cawood et al. 1994), and the onset of Windermere Supergroup sedimentation in the English Lake District (Cooper et al. 1993), also suggests that the SNSLT should be considered as part of Avalonia. If so, the Heligoland-Pomerania Deformation Belt (HPDB), which largely comprises a zone of overthrusting, marks the collision zone between Avalonia and the EEC. It shows little evidence of contemporary magmatism in boreholes, but geophysical evidence may indicate the presence of buried arc volcanic rocks (Williamson et al. 2002; Pharaoh et al. 2006), suggesting that convergence was accompanied by southdirected subduction beneath the Avalonian margin.

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(Fig. 2). Recognizing the Avalonian affinities of such detached extensions, where later metamorphism may have destroyed faunal evidence, is problematic; discrimination between Avalonia and the ATA must rely instead on the presence (as in Avalonia), or absence (as in the ATA) of mid-Proterozoic (1.45 Ga) inherited zircon dates, related to the previously adjacent (pre-rifting from Gondwana) Rondonian event in the Amazonian Craton.

Avalonian easternmost extremities Although totally concealed by thick Mesozoic and Cenozoic sequences in the Polish Trough, the easternmost end of Avalonia appears to abut the Lysogory and Matopolska blocks of the Holy Cross Mountains, which form part of the Bruno-Silesian Promontory (BSP). Further south, tectonic structures along the western margin of the Bruno-Silesian Block show highly oblique (dextrally transpressive), complex overthrusting to the east (Moldanubian and Drinova thrusts) in the early Carboniferous, between 350 and 330 Ma (Schulmann & Gayer 2000). This junction is traceable northwards beneath the thick sedimentary cover of the Polish Trough, using seismic profiling. Both the Polonaise PI and TeisseyreTornquist Zone (TTZ) profiles (Grad et al. 1999; Jensen et al. 1999) show a clear change of mid-crustal structure north of the Moldanubian Thrust, suggesting that it continues northward as a major feature (Moravian Line of Winchester et al. 2002). The TTZ profile shows the mid-crustal break to be displaced eastwards compared with Polonaise P 1, indicating dextral displacement of the Moravian Line by strike-slip faulting between the two profiles, perhaps along the Dolsk Line (Grad et al. 2002). On accretion, Avalonia was unlikely to fit exactly into the position against the EEC margin that it now occupies, bounded to the east by the BSP. Therefore, former eastern extensions are likely to have been detached by sheafing initiated by collision with the BSP

Fig. 2. (a) Sketch map illustratingthe supposed configurationof Avalonia on its initial impact with the Bruno-SilesianPromontory. (Note the detachment and displacementeastward of its eastern extremity). Abbreviationsas in Figure 1. (b) Sketch illustrating the likely configurationof the Armorican Terrane Assemblageon initial impact with the Bruno-SilesianPromontory in the early Carboniferous. GWSO, Giessen-Werra-Sudharz Ocean; MGCH, Mid-German CrystallineHigh. Other abbreviationsas in Figure 1. (c) Likely configuration of crustal blocks followingthe main Variscan Orogeny. (Note the eastward displacementof the eastern Variscides.)

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J.A. WINCHESTER ETAL.

The M o r a v i c u m Nappe

In NE Austria and Moravia, the Dobra Gneiss (Gebauer & Friedl 1993) and Bittesch Gneiss (Friedl et aL 2000) both yield midProterozoic inherited zircon dates of c. 1.5 Ga, contemporary with the Rondonian orogeny affecting the NW side of the Amazonian Craton (Tassinari et al. 2000; Cawood et al. 2003). They also show clusters of inherited dates at 1.2 and 1.78 Ga, which have also been noted both from Ganderian (Avalonian) basement in southern New Brunswick and from the Amazonian Craton (where they are termed the Sunsas and Rio Negro provinces; Tassinari et al. 2000). Because these rocks were thought to be linked to the Bruno-Silesian massif (BSM), these dates were interpreted to assign an Avalonian affinity to the BSM (Finger et al. 2000). However, if the latter was attached to the EEC, at least since the end of the Cambrian, it cannot also have been part of Avalonia. However, both the Dobra and Bittesch gneisses are situated in the deformed Moravicum nappe, emplaced onto the western margin of the Bruno-Silesian block by movement on the Drinova Thrust (Hock et al. 1997; Melichar & Kotkova 2003). They therefore need not belong to the BSM, and could instead represent small detached slivers of Avalonian crust thrust obliquely onto the margins of the BSM. The Danubian basement o f the southern Carpathians in Romania

Exposed in nappes towards the western end of the southern Carpathians of Romania are rocks of the Lower and Upper Danubian basements (Berza et al. 1983, 1994, 2004; Iancu & Berza 2004). Late Neoproterozoic magmatic rocks (Liegeois et al. 1996) including calc-alkaline granitoids such as the Tismana Pluton (567 + 3 Ma, U - P b ) are unconformably overlain by a clastic sedimentary succession of Late Ordovician-Early Silurian age, apparently devoid of recorded glaciogenic diamictites. These rocks underwent a Devonian (?Acadian) deformation and, although clear faunal or palaeomagnetic evidence remains lacking, and some claim that the older rocks may be exhumed Moesian basement (Sandulescu 1984, 1994), these characteristics make an Avalonian affinity an alternative possibility. The Istanbul Block

Further east, in the Istanbul Block of NW Turkey, lithologically similar Ordovician rocks have yielded 'Celtic' (e.g. Avalonian) faunas (Kozur & G6nctio~lu 1998; Dean et al. 2000). Some studies have distinguished, in this area, separate Istanbul and Zonguldak terranes, based on differences of facies between Palaeozoic rocks close to Istanbul and those further east (Goncuoglu & Kozur 1998, 1999; Kozur & Gincuoglu 2000). Also, whereas sedimentation near Istanbul seems to have continued uninterruptedly from the Early Ordovician to the MidCarboniferous, further east there are increased signs of Early Devonian uplift and deformation. However, no clear boundary between these terranes can be mapped, and it seems more likely that these differences relate to facies changes within a single terrane, analogous to those seen within Eastern Avalonia, in which Central English and Welsh successions can be contrasted. According to this analogy, the continuous sequence in the Istanbul area corresponds better to the Lower Palaeozoic shelf deposits in Central England, whereas the overlying Upper Palaeozoic rocks, containing mixed shales, cherts and limestones and their benthic fauna (Tokay 1955) have more in common with the marine Devonian and Carboniferous units of the Rheno-Hercynian zone, which in both SW England and Germany overlie Avalonian basement. By contrast, the Zonguldak area, after the deposition of an initial pebbly quartzite, reveals a shale-dominated Ordovician sequence more characteristic of the Welsh Basin, and this

analogy is enhanced by the presence of a Lower Devonian unconformity. Ensuing Late Devonian and Early Carboniferous sedimentation is dominated by limestone, which is in turn succeeded by regressive flood plain deposits containing coals (Yanev et al. 2006). Both this Upper Palaeozoic sequence and the major unconformity above which Permo-Triassic continental clastic rocks occur are analogous to the sequence seen in England overlying Avalonian basement. The Avalonian link for the Istanbul Block, suggested by both the lithological sequences and the faunal evidence, is further strengthened by a mid-Proterozoic discordia date of 1 4 4 5 _ 24 Ma obtained from a late Neoproterozoic granite in the Karadere basement (Chen et al. 2002). A further, less well-constrained inherited age of 1189 _+ 110 Ma from a metatonalite in the same area (Chen et al. 2002) also resembles some of the ages obtained from the Moravian Nappe, and from the Ganderian part of Avalonia in southern New Brunswick. This information, combined with evidence of Silurian deformation in the northern part of the block near Zonguldak (although not nearer to Istanbul itself) also is consistent with Siluro-Devonian docking with Baltica of an Avalonian fragment, and post-Carboniferous (Variscan) deformation, as in the Rheno-Hercynian zone of Western Europe, may attest to the deformation of the southern Avalonian margin caused by accretion of terranes of the ATA. Moesia and Dobrogea

In contrast, there is little evidence to link the basement rocks of any part of either the Moesian Platform or the Dobrogea with Avalonia. Neoproterozoic and Lower Palaeozoic rocks of the central and southern Dobrogea, in eastern Romania, appear to have similarities to those in the Holy Cross Mountains of Poland rather than the Istanbul Block, and borehole sections indicate that this is also true of the northeastern part of the Moesian Platform, NE of the Intra-Moesian Fault (IMF). Metamorphic basement to the southern Dobrogea has also yielded Mid-Proterozoic ages (Giusca et al. 1967), and is considered to have affinities with rocks in the Ukrainian Massif of the EEC. Residual gravity and magnetic anomalies also indicate a link with the EEC (Ioane & Atanasiu 2000). However, as the Bruno-Silesian Massif itself may also have originated adjacent to the southern EEC, it is possible that this basement may also underlie the Central Dobrogea at greater depth. West of the IMF, Neoproterozoic granitoids have also been recovered from boreholes (Savu & Paraschiv 1985), and Cambrian rocks from deep boreholes (Mutiu 1991) have yielded numerous fragmentary specimens of the trilobites Paradoxides (species undetermined) and Peronopsis fallax (Linnarsson). A. Rushton (pers. comm.) considered that the latter fossil, although apparently showing some affinities with species associated with the Baltican margin (Rushton & McKerrow 2000), is a widely recorded member of an outer shelf fauna, which may have been able to cross geotectonic boundaries. Thus, although neither trilobite can be used to prove conclusively a periBaltican affinity of the Moesian Platform, similar to that proposed for the Dobrogea or Bruno-Silesian regions, this remains its most likely affinity. However, the presence of a widespread unconformity beneath Silurian rocks (Iordan 1984) may indicate a Late Ordovician uplift, which could record deformation associated with collision of Avalonian fragments immediately to the south.

Avalonian eastern extremities: mechanism of emplacement Lower Palaeozoic rocks deposited directly on the EEC margin are all consistent with a passive margin setting: the narrowing of the Tornquist Sea appears to have occurred exclusively by subduction under the Avalonian margin up to the time of collision in the late

DETACHED TERRANE FRAGMENTS IN EEC Ordovician. Further east, however, the southern margin of the BSP is now concealed beneath younger rocks associated with the much later formation of the Carpathians. Yet, for fragments of Avalonia to migrate east with dextral transpression, continued subduction was probably needed, but the Silurian rocks in the Istanbul Block do not include magmatic rocks indicative of continued subduction. It therefore seems likely that, on collision of easternmost Avalonia with the BSP, a change of polarity of subduction occurred analogous to that recorded in New Brunswick with closure of the Iapetus Ocean in that sector (van Staal et al. 1991). With continued subduction, but this time northwarddirected, the buoyant continental fragments of Avalonia could be transported eastwards with sinistral transpression, along the southern margin of the EEC, until 'trapped' in re-entrants of the continental margin (Fig. 2a). Smaller fragments, such as the Moravicum Nappe and the Danubian Terrane, may be interpreted as slivers on the continental margin, abandoned during the eastward progress of the Istanbul Block. The present position of the Istanbul Block partly arises from its southward displacement during the opening of the Black Sea basins, since the Cretaceous. There seems to be no clear westward continuation into Moesia or the Dobrogea, which suggests that almost the entire Avalonian fragment was displaced southwards as the Istanbul Block.

Accretion history of the Armorican Terrane Assemblage: mechanisms of migration and ocean closure The Armorican Terrane Assemblage (sensu Franke 2000; Tait et al. 2000), also previously referred to as 'Peri-Gondwanan Terranes' or 'Northern Gondwana terranes' (e.g. Erdtmann & Kraft 1999), is exposed in a series of massifs across much of SW to Central Europe from Iberia to Poland. In Western and Central Europe, these terranes were accreted to Laurussia during the Late Palaeozoic. The term 'Variscan Orogeny', which has been used to describe the deformation and magmatism associated with the closure of the Rheic Ocean, its successor basins, and basins separating constituent terranes within the ATA, does not fully convey the complexity of these multiple accretions: a revised overview following intensive study of the constituent terranes in Central Europe and their accretion histories has been given by Pharaoh et al. (2006). In summary, early Devonian metamorphism and magmatism (sometimes called 'Caledonian', but historically and collectively termed Eo-Variscan elsewhere in Hercynian Europe; e.g. Faure et al. 1997; Shelley & Bossi~re 2000) was confined in the northern Bohemian Massif to isolated high-grade metamorphic rocks in the Gdry Sowie Block (GSB; Brueckner et al. 1996; O'Brien et al. 1997) and the Mtinchberg klippe (395-390 Ma; Kreuzer et al. 1989; Stosch & Lugmair 1990). It may record local tectonothermal and hence collisional activity between migrating platelets of the ATA, with subsequent exhumation. Whereas high-P metamorphism was initiated somewhat earlier in the GSB than further west, as indicated by growth of metamorphic (granulitefacies) zircon at 402 __ 0.8 Ma (O'Brien et al. 1997), subsequent late Devonian H T - M P metamorphism in the GSB is well constrained by U - P b monazite ages (van Breemen et al. 1988; Br6cker et al. 1998; Timmermann et al. 2000) and appears to be contemporary with H P - L T metamorphism along the contact zone of the Saxo-Thuringian and Teplfi-Barrandian blocks between 380 and 365 Ma. Further west, recently obtained mid- to late Devonian dates for the emplacement and metamorphism of the Lizard Peridotite and associated rocks at the southern margin of the Cornubian Massif (Sandeman et al. 2000; Nutman et al. 2001) reinforce the parallel with the Giessen-Werra-Sudharz Ocean (Franke 2000), as the latter also underwent contemporary metamorphism and deformation. The latter has been interpreted as an obducted successor basin to the Rheic Ocean.

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In the Karkonosze-Izera complex (central West Sudetes) tectonic exhumation was earlier and greater in the SE. This is shown by: (1) early kinematic indicators in mylonitic ductile shear zones (Mazur 1995; Seston et al. 2000); (2) decrease in metamorphic grade from garnet zone in the SE to chlorite zone in the NW (Baranowski et al. 1990; KachlN & Patorka 1998; Collins et al. 2000); (3) northwestward decrease of 4~ cooling ages (Marheine et al. 1999); (4) progressively later flysch sedimentation onsets towards the NW. Also, late Devonian unconformities in the central West Sudetes occur between the Ktodzko metamorphic complex and the overlying Bardo Unit (Hladil et al. 1998; Kryza et al. 2000), while Late Devonian and Carboniferous coarse-grained clastic sedimentary deposits, derived from exhumed metamorphic complexes to the east, were deposited in syntectonic basins (Aleksandrowski & Mazur 2002). Deformation and metamorphism, which started in the central West Sudetes in pre-late Devonian times (e.g. Hladil et al. 1998) continued until the Tournaisian in both the northwesternmost frontal parts of the West Sudetic orogenic wedge, where m~langes formed in the Kaczawa Complex (Collins et al. 2000), and in the metamorphic core of the complex, as in the OrlicaSnieznik area, where HP metamorphism produced eclogites. This range of dates suggests that a series of small-scale collisional events occurred, consistent with a progressive aggregation of the constituent terranes of the ATA. In the West Sudetes Carboniferous metamorphism was followed by tectonic exhumation of deeply buried crustal slices (353-350 Ma) and the superimposition of a greenschist- to lower amphibolite-facies overprint dated at 345-340 Ma). 4~ dating (325-320 Ma) suggests that metamorphism was complete by the mid- to late Carboniferous (Marheine et al. 2000), a timing supported by the age of deposition in adjacent intramontane basins. These Carboniferous events are generally considered to reflect the docking of the amalgamated ATA with the Avalonian and Bruno-Silesian margin of the growing Laurussian supercontinent. The range of dates suggests that collision was not a simple process: it probably began earlier where the accreting ATA first impinged on promontories, such as that of the Bruno-Silesian Massif, and occurred later further west. Deformation of Devono-Carboniferous sedimentary sequences on the Laurussian passive margin in the Cornubian, Rhenish and Bruno-Silesian massifs, as a result of this collision, produced the only significant late Palaeozoic deformation to affect both Avalonia and Bruno-Silesia. As the ATA approached Laurussia, subduction was south-dipping beneath its leading edge, causing the formation of an arc edifice preserved as volcanic rocks of the Mid-German Crystalline High (MGCH), with its associated oceanic back-arc basin, the Giessen-Werra-Siidharz 'ocean'. Subduction of this successor back-arc basin, which developed on the south side of the Rheic Ocean, occurred in DevonoCarboniferous time, with obduction of fragments of it, originally developed on the southern side of the ocean, eventually thrust northwards across the Rheic Suture, so that they are now preserved as ophiolitic outliers assigned to the Giessen-Werra-Stidharz or Selke Nappe (e.g. Franke 2000), north of the Rheic Suture. Thus, the MGCH marks the superimposition of both late Silurian-Devonian arc magmatism on the Avalonian margin below the south-dipping Rheic Suture, and Carboniferous age volcanism above it (Oncken 1997). Small magnetic highs seem to indicate a continuation of the volcanic centres within the MGCH eastwards into Poland as far as a point just NE of the LesznoWolsztyn High, corresponding to the location of the Moravian Line.

Eastern extremities As with Avalonia, the eastern extremities of the ATA abut the Bruno-Silesian Massif (BSM), which must have still formed a

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J.A. WINCHESTERETAL.

promontory on the Laurussian margin at the time of ATA accretion. Without a perfect fit, the ATA presumably included crustal blocks that converged with Laurussia further east, and that might be expected to be accreted to the southern margin of the BSM. However, because the latter margin is overthrust by the Carpathian-Alpine Front, the mechanism for distribution of ATA-related blocks further east is obscured. However, rocks apparently subjected to Variscan-age metamorphism, often intruded by mid-Carboniferous post-orogenic granitoids, occur as basement inliers in the Carpathians, such as the Tatra Mts. In the western Tatra Mts, metamorphic rocks containing amphibolites with similar chemistry to those in the West Sudetes (Gaw~da et al. 2000) are cut by post-metamorphic Variscan granitoid rocks, dated by both 4~ and Rb-Sr methods at 300-330 Ma (Burchart 1968; Janak 1994). In the Romanian Carpathians, the Getic-Supragetic basement (Iancu & Berza 2004) contains similar lithologies subjected to Variscan deformation and metamorphism. Further SE, the Balkan terrane exposed in western Bulgaria (Fig. 1), and also sampled north of Sofia in the Svoge borehole, contains mid-Ordovician faunas similar to those of Bohemia and North Africa (Gutteriez-Marco et al. 2003), and typical of a cold, peri-Gondwanan environment (Haydoutov & Yanev 1997). Mid-Ordovician trilobites ( C y c l o p y g e p r i s c a ) occur in shales overlain by Ashgill diamictites, indicating that the Balkan Terrane remained attached to Gondwana in high latitudes long after Avalonia had rifted off and migrated to lower latitudes. Built upon a basement of Neoproterozoic ophiolites and Cambrian calc-alkaline volcanic rocks, the thick Palaeozoic sequence also includes Silurian argillites, Devonian clastic deposits and an unconformity above the Lower Carboniferous units. All these indicators point to an 'ATA' Gondwana affinity, with collision with Moesia during the Carboniferous. However, the presence of a late Cambrian subduction-related sequence (493 Ma, Carrigan et al. 2003) also needs explanation. Although this could be interpreted as the product of intercontinental collision, it could also be the result of a Cambrian arc-continent collision, closing the intervening oceanic back-arc basin that had been formed in the late Neoproterozoic. If so, the Balkan terrane could represent yet another portion of the NeoproterozoicCambrian supercontinent-fringing series of arcs and back-arc basins. To the SE, the Balkan Terrane is structurally juxtaposed with the Rhodope (Thracian) and Strandja terranes, which nevertheless seem to share its Palaeozoic continental affinities. Still further east, in NW Turkey, the basement to the Sakarya Zone shares a similar Palaeozoic history to blocks comprising the ATA, in that it underwent Carboniferous metamorphism, followed by intrusion of late Carboniferous post-orogenic granites (Yilmaz et al. 1997). A rupture of the ATA, similar to that experienced by Avalonia on collision with the Bruno-Silesian Promontory, might explain, in the same way, the eastward migration of displaced ATA-related blocks.

Why did Avalonia and the ATA separate from Gondwana? The composition of Palaeozoic magmatic rocks provides clues to the causes of the separation of Avalonia and the ATA from the Gondwana margin. In the northern Bohemian Massif extensive bimodal magmatism occurred in the early Ordovician, with bursts of magmatism continuing until the Devonian. Early, mainly acidic magmatism of Cambro-Ordovician age (e.g. Korytowski et al. 1993; Krrner et al. 1994; Philippe et al. 1995; Hammer et al. 1997) shows calc-alkaline chemistry, which some interpreted as evidence for an arc or active continent margin tectonic setting (e.g. 0liver et al. 1993; Krrner & Hegner 1998). Others suggested that the absence of supporting geological evidence for an arc edifice at the time suggested that chemical

characteristics of the intrusions were inherited from extensive melting of the calc-alkaline Panafrican basement (Kryza & Pin 1997; Aleksandrowski et al. 2000; Floyd et al. 2000). Subsequent, dominantly basic volcanism was associated with clastic basin-fill metasedimentary rocks, typical of magmatism associated with an extensional tectonic setting. Minor associated felsic volcanic rocks were shown by Sm-Nd systematics and their REE distribution to result from continued melting of continental crust (Fumes et al. 1994; Patofika et al. 1997; Dostal et al. 2000), whereas the compositional range of the basic rocks (e.g. Floyd et al. 1996, 2000; Winchester et al. 1995, 1998) indicated magma production resulting from the interaction of an enriched plume with both asthenospheric and sediment-contaminated lithospheric mantle sources (Floyd et al. 2000). Although the preserved volume of magmatic rocks is smaller than younger plume-influenced magmatic provinces, it has widespread correlatives in many parts of Western Europe, including the Massif Central (Briand et al. 1991, 1995) and Massif des Maures (B. Briand, pers. comm.) in France and NW Spain (e.g. Peucat et al. 1990). Floyd et al. (2000) suggested that plume-induced magmatism could also explain the amount of heat needed to melt substantial volumes of lower crust to produce the major granitoid bodies, this providing a possible mechanism for the fragmentation of the Armorican Terrane Assemblage (ATA) as it separated from Gondwana, and the repeated rifting of crustal fragments from the Gondwana margin, including Avalonia and the ATA.

Palaeozoic palaeogeographical evolution and accretions to the EEC Recent reconstructions show that the main pre-Alpine, Central European and related microcontinents formed an active continental margin (ACM) to the Pannotian supercontinent, with Avalonia adjacent to the Amazonian Craton, based on the presence of inherited 1.5 Ga 'Rondonian' ages obtained from rocks in Nova Scotia (Nance & Murphy 1994) and central England (Tucker & Pharaoh 1991). To the east (present co-ordinates) the ACM extends through the ATA (shown adjacent to the North African Craton as it lacks inherited 'Rondonian' ages) and other blocks that are thought to have separated from their peri-Gondwanan positions later, notably the basements of Italy, the Pannonian blocks, and the Tauride basement of southern Turkey. The presence of late Neoproterozoic ophiolitic fragments within this ACM (e.g. Yifgitba~ et al. 1999; Scarrow et al. 2001) attests to the obduction of successor basins and suggests that the continental margin was originally of West Pacific rather than Andean type. Shared end-Proterozoic calc-alkaline magmatism and deformation affecting all the accreted blocks records their former location along an active margin to the end-Proterozoic supercontinent Pannotia. During the Cambrian, subduction along this margin appears to have ceased or been greatly reduced, whereas during the Tremadoc, renewed subduction resulted in calc-alkaline magmatism and the formation of large back-arc basins (the Gander Arc and associated ophiolites) in the western part of the margin, now preserved in Atlantic Canada. During the Llanvirn Stage, bimodal acid-basic magmatism marks the detachment of Avalonia, possibly as more than one block (Pharaoh 1999; Banka et al. 2002; Winchester et al. 2002), marking the break-up of the Gondwana margin, and renewed arc magmatism in the 'Caradoc' Stage (Exploits and Lake District arcs) marks its rapid northward migration, narrowing the Iapetus Ocean. By this stage, a widening Rheic Ocean opened between Avalonia and the Gondwana margin, from which parts of the ATA were already starting to rift as a series of linked blocks. By the early Silurian, Avalonia had moulded itself onto the TESZ margin of Baltica, with its easternmost extremity detached and displaced eastwards along the southern margin of the new

DETACHED TERRANE FRAGMENTS IN EEC

supercontinent of Laurussia, comprising Avalonia, Baltica and Laurentia. By this time also, many blocks of the ATA, already rifted into an archipelago or related microcontinents, had separated from Gondwana, narrowing the Rheic Ocean, although the widespread occurrence of late Ordovician glacial deposits (lacking in Avalonia) indicates that significant separation from Gondwana by even the earliest blocks occurred only after the end of the Ordovician. However, the contrast between Silurian microfaunas of the French Armorican terranes and those of the Brabant Massif (Verniers 1982), suggests that the Rheic Ocean remained broad. Subduction was initiated along the southern margin of Avalonia, marking the earlier stage of volcanism in the Mid-German Crystalline High. As terranes of the ATA moved away from the Gondwana margin, the new seaway being formed was the Proto-Tethys Ocean. During the Devonian (Emsian), high-P, low-T metamorphism, recording subduction and closure of intervening seaways within the ATA, suggests that amalgamation of individual ATA terranes had begun, eventually resulting in the production of a single ATA microcontinent. Southward subduction, marked by renewed volcanism in the Mid-German Crystalline High, recorded the final stage in the approach of the now-amalgamated ATA to Laurussia, also impelled by Gondwanan convergence. Contact with the BSP, still not firmly enough attached to Baltica to prevent some displacement and relative rotation, was marked by dextral strike-slip faulting along its western margin. This was followed by the docking of most ATA blocks along the southern margin of Laurussia. Easternmost parts of the ATA were detached on collision with the BSP and displaced eastwards by sinistral faulting to form the Variscide basement seen in Carpathian inliers, the Balkan and Thracian terranes of Bulgaria, and the Sakarya and Eastern Pontide crustal blocks of N W Turkey. These investigations, and the collation of information, were supported by the EU-funded PACE (Palaeozoic Amalgamation of Central Europe) TMR Network, No. ERBFMRXCT97-0136. Part of the study is sponsored by the FWO Research Project No. G.0094.01 'Tectonics of the Early Palaeozoic basin development in NW Europe: basin analysis and magnetic fabric analysis in the Belgian Caledonides'. The contribution of T.C.P. appears with permission of the Executive Director, British Geological Survey (NERC). Particular thanks are expressed to M. C. G6nctio~lu (Turkey), I. Haydoutov and S. Yanev (Bulgaria), A. Okay (Turkey), and A. Rushton (UK), who all provided valuable additional data for inclusion.

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VECOLI, M. & SAMUELSSON,J. 2001. Quantitative evaluation of microplankton palaeobiogeography in the Ordovician-Early Silurian of the northern TESZ (Trans-European Suture Zone): implications for the timing of the Avalonia-Baltica collision. Review of Palaeobotany and Palynology, 115, 43-68. VERNIERS, J. 1982. The Silurian Chitinozoa of the Mehaigne area (Brabant Massif Belgium). Professional Paper of the Belgian Geological Survey, 1982(6), 192. VERNIERS, J., PHARAOH,T. C., ANDRI~,L., ETAL. 2002. The Cambrian to Mid-Devonian basin development and deformation history of Eastern Avalonia east of the Midlands Microcraton: new data and a review. In: WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) Palaeozoic Amalgamation of Central Europe. Geological Society London, Special Publications, 201, 47-93. WILLIAMSON,J. P., PHARAOH,T. C., BANKA, D., THYBO, H., LAIGLE,M. & LEE, M. K. 2002. Potential field modelling of the BalticaAvalonia (Thor-Tomquist) suture beneath the southern North Sea. Tectonophysics, 360, 47-60. WINCHESTER, J. A., FLOYD, P. A., CHOCYK, M., HORBOWY, K. & KOZDROJ, W. 1995. Geochemistry and tectonic environment of Ordovician meta-igneous rocks in the Rudawy Janowickie Complex, SW Poland. Journal of the Geological Society, London, 152, 105-115. WINCHESTER, J. A., FLOYD,P. A., AWDANKIEWICZ,M., PIASECKI,M. A. J., AWDANKIEWICZ,H., GUNIA, P. & GLIWICZ, T. 1998. Geochemistry and tectonic significance of metabasic suites in the G6ry Sowie Block, SW Poland. Journal of the Geological Society, London, 155, 155-164. WINCHESTER, J. A. & PACE TMR NETWORK 2002. Palaeozoic Amalgamation of Central Europe: new results from recent geological and geophysical investigations. Tectonophysics, 360, 5-22. YANEV, S., GONCUOGLU,M. C., GEDIK, I., E T AL. 2006. Stratigraphy, correlations and palaeogeography of Palaeozoic terranes in Bulgaria and NW Turkey: a review of recent data. In: ROBERTSON, A. H. F., MOUNTRAKIS, D. & BRUN, J.-P. (eds) Tectonic Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 260, 51-67. YIdlTBA~, E., ELMAS, A. & YILMAZ, Y. 1999. Pre-Cenozoic tectono-stratigraphic components of the Western Pontides and their geological evolution. Geological Journal, 34, 55-74. YILMAZ, Y., TOYSI)Z,O., YIGITBA$,E., GEN(~,S. C. & ~ENGOR,A. M. C. 1997. Geology and tectonic evolution of the Pontides. In: ROBINSON, A. G. (ed.). Regional and Petroleum Geology of the Black Sea and Surrounding Region. American Association of Petroleum Geologists, Memoirs, 68, 183-226.

The Variscan orogen in Central Europe: construction and collapse WOLFGANG FRANKE Geologisch-Paliiontologisches Institut, Johann Wolfgang Goethe-Universitiit, Senckenberg-Anlage 32, Frankfurt, Germany (e-mail: [email protected])

Abstract: On the basis of a brief survey of the subdivision and evolution of the Variscides, this paper addresses controversial issues relating to the plate kinematic assembly and the 'collapse' of the orogen. A widespread phase of Devonian extension and basaltic magmatism is at variance with overall convergence. This episode either reflects subduction of the Rheic mid-ocean ridge, or else relates to a set of mantle plumes that also produced the Dniepr-Donets aulacogen. Another controversy regards the position of Gondwana in Devonian and Early Carboniferous time. Contrary to recent proposals of a wide Palaeotethys ocean, biogeographical and palaeomagnetic data suggest, until the Late Carboniferous, a Pangaea B model with Gondwanajuxtaposed against Southern Europe. Contrary to the concept of Late Carboniferous-Permian 'collapse' of a central Variscan high plateau, major crustal thickening occurred only in relatively narrow belts, and parts of the central Variscides were close to sea level from the Late Devonian onwards. Collision occurred in a hightemperature regime from c. 350-340 Ma onwards. Heating by several independent mechanisms effected the reduction of orogenic roots by buoyant rise and lateral spreading of thermally softened crust. However, major flysch wedges reflect the importance of erosion and uplift. Late Carboniferous-Permian magmatism and extension associated with strike-slip zones affected a largely equilibrated crust. These events probably relate to the westward displacement of Gondwana and the opening of the Palaeotethys embayment (Pangaea B to Pangaea A).

It is generally agreed that the Variscan crust is a collage of Gondwana-derived microplates (Avalonia, Armorican Terrane Assemblage; ATA), which were sequentially accreted to Baltica and eventually caught up in the collision of Gondwana and the 'Old Red Sandstone Continent' (Laurussia plus Avalonia). Late Devonian and Carboniferous subduction and collision created a large and heterogeneous orogen, with two zones of subduction on the northern flank and one on the southern flank of the belt (Figs 1 and 2). Shortening of continental crust involved in the Variscan collisions (Fig. 3) amounts to at least 800 km. Much higher values are probable. As discussed by Franke et al. (1995) and Franke (2000), the Variscan sutures have been overprinted by important dextral strike-slip movements that are difficult to constrain. The evolution of major parts of, or the entire Variscides has repeatedly been summarized (see, e.g. Martin & Eder 1983; Matte 1986, 1991; Matte et al. 1990; Franke et al. 1995; Franke 2000). Pharaoh et al. (2006) has provided a summary that includes the broader plate-tectonic context. However, the plate kinematic evolution and its geodynamic background as well as build-up and destruction of the orogen still present many unsolved questions. The present paper attempts a brief review of the main facts and highlights major open problems encountered in the German segment of the orogen, with references to some neighbouring areas. Main issues concern Devonian extension and its geodynamic causes, the existence of a Palaeotethys ocean, and the processes that destroyed the orogen. For a more detailed assessment and a survey of earlier literature, the reader is referred to Franke (2000) and, for eastern parts of the Bohemian Massif, Franke & Zelainiewicz (2000, 2002). Correlation of exogenic and endogenic events is based upon the time scales of Gradstein et al. (2004) and, for the Carboniferous, Menning et al. (2000). Plate kinematic scenario Major oceanic basins and constraints on their closure (Figs 2 and 3) Rheic ocean. Avalonia had rifted off from Gondwana during the Ordovician, opening the Rheic ocean in its wake (e.g. Tait et al. 2000). Avalonia made contact with Baltica in Late Ordovician time, thus producing a narrow tectonic belt known only from drillholes (see the review by Pharaoh et al. 2006). In Silurian-Early

Devonian times, the Rheic ocean was closed by intra-oceanic subduction, giving rise to an island arc now preserved in the Mid-German Crystalline High and at the southern margin of the Rhenish Massif (Rheno-Hercynian belt). Felsic members of the arc have been dated to the latest Ordovician to Early Devonian (444 _+ 22 to 398 _+ 3 Ma; U/Pb, Pb/Pb zircon and a few Rb/Sr whole-rock ages, see Franke 2000). By the earliest Devonian, the ATA must have been juxtaposed against Avalonia, as both microplates share the same Lochkovian non-marine fish (around 415 Ma, Young 1990; see biogeographical summary by McKerrow et al. 2000). Silurian and Devonian sedimentary sequences of the Rheno-Hercynian belt do not show any evidence of deformation and synorogenic clastic sedimentation during the relevant time span. Instead, important Gedinnian to Siegenian subsidence and sedimentation suggest crustal extension, probably effected by subduction toward the north and resulting back-arc spreading. Rheno-Hercynian narrow ocean. Shortly after, in Emsian time, a new spreading episode started to open the Rheno-Hercynian (Lizard-Giel3en-Harz) basin. The age of the oceanic crust is constrained by the oldest pelagic sediments overlying the pillow lavas (Emsian and Eifelian near Giegen, Birkelbach et al. 1988) and by a U - P b zircon age of 397 +__2 Ma from the Lizard allochthon in SW England (Clark et al. 1998), which again falls into the Emsian or early Eifelian, thus matching the biostratigraphic evidence in Germany. During Variscan collision, nappes in SW England and Germany transported oceanic fragments over the Avalonian foreland. At their base, these thrust sheets contain fragmented sedimentary sequences of Ordovician to Early Devonian age with Armorican faunas (see discussion by Franke & Engel 1982; Franke & Oncken 1995; Plusquellec & Jahnke 1999). Hence, opening of the Rheno-Hercynian ocean must have sprit off a fragment from the ATA adjacent to the south, which was left stranded on the N W shore of the nascent Rheno-Hercynian basin. Because of this Rheno-Hercynian reworking, the Rheic suture does not correspond exactly to the present-day fault zone at the southern margin of the Rhenish Massif, but is contained within the Northern Phyllite Zone, a narrow belt of Variscan pressuredominated metamorphic rocks at the southern margin of the Rhenish Massif and Harz Mts. (Anderle et al. 1995; Franke 2000). Rheno-Hercynian extension is reflected in three volcanic episodes. Early Devonian calc-alkaline rhyolites probably relate to incipient rifting (Jones & Floyd 2000). Younger, intraplate basaltic volcanism occurs in two peaks at about the

From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 333-343. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Structural map of the European Variscides (Permian plate configuration), with foreland basins marking the major sutures (yellow, Rheno-Hercynian-Moravo-Silesian; orange, Saxo-Thuringian; blue, retro-arc basin of the Moldanubian zone; brown, Mediterranean-Alpine, resulting from the Massif Central-Moldanubian collision). After Franke (2000).

Fig. 2. Plate kinematic model for the assembly of minor and major plates in the German segment of the Variscides. North is to the left. After Franke (2000).

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Fig. 3. Structural map and diagrammatictectonic cross-section of the German Variscides. (Note the distinctionbetween terranes (colours) and tectonic zones (black & white lettering).) Section is combined from a northwestern part (RheinischesSchiefergebirge-Spessart)and a southeastern part (SW part of the Bohemian Massif). KTB, site of deep continentaldrilling (KontinentalesTiefbohr-Programm). After Franke et al. (2004).

Givetian-Frasnian boundary and in the Tournaisian to early Vis~an (see summaries by Floyd 1995; Nesbor 2004) These volcanic episodes are also recorded in lateral equivalents of the Rheno-Hercynian in Germany, such as the Moravo-Silesian belt (Dvo}fik 1995) and SW England (see Holder & Leveridge 1986, and references therein). The Carboniferous volcanic episode is also represented in the Pyrite Belt of southern Portugal, with a predominance of felsic lavas (Oliveira & Quesada 1998; Boulter et al. 2001). Because the intra-plate basalts post-date the early Devonian onset of oceanic spreading, and the passive, northern margin of the Rheno-Hercynian ocean is clearly non-volcanic (Franke 2000), they require a separate geodynamic cause. The Rheno-Hercynian ocean is not detectable in the palaeomagnetic and biogeographical records. This is understandable, as flysch greywackes deposited on the oceanic crust suggest that subduction was already active in mid-Frasnian time (c. 380 Ma), which leaves only c. 30 Ma (Emsian to Givetian) for the drift stage. The Frasnian to Namurian flysch sediments are derived from the active, southern margin of the Rheno-Hercynian basin (Mid-German Crystalline High, MGCH; see Kopp & Bankwitz (2003) for the latest review of regional geology). The Crystalline High evolved from a north Armorican microplate (Franconia), which is documented only in Neoproterozoic detrital micas from Late Devonian greywacke turbidites (Huckriede et al. 2004). Magmatic activity in the arc of the MGCR is detectable from c. 360 Ma onwards. Retrogressed eclogites in the eastern Odenwald have been dated at 357 _+ 7 and 353 + 11 Ma ( L u - H f garnet-whole

rock, Scherer et al. 2002). Collision is documented by the crossover of greywacke turbidites onto the Avalonian foreland from the Devonian-Carboniferous boundary onwards (see Franke 2000, p. 50). During the Early Carboniferous to Namurian B, the front of synorogenic clastic sedimentation migrated across the foreland (Kulick 1960; Engel & Franke 1983; Franke & Engel 1986). From the Namurian C to the Westphalian C, sedimentation continued in a paralic molasse basin with coal seams, which can be traced from the Ruhr district of the northern Rhenish Massif westwards through northern Belgium as far as south Wales, and eastwards in Silesia, on the SE flank of the Bohemian Arc. S a x o - T h u r i n g i a n n a r r o w ocean. The Saxo-Thuringian basin originated from Cambro-Ordovician tiffing, which separated Bohemia from the north Armorican microplates of Saxo-Thuringia and Franconia. The Vesser Rift (Fig. 2; Kemnitz et al. 2002) between the latter terranes probably failed and did not evolve into a separate orogenic belt. Prolonged Saxo-Thuringian extension is documented in episodes of basaltic intra-plate volcanism in Silurian, Early Devonian and early Frasnian time (see summary by Falk et al. 1995). Southward subduction of Saxo-Thuringian crust commenced no later than c. 400 Ma (a summary of metamorphic events has been given by Franke et al. 1995; Scherer et al. 2002) and produced eclogites now preserved at the deformed northwestern margin of the Bohemia terrane (Tepl~-Barrandian unit) and in allochthons emplaced on the Saxo-Thuringian foreland (Franke 1984a,b).

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Collisional closure of the basin is recorded in widespread medium-pressure metamorphism around 380 Ma; in rocks on both sides of the suture and is also constrained by the onlap of early Famennian flysch greywackes on the foreland (c. 375 Ma; Sch~ifer et al. 1997). Massif Central-Moldanubian (MCM) narrow ocean. Oceanic separ-

ation between the ATA and mainland Gondwana is inferred from allochthons in the Moldanubian belt of the Bohemian Massif, which contain rocks that have undergone metamorphism at pressures up to 4 GPa derived from Palaeozoic ultramafic-mafic as well as continental protoliths (e.g. Becker & Altherr 1992; O'Brien 2000; VrLqa & Fryda 2003). Tectonic transport was directed generally southwards. As in the Rheno-Hercynian and Saxo-Thuringian cases, this southern ocean is not documented in the palaeomagnetic and biogeographical records (McKerrow et al. 2000; Robardet 2003). However, the presence of mantle rocks in the (ultra-)high-pressure metamorphic assemblage of the Gfrhl Moldanubian (Medaris et al. 1995) requires a zone of major crustal extension or narrow ocean, which guided continental subduction. Several lines of evidence suggest that closure of the MCM ocean occurred in latest Mid-Devonian to Late Devonian time. Late Givetian (c. 380 Ma) clastic sediments in the Barrandian syncline near Prague (Chlup~i6 1993) may be taken as a first sedimentary signal of orogeny preserved in the upper plate. Metagranites at the southern margin of the Tepl~i-Barrandian block were intruded at c. 370 Ma (Ko~ler et al. 1993; Ko~ler & Farrow 1994). In an eastern part of the Tepl~i-Barrandian unit, now concealed under Cretaceous cover (Fig. 3), limestones of late Famennian age unconformably overlie folded and cleaved Palaeozoic rocks, which indicates deformation prior to c. 360 Ma (Chlup~i~ 1994). This post-tectonic sedimentary cover extends into the Bardo basin of the West Sudetes, where limestones of Late Devonian age unconformably overlie a greenschist-grade basement (Ktodzko unit), whose protoliths include Neoproterozoic and early Ordovician felsic magmatic rocks (Mazur et al. 2003) as well as Givetian carbonates (Hladil et al. 1999); see also discussion by Franke & Zelainiewicz (2000, 2002). The suture zone at the southern margin of the Tepl~i-Barrandian block is sealed by the largely undeformed Central Bohemian batholith, which was intruded between 354 and 337 Ma (Drrr et al. 1997; Holub et al. 1997; Janousek & Gerdes 2003). In the French Massif Central, a western equivalent of the Moldanubian Zone, collision likewise occurred in Mid- to Late Devonian time (before c. 380 Ma; e.g. Lardeaux et al. 2001; Cartier & Faure 2004). High-pressure granulites from the Moldanubian allochthon further south (Gfrhl unit) have consistently yielded U-Pb zircon ages around 340 Ma (e.g. Krrner et al. 2000). Overthrusting of the Gfrhl unit over the less allochthonous Drosendorf unit led to widespread metamorphism around 335 Ma (see compilation by Franke 2000). In the SE part of the Moldanubian unit, post-metamorphic durbachites with U-Pb zircon ages between 338 and 335 Ma reflect HP melting in mantle or slab rocks (Kotkowi et al. 2003). Tectonometamorphic and magmatic events in the Moldanubian unit are difficult to interpret, as there is evidence for two subduction-collision events: the earlier one (>340 Ma) probably records collision between Bohemia and Gondwana (or some other Gondwana-derived fragment), with the suture extending westwards into the Massif Central (Figs 1 and 2). Later, the Moldanubian, Saxo-Thuringian and Rheno-Hercynian belts were dissected by NW-trending dextral shear zones (the Elbe and Intrasudetic fault zones) and rotated clockwise to form the Bohemian Arc. Shortly after, the rotated tectonic belts were truncated by dextral transpression along the Moldanubian Thrust, and juxtaposed against the Moravo-Silesian block (Figs 1 and 3; Franke & Zelalniewicz 2000, 2002; Gayer & Schulmann 2000). This latter process is documented by Vis~an flysch sediments deposited on the Moravo-Silesian foreland from the early Vis~an

(c. 340 Ma) to the early Namurian, which was followed by Late Carboniferous fluvio-lacustrine molasse with coal seams. The latest Visran Moravice Formation of the Moravo-Silesian flysch contains granulite pebbles derived from the Moldanubian allochthon and transported across the Moldanubian Thrust (Hartley & Otava 2001). This indicates that, by c. 325 Ma, the transpressional event must have been completed. The Moldanubian belt is conventionally correlated with the southern Black Forest and Vosges in SW Germany and, beyond, the French Massif Central (e.g. Matte 1986, 1991). At the southern margin of the Massif Central, in the Mouthoumet Massif and in the Pyrenees, synorogenic clastic sediments record the southwestward advance (in present-day coordinates) of the orogenic front in late Vis~an to Namurian time (Engel 1984; Franke & Engel 1986). The foreland is taken to represent Gondwana, as thick sequences of Cambro-Ordovician shelf sediments cannot be derived from a microplate, but require a large catchment area. The Variscan basement fragments dispersed over the Alps and the Mediterranean realm cannot be treated in this paper. However, it is important to note that the Carboniferous flysch of southern France is generally correlated with the southwarddriving flysch wedge exposed in the Southern Alps (see Fig. 1 and, e.g. Franke & Engel 1986; for the Carnic Alps and Karawanken Mts, see L~iufer et al. 2001). Therefore, it can be expected that evolution of the Variscan basement units now incorporated in the Alps was similar to that of France. The south Alpine Hochwipfel Flysch is unconformably overlain by shallow-marine deposits of Late Carboniferous to Permian age (Krainer 1993; Sch6nlaub & Histon 2000), which probably indicate incipient rifting at the tip of the westward-propagating Palaeotethys. O p e n questions Devonian plate divergence and basaltic magmatism. It is difficult to

understand why Devonian sea-floor spreading should occur within an assembly of plates that, during this time interval, underwent large-scale convergence. Back-arc extension (as proposed by Ziegler; e.g. Ziegler & D~zes 2006) can account for only the late Silurian-earliest Devonian part of Rheno-Hercynian extension in areas to the north of the intra-oceanic arc (Fig. 2). However, formation of Rheno-Hercynian ocean crust occurred within the northern part of the ATA and would, therefore, require southward subduction under the northern margin of the ATA (for which there is no evidence), or else subduction of the Saxo-Thuringian narrow ocean towards the north (which, in fact, was towards the south under Bohemia). Instead, it might be speculated that the narrow Rheno-Hercynian ocean was formed, when the northward moving ATA overrode the mid-ocean ridge of the Rheic ocean, much like the Bay of California is being opened, today, because the North American plate overrides the East Pacific Ridge. A back-arc model is feasible only for the early Frasnian intra-plate volcanism in the Saxo-Thuringian basin adjacent to the south, which might have been caused by southward subduction of Rheno-Hercynian ocean crust under the nascent Mid-German Crystalline High. A separate explanation is required for the Givetian to Frasnian and Early Carboniferous intra-plate basalts in the RhenoHercynian autochthon of the Rhenish Massif and Harz Mts (and equivalents from Portugal to Moravia). The geometry of the subduction zones active during this time interval (Fig. 2) precludes back-arc spreading. It is interesting to note that the majority of Devonian alkaline magmatic rocks in the Kola Peninsula were emplaced between 382 and 362 Ma (Sindern et al. 2003). Devonian to Early Carboniferous extension and magmatism are also important within the East European Craton (see Stephenson et al. 2006). It may be speculated that all these magmatic provinces represent a large-scale cluster of mantle plumes, whose activities were independent of the convection systems driving Variscan plate tectonics. Late Devonian-Early Carboniferous extension and magmatism in the Brevenne unit of the French

VARISCAN OROGEN IN CENTRALEUROPE Massif Central and Early Frasnian basaltic magmatism in the Saxo-Thuringian belt (see above) possibly record back-arc extension, but could equally well be part of the plume scenario. Position o f Gondwana and evolution of the Palaeotethys ocean. The

position of Gondwana during the Devonian and Early Carboniferous is still controversial. Most palaeomagnetic scenarios (e.g. Tait et al. 2000; Cocks & Torsvik 2002) propose an oceanic separation between Gondwana and the ATA, which corresponds to the Massif Central-Moldanubian ocean. This ocean is inferred to have opened from the Early Devonian and to have widened during the Carboniferous. A detailed assessment of the biogeographical and palaeomagnetic findings involved would go beyond the scope of this paper. It should be pointed out, however, that recent biogeographical reviews by McKerrow et al. (2000) and Robardet (2003) do not reveal indications of the Palaeotethys. Differences in the floras and faunas in post-Early Devonian times are attributed, by those workers, to climatic belts. In addition, the U - P b signatures of zircons from Ordovician to Devonian clastic sediments in Iberia suggest derivation from the West African Craton and the surrounding Pan-African belts (Martfnez-Catal~in et al. 2004). Lastly, the tectonic and sedimentary records of Variscan collision in southern France, the Southern /kips, and the Moldanubian part of the Bohemian Massif consistenfly document deposition of flysch and emplacement of thrust sheets on continental forelands. In the best-preserved section (southern France), the pre-flysch sedimentary and faunal records clearly suggest that this foreland was part of Gondwana (Robardet 2003). As noted above, the oldest marine ingression attributable to the opening of Palaeotethys occurs in the Late Carboniferous of the Carnic Alps. This is consistent with the findings of Muttoni et al. (2003), who have documented that the change from Pangaea 'B' (with Gondwana juxtaposed to Europe) to Pangaea 'A' (with the Palaeotethys to the south of Europe) occurred not before the Permian, and was accommodated by large-scale dextral shear zones (as proposed already by Arthaud & Matte 1977; Matte 1986). These findings rule out the concept of an oceanic subduction zone dipping to the NW under the ATA in Carboniferous time (as shown, e.g. by Cocks & Torsvik 2002), for which there is no evidence in any part of the southern Variscides. Likewise, the southern Variscides do not show any indications of Late Devonian-Early Carboniferous extension and magmatism. To avoid these difficulties, Tait et al. (2000) have proposed that the continental foreland found in the southern Variscides (southern France, Southern Alps) does not represent Gondwana mainland, but another Gondwana-derived microplate. This explanation, however, just shifts the problem to unknown areas further south, so that evidence of Late Devonian-Early Carboniferous rifting and Carboniferous northward subduction remains elusive.

Destruction of the orogen: where, when and why? Destruction of an orogen is brought about by reduction of the thickened crust either by plate boundary forces leading to lithospheric extension or else by buoyancy forces. In the latter case, crustal roots may be reduced by erosional or tectonic removal of orogenic topography. Alternatively, topography may be reduced by gravitational spreading of hot, low-viscosity lower crust. In reality, these processes will often work together. Their evolution in the Central European Variscides is assessed in the following paragraphs. P a l a e o - t o p o g r a p h y , vertical crustal m o v e m e n t s a n d thermal regime Late Devonian and Tournisian events ( 3 6 0 - 3 4 0 Ma). During the Late

Devonian, subduction of oceanic and continental materials

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occurred both on the northwestern margin (Saxo-Thuringian belt) and southeastern margin (Moldanubian belt) of Bohemia. Major erosion and some uplift is documented at the active, northwestern margin of Bohemia, where 400 Ma high-pressure and 380 Ma medium-pressure metamorphic rocks were already being eroded in Fammenian time and deposited in the marine Saxo-Thuringian foreland basin (Sch~ifer et al. 1997). Clues to the palaeo-topography have survived only in eastern parts of the Tepl~i-Barrandian block, where the transgression of Late Devonian to Toumaisian marine sediments on deformed Palaeozoic rocks indicates zero elevation. Also, the preservation of very low-grade sedimentary rocks and moderate tectonic shortening argue against major crustal thickening and uplift. These findings clearly rule out the concept of a Tepl~i-Barrandian high plateau proposed by Zulauf (1997, 2002). The absence of siliciclastic debris from these deposits argues against major elevation also in the neighbouring regions in the time between c. 360 and 340 Ma. Coarse-grained clastic sediments of Late Devonian to Early Carboniferous age do occur in the SwiCbodzice pull-apart basin of the West Sudetes (PorCbski 1990), but late Frasnian and Famennian intercalations of marine mudstones and limestones indicate that the sediments of the intra-orogenic basin were deposited close to sea level. These findings demonstrate that, < 20 Ma after the closure of the Saxo-Thuringian and Massif Central-Moldanubian oceans, parts of the central Variscides were inundated by the sea, and orogenic topography existed only in marginal parts of Bohemia. This clearly precludes an areally extensive high plateau in the area. Similar considerations apply to the Morvan in the northern part of the French Massif Central, where marine shales have been dated as Tournaisian (Weyer 1976), and to the Beaujolais and Brevenne units further south, which contain Devonian to Early Carboniferous marine sedimentary and volcanic rocks of very low to low metamorphic grade (Leloix et al. 1999; Lardeaux et al. 2001). A relatively cool thermal regime is indicated by the observation that most of the metamorphic rocks of the Bohemian Massif dated at c. 400-380 Ma were formed in, or else exhumed through, the amphibolite facies. H P - H T metamorphism is restricted to one locality in the Saxo-Thuringian region of NW Bavaria (Kleinschrodt & Gayk 1999) and to the Grry Sowie of the West Sudetes (Zelalniewicz 1990; O'Brien et al. 1997; Timmermann et al. 2000). The 340 Ma event. Areas surrounding the Tepl~i-Barrandian unit (i.e. the Bohemia microplate and rocks accreted to it) are characterized by high-temperature metamorphic rocks and granitoids dated at c. 340 Ma. The Saxonian Granulites of the SaxoThuringian belt were equilibrated at c. 22 kbars and 1050 ~ (Rrtzler et al. 2004), the highest metamorphic temperatures hitherto documented in the Variscan belt. Ultrahigh pressures are documented in metamorphic diamonds in continental rocks of the Erzgebirge dated at c. 340 Ma (Massonne 2001, 2003; Massonne et al. 2001). Shortly after their formation, granulites and eclogites were emplaced in the continental crust of the foreland (e.g. Reinhardt & Kleemann 1994). Emplacement was probably driven by hydraulic forces and occurred under the floor of the Saxo-Thuringian foreland basin (DEKORP & Orogenic Processes Working Groups 1999; Franke & Stein 2000; Henk 2000), a process requiring low viscosity (Zulauf et al. 2002a). In the southeastern, internal part of the Saxo-Thuringian belt, the exhumed high-pressure rocks were subsequently involved in the accretion of the foreland (Erzgebirge: Franke & Stein 2000; Konopfisek et al. 2001). Eclogites in the Orlica-SnieZnik unit of the West Sudetes (Fig. 3) probably represent an eastern continuation of the Erzgebirge (see discussion by Franke & Zelainiewicz 2000). On the southern (Moldanubian) flank, high-pressure granulites are widespread in the allochthonous Gfrhl unit. These granulites have consistently yielded U - P b zircon ages around 340 Ma

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W. FRANKE

(compilation by Franke 2000; Kr6ner et al. 2000). However, deformed granites within allochthonous Bohemian granulite gneisses show U-Pb zircon minimum ages of 354 Ma (Svojtka et al. 2002). Zircon growth possibly occurred during decompression melting in felsic granulites, so that the time of deepest burial might be older (Finger et al. 1996; Roberts & Finger 1997). This is compatible with the findings of Kr6ner et al. (2000), who have demonstrated that zircons from granulites with U-Pb ages of c. 340 Ma were repeatedly formed at various stages of decompression. Zulauf (1997, 2002b) and Zulauf et al. (2002) have shown that the Teplfi-Barrandian is surrounded, today, by rocks whose metamorphic grade implies subsidence of the Tepl~i-Barrandian (or uplift of the surrounding rocks) by c. 10-15 km. Overthrusting of the Gf6hl unit over the less allochthonous Drosendorf unit led to widespread metamorphism around 335 Ma (see compilation by Franke 2000). In the southeastern part of the Moldanubian unit, post-metamorphic durbachites with U-Pb zircon ages between 338 and 335 Ma reflect HP melting in mantle or slab rocks (Kotkovfi et al. 2003), probably belonging to the underthrust Moravo-Silesian belt. 3 4 0 - 3 2 0 Ma. Orogenic uplift and erosion from 340 to 320 Ma is

documented by flysch sediments in the foreland basins of the Rheno-Hercynian, Saxo-Thuringian and Moravo-Silesian belts. Clastic sediments were derived from the Mid-German Crystalline High, the northwestern margin of Bohmia, and from Moldanubian sources (Fig. 1). The Variscan topography, during this time span, was characterized by an alternation of foreland basins and flyschproducing collisional uplifts. Whereas sedimentation in the Rheno-Hercynian and Moravo-Silesian basins continued into Late Carboniferous coal-bearing molasse deposits, sedimentation in the Saxo-Thuringian was terminated by folding around 330 Ma. A short time later, there is again evidence of low elevations. A eustatic sea-level rise occurred in the Goniatites crenistria zone of the traditional European Culm zonation (Go oL of the classical goniatite stratigraphy, Late Asbian; see Herbig 1998), which corresponds to an isotopic age of c. 327 Ma (Menning et al. 2000). Marine sediments of this age were deposited on folded Cambrian rocks in the the northwestern part of the Saxo-Thuringian basin (Doberlug-Kirchhain, Vis~an 3b, Weyer 1965) and in the strike-slip, related Intra-Sudetic Basin (Fig. 3: SW of the Gdry Sowie; Zakowa 1963; Herbig 1998; Tumau et al. 2002). In both these basins, the marine beds represent an early phase of sedimentation, so that the orogenic topography must have been low from the beginning of basin evolution. During the time interval between 340 and 320 Ma, there is no evidence of pressure-dominated metamorphism, and P - T conditions suggest collisional stacking and heating. The Moldanubian zone contains large volumes of granites intruded between c. 335 and 325 Ma (see references in Franke et al. 2000; discussion by Gerdes et al. 2002, 2003). A narrow, NW-trending belt along the SW margin of the Bohemian Massif cuts across the tectonic zonation. It is characterized, between c. 325 and 320 Ma, by an especially high-temperature regime. Low-pressure-high-temperature metamorphism with anatexis (Tanner & Behrmann 1995; Tanner 1999) dated to a narrow interval of 327-320 Ma (see reviews by Franke 2000; Kalt et al. 2000) was immediately followed by the intrusion of post-tectonic granitoids (e.g. Siebel et al. 2003; Chen& Siebel 2004). This transverse zone extends northwestwards into the very low grade rocks of the Saxo-Thuringian foreland (Kosakowski et al. 1999). Younger granites (315-290 Ma) occur in this SW Bohemian Transverse Zone, but also in the Fichtelgebirge-Erzgebirge antiform of the southeastern Saxo-Thuringian zone and in the Moldanubian zone. Evolution in SW Germany, France and Iberia. A detailed assessment

of metamorphism and granitoid magmatism in other parts of the

Variscan intemides would go beyond the scope of this paper. It should be noted, however, that the time interval of granitoid intrusion observed in the Bohemian Massif is the same as that in Iberia (352-297 Ma, maximum at 335-305 Ma; Montero et al. 2004). Intense metamorphism and granite emplacement at about 340 Ma have also been recorded from the Black Forest, the Vosges and the Massif Central (Costa 1992; Boutin et al. 1995). In the intramontane basins of the French Massif Central, there are no marine sediments, which would suggest moderate elevations also in neighbouring regions. Becq-Giraudon & Van den Driessche (1994) and Becq-Giraudon et al. (1996) even claimed to have found petrographic, sedimentological and palaeobotanical evidence of sediments deposited >5000 m above sea level. However, their sedimentological findings are equivocal. Also, the floras of the Permo-Carboniferous basins in the Massif Central do not reveal a cold environment, but warm and humid conditions (H. Kerp, Mtinster Univ., pers. comm.).

Geodynamic model

Whereas the closure of the Rheno-Hercynian basin conforms to a classical model of subduction-collision, the tectonothermal evolution of the internal Variscides (Bohemian Massif, Saxo-Thuringian and Moldanubian belts) is problematic. In Mid- to Late D e v o n i a n time, the Saxo-Thuringian and Moldanubian narrow oceans were closed by subduction from the SE and the NW under the Bohemian microplate (now largely represented by the Teplfi-Barrandian tectonic unit). Eclogites in the Saxo-Thuringian belt were formed and obducted in a mediumtemperature regime. Obduction may be explained by buoyant rise of subducted continental material according to the model of Chemenda et al. (1995; see Franke & Stein 2000). Late Devonian subduction on the Moldanubian flank cannot be excluded, but isotopic evidence has been obliterated by later high-temperature metamorphism. Crustal thickening, resulting in uplift and erosion, is documented only for the northwestern margin of Bohemia (the active margin of the Saxo-Thuringian basin). The central part of Bohemia (Teplfi-Barrandian) was only moderately thickened and had already been inundated by the sea by Late Devonian time. B e t w e e n c. 350 a n d 340 Ma, granitoids were emplaced along the southeastern and western flanks of the Teplfi-Barrandian block and remained largely unaffected by later ductile deformation. During the same time, or shortly thereafter (c. 340 Ma), large volumes of low-viscosity, high-pressure granulites and some eclogites rose on both flanks of the Bohemian 'median massif' (Fig. 4). The buoyancy of felsic material will certainly have contributed to uplift (Reinhardt & Kleemann 1994; Gerya et al. 2002a,b; see also Lardeaux et al. (2001) for a similar situation in the French Massif Central). However, expulsion by hydraulic or compressional forces probably also played a major role in their emplacement (Franke & Stein 2000; Henk 2000; Stfpskfi et al. 2004) Ultrahigh pressures in metamorphosed continental rocks on both flanks of the Teplfi-Barrandian suggest that at least part of the heat was derived from contact with the asthenosphere. In addition, crustal shortening on the NW and SE flanks of Bohemia amounts to at least 500 km. This implies subduction of the same amount of lithospheric mantle, which cannot have been accommodated under the narrow 'median massif' of the Tepl~-Barrandian block (Franke 2000). Loss of parts of the subducted lithospheric mantle slabs and subsequent rise of hot asthenosphere probably added to the high-temperature regime. Because the density contrast between lithospheric and asthenospheric mantle is rather low (c. 0.05 g cm-3; e.g. Grow & Bowin 1975), neither subduction nor break-off of the lithospheric mantle slabs can be expected to have important isostatic consequences. This is in accord with Late Devonian to Tournaisian shallow-marine environments in the West Sudetes.

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Fig. 4. Tectonic model (not to scale) of continental subduction and subsequent exhumation at the northwestern and southeastern margins of Bohemia, around 340 Ma. Downward thinning of subducted lithospheric slabs is intended to indicate transition into oceanic lithosphere (already subducted).

B e t w e e n 340 and 325 Ma, accretion and thickening of Moldanubian crust propagated toward the SE, in a medium- to low-pressure metamorphic regime with widespread migmatization and intrusion of large volumes of granitoids (not depicted in Fig. 4). Both heating by radiogenic decay and transfer of mantle heat through a thinned mantle lithosphere may be responsible for the thermal environment and have caused extensional spreading of thickened crust. However, shedding of flysch sediments into the MoravoSilesian foreland basin indicates that crustal thinning was also effected by erosion and uplift. These considerations also apply to the active, southeastern margins of the Rheno-Hercynian and Saxo-Thuringian basins. In any case, the late Asbian (c. 328 Ma) marine sediments in intramontane Saxo-Thuringian basins preclude, also for this time interval, the existence of a central Variscan 'Tibetan' plateau. A r o u n d 325 Ma, low-pressure metamorphism and granitoid intrusion in the SW Bohemian transverse zone cut across the collisional zonation, from the Moldanubian belt in the SE to the Saxo-Thuringian foreland in the NW. This argues against causes acting 'along-strike', such as crustal thickening or delamination of mantle lithosphere. Franke et al. (1995), Franke (2000) and Kalt et al. (2000) have proposed advective heating by melts, probably triggered by processes in the asthenospheric mantle independent of the mechanics of the orogen. Younger granitoids ( 3 2 0 - 2 9 0 Ma) are widespread in Western and Central Europe. This magmatic pulse was associated with re-equilibration of the Moho (e.g. Ziegler et al. 2004). In particular, the youngest plutonic and volcanic rocks (_<300 Ma) and timeequivalent basin formation clearly extend beyond the limits of the Variscan orogen (see the contributions in the volume edited by Wilson et al. 2004).

incorporated numerous felsic magmatic rocks, which were at least partly derived from the melting of fertile sediments (NeoProterozoic arc magmatism, Cambro-Ordovician rift magmatism, and Devonian-Early Carboniferous arc magmatism). However, Henk et al. (2000) demonstrated that the Variscan high-temperature regime observed between 340 and 325 Ma also requires heat advection by the rise of asthenospheric mantle. This may have been brought about, in several ways, by reduction of the thickness of the insulating mantle lithosphere (by loss of the thermal boundary layer, delamination or slab break-off). It appears, however, that there was no break-through of primitive asthenospheric melts, as the composition of the granitoids is dominated by crustal protoliths, with only a minor component of lithospheric mantle (Henk et al. 2000). The large areal extent of Late Carboniferous-Permian magmatism, which is present also outside the Variscan orogen, argues against geodynamic explanations relating to preceding orogenic processes such as slab detachment. This contradicts proposals of, for example, Ziegler et al. (2004) and Ziegler & D~zes (2006). The important mantle component in the magmatic rocks is much better explained in the context of the dextral displacement of Gondwana with respect to Europe and the ensuing opening of the Palaeotethys ocean (Arthaud & Matte 1977), which may have occasioned asthenospheric upwelling. According to Henk (1997, 1999), the post-Variscan dextral mega-shear was also essential for subsidence in strike-slip related basins: Henk's numerical modelling suggests that the Variscan orogen 'did not collapse, but was torn apart'.

Heat sources. By comparison with many other orogenic belts, the Variscan orogen was 'hot'; that is, it is characterized by numerous granitoids and HT metamorphic rocks. Gerdes et al. (2000) have modelled the importance of radiogenic heating and concluded that this process is sufficient to produce the large volume of granitoids observed in the central Variscides. This is plausible, as Variscan tectonic stacking affected thick clastic sequences, which are rich in feldspar and mica and, therefore, thermally fertile (Devonian and Early Carboniferous synorogenic greywackes, Cambro-Ordovician and Devonian shelf sediments, Late Proterozoic synorogenic greywackes, for example (Gerdes et al. 2002; Bea et al. 2003). This situation differs from that in the Alps or Scandinavian Caledonides, where warm climates during the rift and drift stages favoured deposition, on the shelves, of carbonates rather than siliciclastic sediments. Variscan stacking also

(1) Tectonic, stratigraphic and biogeographical datasets consistently suggest that the Variscan plate collage was largely assembled during the Mid-Late Devonian, so that the tectonic regime from the earliest Carboniferous onwards was collisional. Apart from palaeomagnetic data, there is no indication, in Late Devonian to Early Carboniferous time, of a Palaeotethys ocean separating the Armorican Terrane Assemblage from Gondwana. (2) Contrary to the general trend of plate convergence, Devonian extension is documented in the Rheno-Hercynian, Saxo-Thuringian and Moldanubian belts. Although some of these extensional zones and episodes may be explained by back-arc spreading, opening of the Rheno-Hercynian narrow ocean requires a separate explanation (subduction of the midocean ridge of the Rheic ocean, or, alternatively, a set of mantle plumes active under parts of Eastern and Central Europe).

Conclusions

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W. FRANKE

(3) The stratigraphic and tectonometamorphic records clearly reveal that destruction of the Variscan orogen was a multistage process. An areally coherent, Tibetan-style plateau has never existed: some central parts of the Bohemian Massif were repeatedly inundated by the sea from the Late Devonian onwards. (4) While subduction-exhumation processes during the Devonian occurred in a medium-temperature regime, a widespread tectonometamorphic episode around 340 Ma created high- to ultrahigh p r e s s u r e - h i g h temperature metamorphic rocks derived from continental protoliths. Heat was probably derived from contact with the asthenosphere, probably during subduction and possibly also after partial loss of the mantle lithosphere. Uplift and exhumation of these rocks were caused by buoyant rise through the mantle and hydraulic or tectonic 'squeezing' upwards and outwards. (5) After 340 Ma, crustal thickening and magmatism occurred in three active margins feeding the Rheno-Hercynian, Saxo-Thuringian and Moravo-Silesian foreland basins. Reduction of these individual roots was at least partly caused by uplift and erosion, although 'collapse' of thermally softened crust cannot be ruled out. (6) From c. 325 Ma onwards, magmatic and metamorphic events no longer reveal clear spatial relations with the plate tectonic zonation of the Variscan belt. Around 300 Ma, magmatism and basin formation even occur outside the orogen. These processes affected an orogenic crust, which had been largely equilibrated already by preceding events. Both the formation of strike-slip-related Permo-Carboniferous basins and timeequivalent, mantle-derived magmatism may be explained by the westward displacement of Gondwana with respect to Europe, and subsequent opening of the Palaeotethys.

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YOUNG, G. C. 1990. Devonian vertebrate distribution patterns and cladistic analysis of palaeogeographic hypotheses. In: MCKERROW, W. S. & SCOTESE, C. R. (eds) Palaeozoic Palaeogeography and Biogeography. Geological Society, London, Memoirs, 12, 243-255. ZA~ZOWA,H. 1963. Stratigraphy and facial extents of the Lower Carboniferous in the Sudetes. Kwartalnik Geologiczny, 7, 73-94. ZELAZNIEWICZ, A. 1990. Deformation and metamorphism in the Gdry Sowie gneiss complex, Sudetes, SW Poland. Neues Jahrbuch fiir Geologic und Paliiontologie, Abhandlungen, 179, 129-157. ZmCLER, P. A. & DEZES, P. Crustal evolution of Western and Central Europe. In: GEE, D. G. & STEPHENSON, R. A. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 43 -56. ZIEGLER, P., SCHUMACHER, M. E., DgZES, P., VAN WEES, J.-D. & CLOETINGH, S. 2004. Post-Variscan evolution of the lithosphere in the Rhine Graben area: constraints from subsidence modelling. In: WILSON, M., NEUMANN, E.-R., DAVIES, G. R., TIMMERMAN, M. J., HEEREMANS, M. & LARSEN, B. T. (eds) Permo-Carboniferous Magmatism and Rifting in Europe. Geological Society, London, Special Publications, 223, 289 - 3167. ZULAUF, G. 1997. Von der Anchizone bis zur Eklogitfazies: Angekippte Krustenprofile als Folge der cadomischen und variscischen Orogenese im Teplfi Barrandium (B6hmische Masse). Geotektonische Forschungen, 89, 302 S. ZULAUF, G. 2002. Elevator-style tectonics and tilted crustal sections in the Bohemian Massif. Post-Conference Field Trip, TSK 9, March 23-24, 2002. In: Abstracts TSK 9 (9. Symposium TektonikStrukturgeologie-Kristallingeologie). Erlanger Geologische Abhandlungen, Sonderband 3, 101-121. ZVLAtJF, G., BUES, C., DORR, W. & VEJNAR, Z. 2002a. 10 km minimum throw along the West Bohemian shear zone: evidence for dramatic crustal thickening and high topography in the Bohemian Massif (European Variscides). International Journal of Earth Sciences (Geologische Rundschau), 91, 850-864. ZULAUF, G., D6RR, W., FIALA, J., KOTKOVA, J., MALUSKI, H. & VALVERDE-VAQUERO, P. 2002b. Evidence for high-temperature diffusional creep preserved by rapid cooling of lower crust (North Bohemian shear zone, Czech Republic). Terra Nova, 14, 343-354.

Transpressional collision tectonics and mantle plume dynamics: the Variscides of southwestern Iberia J. F. S I M A N C A S l, R. C A R B O N E L L 2, F. GONZ/~LEZ L O D E I R O l, A. PI~REZ E S T A I ) N 2, C. J U H L I N 3, P. A Y A R Z A 4, A. K A S H U B I N 3, A. A Z O R 1, D. M A R T I N E Z POYATOS 1, R. S,~EZ 5, G. R. A L M O D O V A R 5, E. P A S C U A L 5, I. F L E C H A 2 & D. MARTI 2

1Departamento de Geodindmica, Facultad de Ciencias, Universidad de Granada, Campus de Fuentenueva, 18071 Granada, Spain (e-mail: simancas @ ugr. es) 2Instituto de Ciencias de la Tierra 'Jaume Almera', Consejo Superior de Investigaciones Cientfficas, c/Lluis Sold i Sabarfs s/n, 08028 Barcelona, Spain 3Department of Earth Sciences, Uppsala University, Villavagen 16, SE-752 36 Uppsala, Sweden 4Departamento de Geologfa, Universidad de Salamanca, 37008 Salamanca, Spain 5Departamento de Geologfa, Universidad de Huelva, 21071 Huelva, Spain

Abstract: In southwestern Iberia, three continental domains (the South Portuguese Zone (SPZ), Ossa-Morena Zone (OMZ) and Central Iberian Zone (CIZ) collided in Devonian-Carboniferous time. The collision was transpressional, with left-lateral kinematics, and was interrupted by extensional tectonics during the earliest Carboniferous, when bimodal magmatism (with associated mineral deposits) and basin development were the dominant orogenic features. Transpression was renewed in Visran time, and persisted until the end of the Carboniferous. The IBERSEIS deep seismic reflection profile helps to define the 3D geometry of transpressional structures: out-of-section displacements concentrate in bands, which bound wedges of upper crust; this crustal wedging strongly modifies the geometry of the sutures between continental blocks. A mid-crustal strongly reflective thick band (the lberseis Reflective Body, IRB) is interpreted as a huge body of basic rocks. The IRB magma trapped in the middle crust was linked to the Early Carboniferous mantle-derived magmatism that crops out in the SPZ, OMZ and CIZ. Magmatism at the surface and trapped in the crust, high thermal gradients and basin development reflect a thermal anomaly in the underlying mantle, influencing both the thermal and the stress state of the orogen at that time. A mantle plume is inferred to have existed in the Early Carboniferous, the transpressional tectonic regime dominating again after its decay.

New knowledge of the orogenic evolution of southwestern Iberia has come from the completion of the SW Iberia EUROPROBE project (Ribeiro et al. 1996), which resulted in the recent acquisition of the IBERSEIS deep seismic reflection profile and its interpretation in the light of current geological research. We document here the IBERSEIS crustal seismic section and interpret its orogenic evolution. Two outstanding features of this collisional evolution are the following: (1) its general oblique left-lateral kinematics; (2) the existence of an extensional intra-orogenic stage associated with abundant magmatic activity, which has been detected both at the surface and trapped in the crust. The general left-lateral transpressional regime accords well with the location of SW Iberia in the broader plate-tectonic context of the Variscides. The extensional and magmatic stage indicates a lithospheric anomaly interpreted as a mantle plume underlying the orogen during the Early Carboniferous and strongly influencing the tectonic scenario at that time. The mantle plume decay, or its migration elsewhere, gave rise to a renewal of the continental transpressional collision.

Collisional Variscan terranes in SW Iberia The Iberian Massif forms a large part of the Variscan orogen (Fig. la). Besides the peculiarities of its orogenic evolution, the Iberian Massif deserves attention because it shows a complete section across the Variscan belt. The Iberian Massif has traditionally been divided into a number of zones (Fig. lb; Farias et al. 1987): the Cantabrian Zone, West Asturian-Leonese Zone, Central Iberian Zone, Galicia Tras-Os-Montes Zone, O s s a Morena Zone and South Portuguese Zone. Geological research indicates that some of the boundaries between zones are sutures of the orogen whereas others have much less tectonic significance.

The Cantabrian Zone to the north and the South Portuguese Zone to the south represent opposing foreland fold and thrust belts (Fig. lb and c). In the northwestern part of the Iberian Peninsula, the Galicia Tras-Os-Montes Zone consists of a pile of allochthonous tectonic units, some of them with high-pressure metamorphic and/or ophiolitic rocks (Arenas et al. 1986, 1997; Ribeiro et al. 1990). The suture defined by these allochthonous terranes is rooted offshore in the Atlantic margin, NW of Iberia (Fig. lc, section C1; Matte 1986; Martfnez Catal~in et al. 1997, 2002). However, the curved architecture of the orogen (Fig. l a) suggests that the suture must crop out on land in southwestern Iberia. In fact, two tectonic contacts can be interpreted as Variscan sutures in southwestern Iberia, corresponding to the northern and southern boundaries of the Ossa-Morena Zone. Nevertheless, the connection between the northwestern and the southwestern sections of Iberia (Fig. lb and c) is not simple; a transform fault has been suggested to exist between them, to explain the envisaged opposed subduction polarity (Simancas et al. 2002). In this paper, we are concerned with the geodynamic evolution of southwestern Iberia, which includes the South Portuguese (SPZ) and O s s a - M o r e n a (OMZ) zones in conjunction with the southern part of the Central Iberian Zone (CIZ). The boundary between the OMZ and the SPZ is interpreted as a suture, considering that it is marked by a continuous strip of amphibolites with oceanic geochemical signatures (the Beja-Acebuches ophiolite; Bard 1977; Andrade 1983; Munhfi et al. 1986; Crespo Blanc 1989; Fonseca & Ribeiro 1993; Quesada et al. 1994; Castro et al. 1996), and an accretionary prism with slices of oceanic metabasalts (the Pulo do Lobo Unit; Silva et al. 1990). Small klippen of oceanic rocks imbricated with high-pressure continental rocks, which crop out in the southwestern part of the OMZ (Fonseca et al. 1999), may also be rooted in the S P Z - O M Z boundary. The boundary between the OMZ and the CIZ is marked by a complex tectonic unit, the Badajoz-C6rdoba Shear zone or

From: GEE, D. G. & ST~PHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 345-354. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

345

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J.F. SIMANCASETAL.

Fig. 1. (a) The Iberian Massif in the context of the Variscan orogen. (b) General view of the Iberian Massif, including the main tectonic contacts. (e) Generalized cross-sections of the Variscan orogen in Iberia (see location in (b)). Cross-section CI is after P~rez Estafin et al. (1991).

Central Unit (Burg et al. 1981; Azor et al. 1994), which includes some retro-eclogites (Abalos et al. 1991; Azor 1994; L6pez Sfinchez-Vizcafno et al. 2003) and amphibolites, with oceanic geochemical signatures (G6mez Pugnaire et al. 2003). Large-scale upper crustal structures root in this boundary (Fig. lc, section C2; Simancas et al. 2001). The Iberian Massif results from the amalgamation in Devonian to Carboniferous times of three continental blocks: the SPZ, the OMZ, and the ensemble of the CIZ, West Asturian-Leonese and Cantabrian zones (Fig. l b). The last three of these must have belonged to the same continental block, most probably the Gondwana continental margin. The SPZ seems to be part of the Avalonian border on the opposite continent, so that its boundary with the OMZ would be the suture of the Rheic ocean, which has been proposed in palaeomagnetic reconstructions (Crowley et al. 2000; Tait et al. 2000; Matte 2001). The OMZ, placed between the two sutures, was an independent terrane, although it must have always been closely tied to the Gondwana margin,

given that palaeobiological data do not support the existence of a wide Palaeozoic ocean at this boundary (Nysa~ther et al. 2002; Robardet 2002).

Crustal image provided by the IBERSEIS deep seismic reflection profile A detailed description of the IBERSEIS seismic profile and its geological interpretation has been provided by Simancas et al. (2003) and Carbonell et al. (2004), and in this paper we focus on particular important features. The IBERSEIS crustal image shows a general detachment between strongly deformed upper crust and less deformed lower crust (Fig. 2). The lower crust that was subducted in the suture zones, as a consequence of the convergence and crustal detachment, is not observed in the seismic image. In fact, the present

TRANSPRESSION AND MANTLE PLUME IN IBERIA

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Moho is a post-orogenic structure (Permian?) with a fiat topography, found at 10.5 s, which has obliterated the previous orogenic root. Besides the mid-crustal detachment, two outstanding features of the crustal section need to be emphasized: (1) the existence of upper crustal wedges corresponding to left-lateral (out-of-section) displacements during the continental collision; (2) the presence of a mid-crustal reflective band, interpreted as a voluminous intrusion of basic magma (Fig. 2), namely, the Iberseis Reflective Body (IRB). The suture zones in southwestern Iberia have been imaged in the IBERSEIS section with markedly different features from those typical of orogenic sutures. Neither the OMZ-SPZ nor the CIZ-OMZ boundary depicts well-defined seismic fabrics imaging one continental block beneath the other. Instead, the seismic fabrics approximate to a face-to-face arrangement, being cut by steeply dipping reflectors that correspond to faults with variable strike-slip component (Fig. 2). At the OMZ-SPZ boundary, a set of steeply dipping reflectors delineate a fan-like structure, converging at mid-crustal level. The fan-like geometry of the faults complements the tapering and sigmoidal shape of the units as seen on geological maps, thus providing a 3D image of the upper crustal wedges characteristic of this orogenic region. At the CIZ-OMZ boundary, similar wedge features can be observed both in the crustal section (Fig. 2) and at the surface. Crustal wedging as a result of strike-slip tectonics strongly modifies the initial geometry of the sutures at the boundaries of the continental blocks. The resulting geometry is characteristic of transpressional suture zones, where out-of-section displacements caused by oblique collision are seen to be concentrated along shear zones or faults, leaving intervening domains where shortening is dominant. The extreme scarcity of arc volcanic rocks at the OMZ-SPZ boundary, related to the SilurianDevonian subduction of the Rheic ocean, could be partially explained by these out-of-section displacements, which may have transported fragments of the volcanic arc away from southwestern Iberia. The lower crust under the suture zones shows another type of wedging (Fig. 2), consisting of indentation structures geometrically similar to those described, for instance, in the lower crust of northern Iberia (P~rez Estatin et al. 1994), or in the western Pyrenees (Teixell 1998). Based on geometric arguments, the indentation wedging observed in the lower crust of southwestern Iberia is likely to have developed at an advanced stage of the collision, thus contributing to its blocking. Perhaps the most outstanding feature in the IBERSEIS seismic profile is a strongly reflective band, irregular in thickness, located in the middle crust of the OMZ: the Iberseis Reflective Body (Fig. 2; Simancas et al. 2003; Carbonell et al. 2004). The IRB shows relatively high velocities when compared with the background velocity field, and contains small domains of layered reflectivity. It has been interpreted as a layered maficultramafic igneous body with some assimilation of surrounding rocks. The surface geology favours this interpretation for the IRB and provides evidence for the timing of emplacement of the magma in the middle crust. The following data are relevant: (1) the relative abundance of Early Carboniferous mafic intrusions in the OMZ (Capdevila et al. 1973; Sfinchez Carretero et al. 1990; Dallmeyer et al. 1993, 1995; Pin et al. 1999; Montero et al. 2000), which we infer to be derived from the IRB magmatic body; (2) the Early Carboniferous high-temperaturelow-pressure, amphibolitic-granulitic metamorphic belt cropping out in the southernmost OMZ (Bard 1969; Crespo Blanc 1989; Castro et al. 1999), which we interpret as tectonically exhumed rocks, originally located adjacent to the IRB mid-crustal magma; (3) some Ni-Cu ore deposits found in the OMZ, which are believed to be related to mantle-derived Early Carboniferous gabbroic magmas, contaminated by pelitic crustal material after some period of residence in the crust (Casquet et al. 2001; Tornos et al. 2001).

The emplacement of this mantle-derived magma in the middle crust of the OMZ would have taken advantage of the orogenic detachment level, which played the role of a major mechanical discontinuity.

Chronology and kinematics of the Variscan collision in southwestern Iberia A schematic representation of major structures of various ages, mapped in southwestern Iberia, is shown in Figure 3. The timing of the tectonic events giving rise to these structures and the tectonic regime in which they were developed are discussed further below. The earliest stages of the Variscan collision are poorly known. In northwestern Iberia, stratigraphic and radiometric data indicate that the convergence started in Early Devonian times (Gonz~ilez Clavijo & Martfnez Catal~in 2002). In southwestern Iberia, a similar situation is envisaged, as discussed below. In the OMZ, above Lower Cambrian carbonates and Cambro-Ordovician terrigenous deposits and volcanic rocks, the Silurian is represented by pelagic sediments (black siliceous slates, radiolarian cherts), passing up, in the Lower Devonian sequence, into siltstones overlain by flysch deposits (Robardet & Guti~rrez Marco 1990; Guti6rrez Marco et al. 1998; Pi~arra 1998). This stratigraphic record is interpreted as providing evidence of an Early Cambrian platform and Mid-Cambrian rifting (Lift,in & Quesada 1990). The Silurian deposits, characterized by the lack of clastic influx, suggest the isolation of the OMZ crust as a result of the development of open sea domains at its boundaries. Thereafter, the renewed clastic influx in Early Devonian times points to the closing of the open sea domains, that is, the beginning of the Variscan collision. High-pressure metamorphism recorded in the Central Unit (CIZ-OMZ boundary) is related to underthrusting of an oceanic and/or thinned continental crust beneath the southern border of the CIZ, during an early compressional stage whose structural record has been completely obliterated by shearing associated with the later exhumation of the unit (Azor et al. 1994). The approximate radiometric age of 427 + 45 Ma (Sm/Nd on garnet), provided by Sch~ifer et al. (1991), can be taken to support a Silurian to Early Devonian age for the earliest Variscan convergence. The kinematics of this subduction-to-collision, convergent stage is still unknown (Fig. 4a). SW-vergent recumbent folds and thrusts are the first collisional structures that propagate throughout the OMZ (Expdsito et al. 2002); at the same time, NE-vergent recumbent back-folds were formed in the southern border of the CIZ. Because the train of recumbent folds and thrusts migrated from the CIZOMZ boundary, it is inferred that these structures must be a consequence of the OMZ underthrusting beneath the CIZ (Fig. 4b; Simancas et al. 2001). The Devonian age of these structures is indicated by the unconformable deposition of Late Devonian to Vis6an sediments over the already folded rocks of Late Proterozoic and Early Palaeozoic age (van den Boogaard & V~izquez 1981; Ribeiro 1983; Giese et al. 1994). Some of the flyschoid deposits are dated to the Early Devonian (Piqarra 1998; Pereira et al. 1998) and probably indicate the very beginning of the collision. Although there are no definitive data on the kinematics of the collision at this stage, the obliquity between the traces of Devonian thrusts and the boundaries of the OMZ (Fig. 3) suggests that some left-lateral component of displacement may have existed. During the Early to Mid-Devonian, when continental collision between the OMZ and the CIZ started, ophiolite obduction (the Beja-Acebuches ophiolite) and accretionary prism development (the Pulo do Lobo Unit) were occurring at the OMZ-SPZ boundary, suggesting that subduction rather than collision was

TRANSPRESSION AND MANTLE PLUME IN IBERIA

349

Fig. 3. Structural sketch of SW Iberia, showingmajor structures (for location, compare with Fig. lb). In the OMZ, Carboniferousstructures and some of the Devonian structures are aligned obliquely to the tectonic boundaries of the zone. In the SPZ, an en echelon pattern of early Carboniferousfolds and thrusts is envisaged. At both boundaries of the OMZ, conspicuous early Carboniferousleft-lateral ductile shear zones developed, and late Carboniferous strike-slip faults are very prominent. All these geometric and kinematic data support the hypothesisthat in SW Iberia the collisional convergence was very oblique (left-lateral).

happening there (Fig. 4b). Subduction polarity, if indicated by the Pulo do Lobo accretionary prism, would have been towards the NE (present-day coordinates). The subducted oceanic crust apparently belonged to the Rheic ocean, according to palaeomagnetic reconstructions (e.g. Nys~ether et al. 2002). However, a more complex history, similar to that invoked in Central Europe, involving Early Devonian closure of the Rheic ocean and the subsequent opening of a small oceanic domain that finally closed in late Devonian times (Berthelsen 1992; Franke 2000; McKerrow et al. 2000), cannot be completely discarded. In this respect, the lack of outcrops in the SPZ of rocks older than Late Devonian precludes the possibility of obtaining reliable data about the Devonian history of this collisional border. The most unusual features of the orogenic evolution of southwestern Iberia occurred in latest Devonian to Early Carboniferous times, when extension and voluminous magmatism took place (Fig. 4c). Early Carboniferous terrigenous deposits and interlayered volcanic rocks unconformably overlie the older Palaeozoic and Late Proterozoic rocks in the southern CIZ and in the OMZ. In the SPZ, Late Devonian to Carboniferous

sedimentary deposits and volcanic rocks constitute the total outcrop in this zone (Oliveira 1990), with early Carboniferous volcanism reaching its maximum manifestation here (Munh~i 1983; Thirblemont et al. 1998) in connection with pervasive hydrothermal activity and giant sulphide deposits (the Iberian Pyrite Belt; Leistel et al. 1998; Sfiez et al. 1999). In the OMZ and southern CIZ, Early Carboniferous volcanism is much less voluminous, despite the presence of a number of contemporaneous plutonic bodies of that age (e.g. the Beja gabbro; Pin et al. 1999). However, the IBERSEIS seismic crustal image strongly suggests that a huge volume of magma was trapped in the OMZ middle crust (IRB, Fig. 2). Accordingly, we conclude that an extensional stress regime accompanied by voluminous magma production characterized the latest Devonian to early Visran times in southwestern Iberia, straddling previously amalgamated continental blocks (Simancas et al. 2002). Strong evidence of oblique-slip (left-lateral and normal) kinematics during this extensional stage comes from the important shear zone at the C I Z - O M Z boundary, which is dominated by left-lateral displacement (Fig. 3); this shearing had a thrust

350

J.F. SIMANCASETAL.

Fig. 4. Proposed Variscan evolution of SW Iberia based on geologicalstudies and the interpretation of the IBERSEIS seismic profile. The transtensional,hotspot stage in earlier Carboniferoustimes is probably related to an underlyingmantle plume, as suggested in the figure. The major left-lateraldisplacementsare indicated by appropriate symbols, but cannot be fully appreciated in this 2D view.

component in Devonian time and a normal fault component during the Early Carboniferous (Burg et al. 1981; Azor et al. 1994; Simancas et al. 2001). Thus, the earliest Carboniferous stage may have been transtensional. Compressional tectonics restarted to dominate southwestern Iberia at the beginning of the late Vis6an, as evidenced by several structural observations. Thus, throughout the OMZ and in the southern CIZ, upright to inclined folds developed at this time, superimposed on previous recumbent folds and thrusts and folding, for the first time, sediments and volcanic rocks in the Early Carboniferous basins (Figs 2c and 4d). At the O M Z - S P Z boundary, a conspicuous early structure of this renewed compression is a ductile shear zone affecting the oceanic amphibolites (Crespo Blanc & Orozco 1988), which has been dated to c. 340 Ma (i.e. Early Vis~an; 4~ on amphiboles, Dallmeyer et al. 1993). In the SPZ, terrigenous flysch deposits heralded the arrival of shortening directed from the southern OMZ and propagating southward, as indicated by the southward younging of the flysch, from Late Vis~an to earliest Westphalian (Oliveira 1990). Deformation in this area built the southward-vergent fold-and-thrust belt depicted in Figs 2c and 4d. The kinematics of the Carboniferous shortening in SW Iberia is well known from a number of lines of structural evidence (Fig. 3): (1) at the O M Z - S P Z boundary, the shear zone affecting the oceanic amphibolites shows an oblique

left-lateral-thrust displacement (Crespo Blanc & Orozco 1988); (2) folds in the OMZ run in a direction markedly oblique to the boundaries of the zone; (3) folds and thrusts in the SPZ show an en echelon arrangement, typical of left-lateral oblique shortening. The latest stages of the convergence (late WestphalianStephanian) are characterized by sinistral strike-slip faulting (Figs 3 and 4d; Simancas 1983; Crespo Blanc 1992; Jackson & Sanderson 1992), apparently indicating a progressive increase of the left-lateral component of the collision during Carboniferous. Three-dimensional crustal wedges observed both in the seismic section of the IBERSEIS profile and as tapering units on geological maps can be taken to be the result of these lateral displacements. The orogenic manifestations end with moderate extensional readjustments, giving way to basaltic dykes and minor basins filled in with Autunian molasse and volcanic flows (Simancas 1983; Sierra & Moreno 1998). There is little doubt that the left-lateral kinematics found in this region is connected with the development of the Ibero-Armorican arc, with SW Iberia being placed in the southern branch of the arc (Fig. la). Geological and palaeomagnetic data have demonstrated that the arc is mostly a secondary orogenic feature (P~rez Estafn & Bastida 1990; Weil et al. 2001), which has been explained by several indentation models (Fig. 1C; Matte 1986; Ribeiro et al. 1995).

TRANSPRESSION AND MANTLE PLUME IN IBERIA

A mantle plume interacting during continental collision The geological and seismic data, summarized in the previous sections, indicate that since the onset of the continental convergence in the Early Devonian until the Late Devonian, continental collision was taking place in southwestern Iberia. The first collisional stage is recorded by a number of structural and metamorphic features, namely recumbent folds and thrusts deforming a detached upper crust, eclogite-facies metamorphism and ophiolite obduction. However, in the latest Devonian the tectonic scenario passed to be dominated by crustal extension (normal faulting), basin formation and conspicuous magmatism. In the SPZ, the latest Devonian-Early Carboniferous, rift-related volcanism is a conspicuous feature, but it should be stressed that an additional quantity of magma must have been trapped deeper in the crust, if one is to explain the thermal input and the mantle signature of the huge hydrothermally driven sulphide deposits in the Iberian Pyrite Belt (Leistel et al. 1998; Sfiez et al. 1999; Nieto et al. 2000). In this respect, many prominent reflectors in the SPZ portion of the IBERSEIS seismic profile may be extensional fault zones filled with igneous rocks; these extensional faults have been later inverted as thrusts (Fig. 2). The OMZ and the southern CIZ suggest a different scenario: the Early Carboniferous magmatism observed at surface is less abundant than in the SPZ, but there is seismic and geological evidence for the storage of a great volume of magma in the middle crust (the above-mentioned IRB; Simancas et al. 2003; Carbonell et al. 2004). Thus, taking together the geological (structural, petrological and metallogenic) and seismic (IBERSEIS deep reflection profile) data, an extraordinary production of mostly mantle-derived magma can be inferred, along with a sudden change in the regional stress regime in southwestern Iberia during the Early Carboniferous. These important tectonic and magmatic changes are compelling evidence for a mantle perturbation at that time. Most of the extensional and/or magmatic intra-orogenic stages described so far in old and modern orogens have been tentatively explained in connection with the behaviour of the subducting lithospheric slab, either as slab retreat (Royden & Burchfiel 1989; Huismans et al. 2001; Jolivet et al. 2001) or as some other mechanism (delamination, slab break-off) that was able to remove the unstable lithospheric root (Chemenda et al. 2000; Houseman & Molnar 2001). These tectonic hypotheses imply that the observable effects have to be located above the subducted slab; that is, in the overriding plate. However, this is not the case in southwestern Iberia, where the extensional and magmatic intra-orogenic event is recorded simultaneously in different continental blocks, separated by sutures. Specifically, the SPZ is a very unusual external zone of the Variscan orogen: it shows a high-thermal regime at this stage, expressed by widespread volcanism (Munhfi 1983; Mitjavila et al. 1997; Thi~blemont et al. 1998), giant sulphide deposits (Leistel et al. 1998; Sitez et al. 1999) and low- to very low-grade metamorphism related to a high geothermal gradient (Munhfi 1979; Simancas 1983). A mantle plume hypothesis is a tectonic scenario that considerably better fits the features of the Early Carboniferous evolution in southwestern Iberia. The mantle plume would have been most active during the latest Devonian-Visran, the age interval of most Variscan basic rocks and the time when extensional regional stress prevailed. Numerical models provide support for the view that mantle plumes can strongly influence the state of regional stress (Ratcliff et al. 1998). In late Visran times, the effects of the mantle plume are inferred to have faded out in southwestern Iberia, as a result of either the decay of the plume or its migration elsewhere. Significant motion (both drift and rotations) of the Variscides during the Carboniferous collision has been derived from palaeomagnetic studies; specifically, Edel (2001) suggests a southward drift during the late Visran, which could account for the disappearance

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of the plume manifestations. In this respect, we note that basin opening and volcanism of late Visran age have also been described in the Variscan Massif Central of France as an intra-orogenic extensional event (Bruguier et al. 1998). Nevertheless, current knowledge of the Devonian and Carboniferous plate motions is not accurate enough to undertake the task of reconstructing the track of hypothetical mantle plumes.

Conclusions The IBERSEIS deep seismic reflection profile is a high-quality crustal image that has shed light on the Variscan orogenic evolution of southwestern Iberia (Fig. 2). The crustal architecture of southwestern Iberia is now well constrained and, therefore, better-founded suggestions can now be made about its lithospheric evolution, despite the limitations inherent in most of the seismic images of old orogens, mainly the development of post-orogenic flat Mohos. The southwestern Iberia crust was rebuilt during the Variscan orogeny, recording the amalgamation of three continental domains: the SPZ, the OMZ and the CIZ. The major suture, which would correspond to the closing of the Rheic ocean, seems to be located at the S P Z - O M Z boundary, but some kind of oceanic domain may also have existed between the OMZ and the CIZ. Geological mapping and kinematic data (indicating left-lateral displacements) demonstrate the transpressional evolution of the continental collision in southwestern Iberia. Furthermore, the IBERSEIS seismic image provides a new section that completes the 3D view of the upper crustal wedges in which transpression is resolved. This wedging strongly modifies the geometry of the sutures in the upper crust, and out-of-section displacements can explain the crucial problem of the lack of volcanic-arc rocks, formed during the subduction of the Rheic ocean, at the surface. The lower crust of both SPZ and OMZ seems to have evolved independently (i.e. detached from the upper crust), although some shortening is also visible at this deep level. The continuation in the lower crust of the S P Z - O M Z suture is probably located along gently NE-dipping reflections, thus placing the SPZ lower crust under the southern OMZ (Figs 2c and 4d). The record of an extensional and magmatic Early Carboniferous intra-orogenic stage is a characteristic feature of the orogeny in southwestern Iberia. The IBERSEIS seismic image has strongly enhanced the significance of this stage by revealing the existence of a great, more than 100 km long, reflective body in the middle crust of southwestern Iberia, which has been interpreted as a voluminous accumulation of basic magma. Accordingly, in the OMZ, Early Carboniferous magmatism is as important as in the SPZ, although, in the latter, surface manifestations are prominent and intrusions are dispersed in the crust (Figs 2c and 4c). The trapping of such a volume of magma in the OMZ middle crust, with the subsequent crustal contamination, has proved to be of crucial importance in explaining the contrasted petrological and metallogenic evolution of the SPZ and the OMZ. In the former, giant sulphide deposits are found, exhibiting a mixed crustal-mantle signature and indicating widespread hydrothermal activity in the upper crust; in the latter, the most noteworthy mineral deposits are Ni-Cu, mantle derived but showing significant contamination with pelitic carbonaceous material (most probably, the Serie Negra rocks of Late Proterozoic age). Despite these differences, the same basic phenomenon seems to have been active under southwestern Iberia in Early Carboniferous times: a mantle anomaly producing abundant magmatism and extension. We can refer to this period of time as a hotspot stage, and we relate this anomaly to a mantle plume influencing southwestern Iberia in Early Carboniferous times. Therefore, the geology of this region can be taken to reflect the interaction of a transpressional orogenic evolution and a mantle plume.

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Funding for the IBERSEIS deep seismic reflection profile has been provided by CICYT-FEDER (1FD1997-2179/RYENI), Junta de Andalucfa, ENRESA, the Swedish Research Council and the Instituto Geol6gico y Minero de Espafia. The research has also been supported by the Spanish Ministery of Science and Technology (grants: BTE2000-0583-C02-01, BTE2000-3035-E and BTE2000-1490-C02-01). We thank C. Biermann and P. Matte for their constructive criticism. F. Gonz~ilvez Garcla has improved our English text.

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Post-Variscan (end Carboniferous-Early Permian) basin evolution in Western and Central Europe T. M c C A N N 1, C. P A S C A L 2'3, M. J. T I M M E R M A N 4, P. K R Z Y W I E C 5, J. L O P E Z - G O M E Z 6, A. W E T Z E L 7, C. M. K R A W C Z Y K 8, H. RIEKE 9 & J. L A M A R C H E 1~

l~

lGeologisches Institut, Bonn University, Nuf3allee 8, 53115 Bonn, Germany (e-mail: [email protected]) 2Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands 3Present address: NGU, Geological Survey of Norway, N-7491 Trondheim, Norway 4Institut fiir Geowissenschaften, Universitiit Potsdam, Karl-Liebknecht-Strasse 24, 14476 Potsdam, Germany 5Polish Geological Institute, ul. Rakowiecvka 4, 00-975 Warsaw, Poland 6Instituto de Geolog{a Econdmica, CSIC-UCM Facultad de Geolog{a, 28040-Madrid, Spain 7Geologisch-Paliiontologisches Institut, University of Basel, Bernoullistrasse 32, CH-4046, Basel, Switzerland 8Geoforschungszentrum, Telegrafenberg, 14473 Potsdam, Germany 9panTerra Geoconsultants B.V., Weversbaan 1-3, 2352 BZ Leiderdorp, Netherlands de Provence, Unitd CNRS Gdologie des Carbonatds, Place Victor Hugo, Case 67, 13331 Marseille, France

Abstract: The Variscan orogeny, resulting from the collision of Laurussia with Gondwana to form the supercontinent of Pangaea, was followed by a period of crustal instability and re-equilibration throughout Western and Central Europe. An extensive and significant phase of Permo-Carboniferous magmatism led to the extrusion of thick volcanic successions across the region (e.g. NE German Basin, NW part of the Polish Basin, Oslo Rift, northern Spain). Coeval transtensional activity led to the formation of more than 70 rift basins across the region. The various basins differ in terms of their form and infill according to their position relative to the Variscan orogen (i.e. internide or externide location) and to the controls that acted on basin development (e.g. basement structure configuration). This paper provides an overview of a variety of basin types, to more fully explore the controls upon the tectonomagmatic-sedimentary evolution of these important basins.

The Permo-Carboniferous collision of the continents of Gondwana and Laurussia, termed the Variscan orogeny, led to their amalgamation and the formation of the late Palaeozoic supercontinent of Pangaea. Oblique convergence resulted in collisional processes in the Appalachians and the Urals, whereas sinistral wrench faulting caused widespread rifting of the Northern European crust (Pegrum 1984a,b; Ziegler 1990). In addition, the collapse of the thickened Variscan orogenic crust resulted in late orogenic crustal extension. The post-Variscan period was one of intense crustal re-equilibration and reorganization under an alternating transtensional and transpressional tectonic regime, and the combined effect of re-equilibration and tectonic activity controlled the kinematic patterns and subsidence of approximately 70 basins, all of which are characterized by a major strike-slip component in their deformational history. Western and Central Europe was already thermally weakened by the preceding orogeny and most of the Carboniferous-Permian basins trace long-lived Variscan fault systems (Henk 1993). Subsequent basin evolution involved extensive, predominantly clastic sedimentation (e.g. Glennie 1990; Maynard et al. 1997), and some of the newly formed rifts became the loci of extensive intraplate magmatism (e.g. Neumann et al. 2004). In recent times, there have been a number of volumes published on the Variscan history of Europe (e.g. Dallmeyer et al. 1995; Franke et al. 2000). Although these books provide much-needed information conceming the nature and style of deformation during the Variscan orogeny, they are not really concerned with the postVariscan evolution of the region. One of the main problems with unravelling the post-Variscan deformational history of Western Europe is the localized nature of deformation. Small isolated grabens were gradually filled by predominantly locally derived sediments and/or associated volcanic and volcaniclastic rocks. The nature of the sedimentation (predominantly alluvial) constitutes another problem, as correlation is much more difficult, both within and, more importantly, between grabens. The nature of the

sedimentary record within the basins that formed as a result of Permo-Carboniferous wrench-fault activity, therefore, frequently precludes detailed investigation, as it is difficult to correlate the sedimentary record from one basin to another. In those Early Permian basins that are rich in biogenic remains it may be possible to correlate within the basin (e.g. Schneider 1989, 1996). However, the general level of biostratigraphic uncertainty leads to problems with correlation in this stratigraphic interval. It is only when basinwide crustal subsidence gave rise to the widespread Northern and Southern Permian basins, with the establishment of a unified depositional pattern across Northern and Central Europe, that areally more extensive correlation becomes possible (e.g. McCann 1998a). There are, however, many Permo-Carboniferous basins that are outside the Southern and Northern Permian basins region, and such basins provide much-needed understanding of the broader evolution of the Permo-Carboniferous transition in terms of magmatic, tectonic and sedimentary activity outside the area of the former Carboniferous foreland basin. The following summary, by nature selective, reviews the evidence for basin formation at this time and investigates a selection of these basins. Of marked interest is the similarity in terms of basin infill (sedimentary and magmatic) between the basins, despite differences in both timing and location. The aim of this study, therefore, is to provide an overview of the main areas of basin initiation that followed the cessation of Variscan orogenic activity in Western Europe. The following sections will outline the regional framework of Western Europe and provide an introduction to the most recent research in these areas, including sedimentology, tectonics, magmatic history and basin modelling.

Background geology and palaeogeographical setting The Variscan belt is a broad (c. 1000 km) complex curvilinear feature extending across Europe and marking the zones of Vafiscan-age deformation (Fig. 1). Variscan orogenic activity was

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 355-388. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. The main tectonic elementsof Central and Western Europe (modified after Ziegler 1990; Berthelsen 1992). Inset map shows the tectonic zones corresponding to volcanic belts: MZ, MoldanubianZone; OMZ, OssaMorena Zone; RHZ, Rheno-Hercynian Zone; SPZ, South Portuguese Zone; STZ, Saxo-ThuringianZone; TTZ, Tornquist-Teisseyre Zone. Massifs: a, Armorican; b, Bohemian;bf, Black Forest; c, Massif Central; h, Harz Mountains;ic, Iberian Cordillera;rh, Rheinische Schiefergebirge;v, Vosges (after Francis 1988).

associated with the convergence of the southern continent of Gondwana with the northern continent of Laurussia (i.e. Laurentia, Baltica and Avalonia), to form the supercontinent of Pangaea and leaving a relict Palaeotethys to the east (Scotese & Langford 1995; Fig. 2). The reconstructed geometry of the Variscan orogen can be divided into a number of distinct geotectonic zones, many of which are separated by steeply dipping faults or shear zones, that exhibit a general continuity around the belt. Some of the boundaries may represent suture zones and these are marked by the occurrence of ophiolite assemblages and calc-alkaline continental arc volcanic rocks of Devonian or Carboniferous age. The high-grade core of the Variscan orogen runs through central and NW Spain, France, Germany and the Bohemian area of the Czech Republic (Figs 1 and 3). The Devonian to Early Carboniferous evolution of the Variscan orogenic system was governed by alternating tensional and compressional tectonic cycles, reflecting the development of an essentially intra-continental Pacific-type, back-arc system (Ziegler 1990). Four principal phases of deformation and exhumation may be recognized, each of which lasted between 20 and 39 Ma and was restricted in geographical extent. These phases probably resulted from the successive docking of continental lithospheric fragments along the southern margin of Laurussia. The first of these periods, the Ligerian Phase (Late Silurian-Early Devonian), was contemporaneous, but not cogenetic, with the Acadian phase of the Caledonian orogeny. The remaining three phases, Bretonian (Late Devonian-Early Carboniferous), Sudetian (Vis6an-Early Namurian) and Asturian (Westphalian-Early Permian), are assigned to the Variscan orogeny (Warr 2000). During the late Vis6an the Variscan orogenic system entered the Himalayan-type continent-continent collision stage. Continued dextral-oblique convergence of Gondwana and Laurussia and the progressive closure of the Proto-Atlantic and western

Proto-Tethys oceans was accompanied by the propagation of their collision front into the Appalachian and central Mediterranean domains. At the same time, major crustal shortening was achieved in the Variscan fold belt as indicated by the occurrence of major nappe structures, in part involving basement in the Moldanubian area, the southern Massif Central and the Variscan externides (Behr et al. 1982; Burg et al. 1984; Cazes et al. 1986; Behr & Heinrichs 1987). During the late Vis6an and Namurian, the collisional front between Africa and the Southern European margin propagated rapidly westwards and eastwards. Following the Late Westphalian-Early Stephanian consolidation of the Variscan fold belt, convergence between Gondwana and Laurussia apparently changed from an essentially northsouth-directed collision to an east-west convergence. During Late Carboniferous times a subduction zone developed along the Variscan Deformation Front, with oceanic crust being subducted beneath the continental Variscan system (Gast 1988). Destruction of the subducted crust led to stretching (dominantly east-west) and associated block faulting. Coeval counter-clockwise rotation of the southern African plate against the stable northern European craton caused wrenching along NW-SE-trending faults (Ziegler 1990). Continental-scale dextral shears (e.g. the TornquistTeisseyre fracture zone) were linked by secondary sinistral and dextral shear systems (Ziegler et al. 2006). The westward translation of Gondwana relative to Laurussia along a 3500 km long dextral megashear has been proposed, but recent work from northern Spain does not support the existence of such a major structure (Weil et al. 2001). It is, however, possible that a series of smaller shear structures were active (see the following section). The development of the Variscan orogen involved major crustal shortening and subduction of substantial amounts of supra-crustal rocks, continental and oceanic crust, and mantle-lithosphere

POST-VARISCANBASINEVOLUTION,EUROPE

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From late Early Carboniferous times, compression within Central Europe was continuous for c. 45 Ma, with Variscan deformation advancing northward at a rate of c. 0.5 cm a -1 (Ahrendt et al. 1983). This is only slightly less than the value of 0.79 cm a -1 that characterized Alpine orogenic compression (Schmid et al. 1996). In western Germany, for example, a series of major north- and southverging, basement-cored nappes were emplaced during Namurian and Westphalian times. The widespread occurrence of low-pressure metamorphic rocks and late to post-orogenic calc-alkaline intrusive rocks in the internal parts of the Variscan orogen suggests that a significant amount of crustal shortening, accompanied by crustal delamination, subduction and anatectic remobilization of lower crustal material, and partial melting of upper mantle material, occurred (Ziegler 1984, 1986). Rough estimates of total crustal shortening yield a value of c. 150-200 km (Ziegler 1990). This is less than half of the crustal shortening recorded from the Alpine Realm during the Tertiary, which was estimated at c. 500 km (Dewey et al. 1989; Schmid et al. 1996). By latest Westphalian times, the Variscan fold-and-thrust belt of Westem and Central Europe was consolidated and inactive (Ziegler 1990). However, localized tectonic activity is evidenced by diverse angular unconformities between the Westphalian and Permian deposits (e.g. NE German Basin; McCann 1999). In Germany, post-orogenic uplift of the Rheno-Hercynian thrust belt, preceding deposition of the Permian sediments, amounted along its northern margin to 2-3 km, increasing southwards to 6 km (Littke et al. 2000), with values of up to 10 km in the area of the Saar-Nahe Basin (Oncken et al. 2000).

The e a s t w a r d s extent o f Variscan d e f o r m a t i o n

Fig. 2. Early(a) and Late (b) Permian palaeogeography (after Scotese & Langford 1995).

(Ziegler et al. 1995). The convergence rate between Gondwana and Laurussia, however, was not constant during Devonian and earliest Carboniferous times, as suggested by the Early Devonian development of an extensive back-arc rift system in Western and Central Europe, which remained intermittently active until the late Visran (Zieger 1990). Full-scale collision between Gondwana and the southern extension of Laurussia occurred during the late Visran at a time when the back-arc extensional system that had governed the early evolution of the Variscan orogen in Western and Central Europe finally became replaced by a compressional stress regime (Ziegler 1990). Gondwana and Laurussia became increasingly coupled during the Early Carboniferous, as is evident from their joint clockwise rotation and northward drift (see Ziegler 1990, Fig. 8).

As noted above, the collision between Gondwana and Laurussia continued to develop until late in the Carboniferous-Early Permian, at which time the intercontinental collision began to affect the northwestern part of Africa (e.g. the West African orogens have a Late Carboniferous to Early Permian age; Lrcorch6 et al. 1989). As a result of the extensive collision episode, a central Pangaean mountain range was formed, extending from Mexico to Poland (Golanka & Ford 2000) and southwards to Morocco (Pique & Michard 1989). Late Carboniferous events were also marked in the Alps (e.g. Matter et al. 1987), the Carpathians (Dallmeyer et al. 1995) and the Rhodope area (Yanev 1992). Further eastwards, the situation was a more complex one, as it involved the rifting (from Late Carboniferous times onward) from Gondwana of several Cimmerian continents (e.g. parts of Turkey, Iran and Afghanistan; Seng6r et al. 1984; Scotese & Langford 1995; Fig. 4). During the main phase of the Variscan orogeny (late Visran to Westphalian), the collision front between Gondwana and Laurussia propagated eastward and southwestward in conjunction with the progressive closure of the Palaeotethys and Protoatlantic oceans. By Westphalian times, Gondwana had collided with North America whereas to the east Palaeotethys remained open. Therefore, the westem parts of the Variscan megasuture were characterized by a Himalayan-type (continent-continent collision) setting, whereas its eastern parts remained in an Andean-type (continent-ocean collision) setting (Stampfli 2000; Ziegler & Stampfli 2001). The setting of the West European segment of the Variscan orogen was, therefore, transitional between a Himalayan- and an Andean-type setting (see Ziegler et al. 2006, for details). During the Stephanian and Autunian orogenic movements continued in the AppalachianMauretanides (Rodgers 1970; Michard & Sougy 1977) and in the Urals (Ivanov et al. 1977), whereas the Variscides region remained largely inactive. These latter movements were accompanied by the emplacement of a right-lateral transform fault system, which linked the southern Uralides and northem Appalachians, and which crossed Europe, where it caused the development of a complex pattern of conjugate shear faults and related pull-apart structures (Arthaud & Matte 1977; Ziegler 1978a).

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Fig. 3. Structural sketch map of the European Variscan orogen. Three external coal basins are shown in black. C.C.S., Coimbra-Cordoba suture; L.R.S. Lizard-Rhenish suture; N.V.F., North Variscan Foreland; M.C.S., Massif Central suture; M.T.S., M~nchberg-Tepla suture; O.M.S., Ossa-Morena suture. The Iberia and Corsica-Sardinian blocks are represented in their possible Permian positions relative to Europe (after Matte 1991).

An important problem is the continuation of the Variscan orogen in SE Europe and Turkey. In Western Europe, subduction ended with the collision of NW Gondwana (i.e. Morocco, Algeria) with SW Europe (i.e. Iberia, Corso-Sardic block). Further to the east (i.e. Tunisia, Libya, Egypt) there is no evidence of subduction-related activity (e.g. intense folding, subductionrelated magmatism) at the northern margin of Gondwana during the Late Palaeozoic (Dornsiepen et al. 2001). In contrast, orogenic activity can be traced along the southern margin of Laurussia from the Alps to the Caucasus, although it appears that the closure of Palaeotethys was asymmetrical, with subduction occurring on the northern (i.e Laurussian) margin (Vavassis et al. 2000). It is probable that the active spreading ridge of Palaeotethys arrived at the southern margin of Laurussia around the Permo-Carboniferous boundary, leading to a cessation in active subduction, and the convergent margin was then transformed into a dextral strike-slip movement zone (Dornsiepen et al. 2001; Fig. 4). Stampfli et al. (2001) noted that there is a general pattern of rifting across the region of Tethys. The dynamic evolution of the region of Neo-Tethys (India, Arabia, Turkey) reveals that the Early Carboniferous rifting phase was followed by a second rifting phase in the Early Permian and the late Early Permian-Late Triassic opening of Neotethys. A similar pattern of rift activity is recorded in the East Mediterranean area, where the initial Carboniferous rifting phase affected the southern margin of the area between Syria and Tunisia. This was followed by a Permian rifting phase and protracted Late Permian subsidence (extending to the present day). Following late-stage subduction of Palaeotethys, and in conjunction with slab roll-back (since Early Permian times) and slab detachment beneath the Variscan orogenic belt, the Cimmerian Terrane was detached from the Gondwana margin and the Mediterranean-Neotethys Basin began to open along the Eurussian margin. The post-collisional Permo-Carboniferous rifting phase accompanied the transtensional collapse of the overthickened Variscan lithosphere in a basin-and-range fashion (Malavielle 1993). This analogy implies the presence of a still-active or transform margin in areas where the Palaeotethys was not yet closed during the Permian. The mid-ocean ridge of

the Palaeotethys was obliquely subducted beneath the active Eurasian margin during the Permo-Carboniferous (Stampfli 1996). This was responsible for the widespread Late Carboniferous and Early Permian magmatism and volcanism characterizing the southern Variscan domain.

L a t e to p o s t - V a r i s c a n s t r u c t u r e s

During the main phase of the Variscan orogeny, lateral escape tectonics was accompanied by the development of synorogenic transtensional basins in the Armorica-Iberia-Avalonia domain. During the latest Carboniferous and Early Permian the convergence of Gondwana and Laurussia changed from oblique collision to dextral translation controlling the Alleghenian orogeny of the Appalachians. At the same time the Variscan orogen was transected by conjugate wrench faults, partly terminating in pull-apart basins (Ziegler 1990; Matte 1991). Along the northern margin of Gondwana, rifting resumed during the Early Carboniferous, reflecting an intraplate stress reorganization following the docking of the composite Hun terrane to the Laurussian margin (see Stampfli et al. 2001, for details). A number of major late to post-Variscan tectonic structures have been recognized. The Tornquist Zone, comprising a fan-shaped series of fault and fault-zone splays was activated in the western Baltic area (Berthelsen 1992). Indeed, right-lateral wrench movements along the Tornquist-Teisseyre lineament led to the formation of the Oslo-Skagerrak Graben system (Ziegler 1978b; Neumann et al. 1992). According to Ziegler (1990), a major postVariscan wrench fault was active along the present Elbe Fault System (northern Germany) as part of a complex megashear zone that developed in North-Central Europe between the Appalachians and Uralides as a result of dextral translation of the Armorican-European plate with respect to the African plate (Arthaud & Matte 1977). Along the SE end of this wrench fault, geological evidence for Late Carboniferous-age dextral ductile shearing has been found in outcrops of the Elbe Zone (Mattern 1996). Scheck et al. (2002) have noted that maximum deformation occurred along the Elbe Fault System during late Carboniferous

POST-VARISCAN BASIN EVOLUTION, EUROPE

Teisseyre lineament. Displacements along the Bay of Biscay fault system were in part taken up in the Arctic-North Atlantic rift, in which crustal distension continued during the latest Palaeozoic, as illustrated by the East Greenland graben system (Haller 1971). Indeed, from the late Early Carboniferous onwards, Laurussia was transected by the Arctic-North Atlantic rift system, which was partially superimposed on the Caledonian suture zone (Ziegler 1990).

A'ER'J EUROPEIf SOUTH AMERICA

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Palaeogeography and palaeoenvironments

ARABIA

NA EU SIB

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(c) Fig. 4. Schematic plate configuration and plate boundary geometry between Europe and Gondwana. (a) Early Permian. (b) Mid-Late Permian. (e) Mid-Late Triassic. (After Dornsiepen et al. 2001.) wrenching and volcanism (and subsequent to these events). Late Palaeozoic dextral transtensional movements with estimated displacements of 60-120 km have been postulated. The major Variscan faults in Germany and Poland (i.e. Elbe Zone, Main Intra-Sudetic Fault, Odra Fault Zone) are considered to have been dextral shears trending N W - S E during the late Devonianearly Carboniferous with maximum offsets of 5 0 - 3 0 0 k m (Aleksandrowski 1995; Alexandrowski et al. 1997) and dextral offsets of up to 400 km along the Teisseyre-Tornquist Zone (Lewandowski 1993). In late Carboniferous-early Permian times, many local and regional W N W - E S E - to NW-SE-oriented fault or shear zones in the Western Sudetes were reactivated under a semi-brittle regime. The main elements of the Late Variscan fracture system of northern Africa and Europe are the Agadir-Kelvyn, Gibraltar and Bay of Biscay fault systems, as well as the Tornquist-

Lithologically, the Upper Carboniferous successions across the region are relatively monotonous. This is partly due to the similarity in terms of lithological and stratigraphical evolution between the Variscan fold-and-thrust belt and that of its foreland basins, related both to the infilling of the basins and the gradual cessation of Variscan-age tectonic activity. Only the Namurian A still shows the orogen-foreland differentiation. During Late Carboniferous times, paralic coalfields extended across Laurussia from North America across Britain and the Rheno-Hercynian region of Germany to Silesia and the Donets region, Ukraine. Coal-forming environments were present south of the Rheic suture in the Saar-Lorraine Basin during late Westphalian times (possibly earlier) and continued until early Stephanian times. After a short hiatus, these conditions were re-established in mid-Stephanian times (Cleal 1984). This extensive region experienced several marine incursions and clearly most of the Westphalian deposition (largely non-marine) occurred close to sea level (Paproth 1991) on the northern (passive margin) side of the Rheic suture. In contrast, to the south of the suture mainly intermontane limnic basins (of Stephanian age) developed on the various components of the Armorican Terrane Assemblage (McKerrow et al. 2002). During the Permian, continental size and aggradation were at a maximum, as few continental fragments escaped incorporation into Pangaea. The new continent, stretching almost from pole to pole, had a significant amount of terrestrial area. A single world ocean, Panthalassa, with a semi-enclosed Tethyan Sea, dominated the marine environment. This singular continental configuration led to the development of extreme climatic conditions. Furthermore, the Permian climate was not a stable one, with evidence of profound changes throughout the duration of the period (Veevers & Powell 1987; Barton & Fawcett 1995; Parrish 1995). Climate was not the only unstable factor. Given the sheer size of Pangaea, the supercontinent was unstable from the outset. The timing of break-up was varied, beginning within the late Carboniferous immediately north of the Variscan orogen. In many places (e.g. Germany and Belgium), extension to the north of the Variscides occurred coevally with thrusting and strikeslip faulting further south (Ziegler 1990). However, the timing and extent of individual phases of extension and rifting throughout the North Atlantic (and associated) rift systems are still subject to some debate, as the dating of Late Carboniferous and Early Permian red beds is imprecise. Following the main phases of Variscan compression, thermal relaxation of the crust occurred in Early Permian times, creating the rifts and graben that allowed accumulation of the first phase of sedimentation.

Stephanian-Autunian magmatic activity During the Permo-Carboniferous, the Variscan foreland in Europe was subjected to extensive rift-related tectonism and related extensional magmatic activity (Fig. 5). The rare occurrence of Autunian-age sediments suggests that much of the area later occupied by the Southern Permian Basin was regionally uplifted and subjected to profound erosion. This uplift, probably induced

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Fig. 5. Distribution of late Carboniferous and early Permian magmatic rocks in NW Europe (after Ziegler 1990). NGB, North German Basin; OS, Oslo Rift; WSC, Whin Sill Complex; VF, Variscan Front; RFH, RingkObing-FynHigh; OR, Oslo Rift; MV, Midland Valley. by a combination of wrench-related lithospheric deformation, magmatic inflation of the lithosphere and thermal erosion of the mantle lithosphere, was coupled with periods of significant magmatic activity. Melt generation was probably related to localized divergent wrench-induced decompressional partial melting of the uppermost aesthenosphere and the lithospheric thermal boundary layer, possibly combined with the impingement of a not very active mantle plume on the base of the lithosphere (Ziegler 1996). Although the most voluminous and evolved magmatism occurred in the North German Basin (Benek et al. 1996), dated at 297-302 __+3 Ma (Breitkreuz & Kennedy 1999), the Oslo Rift contains the most extensive and best-preserved sequences of basaltic lavas associated with this event. In the latter, Rb-Sr age determinations indicate that the total period of magmatic activity extended from c. 305 to 245 Ma (Sundvoll et al. 1990), and the earliest basaltic magmatism appears to have been restricted to a relatively short period (305-290Ma) (Neumann et al. 2002). However, recent dating suggests that these ages are slightly too young and that the duration of the various magmatic periods in the Oslo Graben could be shorter than proposed by Sundvoll et al. (1990) (see Neumann et al. 2004). The early basalts from the Oslo Graben contain the least evolved magmatic products in the region and provide information about the primary magmas and the magma sources involved in the magmatic activity. Volcanic activity, however, was widespread throughout Europe, with thick sequences being deposited elsewhere; for example, in Spain (e.g. Iberian Range, Lago & Pocovi 1984), dated at 282 + 12 Ma (Hern~indo et al. 1980), and in the Italian Alps (e.g. Collio Basin, Breitkreuz et al. 2001), dated at 283 + 1 Ma and 2 8 1 _ 2Ma (Schaltegger & Brack 1999), Liguria and Sardinia (e.g. Cortesogno et al. 1998). Autunian-age magmatic activity also included the intrusion of the Whin Sill dolerite (northern England), some dolerite dyke swarms in the Midland Valley of Scotland and the Northumberland Basin (Francis 1978; Coward 1995), and in the Argyll area of Scotland

(Speight & Mitchell 1979), and episyenite dykes along tensional structures in the Central Range of Iberia (Gonz~ilez-Casado et al. 1996). In these areas volcanism followed Westphalian-age inversion and predated much of the Permo-Triassic succession, suggesting that the region was underlain by hot asthenosphere. The widespread Stephanian-Early Permian (305-285Ma) alkaline intrusive and extrusive magmatism of the Variscan area and its northern foreland is mantle derived and shows evidence of strong crustal contamination (e.g. Marx et al. 1995; Benek et al. 1996; Breitkreuz & Kennedy 1999). Melt generation by partial melting of the uppermost aesthenosphere and the lithospheric thermal boundary layer was probably triggered by a rise in the potential temperature of the aesthenosphere and localized divergent wrench-induced decompression. Aesthenospheric upwelling was presumably triggered by detachment of subducted lithospheric slabs, resulting in a reorganization of the mantle convection system and the impingement of a not very active system of mantle plumes on the base of the lithosphere. There are at least two discrete tectonic settings for sedimentary basins that formed simultaneously during the Late Palaeozoic Variscan orogeny in Western Europe. Immediately adjacent to, and to the north of, the Variscan deformation front, narrow foredeep basins, termed the external basins (Variscan externides), are interpreted to have formed as a result of subsidence related to thrust loading of the crust. The related internal basins (Variscan internides) comprise those basins formed within the orogenic belt and that were structurally controlled. The following discussion of the magmatic activity within the Variscan-influenced regions of Europe will examine both basin types.

Extrusive activity (Variscan f o r e l a n d - e x t e r n i d e s )

In the foreland of the Variscan region, sedimentation and magmatism generally occurred in and close to the north- and NW-trending grabens. Widespread magmatism occurred as

POST-VARISCANBASINEVOLUTION,EUROPE felsic to mafic volcanism, and emplacement of mafic dykes and sills. In the Oslo Graben and Central North Sea, extension, block faulting, tilting, uplift and erosion of structural highs occurred simultaneously with, and subsequent to, the initial phase of volcanic activity. As noted above, the most voluminous magmatic activity occurred in the Oslo Rift, with more than 100 000 km 3 of produced melts (Neumann et al. 2004), and the NE German Basin, where more than 48 000 km 3 of magmatic material was extruded. Five stages of magmatism have been recognized and dated in the Oslo Graben (Olaussen et al. 1994; Neumann et al. 2004). The pre-rift stage (i.e. between 304 and 294 Ma) comprises felsic sill emplacement in underlying Westphalian sediments. The second stage corresponds to rift initiation and is characterized by widespread basaltic volcanism (i.e. B 1 basalts) representing the most primitive magmatic rocks in the Oslo Graben. The main rift stage (i.e. Stage 3) resulted in the fissure eruption of thick sequences of porphyritic trachyandesite flows (rhomb porphyries), which was accompanied by the intrusion of large amounts of syenitic magmas that have yielded ages of 292-298 Ma according to recent U-Pb dating (Neumann et al. 2004). During Stage 4 rifting abated and the magmatic style changed to the development of central volcanoes and caldera collapse. Rb-Sr dating of associated magmatic rocks gives ages between 280 and 240 Ma (Sundvoll et al. 1990). The final magmatic stage was dominated by the intrusion of composite batholiths of trachyandesite to rhyolitic compositions. In the subsurface of Northern Germany and Poland extensive Late Carboniferous to Early Permian felsic to intermediate volcanic rocks have been cored by exploration wells (e.g. Hoth et al. 1993). In this region, centres of volcanic activity appear to coincide with the intersection of fault systems (Plein 1978) and are probably related to pull-apart structures at the termination of subsidiary wrench faults that parallel the Tomquist-Teisseyre lineament. In Northern Germany the magmatic succession is up to 2000 m thick, and comprises predominantly calc-alkaline SiOz-rich volcanic rocks (Kramer 1977; Marx et al. 1995; Benek et al. 1996). The volcanic and pyroclastic succession of Northern Poland is broadly similar to that of NE Germany and comprises mainly acid volcanic rocks (rhyolites, rhyodacites) and locally abundant trachytes and trachybasalts (Pokarski 1988, 1989). In the northern part of the Fore-Sudetic Monocline, there is an Early Rotliegend-age volcanic succession composed of lava flows, pyroclastic deposits, and hypabyssal and intrusive rocks (Jakowicz 1994). The rocks, however, are not comagmatic, as they were derived from two sources, the mantle and crust. The parent magma was a high-aluminium basalt, but anatectic crustal fusion produced acid melts that mixed with the derivatives of the parent magma to produce poorly homogenized hybrids with intermediate compositions. The distribution of these rocks is parallel to the trend of regional tectonic units of Early Rotliegend age, and their thickness varies from a few metres to over 1500 m. In general, the volcanic rocks of the Polish Basin are limited to its western margin and overall only attain a thickness of c. 100-200 m (with exceptions, see above), and are, thus, an order of magnitude less than in the adjacent NE German Basin (Karnkowski 1999). In the offshore regions of Northern Europe it is more difficult to determine the extent of magmatic activity. For the Central North Sea, Denmark and Skagerrak Graben there is geophysical evidence of mafic intrusions in the middle and lower crust. Furthermore, in the Central North Sea area, the thick basalts and tufts (< 160 m) of the Inge Volcanics Formation (c. 299 Ma, Heeremans et al. 2004) are interbedded with Rotliegend-age mudstones and sandstones, whereas bimodal suites of basalt, trachyandesite and rhyolite flows occur in the eastem North Sea. However, in other parts of the Central and Northern North Sea and along the Scottish-Irish Atlantic seaboard, the importance of the Stephanian-Autunian tectonism is difficult to assess.

361

Many of these latter areas, however, were subjected to uplift (e.g. Dunlap & Fossen 1998). Intrusive activity (Variscan f o r e l a n d - externides)

Stephanian to Early Permian-age intrusive activity in the Variscan foreland occurred largely in the form of dolerite sills and dyke swarms. For example, the Whin Sill Complex in northern England comprises a series of sills ( < 9 0 m thick) and four major ENE-trending dyke echelons of high-Fe, sub-alkaline basalt (Johnson & Dunham 2001). The total volume of the dyke complex has been estimated at 120-215 km3; however, as it extends eastwards under the North Sea, this figure is a minimum. In the Midland Valley of Scotland, earthwest-trending dykes (up to 50 m wide and 130 km long) and a sill (< 180 m thick) are composed of sub-alkaline to transitional basalt, with a total volume of almost 500 km 3. This dyke swarm also extends c. 200 km eastwards into the North Sea (Smythe 1994). The Midland Valley of Scotland also contains Westphalian-Early Permian-age pyroclastic volcanic rocks, vents, necks, alkaline dolerite sills, dykes and plugs of basaltic to trachytic and phonolitic composition (Francis 1992). Many of the vents contain mantle and crustal xenoliths of high-grade mafic and felsic gneisses. The mantle xenoliths probably reflect derivation from variously metasomatized mantle sources, but biotite, amphibole and/or alkali feldspar-bearing ultramafic xenoliths and some of the megacrysts may be fragments of coarsegrained intrusions that fractionated under high-pressure conditions in the upper mantle. The mafic crustal xenoliths may represent metamorphosed cumulate or underplated mafic mantle melts, whereas the felsic gneisses were derived from Precambrian to Palaeozoic crust. In the western Scottish Highlands at least 3000 Early Permian, SW- to NW-trending camptonite and monchiquite dykes, many containing mantle xenoliths, intruded the Caledonian basement (although some of these dykes may be of Tertiary age; Rock 1983). In southern Sweden a c. 70 km wide swarm of NW-trending sub-alkaline basalt and basaltic andesite dykes occurs. Individual dykes can be up to 100 m wide but the majority vary between 1 and 50 m; their total volume has been estimated at c. 4000 km 3 (Obst et al. 2004). Of similar age may be the isolated, NNWtrending dolerite dykes on the SW coast of Sweden and the dolerite sills that intruded Cambrian alum shales in south-central Sweden (V/istergrtland). These dykes and sills may be related to volcanic rocks obtained from drill core in the Kattegat and offshore eastern Denmark and interpreted from seismic data as volcanic edifices (Mogensen 1994; Marek 2000). Drill core from the island of Rtigen (NE Germany) contains a few basalt sills that intrude lower Palaeozoic sediments, and that are also interpreted as having an age of late Carboniferous-early Permian (Korich & Kramer 1994).

Magmatic activity in the Variscan internides

Compared with the foreland and internal Variscides, post-tectonic magmatism in the Rheno-Hercynian foreland fold-and-thrust belt was largely restricted to the intrustion of granites (c. 295-293 Ma) in SW England and granitoids (c. 290-307 Ma), gabbros, and rhyolitic to basaltic sills and dyke swarms in the Mid-German Crystalline Rise and Harz area in Germany. StephanianAutunian-age magmatism in the internal parts of the Variscan orogen is dominated by granitoid intrusions exposed in midcrustal metamorphic terranes. Volcanic activity, on the other hand, consisted of widespread ash-fall tufts of distal provenance, with evidence of more pronounced activity being locally restricted. For example, the Ilfeld Basin (southern Harz region) is a small dextral pull-apart basin that contains up to c. 400 m of

362

T. McCANN ETAL.

Autunian-age tufts, ignimbrites and latitic, trachytic and rhyolitic volcanic rocks that erupted from a large number of small centres to overlie Stephanian-age sediments. To the west, the Saar-Nahe Basin contains Autunian-age sub-alkaline basalts and andesite flows that are associated with sub-volcanic dacite and rhyolite domes and pyroclastic deposits, and that are interbedded with fluvio-lacustrine sediments (Stollhofen 1994, 1998; Schmidberger & Hegner 1999). In the Western Mediterranean (Iberia, Pyrenees, Balearic Islands, Sardinia, Corsica, Provence), Stephanian-Autunian-age sedimentation occurred in deep basins. Here, volcanic activity took place in at least two stages, separated by a period of strikeslip activity and granite intrusion (270-290 Ma). The initial magmatic stages are of predominantly calc-alkaline character, whereas later, mid-Permian-age, magmatic rocks are alkaline. In the Iberian Peninsula and the Pyrenees, deep basins were formed in which sequences of late Westphalian-C to Stephanian-age clastic sediments are overlain by Autunian-age volcanic rocks. In some half-grabens in the Pyrenees, several cycles of calc-alkaline rhyolitic to andesitic and alkaline basaltic volcanic rocks have been recognized, ranging in age from Stephanian to Early Permian. Late Carboniferous magmatic activity is evidenced by thick sequences of lavas and ignimbrites, but Permian-age volcanicity has also been established (Mart/& Barrachina 1987; Marti 1996). The Permo-Carboniferous volcanic rocks are represented by a suite of basaltic andesite to rhyolite of calc-alkaline orogenic type, with a predominance of rhyolitic and rhyodacitic volcaniclastic deposits. Volcanic activity was essentially continuous, and without any significant change in terms of its character, from the deposition of the first Mid-Stephanian sediments to the beginning of the Late Permian sedimentation (c. 18 Ma; Marti 1996). During the Late Carboniferous, volcanism is represented by lavas and pyroclastic deposits mainly associated with caldera-forming events. Indeed, at this time it would appear that each basin had an independent volcanic history, with several volcanic centres located around the basin margins. Most of these deposits appear to be remnants of intra-caldera fill. During the Early Permian only rhyodacitic and rhyolitic magmas were extruded, and volcanic activity was concentrated in one basin (i.e. Castellar de N'Hug Basin), although it is possible that other volcanic centres were also active in the Central Pyrenees at this time (Martf 1996; Fig. 3). Eruptions were explosive, and generated widespread pyroclastic flows and associated pyroclastic surge and fall deposits. In the northern and central parts of the Iberian Peninsula, post-collisional magmas of mafic composition are rare. The magmatic rocks of the Massif Central area of France are limited to centimetre- to decimetre-thick layers of strongly altered air-fall tufts (tonstein) in both the Stephanian and Autunian sequences, and minor amounts of felsic to intermediate volcanic rocks. In the Rodez Basin, c. 5 m of andesite lavas and c. 30 m of trachytic tufts occur at the base of Autunian sandstones and are probably of early Autunian age. The Variscan basement south of the Decazeville Basin is intruded by a topaz-bearing rhyolite dome and a later leucogranite, both of crustal origin (Badia & Fuchs 1989). However, U - P b zircon dating of a tuff (332 • 4 Ma) and a rhyolite flow (333 • 2 Ma) at the base of, respectively, the Bosmoreau and Decazeville basins showed that some of the sediments and volcanic rocks, previously thought to be of Stephanian age, were actually deposited in the late Visran (Bruguier et al. 1998).

Permian basins in Europe Following the end of Variscan contraction in latest Westphalian times, subsequent Stephanian-Autunian magmatic activity and basin formation in both the internal Variscides (i.e. the area within the Variscan fold-and-thrust belt and to the south of it)

and the external Variscides (i.e. the area of the northem foreland basin, to the north of the Variscan orogenic belt) took place within a relatively short time span (c. 300-290 Ma) in a dextral wrenching setting. In the internal parts of the orogen the basins tend to be small, deep and isolated. Basin infill was a relatively rapid process with the accumulation of thick sequences of Rotliegend-age aeolian and fluviatile-lacustrine sediments. Tectonic activity coincided with an overall change in climate to semi-arid (Stephanian-Early Permian). As noted above, many of the basins contain volcanic rocks, and the presence of these magmatic successions aids in distinguishing the Permian sequences from the underlying Stephanian (see Plein 1995). The following section will examine a number of basins and regions across Europe to provide an overview both of the main basin types that formed at this time, and of the differences between the basins that formed in front of and within the Variscan orogen. The section is subdivided into those basins (Northern and Southern Permian basins) that were located to the north of the Variscan orogen, and those that were located within or to the south of the orogenic belt (Fig. 6).

Basins occurring within the Variscan f o r e l a n d Southern Permian Basin (SPB). The Southern Permian Basin comprises a series of connected basins extending across Northern Europe from England to Poland (Figs 1 and 6). The SPB, with a north-south extension of 300-600 km and an east-west extension of c. 1700 km, developed between the northern foreland of the Variscan mountain belt and the Ringkr High. The south-central parts of the basin are superimposed on the Variscan fold-and-thrust belts and intervening Gondwana-derived blocks (e.g. East Silesian Massif, SW part of the Malopolska Massif) and on the Precambrian of the East European Craton. Despite this, the geometry of the SPB shows no direct relationship with either the different crustal domains or the suture zones on which it subsided. The SPB is of great economic importance, containing a number of significant hydrocarbon finds (e.g. Groningen, SalzwedelPeckensen). By the early 1990s almost 180 wells had been drilled into the Rotliegend of Western Germany, and in Eastern Germany more than 1500 wells were drilled (see McCann et al. 2000, for details). The succession is subdivided into two distinct units, separated by the Saalian unconformity (Schneider et al. 1995). The Lower Rotliegend is characterized by acid and intermediate volcanic rocks with only minor sediments, whereas the Upper Rotliegend is essentially sedimentary and contains only rare volcanic rocks, of a more basaltic composition (Fig. 7). In the area of the Southern Permian Basin, up to 800 m of Stephanian continental red beds, containing correlative marine bands, were deposited in a broad successor basin to the Namurian-Westphalian Variscan foreland basin. During the late Stephanian-Early Permian, this basin was disrupted by predominantly transpressive wrench tectonics, as evidenced by the deep truncation of Late Carboniferous series and the conspicuous absence of deep Early Permian basins (Ziegler 1990; McCann 1999). These wrench tectonics were associated with the development of extensive volcanic fields (see above). Ziegler et al. (2006) noted that crustal thinning within the North German part of the SPB may be interpreted in terms of a 'stretching' factor of 1.45. This thinning is attributed to late Stephanian-Early Permian magmatic destabilization of the crust-mantle boundary that was paralleled by major thermal attenuation of the mantle-lithosphere (Van Wees et al. 2000). Southern Permian Basin: North German Basin. The North German

region, forming part of the later SPB, comprises crystalline basement of varying Precambrian ages (see McCann 1998b, for details), which is covered by a thick ( > 1 2 k m ) sedimentary

POST-VARISCAN BASIN EVOLUTION, EUROPE

I

Fenn~ ~Shield

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"~

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363

cb

Fig. 6. Outline of the Northern and Southern Permian basins of NW Europe (after Stemmeriket al. 2000).

,--'i'~'

succession of early and late Palaeozoic-, Mesozoic-, and Cenozoic-age strata. Basin evolution commenced with the destruction of the subducted oceanic crust of the Rhenohercynian Ocean, which led to lithospheric stretching (dominantly eastwest) and associated block faulting (Late Carboniferous-Early Rotliegend, and Early-Late Rotliegend; e.g. Gast 1991;

R e g i o n a l Scale NE - G e r m a n y

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...... ~ 9

I

Fig. 7. Permian time scale (after Menning 1995). In addition, Menning (2001) has provided an integrated Permian time scale based on absolute dates and field indicators for the duration of permian stages. More recently, the International Stratigraphic Chart (Gradstein et al. 2004) set the CarboniferousPermian boundary at 299 +__0.8 Ma (see also StratigraphischeTabeUe yon Deutschland, STD 2002; Menning et al. 2005).

Kiersnowski e t al. 1995; Stemmerik e t al. 2000; Kockel 2002), accompanied by calc-alkaline magmatic activity (which was particularly pronounced in NE Germany), which resulted in the formarion of a series of north-south-striking grabens. These acted as a feeder rift system for sediment transport from the Variscan hinterland to the areas to the north. The Lower Saxony rift system and associated basins (e.g. Hessen Basin) were part of a major continent-separating suture parallel to the subsequent MidAtlantic Rift (Gast 1988, 1991), which can be extended northwards via the Glfickst~idter Trough (Schleswig-Holstein), Hom Graben, and the Skagerrak Graben into the Oslo Graben (Fig. 1). Analysis of the form and sedimentary infill of a series of half-grabens along an east-west transect in Northern Germany reveals variations in both, suggesting that there were differences in both the rates and amounts of stretching. In the NE German Basin (NEGB) the angular unconformities between the the Westphalian-Stephanian and the overlying Permian sequences reflect a period of tectonic activity (McCann 1998a). This was followed by the major Permo-Carboniferous volcanic event that heralded the onset of Permian basin evolution. Initial sedimentation (i.e. Autunian) was restricted and mostly confined to isolated basins. At the onset of the latest Rotliegend (i.e. Rotliegend II), there was a clear change in terms of the basin geometry (e.g. Rieke e t al. 2001, 2003). Initially, the basin comprised two distinct sub-basins (Havel-Miiritz and West Mecklenburg basins), although it is not clear to what extent these were isolated from one another (Fig. 8). However, c. 2 Ma later there was a clearly unified depositional area across the entire NEGB, with sediment being sourced from basin margin highs, and adjacent orogenic piles, and transported towards the basin centre (McCann 1998a). The distribution pattern within the basin itself broadly resembles the models of closed-basin sedimentation as outlined by Leeder & Gawthorpe (1987) albeit with the significant difference that the NEGB was not a half-graben structure. Strongly increasing thermal subsidence modified the facies architecture as well as the basin geometry through the remainder of the Rotliegend II period, resulting in a broadly smooth topography, a decrease in sediment supply and the expansion of a playa lake environment across almost the entire basin. The NEGB shows no evidence of significant synsedimentary tectonism (e.g. Kossow e t aL 2000), as described for the NW German Basin, the Dutch Basin (e.g. Verdier 1996; Geluk 2005) and the

364

T. McCANNET AL.

Fig. 8. Isopachmaps of the Rotliegend-age (a) Parchim, (b) Mirow, (c) Dethlingen and (d) Hannover formations, NE German Basin (after McCann et al. 2000).

English Basin (e.g. George & Berry 1997). This absence presumably reflects the trend of gradually increasing tectonic activity in these latter areas, which may be related to the initiation of the break-up of Pangaea in Late Permian times. Southern Permian Basin: Polish Basin. The Mid-Polish Trough

(MPT) was continuously connected with the North East German Basin to the NW (Ziegler 1990), but not with the Tethyan domain to the south (Dadlez et al. 1998; Figs 9 and 10). Rotliegend development of the main depocentre of the MPT followed a period of Westphalian wrench tectonics and Early Permian volcanic activity (within the NW MPT and the NE Germany Basin) associated with regional wrenching and generalized crustal destabilization. Immediately overlying the volcanic rocks are late Rotliegend continental sediments deposited under arid conditions (Marek 1988). The sediments are broadly similar to those deposited in the NE German Basin to the west, comprising mainly coarser-grained clastic deposits (confined to the basin margins) with finer-grained clastic deposits occupying the central parts of the basin. Depositional environments were mainly fluvial, aeolian and lacustrine (Karnkowski 1999). The extent of the Variscan orogen in Poland is still uncertain (Pozaryski et al. 1992; Dadlez et al. 1994), mainly because of the lack of reliable data documenting the regional extent and styles of orogenic deformation. However, recent studies of well data have provided new clues to the styles of deformation within the external Variscides and the possible location of the front, including the possibly significant role of strike-slip movements within the frontal part of the orogenic zone and source area for foredeep basin infill (Aleksandrowski et al. 2003; Jaworowski

2002; Mazur et al. 2003). Extension, subsidence and basin development, subsequent to Variscan orogenic activity, was parallel to the SW margin of the East European Craton and the TornquistTeisseyre Zone (Figs 7 and 8). The Polish Basin was formed in the Late Palaeozoic to the north of the Variscan orogen (Dadlez 1997, 2006) and belonged to a series of sedimentary basins that developed around the margins of the East European Craton (Nikishin et al. 1996). The main depocentre of the basin is termed the Mid-Polish Trough (MPT), and this was active during the basin's evolution, beginning in the Permian and extending to the Mesozoic, as a zone of maximum subsidence with almost uninterupted sedimentation (Dadlez 1997; Kutek 2001). Indeed, the axial part of the MPT was also the locus of the Late Cretaceous inversion of the Polish Basin (see Krywiec 2002a and Lamarche et al. 1999 for details). The present base of the Permian in the MPT ranges from 3 to 8 km (Dadlez 1998; Znosko 1999). The Polish Basin developed on highly heterogeneous basement. The MPT developed along the Tornquist-Teisseyre Zone, which marks the boundary between the East European Craton and the Palaeozoic Platform (Guterch et al. 1983; Grad et al. 1999). Integration of industry seismic reflection data with gravity and magnetic data has shown that the location of the NE MPT margin was strongly and directly controlled by the southwestern margin of the East European Craton (Krzywiec & Wybraniec 2003; Krzywiec 2004; Lamarche et al. 2003a). Indeed, the form of the MPT is considered to have partly resulted from rheological boundaries within the Trans-European Suture Zone (Stephenson et al. 2003). Furthermore, the MPT is segmented along strike into (from NW to SE) the peri-Baltic, Pomeranian, Kuiavian and

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lithospheric cooling and contraction following the cessation of the Autunian wrench faulting and magmatism.

Baltic Sea

I <4 km

I

Northern P e r m i a n Basin: Denmark. The Rotliegend sediments of

[.=7.. 4-6km 6-8km

Denmark were deposited in the Danish Central Graben and the Danish-Norwegian Basin, located north of the fragmented Mid-North Sea-Ringk0bing-Fyn High, and form part of the succession deposited within the Northern Permian Basin (Fig. 11; Stemmerik et al. 2000). Modelling suggests that early Permian extension resulted in crustal thinning of up to c. 1.35 and was probably related to a major Late Carboniferous-Early Permian heating and extension phase (Frederiksen et al. 2001). In the Danish North Sea the Rotliegend is subdivided into two formations. The lowermost Karl Formation is areally widespread and is defined to include the syntectonic volcanic, volcaniclastic and sedimentary fill of the Permian half-grabens. The succession is dominated by volcanic rocks of alkaline basaltic composition and volcaniclasti sediments. With the exception of some isolated rhyolite flows (e.g. Horn Graben Rhyolite Member), the succession differs in composition from the more acid volcanic rocks that characterize the Lower Rotliegend of the SPB (Aghabawa 1993). The rhyolitic flows were extruded as silicic lava, which solidified as glass and was subsequently devitrified (Aghabawa 1993). The majority of the volcanic rocks are basic in composition, described as alkaline basalts, hawaiites and mugearites (Aghabawa 1993). Associated sediments are mainly volcaniclastic sandstones and conglomerates that were probably deposited in alluvial and fluvial settings. The overlying Auk Formation represents the postrift succession, comprising sandstones, pebbly sandstones and conglomerates with thin interbeds of silty mudstone, which was probably deposited in aeolian, fluvial and sabkha-lacustrine settings.

>8km ...-" ......"'"

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Fig. 9. Structure on the base Upper Rotliegend (below sea level) with the geographical segments of the Polish Basin. Contour lines are at 500 m interval. The position of the cross-section shown in Figure 10 is indicated (after Stephenson et al. 2003). Malopolska segments (Dadlez 1998), which are related to crustal-scale fracture zones (Krrlikowski & Petecki 1995; Grad et al. 1999; Lamarche et al. 2003b; Lamarche & SchechWenderoth, M. 2003). Along the axial zone of the basin the sedimentary succession is 3 - 7 km thick (Marek & Pajchlowa 1997) including about 1500 m of Zechstein salt (Tarka 1994). Deposition of Zechstein evaporites during the Permian was limited to the SE by the Grrjec Fault, which was also active during the Carboniferous (Zelichowski et al. 1983).

Oslo Rift. The Permo-Carboniferous Oslo Rift extends northwards from the Sorgenfrei-Tornquist Zone and comprises three opposing half-graben segments: the southern Skagerrak Graben, the central Vestfold Graben and the northern Akershus Graben (Figs 12 and 13). The basin formed as a result of dextral transtension, which led to the development of a purely extensional regime in the Oslo Rift system and the culmination of rifting and peak magmatic activity in the Oslo Graben (Olaussen et al. 1994; Heeremans et al. 1996; Torsvik et al. 1998). The onshore part of the present-day Oslo Rift consists of a c. 400 km long and c. 35-65 km wide graben containing large volumes of rift-related extrusive and intrusive rocks and minor amounts of rift-related sedimentary rocks (Neumann et al. 1995, 2002). The driving force behind the initiation of rifting in the Oslo region is difficult to ascertain in terms of active (plume-related) and/or passive (lithospheric stretching) end-members (Kirstein et al. 2002). There is no evidence of crustal doming prior to magmatism (Olaussen 1981). Late Carboniferous pre-rift sediments (i.e. Asker Group) were deposited close to sea level and contain evidence of episodic marine incursions (Olaussen 1981). Rifting in the Oslo region appears to have begun in Late Carboniferous time with the formation of a shallow depression close to

Northern Permian Basin. The Northern Permian Basin (NPB) is an

elongate east-west-oriented basin extending across the Central North Sea from Scotland to Northern Denmark (Fig. 6). The basin is bounded to the south by the fragmented Mid-North SeaRingkCbing-Fyn High. To the east and NE the SorgenfreiTornquist Zone separates it from the stable Fennoscandian Shield. As a result of Mesozoic overprinting, the outlines, geometry and facies patterns of the NPB are less well understood that those of the SPB (Ziegler 1990), largely because of the lack of outcrops and limited drilling (Stemmerik et al. 2000). To the west the basin rests on Devonian clastic deposits whereas to the east it overlies Caledonian basement, on Lower Palaeozoic sediments preserved in the Caledonian foreland, and on Precambrian crystalline rocks of the Fennoscandian Shield (Hospers et al. 1985). Ziegler (1990) has noted that it is assumed that the NPB also evolved in response to

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Fig. 10. Geological cross-section based on well and seismic data through the Polish Trough (location shown in Fig. 9). The approximate location of the Trans-European Suture Zone (TESZ) is indicated (after Stephenson et al. 2003).

366

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Fig. 11. Simplified SE-NW cross-section from the eastern North Sea-RingkCbing-Fyn High to the Danish-Norwegian Basin showing the distribution of Rotliegend sediments. (Note the onlap of the Auk Formation towards the Ringkcbing-Fyn High (after Stemmerik et al. 2000).

Fig. 12. Regional structural map of the Oslo Rift and adjacent areas based on maps by Ramsberg et al. (1977), Falkum & Petersen (1980), Buer (1990) and Ro et al. (1990b). FZ, fracture zone; GS, graben. The location of the seismic profile OG-7 (shown in Fig. 13) is indicated.

POST-VARISCAN BASIN EVOLUTION, EUROPE

367

Fig. 13. Seismic line drawing of profile OG-7 across the SkagerrakGraben (after Ro et al. 1990b). (See Fig. 12 for location.) TWT, two-way travel time. sea level, in which deposited sediments are sealed by the firsterupted lava flows (i.e. B1 basalts). This suggests that rifting began with a period of lithospheric stretching and thinning prior to the onset of the main magmatic phase. Coeval intrusion of sills of intermediate to felsic character is interpreted in terms of compressional activity prior to the main phase of extension (Sundvoll et al. 1992). The initial magmatic phase (involving basalt extrusion) and vertical movement along NNW-SSE- to north-south-oriented faults appears to have been contemporaneous in response to E N E - W S W - to east-west-oriented crustal extension (Heeremans et al. 1996). The main phase of rifting and volcanism involved large displacements (up to 3 km) along some of the basin-bounding faults (i.e. the Oslofjorden fault; Neumann et aI. 2004, and references therein). This phase also coincided with the development of the Oslo Graben. Subsequent caldera formation was accompanied by changes in the magma chemistry and the increasing dominance of intrusive activity (Neumann et al. 1995), associated with a change in the orientation of structural deformation to more N E - S W , N W - S E and west-east. The Permian Oslo Rift was located within a broad intracratonic basin during early Palaeozoic time (Ramberg 1976) containing an up to c. 4 km thick sedimentary sequence of Cambrian to late Silurian age. This succession is locally unconformably overlain by Upper Carboniferous and Lower Permian sedimentary rocks which underlie the Permian lavas (Olaussen 1981). Sedimentary rocks also occur as thin layers (<10 m) between lava flows. Significant thicknesses of fanglomerates are preserved close to the Oslofjorden fault zone in the SE part of the Oslo Graben. Here the clasts mainly comprise lavas and these deposits are indicative of synvolcanic tectonic activity (see Larsen et al. 1978, for more details). The thickness of Permian sediments, however, is secondary to that of the magmatic succession, attaining (in the Skagerrak Graben) a maximum thickness of only c. 1 km (Ro et al. 1990a).

The North Swiss Permocarboniferous Basin (NSPB) is found within the subsurface of northern Switzerland and was first discovered early in the 1980s (e.g. Matter et al. 1987; Fig. 15). Initially, it was called the Constance-Frick Trough (Konstanz-Frick Trog, e.g. Laubscher 1987) and later, as its continuation further to the west was recognized, the NSPB (i.e. Nordschweizer Permokarbon Trog). Formed within crystalline basement, the basin is 1 0 - 1 2 k m wide and filled with >1500 m of continental clastic deposits (see Bliim 1989, for overview). Supply was local (Matter 1987; BRim 1989). Carboniferous deposits have been drilled only at Weiach and appear to be restricted to a narrow, graben-like structure. The NSPB is disrupted by a series of NW-SE-trending faults (Fig. 15). At Weiach 572 m of Carboniferous sediments were drilled, but the basement was not reached (Fig. 16). Microfloral remains are dated as Stephanian (Hochuli 1985). Later radiometric dating of zircons from ash layers confirmed the age (303 Ma in the middle of the drilled Carboniferous deposits and 298 Ma at the top; Schaltegger 1997a). The palaeogeographical position within the Variscides, the varying sediment thickness, and the dominance of crystalline basement clasts derived from local sources suggest a pull-apart origin of the NSPB (e.g. Matter 1987; Bliim 1989). The basin is interpreted as a series of en echelon pull-apart basins (e.g. Diebold et al. 1992) that formed in relation to late Variscan wrench tectonics. Areally, the Permian deposits occupy a significantly wider area than those of the Carboniferous. It would appear that only Early Permian deposits were accumulated, but biostratigraphical dating using palynomorphs (Hochuli 1985) may be subject to some uncertainty because of palaeotopographic effects (e.g. Becq-Giraudon 1993). Except for Weiach, the Permian deposits rest on crystalline basement. Permian deposits were mainly correlated according to lithology (e.g. Bliim 1989) and are all continental, ranging from lacustrine to alluvial fan units.

B a s i n s occurring within the Variscan orogen

Massif Central. The Variscan Massif Central shows the character-

Late Palaeozoic basins in northern Switzerland formed after the main Variscan orogeny (dated in Switzerland as pre-late Westphalian), but in the French Alps, post-orogenic sedimentation commenced in the Late Namurian (Trtimpy 1980). Because all these basins formed within the same time span and in close proximity to one another, they exhibit striking similarities, in that the basins, predominantly grabens or graben-like structures, tend to be small, elongate and mainly trending S W - N E (Fig. 14). Furthermore, Carboniferous deposits are restricted to a narrow trough whereas the basins are filled with continental clastic deposits, mainly eroded from the crystalline basement, and containing volcanic and/or volcaniclastic material. Permo-Carboniferous

Basin,

Switzerland. The

istic tectonometamorphic evolution of classic collisional belts, with significant horizontal thrusting and progressive crustal thickening (Echtler & Malavielle 1990). The Montagne Noire forms the external, southernmost segment of the Massif Central, which may be subdivided into a crystalline metamorphic core (Axial Zone) and a tectonically overlying upper unit of strongly deformed sedimentary rocks (Figs 17 and 18). Partly interfolded sediments of late Vis~an and Namurian age occur along the southeastern border of the thrust belt (Engel et al. 1982). Up succession, the unit is increasingly dominated by synorogenic coarse-grained turbidites and a progressively chaotic set of kilometre-sized olistolithic slabs, which are syntectonic (Engel et al. 1982). Deposition of these units was followed by a period of WestphalianStephanian extension.

368

T. McCANN E T A L .

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POST-VARISCAN BASIN EVOLUTION, EUROPE

369

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In the Massif Central the most notable event marking the Stephanian was the appearance of numerous small limnic coalbearing basins, closely related to a network of regional faults (related to Variscan tectonics) that appeared or were reactivated at this time (Arthaud & Matte 1975). From the end of the Westphalian to Stephanian B north-south compression continued in the Massif Central, but the Westphalian regional ductile shear faults progressively gave way to fracture deformation, responsible for the formation and evolution of the Stephanian limnic basins. The anti-clockwise rotation of north-south-

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370

T. McCANN ETAL.

oriented compression to an east-west direction was responsible for the evolution of these regional strike-slip faults and consequently of the associated basins. At the end of the Stephanian, the final stage of east-west-oriented compression interrupted sedimentation and caused intense deformation of the basins located on the approximately north-south-striking faults. In the Pyrenees, the North Pyrenean Fault, which borders the axial zone, probably underwent dextral strike-slip with horizontal displacement estimated at 150 km (Arthaud & Matte 1975). Subsequent to the period of Stephanian compression at the close of the Variscan orogeny, a new period was initiated during that intracontinental basins formed, the thickness and extent of whose deposits differed from those that accumulated during the Stephanian. The Stephanian compressive basins of the Massif Central are long and narrow with varying thicknesses of sediment, whereas the extensional Permian basins contain successions up to several thousands of metres thick and are much more extensive, occupying the sites of the future large Mesozoic basins. In France the Permian includes only continental deposits. The succession is generally subdivided into Autunian and 'Saxonian' strata (both of Early Permian age) and Thuringian deposits (considered to be more or less equivalent to the Late Permian) (Cassinis et al. 1995). The typical Autunian was defined in the Autun Basin, located to the north of the Massif Central. These units comprise lacustrine calcareous and bituminous shales, coarse fluvial deposits and some volcanic ash, which overlie folded Stephanian beds (Cassinis et al. 1995). The many Stephanian-Autunian basins in the Massif Central contain coal measures and alluvial, fluviatile and lacustrine clastic sediments deposited in an active tectonic environment; magmatic rocks are present in only small volumes (Legrand et al. 1994; Djarar et al. 1996; Allemand et al. 1997). Saxonian deposits include those that overlie the Autunian, and these units are generally separated by an unconformity attributed to a period of intra-Permian (i.e. 'Saalian') deformation. The expression of this varies from changes in sedimentation to an actual angular unconformity (Cassinis et al. 1995). Permian-age sediments were deposited on a variety of strata that were deformed and metamorphosed during the Variscan orogeny, and in many basins sedimentation continued into late Permian times (Ch~teauneuf & Farjanel 1989). Sedimentation was controlled by synsedimentary basin margin faults, and faulting caused considerable palaeorelief and the formation of basement horsts that were important sources of sediment, especially alluvial cones. Some of the larger basins began as smaller sub-basins with independent drainage systems that amalgamated during accelerated subsidence (e.g. the Crvennes Basin, Djarar et al. 1996). Those basins located along approximately east-trending faults (e.g. the Saint-Etienne Basin) tend to have pull-apart geometries that are related to strike-slip movements along the basin margin faults. Those basins located along approximately north-trending faults tend to be half-grabens, some of which are strongly asymmetrical (Faure 1995; Mattauer & Matte 1998). Furthermore, the extent of soft-sediment deformation from the Saint-Affrique and Crvennes basins testifies to the syndepositional, extensional to transtensional character of the basin margin faults (Legrand et al. 1994; Djarar et al. 1996). In the southern Massif Central distinct sequences are present within the Autunian deposits. For example, in the Lod~ve Basin (Laversanne 1978) sedimentary sequences between 8 and 15 m thick occur within a unit that is c. 800 m in thickness (Fig. 19). These sequences, which are continental in origin, comprise fluviatile sandstones in the lower part, sandstones and bituminous dolomites of palustrine or lagoonal origin in the middle part, and calcareous silty floodplain mudstones and siltstones in the upper part (which also contains evaporitic precipitates). Each basin developed in a general tensional context. In many cases there is no appreciable lithological change through the entire Upper Permian sequence and, because of a lack of data, many workers

prefer to classify such rocks as Saxonian-Thuringian units. In the Lod~ve Basin the Upper Group is present.

S a a r - N a h e Basin. This Permo-Carboniferous basin extends from

SW Germany into France and is filled with exclusively continental sediments (Fig. 20). The basin has a half-graben geometry, being bordered to the north by the south-dipping Hunsrtick Boundary Fault (HBF), which, according to Henk (1993) is a detachment soling out at mid-crustal levels (at a depth of c. 16 km). Transtensional subsidence of the partly inverted Saar-Nahe Basin, which contains up to 5.6 km of Permo-Carboniferous clastic deposits, accounts for a stretching factor of >1.36. Contemporaneous extrusive activity reflects destabilization of its lithospheric system (Ziegler et al. 2006). The evolution of the Saar-Nahe Basin is closely related to the complex kinematics of the HBF, which is part of a prominent suture zone separating two of the main tectonostratigraphic units of the Variscan fold belt: the Rheno-Hercynian and the Saxo-Thuringian zones. The sedimentary infill of the Saar-Nahe Basin is dominated by continental clastic deposits (Sch~ifer 1989; Sch~ifer & Korsch 1998) with a significant thickness of contemporaneous volcanic rocks. Basin initiation occurred in the latest Namurian or possibly earliest Westphalian, and the absence of older Namurian sediments suggests that this was an elevated area from the late Vis~an to the late Namurian. This period of non-deposition reflects Variscan compression, uplift, exhumation and cooling, as indicated by 320-335 Ma 4~ cooling ages from Mid-German Crystalline Rise (MGCR) basement rocks to the south (i.e. Odenwald and Erzgebirge, Werner & Lippolt 2000; Schubert & Lippolt 2000). There is an angular unconformity between the Westphalian D and Stephanian A, indicative of a reorganization of fault kinematics in the area. Major uplift in the southeastern part of the basin elevated the MGCR basement to provide a new sediment source (Korsch & Sch~ifer 1995). The volcanic rocks are related to the crustal stretching that commenced in the earliest Westphalian, where a regional, rather than local, mechanism was responsible. The cause of this was probably dextral translation between Laurussia and Gondwana, which was about to collide with eastern to southeastern North America. Rotation and translation led to the reactivation of old lineaments (such as the northern boundary of the Saar-Nahe Basin, the Htinsrtick Fault) and the establishment of large-scale, new fault systems in Europe (e.g. Arthaud & Matte 1977; Ziegler 1990). Magmatism led to a thermal anomaly, leading to a subsequent phase of thermal relaxation and crustal re-equilibration. The west-east extensional stress regime that dominated during the post-Westphalian basin formation (Stollhofen 1998) was oblique to the pre-existing fault pattern; the common slip direction was maintained by transtensional strike-slip movements on the transfer faults and oblique-slip motion on the normal faults (Stollhofen et al. 1999). An angular unconformity, indicative of a reorganization of fault kinematics in the area, underlies the synrift megasequence. The Stephanian prevolcanic synrift sequence has a thickness of 3.8-4.7 km and was deposited over a period of c. 14 Ma. It comprises lacustrine, fluvio-deltaic and fluvial sediments with minor limestones, coals and pyroclastic fallout deposits (these last were derived from sources outside of the basin; Stollhofen et al. 1999). This is overlain by the Lower Permian volcanic synrift sequence (1.1km thick, deposited over c. 4 Ma); during this phase widespread bimodal calc-alkaline magmatism occurred with subvolcanic intrusion of rhyoliticdacitic domes and basaltic to andesitic sills and dykes (von Seckendorff et al. 2004b). Magma generation can be attributed to underplating and intrusion of pulses of mantle-derived melts into the crust inducing partial melting, or perhaps a distal subduction zone (Schmidberger & Hegner 1999). The extrusive rocks are interbedded with fluvial sediments and minor lacustrine units (Stollhofen 1994).

POST-VARISCAN BASIN EVOLUTION, EUROPE

371

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VARISCAN BASEMENT MYLONITES AND FAULT REVERSFAULT E Fig. 19. Cross-sections through Permian basins in Central France: Lodbve Basin, south of the Massif Central (after Cassinis et al. 1995), Autun Basin and Blanzy-Le Creusot Basin (after Bibs et al. 1989). (Locations are shown in Fig. 17.)

Fig. 20. Geological interpretation of DEKORP 1C and 9N showing the half-graben form of the Saar-Nahe Basin, Germany (after Henk 1993).

372

T. McCANN E T A L .

Iberian basins. The Iberian microplate was affected during the

(Arche & Ldpez-Gdmez 1996; Gonzfilez-Casado et al. 1996), although transcurrent deformation with a N E - S W orientation continued in the SW of the Iberian Peninsula (Herraiz et al. 1996). There are five regions in Spain where the Permian is well known: the Cantabrian Mountains, the Pyrenees, the Central System, the southern margin of Iberia and the Iberian Range (Cassinis et al. 1995). The succession may be broadly subdivided into a Lower and an Upper Group. The Lower Group, associated with the formation of small basins controlled by strike-slip faults, comprises one or more tectonosedimentary units that unconformably overlie the Stephanian or older Palaeozoic rocks. Its thickness ranges between 200 and 2000 m. Widespread andesitic and rhyolitic volcanism of calc-alkaline type is associated with this group (Lago et aI. 2004). The overlying Upper Group sediments were deposited in an extensional cycle and are areally more widespread. There is a marked unconformity between the two groups. The Permo-Carboniferous succession of the Central Pyrenees has been considered to have originated from strike-slip dynamics that developed during a compressional episode at the end of the Variscan orogeny as established from facies analysis (e.g. Marti 1986, 1991; Besly & Collinson 1991) and from regional palinspastic reconstructions (e.g. Mufioz 1992; Casas et al. 1989). In the Pyrenees, the oldest sediments comprise breccias, sandstones and coal beds, and contain Stephanian flora (Fig. 22). These are overlain by a widespread volcanic unit comprising andesitic pyroclastic deposits and volcaniclastic rocks (see Cassinis et al. 1995; Lago et al. 2004, for details), which is overlain by conglomerates and sandstones that contain a Stephanian-Autunian flora. In the Anayet Massif, for example, there is an extensive

Permian by the final stages of Variscan activity and by the early Atlantic rifting (Fig. 21). These tectonic disturbances led to the development of intracratonic basins. Two phases of Variscan deformation have been described in northern Spain, an early phase consisting of two east-west-oriented compressional events in the period from the Namurian to the Stephanian that resulted in arc-parallel folds and thrusts, and a later northsouth-oriented phase (Sakmarian-Early Permian) that marked the final stage of Variscan deformation in northern Iberia and resulted in plunging fold axes of the early arc-parallel folds (Weil et al. 2000). Subsequent rifting episodes (e.g. Early Permian-Triassic) resulted in the formation of basins with orientations ranging from N E - S W to east-west (Jabaloy et al. 2002). Additionally, magmatic events have been recorded. Well-dated volcanic rocks and coeval granites crop out in the Pyrenees, the Iberian Range and the central part of lberia (Mufioz et al. 1986; Sopefia et al. 1988). The basins are filled with red beds interdigitated with volcanic and volcaniclastic rocks. Thicknesses are variable and can be as much as 2000 m. These units appear as isolated outcrops of generally small extent because they were deposited in complex graben systems. Stephanian deformation involved N E SW compression and horizontal extension orthogonal to the compression direction. In the Early Permian there was major volcanic activity. The fracture patterns associated with this activity in the central part of Iberia indicate extensional deformation in this region (de Vicente et al. 1986; Gonzfilez-Casado et al. 1996; van Wees et al. 1998) and probably also in the Pyrenees. Structural data are scarce, but appear to indicate that the central area was characterized by a NNE-SSW- to NNW-SSE-oriented extension

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POST-VARISCAN BASIN EVOLUTION, EUROPE

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sedimentary record of Stephanian-Permian-age deposits. Here, the succession is composed of four lithological units representing a stratigraphic transition from Stephanian to Permian. In terms of the sediments, these units represent a transition from humid to arid climatic conditions (with fluvial sediments) (Gisbert 1983). Contemporaneous with sedimentation, there were several intrusive and extrusive magmatic episodes favoured by the transtensional tectonic regime, which was also responsible for the development of small listric-fault-bounded basins containing up to 3 km of volcanoclastic and detrital rocks (Bixel et al. 1996). These mixed magmatic-sedimentary successions were deposited in isolated basins (five of which have thus far been identified; Martf 1996). The only outcrops of marine deposits of Stephanian and Permian age to be found in Western Europe (with the exception of the Eastern Alps and Sicily) occur within the Iberian Massif in western Spain and Portugal (Martfnez Garcfa 1990). Compressional tectonics characterizes the evolution of the Stephanian basins, whereas the Permian (and Stephanian C) shows a tensional environment that reveals plate-tectonic conditions suggestive of failed rifting. Two units, corresponding to different basins, have been distinguished in the Permian succession of the Cantabrian and Palential zones by Martfnez Garcfa (1983). The lowermost of these comprises alternating clastic sediments, tufts, volcanic agglomerates, lava flows and shallow marine limestones. Overlying these are volcanic tufts and ash up to 900 m thick and with an increasingly acid character. Thick conglomerate units attest to syndepositional tectonic activity. The unconformably

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Fig. 22. Upper Carboniferous and Permian of the Pyrenees and Iberian Ranges in Spain (after Cassinis et al. 1995). AU, Autunian macro- or microflora; ST, Stephanian macro- or microflora; TH, Thufingian microflora.

overlying succession comprises predominantly continental clastic sediments. The Permian succession of the present-day Iberian Range, eastern Spain, is very well exposed (Figs 22 and 23). These sediments are of continental origin and filled isolated, small basins that were initiated during latest Carboniferous-early Permian times (Sopefia et al. 1988) and that evolved into a single basin (i.e. the Iberian Basin) during late Permian times. This period was one of tectonic readjustment of plates by transtensional faulting in the Chedabucto-Gibraltar and Bay of Biscay areas and the development of a conjugate wrench zone, originating later in the Iberian Basin, following an ancient suture (NESE-oriented) running across the microplate (Arthaud & Matte 1977; Salas & Casas 1993; L6pez-G6mez et al. 2002). During early Permian (Autunian) times, a series of intermontane basins were filled in by alluvial fans, slope breccias and lacustrine deposits associated in their lower part with volcaniclastic rocks of calc-alkaline affinities (Lago et al. 1992). Lacustrine deposits at the top of the volcaniclastic succession record a change from freshwater to saline lakes, indicating a progressive aridity in Iberia. These early Permian sediments were mostly unconformably deposited on rocks of early Palaeozoic age, but also upon late Carboniferous (Stephanian B - C ) sediments. These comprise sandstones and coal-beating successions that filled the Henarejos intermontane small basin contemporaneously with those of the Cantabrian and Palential zones and are related to the final, extensional collapse of the Variscan orogeny, which also resulted in monzogranitic magmatism and uplift in Central Iberia.

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T. McCANN E T A L .

During the late Permian, red bed successions up to 300 m thick accumulated in grabens and half-grabens that formed during a period of widespread extension within the Iberian Basin (L6pez-Grmez et al. 2002). The overall development shows a fining-upwards trend in most of the basins, beginning with transverse alluvial fan deposits and followed by longitudinal braided river system deposits. There was no volcanic activity during this period in the Iberian Basin, where basin boundary faults have shallow listric geometries at depths of only 13-14 km. The outcrops in the Iberian Range are the most widespread and the best studied. The Molina de Aragon series (Ramos 1979; Ramos & Doubinger 1979) is the most characteristic and the one that is best correlated with the Permian of Central Europe (Virgili et al. 1976, 1983; Fig. 23). The lowermost units

,,,

8

PALMACES - RIBA

Geophysical investigations Subsequent to the European Geotraverse (EGT; Blundell et al. 1992), the increase in the number and quality of seismic data

MOLINA DE ARAGON F

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unconformably overlie Lower Palaeozoic rocks and consist of breccias, volcano-sedimentary rocks and thin volcanic intercalations. These are overlain by 300 m of lacustrine and fluvial sandstones and black shales. The succession ends with lacustrine, siliceous dolomites unconformably overlain by the Upper Group (Permian). The onset of sedimentation of the Upper Group is marked by an important and widespread unconformity reflecting an important change in tectonic evolution.

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Fig. 23. The Pfilmaces-Riba and Molina de Aragon sections from the Castilian branch of the Iberian Ranges. AS, Arandilla sandstones; HGC, Hoz de Gallo sandstone; MB, Monstsoro beds; PB, Prados beds; PLC, Pfilmaceslower conglomerates; PMS, P~ilmacesmudstone and sandstone; PS, Pfilmacessandstone; PUC, Pfilmacesupper conglomerates;RGS, Rillo de Gallo sandstone; RSC, Riba de Santiuste conglomerates;RSS, Riba de Santiuste sandstones; VSC, volcano-sedimentarycomplex (after L6pez-G6mez et al. 2002).

POST-VARISCAN BASIN EVOLUTION, EUROPE has greatly improved our understanding of the nature of the crust that underlies many of these Permo-Carboniferous-age basins, allowing us to ascertain the nature and relevance of any preexisting structures (e.g. the Tornquist-Teisseyre Zone and the Sorgenfrei-Tornquist Zone) to the formation of the post-Variscan basins (e.g. Polish Trough and Oslo Rift, respectively). The following section will outline some of the main advances and results that have helped to improve our understanding of the evolution of these basins. Germany

A series of DEKORP profiles were shot in the 1980s to image the western margins of the Variscan deformation front close to the German border with Belgium (Meissner & Bortfeld 1990). Of greater interest, however, was the series of profiles shot in the 1990s, which imaged not only the North German Basin but also the major structures to the north and south of it. Prior to obtaining the deep seismic profiles, the precise nature of the North German Basin (i.e. whether it was an extensional basin or not) had been much discussed in the literature. As is clearly shown on the DEKORP 9601 profile (DEKORP-BASIN Research Group 1999), the basin is intracratonic, and was initiated above a region that had undergone significant rupturing as a result of the wrench tectonics at the Permo-Carboniferous transition. However, there is little evidence of these movements on the seismic profiles. Instead, we have a clear image of the Moho (continuous across the entire profile) and the northern (where Avalonia is obducted onto Baltica) and southern (with clear evidence of Variscan thrusting) margins (Fig. 24). The depth to the Moho is 30 km beneath the entire basin, except at the margins, where it extends to 35 km. More detailed profiles are required to reveal any evidence of Stephanian-Rotliegend tectonic activity. However, tectonic activity in the form of continuous transpression is evidenced by the diverse angular unconformities visible between the Westphalian and Permian deposits on 3D seismic profiles and in drill core (McCann 1998a). In the Saar-Nahe Basin two deep seismic profiles (DEKORP 1C and 9 N) crossed parts of the basin (DEKORP Research Group 1991; Korsch & Sch~ifer 1991). The southernmost part of the basin was also imaged by the ECORS-DEKORP 9S profile (Brunet al. 1991). These profiles reveal the asymmetric geometry of the basin bounded by the South Hunsrtick Fault, which has been interpreted as having either a subvertical (Korsch & Sch~ifer 1995) or listric (Henk 1993) geometry, although the latter interpretation is the favoured one. The upper crust beneath the basin shows a segmentation into three distinctively different reflectivity patterns. The uppermost highly reflective package represents the pre-rift sediments and Permo-Carboniferous basin fill. The underlying wedge-shaped unit lacks major reflectors and is interpreted as crystalline basement rocks of the northern Saxo-Thuringian Zone (i.e. Mid-German Crystalline Rise). Beneath is a thick, highly reflective zone, which may represent remnants of sediments and oceanic crust from the

375

Rheno-Hercynian Ocean (Behr et aL 1984; Franke & Oncken 1990), or an anastomosing pattern of shear zones and duplex structures (DEKORP Research Group 1991). Heat-flow modelling has been carried out in a number of areas, generally in conjunction with other forms of modelling. In NE Germany, for example, Ondrak et al. (1997) integrated the structural models of Scheck (1997) to produce a regional thermal model that allowed the determination of temperature distributions and a depth-dependent estimation of the local heat-flow conditions. Temperature gradients within the model matched the regional trends of heat-flow distributions, although the pattern is more complex. There is a clear relationship between temperature and the Zechstein salt layer, where high temperatures are related to salt margins and regions where sediments with low thermal conductivities cause local elevation of the isotherms. Poland

The crust of central and northern Poland has been extensively studied by deep refraction and wide-angle reflection seismic studies. Older profiles have recently been reprocessed and reinterpreted (see Grad et al. 2003, for details) and integrated with the new high-quality LT7 profile (Guterch et al. 1994), the TTZ profile (Grad et al. 1999) and profiles from the POLONAISE 97 experiment (Guterch et al. 1999; Jensen et al. 1999; Janik et al. 2002). All of these seismic data have provided new information on the deep structure of the transitional area between the East European Craton and the Palaeozoic Platform in central and northern Poland. The depth to the Moho in this region is 32-39 km beneath the two-layered Palaeozoic Platform and 43-45 km beneath the three-layered crystalline crust of the East European (Precambrian) Craton. The Trans-European Suture Zone (TESZ) located between these two main crustal domains is characterized by the presence of thick (up to 20 km) complexes of low-velocity, sedimentary, metamorphic or volcanic rocks, and by a lower crust characterized by high velocities. In this region, the Moho is located at intermediate depths in the NW (30-33 km) and deepens to the SE. According to some interpretations, the lower crust within the TESZ area in central and northern Poland represents the attenuated Baltica margin underthrust towards the SW beneath the Avalonian accretionary wedge (Grad et al. 2002). The results of gravity modelling based on deep seismic refraction data suggest that the crustal structure of the Mid-Polish Trough (MPT), especially of its pre-Zechstein substratum, is more complex than suggested by deep refraction data (Krrlikowski & Petecki 1997, 2002; Petecki 2002). The upper mantle beneath the TESZ is dense, and within the upper crust a highdensity body was also identified, as well as a complex transition zone between the crust and the upper mantle. Long-wavelength gravity anomalies are associated with lateral density variations within the upper mantle and lower crust (Krrlikowski & Petecki 2002; Petecki 2002). Results of magnetic modelling suggest that the average depth of magnetic sources within NW Poland is of the order of 18 km, and could be correlated with the crystalline

Fig. 24. Interpreted line drawing of BASIN 9601 profile and its offshore extensions PQ2-009.1 and PQ2-005 showing the main tectonic and stratigraphic features.

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T. McCANNET AL.

basement as evidenced by seismic data (Petecki 2002). Upper crust fault zones revealed by gravity and magnetic data present within the pre-Zechstein MPT basement played a significant role during the Mesozoic evolution of the MPT (Dadlez 1997) as evidenced by regional interpretation of seismic reflection data (Krzywiec & Wybraniec 2003) and structural modelling (Lamarche et al. 2003a). Geothermal models in Poland have been more concerned with examining the broad structural framework of the region, rather than concentrating on the basin infill succession. For example, Majorowicz et al. (2003) have shown that there is a sharp change in heat-flow data between the East European Craton and the Palaeozoic Platform. Numerical modelling of the crustal temperatures along several deep seismic profiles suggests extensive crustal-mantle warming within the zone located between the Sudete Mountains and the TESZ (Grad et al. 2003; Majorowicz et al. 2003). This anomalous zone coincides with the location of the Dolsk Fault and the Variscan Deformation Front. High heat flow for the Palaeozoic Platform and related high temperatures of the crust coincide with the reduced crustal thickness, whereas the low heat flow of the East European Craton coincides with a higher crustal thickness. Modelling results also suggest that high mantle heat flow is required within the high heat-flow zone located within the Palaeozoic Platform (characterized by the 100 km thermal lithospheric thickness) whereas cold crust and cold mantle are typical of the East European Craton (characterized by the > 200 km thermal lithospheric thickness; Grad et al. 2003; Majorowicz et al. 2003). Thermal modelling of the Polish Basin performed by Karnkowski (1999) suggested that the Rotliegend volcanism began with high geothermal anomalies in the western part of the basin. The anomalies are characterized by higher values ( 1 0 0 - 1 5 0 m W m -2) during the Late Permian-Early Triassic interval, in relation to rifting in the Polish Basin. The Late Triassic and Jurassic were a time of cooling, until the break-point in thermal evolution of the Polish Basin at the Jurassic-Cretaceous boundary as a result of uplift and erosion, after which the heat inflow decreased.

Norway

The Oslo Graben and surrounding areas have been the subject of a number of geophysical investigations (e.g. Ro et al. 1990a,b; Kinck et al. 1991), which have revealed that there is a marked crustal thinning beneath the graben, with the amount of thinning increasing southwards. The depth to the Moho has been estimated to range from 28 to 35 km beneath the Oslo rift system (Cassell et al. 1983; Thybo 1997). Gravity data indicate the presence of a dense 90 km wide body towards the base of the lower crust, which has been interpreted as representing cumulates and gabbroic rocks in a deep crustal magma chamber (Neumann et al. 1992). Furthermore, gravity and seismic data imply a different crustal structure along the rift from in the adjacent Precambrian terrane (e.g. Wessel & Husebye 1987; Kinck et al. 1991). However, more recent data have questioned these interpretations of gravity data in the Oslo region (Ebbing et al. 2006). These data and modelling results show that the high-velocity layer present at the base of the Oslo Graben is similar to the typical high-velocity layers that are commonly found at the base of Proterozoic crusts (Durrheim & Mooney 1991), suggesting that the idea of underplating beneath the Oslo Graben is a false one. It has been noted that although Permian mafic intrusions in the crust may be present, they do not reside in a hypothetical underplate, but most probably in the middle crust (Ebbing et al. 2006). Seismic refraction or wide-angle-reflection data have been used to model variations in crustal thickness and structure in the Oslo rift and adjacent parts of the Precambrian shield (e.g. Bungum et al. 1971; Cassell et al. 1983; Ro et al. 1990a,b; Kinck et al.

1991). The subsequent Moho map shows that the crust is thicker east of the Oslo Graben than west of it (Kinck et al. 1991; Thybo 1997). The refraction or wide-angle reflection data of greatest relevance to the crustal structure of the centre of the Oslo Graben segment include the profiles of Tryti & Sellevoll (1977) (see Neumann et al. 1995, for details). Three distinct crustal layers were noted, including a high-velocity (7.1 km s - i ) lower crustal layer that had been interpreted as resulting from magmatic underplating (see above) and/or a zone of magmatic cumulates and residues (Neumann et al. 1995). Teleseismic studies carried out along the 60th parallel (i.e. parallel to the modelled section AA') suggest that deepening of the base lithosphere occurs abruptly from west to east in the Oslo region (Babuska et al. 1988). These finding are confirmed by a more recent teleseismic study (Plomerovfi et al. 2001).

Modelling post-Variscan basin development Although numerical methods have been applied to tectonic problems for over 30 years, relatively few have been concerned with modelling the post-Variscan (i.e. Permian) basins of Europe. Most Permian basins are buried at considerable depths by Mesozoic and Cenozoic sediments, and their structure has been obscured by younger tectonic events. As a result, the acquisition of accurate data for the modelling is very often difficult. Nevertheless, the increasing collection of subsurface observations (i.e. well and seismic reflection data) has allowed for the modelling of the post-Permian subsidence history in some areas (i.e. Paris Basin, North Sea, NE German Basin, Danish Basin). In other areas, which were uplifted in post-Permian times, direct observation of the Permian basins (i.e. the Oslo Graben) allows relevant data for thermo-mechanical modelling including rheology and decompression melting to be acquired.

Models based on sediment succession-basin

analysis

Models based on subsidence analyses have been the most widely applied to reconstruct the pre-rift configuration and the rifting history of Permian basins (e.g. Van Wees et al. 1998, 2000; Frederiksen et al. 2001). These models are based upon the classical model of McKenzie (1978), and its numerous derivatives, in which rapid synrift subsidence is followed by a more extended phase of thermal subsidence. The pattern of basin subsidence is strongly dependent on the initial configuration of the lithosphere and the amount of stretching. One of the advantages of this method is that it provides information on a previous rift event from the analysis of post-rift sediments. Thus, direct observation of the synrift sedimentary infill can be bypassed (see Allen & Allen 2005, for further discussion on this topic). A variety of basins both within and to the north of the broad zone of Variscan-age deformation have been studied with this method. Basins within the deformation zone include the Paris Basin (e.g. Brunet & Le Pichon 1982; Prijac et al. 2000), the Mid-Polish Trough (e.g. Dadlez et al. 1995), the Iberian Basin (e.g. van Wees et al. 1998), and the Southern Permian Basin (e.g. van Wees et al. 2000); those to the north of it include the Danish Basin (e.g. S0rensen 1986; Frederiksen et al. 2001), the offshore arm of the Oslo Rift (e.g. Pedersen et al. 1991), and the Northern North Sea (e.g. Odinsen et al. 2000). In detail, the results differ somewhat from one basin to another, reflecting the complexity of the structure of the European lithosphere in Permian times and the interplay between different subsidence mechanisms. For example, Permian rifting and post-Permian subsidence in the Paris Basin can be explained by the collapse of the Variscan mountain chain and the slow decay of the associated thermal anomaly (Brunet & Le Pichon 1982; Prijac et al. 2000). In the Variscan foreland such a mechanism cannot be invoked.

POST-VARISCANBASIN EVOLUTION, EUROPE In contrast, Dadlez et al. (1995) and van Wees et al. (2000) argued for strong destabilization of the lithosphere by late Variscan wrenching involving deep fracturing and decompression melting of the underlying mantle. This interpretation explains the large thicknesses of Permo-Carboniferous magmatic rocks and post-rift sediments that accumulated in the absence of major normal faulting of the crust. Modelling results concerning the precise age and timing of the Permian rifting event differ from one study to another. This reflects differences in data coverage between the various studied areas, and differences in the methods used, but also probably different responses of the lithosphere in terms of its structure and past history. Whatever the precise mechanisms associated with the Permian event, it is clear that subsidence modelling studies agree on two major points: (1) latest Carboniferous-early Permian tiffing (290-305 Ma) was a widespread and dramatic event in Europe; (2) the thermal signature of the Permian rifting was a significant control on the subsequent Mesozoic and Cenozoic evolution of the European lithosphere. In particular, the second conclusion has serious implications for oil exploration in the North Sea area (SOrensen 1986; Pedersen et al. 1991; Odinsen et al. 2000) in terms of reassessing the reconstructed thermal history of the region and, by extension, predictions concerning levels of organic maturation. Ziegler et al. (2006) have noted that subsidence curve modelling suggests that there was a period of Permo-Carboniferous 'stretching' from 300 to 280 Ma. This involved decoupled crustal extension and mantle-lithosphere attenuation. Such an assumption is compatible with the concept that during the Permo-Carboniferous re-equilibration of the crust-mantle boundary crustal extension played a significant, but local, role. This can be concluded from the fact that some important Permo-Carboniferous troughs (e.g. Massif Central, Bohemian Massif) do not coincide with major Late Permian depocentres, and depocentres such as the Southern Permian Basin are not underlain by major PermoCarboniferous basins (Ziegler 1990; Ziegler et al. 2006). This would suggest that during the Permo-Carboniferous tectonomagmatic cycle uniform and/or depth-dependent lithospheric extension was, on a regional scale, only a contributing factor and not the dominant mechanism of crustal and mantle-lithosphere thinning. Mechanical stretching of the lithosphere played a subordinate role whereas thermal thinning of the mantle-lithosphere and magmatic and erosional thinning of the crust dominated, providing the principal driving mechanism for the Late Permian and later subsidence of intracratonic basins.

Thermo-mechanical models

Rheological numerical models investigate the way in which materials deform in response to stresses. Rocks are considered to deform in different ways (i.e. by elasticity, plasticity or viscosity) depending on the duration of the load, the petrological composition, the temperature, and the confining pressure. The rocks can display a very complex rheology, as it is possible that various modes of deformation occur at the same time. An additional control is the depth to the brittle-ductile transitions, which can be present within each lithospheric layer and are strongly controlled by the geotherm. Thus, rheological modelling implicitly involves thermal calculations and can include routines to determine possible melt volumes. Application of these theological models is not straightforward, as they require various, and relatively accurate, datasets. The Oslo Rift is one of the few Permian basins in Europe where this requirement is met and, consequently, it has been the primary target for 2D numerical modelling (Ro & Faleide 1992; Pedersen & van der Beck 1994; Pascal & Cloetingh 2002; Pascal et al. 2004). The Oslo Rift presents the paradox of showing little extension (i.e. /3 < 1.3) in association with huge volumes of synrift magmatic rocks (i.e. > 100 000 km 3, Neumann et al. 2004). A

377

possible explanation for this is the presence of an underlying thermal anomaly (i.e. a mantle plume) below the European lithosphere in Permian times, which in turn could also explain the observed widespread rifting and magmatism. However, for various reasons (see discussion by Pedersen & van der Beek 1994; Pascal et al. 2004) this hypothesis is questionable. Melting modelling of the Oslo Graben was carried out by Ro & Faleide (1992) and Pedersen & van der Beck (1994). From a model in which crust and lithospheric mantle are equally stretched, Ro & Faleide (1992) argued for the mantle plume hypothesis. In contrast, Pedersen & van der Beck (1994) showed that the volumes of melts of the Oslo Rift can be accounted for by differential stretching between crust and lithospheric mantle (i.e. the lithospheric mantle is more stretched than the crust) and reduced melting temperatures for the mantle owing to the presence of volatiles (i.e. water and CO2). Based on geophysical observations, Pascal & Cloetingh (2002) proposed a rheological model that considers lithosphere thickness heterogeneities in the Oslo region (Fig. 25). Their modelling shows that such heterogeneities could have resulted in strong localization of deformation in the Oslo Rift. A similar study by Pascal et al. (2004) showed that the introduction of lithosphere thickness contrasts in the models results in pronounced differential stretching between crust and mantle lithosphere, which, in turn, leads to decompression melting of the mantle over relatively short time periods subsequent to the onset of rifting. In summary, the models of Pascal & Cloetingh (2002) and Pascal et al. (2004), in which the mechanical behaviour of the rocks and a more realistic configuration for the lithosphere are included, complement the study by Pedersen & van tier Beek (1994). Although modelling results are very often more suggestive than firmly conclusive and need to be compared with nature, whenever it is possible, they appear here to go against a plume hypothesis for the Permian rift event in Europe. Henk (1999) used rheological modelling of Permian basins of Europe to examine the post-convergence evolution of the region. The purpose of his modelling approach was to explore whether the Variscides simply collapsed following the end of the orogenesis, thus leading to Permian tiffing, or whether the region was also influenced by far-field extension. Various 2D models were presented by Henk (1999), and he concluded that far-field extension superimposed on gravity stresses are required to overcome the strength of the post-Variscan lithosphere. Along the LT-7 deep seismic refraction profile in the NW Polish Basin (Guterch et al. 1994), 1D rheological modelling using a simplified petrological model of lithospheric layering was completed. The results suggest that the lithosphere, except for the East European Craton (EEC), is mechanically decoupled, and that the upper crust is separated from the upper mantle by extremely weak and ductile middle and lower crustal layers (c. 20 km thick). Only within the Tornquist-Teisseyre Zone and the EEC can the lower crust remain strong. The lithosphere of the EEC is probably entirely coupled except for the edge of the craton, where, with the low strain rates, mechanical discontinuity may occur at the middle-lower crust or lower crust-mantle boundaries. Laterally, the cumulative strength of the lithosphere changes by more than an order of magnitude (Jarosinski et al. 2002; Grad et al. 2003).

Tectonic a n d structural m o d e l s

Based on geological and geophysical data, tectonic and structural modelling of an object usually summarizes and tests the admissibility of combined information measured and observed in the field and laboratory. Balanced sections thus provide geologically reasonable constraints (Dahlstrom 1969), a concept that has been widely used in the hydrocarbon industry (Bally et al. 1966; Rowan & Kligfield 1989), but also is used to reveal the nature

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Fig. 25. Numerical modelling of the Oslo Rift, involving rock rheology and heterogeneity in lithosphere thickness (after Pascal et al. 2004). The thickness of the lithosphere in the left and right parts of the model is initially equal to 125 km and 180 kin, respectively. The modelled line is 500 km long at t -----0 Ma. The model is stretched using a velocity of 1.6 cm a- 1. The upper panel presents the horizontal strain distributions (i.e. exx) 1 Ma and 9 Ma after rift initiation. (Note the strong strain localization at the middle of the model and the Earth surface depression simulating basin formation.) The lower panel presents the thermal evolution (i.e. isotherms) of the lithosphere. Note the rise at t ~ 9 Ma of hot mantle rocks below the area that is depressed at the surface. The finite-element grid used for the computations is also shown. U.C., upper crust; L.C., lower crust; L.M., lithospheric mantle.

of tectonic processes and kinematic evolution in the area of interest (e.g. Oncken 1989). In the Central European Variscides, extensive studies were carried out to determine the pre-Variscan and Variscan evolution (see summary by Franke et al. 2000), but only few comprise 2D and 3D geometric and tectonic modelling of late, Variscan (e.g. Plesch & Oncken 1999; Oncken et al. 2000, and references therein; Schtifer et al. 2000) or even postVariscan development (Tanner et al. 1998). In the NE German Basin, the only palinspastic reconstructions available are by Kossow & Krawczyk (2002), based on results from the BASIN96 and commercial seismic surveys (Krawczyk et al. 1999; Kossow et al. 2000). The flexural cantilever model (see Kusznir et al. 1991, for model details) was also applied for forward modelling of the initial phase of NEGB formation in combination with detailed analysis of core material (Rieke et al. 2001; Fig. 26). NE German Basin formation was initiated during the Early Permian and was largely controlled by normal faulting related to deep-seated ductile shearing, with a steep and faulted eastern and a gently dipping western basin margin. A post-rift subsidence phase of 35 Ma immediately followed this east-west extension. The cantilever model predicts a stretching factor of/3 = 1.2 in the basin centre and 1.0 at the margins, which would have only

a slight effect on the crustal structure. The resulting smooth Moho uplift would fit well with the observed seismic data (Krawczyk et al. 1999). Restoration of the subsequent postZechstein kinematic evolution of the NEGB along a 260 km long N E - S W cross-section further indicates two major uplift periods at the Jurassic-Cretaceous and the Cretaceous-Tertiary boundaries (Kossow & Krawczyk 2002). Quantification of geological processes yields a total basement subsidence of 2850 m in the basin centre from end-Zechstein to present, maximum erosion of 860 m during the Cretaceous-Tertiary event at the southern NEGB margin, and at least 9 km of basin shortening. Interestingly, there is a clear correlation between the deformation intensity and the amount of uplift and erosion associated with the Cretaceous-Tertiary deformational period in the NEGB. Deformation intensity decreases from south to north, as do uplift rates, thus suggesting compression from the south, which was probably related to Alpine-induced intraplate deformations (Kossow & Krawczyk 2002). The Permian-Mesozoic development and tectonic inversion of the Polish Basin has been modelled using a 3D structural model combining analysis of 3D depth views and thickness maps (Lamarche et al. 2003a; Lamarche & Scheck-Wenderoth 2005). The model confirms earlier ideas that the Polish Basin and the

POST-VARISCAN BASIN EVOLUTION, EUROPE

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Fig. 26. Schematiccross-section across the northern part of the NE German Basin showingthe faulted basement comprising Permo-Carboniferousvolcanic units, which were subsequentlyoverlain by Rotliegend sediments(after Rieke et al. 2001).

Mid-Polish Swell are genetically related to the Teisseyre-Tornquist Zone, which seems to have tectonically controlled the development of the area through time (e.g. Kutek & Glazek 1972; Dadlez et al. 1995; Kutek 2001). When the Mid-Polish Trough started to form, the Teisseyre-Tornquist Zone constituted a zone of crustal weakness that was prone to extensional deformation. Crustal thinning along the Teisseyre-Tornquist Zone, rifting, and the following Mesozoic subsidence resulted in additional weakening along the zone. As a result, when the stress conditions changed from transtensional to compressional at the end of the Cretaceous, the Teisseyre-Tornquist Zone was preferentially deformed, inducing the inversion of the Mid-Polish Trough and the uplift of a central NW-SE-elongated anticlinorium along the former basin axis, as well as the formation of two bordering marginal troughs (see Krywiec 2002a; Lamarche et al. 2003a for details). This geometry is the surface expression of the tectonic squeezing of the Teisseyre-Tornquist Zone, which played the role of an intra-continental zone of crustal weakness as modelled by Nielsen & Hansen (2000), Hansen et al. (2000) and Gemmer et al. (2002). Although the stress magnitudes may have significantly decreased after the climax of the tectonic inversion, the stress pattern remained compressional, as indicated by the Cenozoic central horst and marginal troughs developed above the Mid-Polish Swell (Lamarche et al. 2003a; Lamarche & Scheck-Wenderoth 2005). The Teisseyre-Tornquist Zone can be considered as a regional weakness zone within which the deformation was localized. A strong tectonic inheritance of Palaeozoic and Precambrian basement structures influenced the deformation during the tectonic inversion (Krzywiec 2004). As a result of the mosaic nature of its Palaeozoic basement, the southwestern flank of the Mid-Polish Trough was tectonically unstable during the Mesozoic, in contrast to the stability of the Precambrian East European Craton beneath the northeastern part of the Mid-Polish Trough. The model of Lamarche & Schech-Wenderoth (2005) and tectonostratigraphic models based on seismic reflection data (Krzywiec 2004) also show the Zechstein salt-beating layer acting as a decoupling level between the pre-Zechstein basement and the Mesozoic cover in the central and northern segments of the Polish Basin, inducing disharmonic deformation during the tectonic inversion. Thus, the idea is that the TTZ was a zone of weakness allowing the Polish Trough to form. Such an idea is supported by the fact that long-lived shear zones (in the crust, but probably also in the mantle) tend to focus strain without regard to the past tectonic context of the area. This is a fact, and is totally independent of theoretical models. For example, the border faults of the Viking Graben are at present the loci of a high degree of micro-seismic activitiy (e.g. Olesen et al. 2004). This observation is in clear contradiction to the idea that crustal thinning implies (following thermal relaxation) lithospheric strengthening (with respect to nearby non-rifted areas). Furthermore, recent advances in fault zone rheology suggest that repetitive deformation of the fault zone results in the development of an in situ mylonitic foliation and concentration of weak phases, which imply a drastic decrease in the coefficient of friction in the

fault zone and potentially a local drop in crustal strength (Bos & Spiers 2002; Holdsworth 2004).

Discussion The Variscan orogen was characterized by a particularly long period of intracontinental deformation, associated with the collision of Gondwana and Laurussia. The post-collisional evolution of Europe (i.e. within the latest Carboniferous-Early Permian time frame) was characterized by the formation of a series of rift, and wrench-induced, basins across the continent, together with significant magmatic events. From the above outline it can be seen that although we have a reasonable understanding of the broad evolution of the late stages of the Variscan orogenic event and the subsequent period of wrench fault activity that was widespread across both the internal and external Variscan provinces, there are many problems relating to our understanding both of the underlying mechanisms that controlled the various observed events and of the detailed integration of the various observations. In particular, there are problems relating the internal and external zones, which have, at times, remarkably similar evolutionary histories (e.g. coeval graben formation and associated volcanism in northern Spain, Italy and northern Germany). Although certain events may be interpreted in terms of plume-related activity, how do we interpret similar successions thousands of kilometres apart? Although it is clear from the above outline that the post-Variscan period in Central, Southern and Western Europe was a period of intense tectonic, magmatic and sedimentary change, any attempt at summarizing these changes must, by necessity, try to assess the various possible driving mechanisms involved in the generation of the post-Variscan basins. It is clear that the coincidence of tectonic activity (both compressional and extensional), magmatic activity and basin formation (with subsequent sedimentation) was very different from the periods immediately before and after. The geological evolution of the region, however, is problematic given the relative lack of significant Early Permian extensional structures. The large amounts of crustal-derived and crustalcontaminated volcanic rocks are also problematic. The processes controlling the post-orogenic modification of the Variscan lithosphere have been variably attibuted to such mechanisms as slab detachment, delamination of the mantle-lithosphere, crustal extension and plume activity during the Stephanian-Early Permian phase of wrench faulting and magmatism that overprinted the Variscan orogen and its foreland (see Ziegler et al. 2006, for references). The following sections will attempt to examine the main controlling mechanisms within the basin, to try to isolate those that are of greatest importance in terms of overall basin evolution.

Rifting history

The examination of a variety of basins across Europe has allowed us to compare and contrast the various successions within the

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basins, as well as features such as basin form, controls on basin formation, and the magmatic, tectonic and infill history. The observed contrasts suggest that the underlying processes that controlled the post-Variscan evolution of Europe were very different between those areas located in the former Variscan foreland basin and those within the thrust front. The various modelling studies carried out on the post-Variscan Permian basins suggest very different mechanisms for each area. Permian rifling in the former Variscan hinterland seems to have been strongly controlled by the collapse of the mountain chain (Brunet & Le Pichon 1982; Prijac et al. 2000) with a possible far-field extension component also being plausible (Henk 1999). This process may also have been modified by the slow decay of the associated thermal anomaly (e.g. Paris Basin). In contrast, rifting in the former Variscan foreland appears to have been dominated by late Variscan wrench tectonics (van Wees et al. 2000), particularly along the boundary between Precambrian and Phanerozoic Europe (Dadlez et al. 1995). Numerical modelling highlights the role of such lithospheric discontinuities in controlling tiffing (Pascal & Cloetingh 2002; Pascal et al. 2003). Butler et al. (1997) have noted that pre-existing heterogeneities in the continental lithosphere are thought to influence its response to subsequent deformation. From the late Early Carboniferous onwards, Laurussia was transected by the Arctic-North Atlantic Rift System, which was partially superimposed on the Caledonian suture zone (Ziegler 1990). Indeed, Variscan exploitation of older Caledonian structures has been reported from other areas (e.g. offshore Ireland; for details, see Shannon 1991; McCann 1996). In Cornwall, early Variscan thrusts were reactivated as late Variscan extensional faults (Shail & Alexander 1997). Additionally, the interaction of the Variscan structures with the pre-Variscan east-west dextral (Badham 1982) transform fault system (running from the Uralides through Europe (Pitra et al. 1999) to the Appalachians) and the NNW-SSE-trending wrench fault system produced a complex series of conjugate shear zones and pull-apart structures in the Cornwall area (Willis-Richards & Jackson 1989) that remained active throughout the early Permian. It is, therefore, highly likely that, within the area under discussion, older structures, both Caledonian and Variscan, were reactivated by later Variscan tectonic activity. However, more recent work (Ebbing et al. 2006) has suggested that even older structure may be involved. In their study of the Oslo Graben they suggested that the rifting in the region is coupled to a reactivation of Precambrian fault systems, and indeed, the very location of the Oslo Graben is more strongly dependent on the pre-rift structure of the area than previously assumed. One factor of note is that Permian wrench activity was not merely limited to 'accreted' Europe, but is also evident in other parts of the craton where there is sufficient stratigraphic evidence. In particular, there is evidence of late Carboniferous-early Permian transtensional tectonic activity in the Dniepr-Donets Basin (Stovba & Stephenson 1999) and even further afield on the margins of the East European Craton (Saintot et al. 2006).

Mantle plume dynamics

Another important issue addressed by modelling of Permian basins, and in particular of the Oslo Rift (Ro & Faleide 1992; Pedersen & van der Beek 1994), is the eventual role played by a mantle plume (although this idea has recently been questioned; see Ebbing et al. 2006, for details). Despite significant differences in the tectonosedimentary setting and the type of magmatic activity within the various basins examined, the StephanianAutunian volcanic rocks in the internal Variscides comprise a high proportion of pyroclastic deposits and are generally of intermediate to felsic composition, of calc-alkaline character, and often have a significant crustal component, as shown by Sr-Nd isotope data and the presence of crustal xenoliths, magmatic garnet, and

(locally) topaz, and (for the volcanic rocks in the NE German Basin) the large amount of inherited zircons necessitating sensitive high-resolution ion microprobe (SHRIMP) dating (Breitkreuz & Kennedy 1999). The calc-alkaline character may reflect the derivation of the melts from a subduction-modified mantle source, extensive assimilation of crustal material, or perhaps inheritance resulting from the melting of older calc-alkaline, crustal sources (such as Cadomian basement). However, the relative scarcity of more primitive mafic melts precludes a more precise interpretation of the mantle source compositions. In addition, numerical studies suggest that huge volumes of magmas can be produced with small amounts of stretching and without the need for any underlying thermal anomaly (Pedersen & van der Beek 1994). Crustal p r o c e s s e s

The Stephanian-Autunian magmatic rocks in the internal Variscides comprise a high proportion of pyroclastic rocks and are generally of intermediate to felsic composition. Their generally calc-alkaline character suggests a subducfion-related origin. With the possible exception of some magmafic rocks in the Alpine basement, this contradicts their intracontinental setting and the fact that the Variscan oceans had closed by mid-Carboniferous times. However, Sm-Nd isotope data and the presence of garnet and crustal xenoliths indicate that many contain a significant crustal component. This is corroborated by the predominantly negative ENd(t) values of the 290-300 Ma volcanic and intrusive rocks of felsic to intermediate composition: - 2.1 to - 6.0 for the Krkonoge Basin (Ulrych et al. 2002), -2.7 to -6.1 for the Intra-Sudetic Basin (Ulrych et al. 2004); -0.8 to - 7 . 0 for the rhyolites of the Halle Volcanic Complex (Romer et al. 2001), -4.3 to -7.5 for the granites in Comwall (Darbyshire & Shepherd 1994), and - 0 . 6 to - 5 . 7 for the Saar-Nahe Basin (Schmidberger & Hegner 1999; von Seckendorff et al. 2004a, and references therein). The parent magmas of the granitoids, rhyolites and andesites may, therefore, have assimilated large amounts of crustal material, or alternatively, be derived from mantle sources that had been modified by earlier subduction events (e.g. Cabanis & Le Fur-Balouet 1989; Schmidberger & Hegner 1999; Innocent et al. 1994; Cortesogno et al. 1998). As in the North German Basin, the granites and rhyolites may be of crustal origin, and their calc-alkaline signature inherited through partial melting of calc-alkaline basement (Schaltegger 1997b; Romer et al. 2001). The possible mechanisms for mantle melting in the internal Variscides may have been the break-off of subducted oceanic crust (e.g. Schaltegger 1997b; Cesare et al. 2002) or even the oblique subduction of the mid-ocean ridge of Palaeotethys beneath the active Eurasian margin (Stampfli 1996). Regional extension leading to lithosphefic thinning and decompressional melting of updoming asthenosphere may have been a contributing factor in the late Carboniferous-early Permian period. Compared with the foreland, Stephanian-Autunian mafic rocks are much rarer in the internal Variscides, which suggests that the mantle-derived parent melts were unable to reach the surface, but stalled at lower to midcrustal levels. This may have been due to the large contrast between the density of the parent melt and a low average density of thinned Variscan crust. Only after fractionation and assimilation of sufficient amounts of crustal material did the melts attain a low enough buoyancy to be able to escape the magma chambers and erupt on the surface. M a g m a t i c - t e c t o n i c activity

The relatively short and widespread pulse of StephanianAutunian magmatism is likely to have taken place in response to changes in the regional stress field at the Westphalian-Stephanian boundary and subsequent thermal equilibration of the lithosphere. The change of stress may have been due to a change in VisranWestphalian crustal shortening and orogen-parallel extension,

POST-VARISCAN BASIN EVOLUTION, EUROPE

and to Stephanian-Autunian gravitational collapse of the Variscan orogen. The latter process was possibly superimposed and aided by a far-field dextral extensional stress-feld that was due to the collision of Gondwana with eastern southeastern North America and concomitant dextral translation (Torsvik & Van der Voo 2002). The invocation of far-field effects is something that has previously been noted in discussions of postVariscan tectonics (e.g. discussion on the origin of the NE German Basin; see DEKORP-BASIN Research Group 1999, for details). In terms of the magmatic history of the post-Variscan there are some indicators that far-field effects might also have played an important role. For example, the alkaline composition, style of volcanism and the presence of abundant megacrysts and mantle xenoliths in the Early Permian mafic rocks in Scotland indicate derivation by low-degree melting of local mantle sources and rapid, vertical transport. In contrast, the sub-alkaline mafic dyke and sills complexes (such as the Whin Sill Complex) indicate higher degrees of mantle melting, and do not necessarily reflect a mantle thermal anomaly of the same extent. The geometry and orientation of the dyke swarms suggest a magmatic focal region in the vicinity of the Denmark-Skagerrak region, which suggests that magma transport may have been horizontal, westwards into the North Sea and Scotland. Thus, the position, trend, number and size of the dykes may have been controlled by the far-field dextral extensional stress field.

Conclusions The end Carboniferous-early Permian history of Europe represents a period of crustal instability and re-equilibration throughout Western and Central Europe. An extensive and significant phase of Permo-Carboniferous magmatism led to the extrusion of thick volcanic successions across the region. Coeval transtensional activity led to the formation of more than 70 rift basins, which differ both in form and infill according to their position relative to the former Variscan Orogenic Front as well as to the controls that acted on basin development. Despite the fact that no unified model for the Permian event can at present be unequivocally proposed from the results of the various modelling studies, recent studies do agree on two fundamental and relevant points: (1) Permian rifting was widespread in Europe with progressively propagated development; (2) its signature strongly influenced the evolution of the European lithosphere during Mesozoic and Cenozoic times (S0rensen 1986). It may not, however, be possible to provide more detailed models for the evolution of the region. Numerical modelling of lithospheric rifting, for example, requires numerous parameters, among which the pre-rifl crust and mantle-lithosphere structure are crucial. Because the pre-Permian lithosphere structure has been obscured by repetitive tectonic phases in most parts of Europe, lithosphere-scale modelling of the Permian event remains difficult and modelling results need to be treated with a high degree of circumspection. The best approach, therefore, to elucidating the tectonosedimentary and magmatic history of the region is to adopt a broad approach, examining the various basins at a range of scales and making use of a variety of techniques. This manuscript was greatly improved by the reviews of two anonymous reviewers. T. Beilfuss is thanked for the production of the excellent diagrams.

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Pre-Alpide Palaeozoic and Mesozoic orogenic events in the Eastern Mediterranean region A. I. O K A Y 1, M. SATIR 2 & W. SIEBEL 2

lJstanbul Teknik (Yniversitesi, Avrasya Yerbilimleri Enstitiisii, Ayazafi, a 80626, Istanbul Turkey (e-mail: okay@ itu.edu.tr) 2Institut fiir Geowissenschaften, Universitiit Tiibingen, Wilhelmstrafle 56, D-72074 Tiibingen, Germany

Abstract: We review the Palaeozoic-Early Mesozoic evolution of the Eastern Mediterranean-Balkan region with special reference to Anatolia, and provide new isotopic data on the Palaeozoic magmatic and metamorphic rocks. The pre-Alpide evolution of the region involves episodic growth of Laurussia by accretion of oceanic terranes and Gondwana-derived microcontinents. Terrane accretion, associated with deformation, magmatism and regional metamorphism, took place in the Late Ordovician-Early Silurian, Carboniferous, Late Triassic-Early Jurassic and Mid-Jurassic. The Late Ordovician-Early Silurian accretion is inferred from stratigraphic and faunal records in the Pontides; other evidence for it is buried under young cover on the northern margin of the Black Sea. The Carboniferous orogeny is related to southward subduction and continental collision on the southern margin of Laurussia. It is marked in the Pontides by high-grade regional metamorphism, north-vergent deformation and post-orogenic latest CarboniferousEarly Permian plutonism. The latest Triassic-Early Jurassic Cimmeride orogeny involved the collision and amalgamation of an oceanic plateau to the southern margin of Laurasia. It is represented by voluminous accretionary complexes with Late Triassic blueschists and eclogites. Late Jurassic regional metamorphism and deformation is confined to the Balkans, and is the result of continental collision between the Rhodope-Serbo-Macedonian and Strandja blocks in the Late Jurassic. The Palaeozoic geological history of the Balkans and the Pontides resembles that of Central Europe, although the similarities end with the Mesozoic, as a consequence of the formation of Pangaea.

Orogenic belts and Mesozoic oceanic basins occupy the Eastern Mediterranean region between the stable areas of the East European Craton in the north, and NE Africa and the Arabian Platform in the south (Fig. 1). The East European Craton, as represented by the Ukrainian Shield north of the Black Sea, is an A r c h a e a n Palaeoproterozoic crystalline terrane. The consolidation of the southern part of the East European Craton was completed by 2300-2100 Ma (e.g. Bogdanova et al. 1996; Claesson et al. 2001). In the Early Palaeozoic, the East European Craton formed part of the Balfica plate, which collided in the west with Laurentia, Avalonia and Armorica, creating Laurussia in the Late Palaeozoic (e.g. Pharaoh 1999; Matte 2001; Wart 2002). In contrast, Africa and the Arabian Platform constituted part of Gondwana, which preserved its unity until the Early Mesozoic opening of the southern Atlantic. Very large areas in the northern margins of Gondwana are characterized by NeoproterozoicCambrian plutonism and metamorphism forming part of the Pan-African-Cadomian orogenic cycle (e.g. Stern 1994), and are therefore readily distinguished from the Palaeoproterozoic basement of the East European Craton. During the Late Palaeozoic and Mesozoic, Tethyan oceanic basins separated Laurussia from NE Africa-Arabia. Parts of the present Eastern Mediterranean Sea represent a Triassic to Jurassic remnant of a Tethyan oceanic crust, whereas the Black Sea is a Late Cretaceous oceanic back-arc basin that opened during the northwards subduction of a Tethyan ocean (e.g. ~eng6r & Yllmaz 1981; Garfunkel 1998). The Anatolian-Balkan region between the Eastern Mediterranean and the Black Sea consists of several small continental fragments or terranes bearing evidence of various periods of deformation, metamorphism and magmatism, the latest and strongest of which is the Alpide orogeny. The Alpide orogeny resulted in the amalgamation of the continental fragments into a single landmass in the Tertiary. Previous to this amalgamation, these continental fragments were situated on the margins of the Tethyan oceans, or formed small edifices within the ocean. The pre-Alpide orogenic history of these terranes forms the subject of this paper.

Terranes in the Eastern Mediterranean-Balkan region As orogenic events are restricted to the plate margins, identification of former plates is important for an understanding of the orogenic evolution. Only the deformed continental parts of the former microplates would be expected to be preserved, and they would be rimmed by sutures marked by linear zones of accretionary complex, blueschist, eclogite and ophiolite, and would show distinctive strafigraphic features, especially if the intervening oceans were large. The main methods used in the differentiation of the former plates include recognition of sutures, palaeomagnetism, faunal provinciality and stratigraphy. A complication in this picture is that the number and configuration of the plates change through time. For example, the Anatolian microplate, which makes up most of the present Anatolian landmass, did not exist before the Miocene (e.g. Seng6r 1979). Furthermore, late-stage strike-slip faulting may lead to the dispersal of a single palaeoplate, as happened during the Cretaceous opening of the Black Sea (Okay et al. 1994). Therefore, a better term for such a palaeo-plate would be 'terrane', and this would be distinguished by its distinctive stratigraphic, palaeomagnetic, faunal, structural, metamorphic and magmatic features. The following terranes are defined in the Anatolian region from south to north: the Arabian Platform comprising part of SE Anatolia, the Anatolide-Taufide Block, the Kir~ehir Massif, the Sakarya Zone, the Istanbul Zone and the Strandja Massif (Fig. 1; Okay & Tfiystiz 1999). Of these, the last three are grouped together as the Pontides. Pre-Alpide orogenic events are especially strong and well documented in the Ponfides (Fig. 2). With the possible exception of its northwestern margin, the Anatolide-Tauride Block was largely free of Palaeozoic and early Mesozoic deformation. Therefore, most of this review concerns the pre-Alpide geological history of the Pontides. In addition to the Anatolian terranes listed above, several other terranes have been defined in the Balkans (e.g. Burchfiel 1980; Stampfli 2000; Stampfli et al. 2001). The major ones include the Moesian Platform, the Rhodope-Serbo-Macedonian Massif, Pelagonia (which comprises the Pelagonian Zone and the

From: GEE, D. G. & STEPHENSON, R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 389--405.0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Tectonic map of the Eastern Mediterranean region showing the major terranes and the bounding sutures. The filled triangles indicate the polarity of the subduction (modified from Okay & Ttiystiz 1999). NAF, North Anatolian Fault; EAF, East Anatolian Fault.

Cyclades) and Apulia, which includes Greece south of the Pindos suture (Fig. 1; Stampfli et al. 2001). Palaeozoic rocks are not exposed on Apulia, so it is not known whether this region was affected by the Variscan orogeny. Assuming that it was not, then Apulia probably forms a single terrane jointly with the Anatolide-Tauride Block. Isotopic data indicate Carboniferous (325-295 Ma) plutonism and regional metamorphism in the Pelagonian Zone and in the Cyclades (see the discussion by Vavassis et al. 2000), which contrasts with the Neoproterozoic magmatic and metamorphic basement ages for the Anatolide-Tauride Block. Their Mesozoic histories are also different, with Late Jurassic-Early Cretaceous ophiolite obduction on the northern margin of Pelagonia contrasting with the Late Cretaceous ophiolite obduction on the Anatolide-Tauride Block. The contact between Pelagonia and the Anatolide-Tauride Block is represented by an Eocene thrust, where the cover sequence of the Cycladic Massif, metamorphosed at blueschist-facies conditions in the Eocene, is thrust on the Menderes Massif (Fig. 3; Okay 2001). This contact, which may represent the extension of the Pindos suture, may link up with the Izmir-Ankara suture, in which case Pelagonia will be correlated with the Sakarya Zone (Fig. 1). Such a correlation is supported by the similar Variscan magmatic and metamorphic ages from Pelagonia and the Sakarya Zone (see below), although their Mesozoic histories are separate. The Alpide orogeny in the Mediterranean area started with the convergence between the Africa-Arabian and Eurasian plates during the Late Mesozoic. The relative movement between these two plates was sinistral strike-slip from Early Jurassic to midCretaceous time (e.g. Savostin et al. 1986; Dewey et al. 1989). Starting with the Cenomanian-Albian (100-90 Ma) the Africa-Arabian and Eurasian plates started to converge, presumably with the initiation of subduction. In the geological record the Albian flysch of the Central Pontides (Ttiystiz 1999), the Turonian (c. 91 Ma) high-temperature-medium-pressure metamorphism in the Klr~ehir Massif (Whitney et al. 2003), and the Campanian (c. 80 Ma) high-pressure-low-temperature metamorphism in the Anatolide Tauride Block (Sherlock et al. 1999) are the first

recognized events of the Alpide orogeny in Turkey. Therefore, the period discussed in this review extends to the Early Cretaceous. We also do not discuss the Neoproterozoic-Cambrian-aged Pan-African orogenic events in Anatolia, which, although important (e.g. Krrner & ~engrr 1990; Yi~itba~ et al. 2004), are poorly preserved and documented.

The Anatolide-Tauride Block The Anatolide-Tauride Block has a Neoproterozoic crystalline basement overlain by a sedimentary succession ranging from Mid-Cambrian to Miocene in age (e.g. Gutnic et al. 1979; Ozgiil 1984, 1997). It was strongly deformed and partly metamorphosed during the Alpide orogeny, and now consists of metamorphic regions in the north (the Anatolides) and a south-vergent Eocene nappe stack in the south (the Taurides). Stratigraphy of several nappe units in the Taurides reveals Palaeozoic to Mesozoic sedimentary sequences with no evidence of pre-Cretaceous deformation or metamorphism (e.g. Gutnic et al. 1979; Ozgtil 1984, 1997). Rare reports of Late Triassic deformation in the Anatolide-Tauride Block (e.g. Monod & Akay 1984) have been questioned and need confirmation (G6nctio~lu et al. 2003). The Anatolide-Tauride Block is here considered as not affected by significant pre-Alpide Phanerozoic contractional deformation, which contrasts with the regions to the west and north that were deformed and metamorphosed during the Variscan orogeny. The largest outcrops of the Precambrian basement in the Anatolide-Tauride Block are found in the Menderes and Bitlis massifs (Figs 3 and 4). In the Menderes Massif, the metagranitoids, which make up most of the crystalline basement, have been dated as to c. 550 Ma using a stepwise Pb evaporation method on zircons (Hetzel & Reischmann 1996; Loos & Reischmann 1999); this is a similar age to those found in the Arabian Platform and in NE Africa. The eclogite-facies metamorphic rocks in the basement of the Menderes Massif are also Neoproterozoic in age (Candan et al. 2001). The Palaeozoic stratigraphy in the Anatolide-Tauride Block is also similar to

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Fig. 2. A chronostratigraphic chart showing generalizedgeological relationships of the Pontic terranes. The main sources of the data are: for the Istanbul Zone, Gedik (1975), Ali~an& Derman (1995), G6rtir et al. (1997) and Dean et al. (2000); for the Strandja Zone, Chatalov (1988), and Okay et al. (2001); for the Sakarya Zone, Altaneret al. (1991) and Okay & Leven (1996).

that of the northern margin of the Arabian Platform. Hence, since Smith (1971), all palaeogeographical reconstructions place the Anatolide-Tauride Block into the Eastern Mediterranean Sea between the Levant and Egyptian margins. Recent evidence for the latest Ordovician (Hirnantian) glaciation in the AnatolideTauride Block, including the presence of striated pebbles and striated basement (Monod et al. 2003), supports this pre-drift position. Stratigraphic evidence from the Levantine (e.g. Bein & Gvirtzman 1977; Garfunkel & Derin 1984) and Gondwana margins in SE Anatolia (Fontaine et al. 1989) indicates that the Anatolide-Tauride Block rifted away from Gondwana during the Triassic or Early Jurassic with the opening of the Tethyan ocean. However, it was never far away from Gondwana, and drifted with Gondwana to the north, as shown by its Jurassic palaeomagnetic record (Piper et al. 2002) and by its Jurassic-Cretaceous stratigraphy, which resembles that of the SE Anatolia.

The Istanbul Zone The Istanbul Zone consists of a Neoproterozoic crystalline basement overlain by Lower Ordovician to Eocene sedimentary rocks (Fig. 2; e.g. Haas 1968; G6rtir et al. 1997). Before the Late Cretaceous opening of the Black Sea, the Istanbul Zone was situated east of the Moesian Platform and adjacent to the Scythian Platform (Okay et al. 1994). These three tectonostratigraphic units have a similar basement and show a similar Palaeozoic-Mesozoic stratigraphic development (e.g. Tari et al. 1997; Nikishin et al. 1998), and formed a single late Proterozoic-Early Palaeozoic terrane, named here the MOIS (Moesian-Istanbul-Scythian) terrane. The crystalline basement of the Istanbul Zone is characterized by voluminous granitoids, which intrude low- to medium-grade metasediments and metavolcanic rocks (Usta6mer & Rogers 1999; Yi~itba~ et al. 1999, 2004). The granitoids have yielded U - P b and Pb/Pb

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Fig. 3. Tectonic map of the western Anatolia illustratingthe geological features discussedin the text. The cross-hatched area shows the extent of the metamorphosed Palaeozoic and Mesozoic rocks of the Menderes Massif.

zircon ages of 590-560 Ma, and the surrounding metasediments have provided similar R b - S r mica ages (Chen et al. 2002). The geochemistry of the granitoids and the metavolcanic rocks indicates a subduction-zone setting in the Neoproterozoic. In terms of age, lithology and geochemistry, the basement of the Istanbul Zone is similar to the Pan-African basement of northern Gondwana, and unlike the East European Craton. Therefore, the Istanbul Zone is generally regarded as a Peri-Gondwana terrane (e.g. Stampfli 2000). The Neoproterozoic basement of the Istanbul Zone is overlain by a thick Palaeozoic sedimentary succession extending from the Ordovician to the Carboniferous (Fig. 2). There are significant stratigraphic differences between the westem and eastern parts of the Istanbul Zone, which led to a suggestion that the Istanbul Zone consists of two terranes, the Istanbul terrane in the west and the Zonguldak terrane in the east (Kozur & G6nctio~lu 1998; Stampfli et al. 2002; von Raumer et al. 2002). The most important difference is in the Carboniferous system, which in the west is represented by Vis~an radiolarian cherts overlain by siliciclastic turbidites, but in the east, in the Zonguldak region, is

represented by Vis6an neritic carbonates overlain by Namurian to Westphalian coal measures (Figs 2 and 5). However, evidence for a Phanerozoic ocean, in terms of pelagic sedimentary rock, m61ange, ophiolite or blueschist, is missing between the western (Istanbul s e n s u stricto) and eastern parts (Zonguldak) of the Istanbul Zone (see Fig. 5), and the stratigraphic differences are a result of facies changes. A similar situation has been reported in the Moesian Platform, where neritic carbonate deposition in the Tournaisian in the north is replaced by radiolarian chert sedimentation in the south in the Elovitza region (Haydoutov & Yanev 1997). During the Carboniferous, the MOIS terrane was part of the southern continental margin of Laurussia. Turbidite deposition in a continental slope setting took place in the western part of the Istanbul Zone and on the southern margin of the Moesian Platform, whereas coal deposition occurred in swamps in the north (Fig. 6). The palaeogeographical situation was similar to that of SW Britain at the same period, when coal measures were being deposited in Wales and siliciclastic turbidites (Culm facies) in Cornwall and Devon (Fig. 6; e.g. Guion et al. 2002).

OROGENS IN THE EASTERN MEDITTERANEAN

393

Fig. 4. Tectonic map of eastern Anatolia illustrating the geological features discussed in the text. (For legend see Fig. 3.) The cross-hatched area shows the extent of the metamorphosed Palaeozoic and Mesozoic rocks of the Bitlis Massif.

Fig. 5. The distribution of pre-Jurassic rocks in the Istanbul Zone (simplified from Aksay et al. 2002; Ttirkecan & Yurtsever 2002). Noteworthy features are the different Carboniferous and Triassic facies in the west and east, and the trend of the facies boundary, which is highly oblique to the Intra-Pontide suture.

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Fig. 6. Carboniferous palaeogeography in the southern margin of Laurussia (a) compared with that of Britain at the same period (b). Both maps are of the same scale. In (a) the Istanbul Zone is restored to its predrift position before the Cretaceous opening of the Western Black Sea basin (Okay et al. 1994). In the Tournaisian-Vis~an, neritic carbonate deposition took place in Moesia and in the eastern Istanbul Zone; this was succeeded by the accumulation of coal during the Namurian and Westphalian. In the same period radiolarian chert sedimentation gave way to siliciclastic turbidite deposition in a continental slope setting in the western Istanbul Zone and the southern margin of Moesia. A similar picture exists in Britain, where, in addition, the Early Devonian Lizard ophiolite in Cornwall provides another indication of the Rheno-Hercynian ocean in the south. It should be noted that the Intra-Pontide suture truncates the facies belts. The Moesia data are from Dachev et aL (1988), Popova et al. (1992), Tenchov (1993), Haydoutov & Yanev (1997) and Tari et aL (1997); the data for Britain are from Guion et al. (2002) and Warr (2002).

The Intra-Pontide suture, which marks the southern boundary of the Istanbul Zone, truncates the Palaeozoic and Triassic facies boundary between the westem and eastern parts of the Istanbul Zone (Figs 5 and 6). This suggests removal of a major section of the Istanbul Zone, possibly by post-Triassic strike-slip faulting. Late C a r b o n i f e r o u s d e f o r m a t i o n a n d p l u t o n i s m

The Palaeozoic sequence in the Istanbul region ends with Visran to Namurian siliciclastic turbidites, whereas in the east, in the Zonguldak region, it extends into the Westphalian coal measures (Fig. 2; G6riir et al. 1997). The Palaeozoic rocks in the Istanbul region are deformed in a contractional mode, with the generation of recumbent folding, local cleavage and minor thrusting, whereas deformation is less intense in the Zonguldak region. The minor folds generally show an east to NE vergence (Seymen 1995; Zapcl et al. 2003), although the timing of deformation, whether Variscan or later, is difficult to constrain. Nevertheless, the observation that the lowermost Triassic red beds step down from Carboniferous to Ordovician (Ttirkecan & Yurtsever 2002) indicates significant deformation and erosion in the Late Carboniferous-Permian interval. The deformed Palaeozoic rocks are intruded by a Permian granite east of Istanbul, which has biotite K - A r and whole-rock R b - S r ages of c. 255 Ma (early Late Permian; Figs 3 and 5; Yllmaz 1977). The age of the undeformed pluton constrains the Variscan deformation in the Istanbul Zone to the Late Carboniferous-Early Permian. The Triassic restoration

The Palaeozoic rocks in the Istanbul Zone are unconformably overlain by Triassic continental clastic rocks with basaltic flows

(Fig. 2). In the Istanbul region, the Triassic sequence continues with neritic to pelagic carbonates, capped by Carnian or Norian siliciclastic turbidites, showing a typical transgressive..passive margin type of development (e.g. Gedik 1975; Yurtta~-Ozdemir 1971), whereas in the east the Triassic is represented mainly by continental clastic rocks and lacustrine limestones (Figs 2 and 5). The change in the Triassic facies closely follows that of the Palaeozoic, suggesting a long-term hinge, possibly controlled by a deep-seated fault (Fig. 5). The termination of deposition in the Carnian or Norian in the Istanbul Zone probably reflects the Cimmeride orogeny, which is particularly strong in the Sakarya Zone farther south.

P a l a e o g e o g r a p h i c a l affinity

The Infra-Cambrian to Cambrian granitoids (590-560 Ma) and Neoproterozoic metamorphism in the basement of the Istanbul Zone suggest a location on the Gondwana margin in the latest Precambrian. This is supported by the Ordovician trilobite faunas, which are similar to those from Central European and Anglo-Welsh successions, and differ from those of Baltica, as well as from those of typical Gondwana realms of the Anatolide-Tauride Block and the Arabian Platform (Dean et al. 2000). Therefore, a location of the MOIS terrane on the western margin of Baltica during the Early Ordovician, as shown in some reconstructions (von Raumer et al. 2002) is not possible. The absence of the latest Ordovician (Hirnantian) glaciation in the Istanbul Zone provides another constraint on its location on the Gondwana margin. However, from the Late Silurian onwards the Istanbul Zone became part of Laurussia, as indicated by its palaeomagnetic record from sediments of Late Silurian, Devonian, Carboniferous and Triassic age (Sarlbudak et al.

OROGENS IN THE EASTERNMEDITTERANEAN 1989; Evans et al. 1991), and by the Devonian-Carboniferous foraminiferal assemblages (Kalvoda 2003; Kalvoda et al. 2003). These data imply that the MOIS terrane separated from Gondwana during the Ordovician, and docked with Baltica in the Late Ordovician-Early Silurian; however, there is little evidence for Ordovician-Silurian collision in the geological record of the Istanbul Zone. Apparently, the zone of collision is hidden under young cover on the northern margin of the Black Sea. The Early Palaeozoic history of the MOIS terrane appears to be remarkably similar to that of Avalonia (Stampfli et al. 2002; Winchester & the PACE TMR Network Team 2002).

The Strandja Massif The Strandja Massif forms part of large metamorphic region in the Balkans, which includes the Rhodope, Serbo-Macedonian and Peri-Rhodope zones (Fig. 1). The relationship between these metamorphic units, and their ages of regional metamorphism are poorly known. The Strandja Massif crops out both in Turkey and in Bulgaria, and is bordered in the west by the Rhodope Massif. It consists of a metamorphic basement of unknown age, intruded by Permian granitoids, and overlain by continental to shallow marine sedimentary rocks of Triassic to Mid-Jurassic age (Fig. 2). During the Late Jurassic, the cover and the basement of the Strandja Massif underwent contractional deformation and regional metamorphism, and Triassic allochthons were emplaced on the Mid-Jurassic metasediments. Cenomanian and younger sediments lie unconformably over the metamorphic rocks (Chatalov 1988; Okay et al. 2001). The basement of the Strandja Massif consists of gneisses and micaschists intruded by voluminous plutonic rocks, several of which have been dated as Early Permian (c. 271 Ma) using stepwise Pb evaporation method on single zircon grains (Okay et al. 2001). The overlying Triassic sequence of the Strandja Massif shows affinities to the Central European Germanic Triassic facies, with a basal continental clastic series overlain by Middle Triassic shallow-marine carbonates (Chatalov 1988, 1991). A hiatus between the Late Triassic and Early Jurassic is probably a distant echo of the Cimmeride deformations farther south (Fig. 2). The shallow marine sedimentation continued into the Mid-Jurassic (Bathonian), and was terminated by the Late Jurassic Balkan orogeny. Late Jurassic deformation and metamorphism in the Strandja Zone

The Triassic to Jurassic sedimentary cover sequence of the Strandja Massif, together with its crystalline basement, underwent deformation and greenschist-facies metamorphism during the Late Jurassic. The age of regional metamorphism is constrained to the Late Jurassic-Early Cretaceous (Callovian-Albian) by the Bathonian age of the youngest metamorphosed strata (Chatalov 1988), and by the Cenomanian post-metamorphic cover (Fig. 2). R b - S r and K - A r biotite ages from the deformed and metamorphosed Permian granites of the Strandja Massif fall in the range of 155-149 Ma (Aydln 1988; Okay et al. 2001), indicating a Late Jurassic age for the regional metamorphism. The Late Jurassic metamorphism in the Strandja Massif was associated with north-vergent thrusting, folding, and the generation of foliation and lineation (Okay et al. 2001). Permian granitoids were penetratively deformed and thrust north over the Triassic to Jurassic mylonitic metasediments and marbles. Large allochthons, composed of Triassic deep-sea metasediments and metavolcanic rocks, were thrust northwards over the epicontinental Triassic-Jurassic rocks of the Strandja Massif (Chatalov 1985, 1988; Dabovski & Savov 1988). A foreland basin, called the Nij-Trojan trough, developed in the Oxfordian between

395

the Strandja-Rhodope massifs and the Moesian Platform. The Nij-Trojan trough migrated northward and persisted until the Early Cretaceous (Barremian; Tchoumatchenko et al. 1990; Harbury & Cohen 1997).

The Sakarya Zone The Sakarya Zone forms a continental sliver, over 1500 km long, south of the Istanbul Zone and the eastern Black Sea (Fig. 1). It consists mainly of Jurassic and younger sedimentary and volcanic rocks, which unconformably overlie a heterogeneous basement. The only sign of the Late Jurassic-Early Cretaceous deformation and metamorphism that is so intense in the Strandja Massif is a parallel unconformity at the base of CallovianOxfordian limestones (Fig. 2; Altlner et al. 1991). The pre-Jurassic basement of the Sakarya Zone includes Devonian plutonic rocks, Carboniferous plutonic and metamorphic rocks, and Triassic accretionary complexes with blueschists and eclogites (Figs 3 and 4). The pre-Jurassic relation between these basement units is strongly overprinted by Alpide deformations. The Devonian and Carboniferous units, and the Triassic accretionary complexes, are described below.

Early Devonian plutonism in the Sakarya Zone

The Devonian was a period of widespread granitoid plutonism in the Caledonides in NW Europe (e.g. Woodcock & Strachan 2002), whereas granitoids of this age were unknown in the Eastern Mediterranean region. Therefore, it was a surprise when a single sample from a granitoid in NW Turkey was dated as Early Devonian (Okay et al. 1996). The ~amllk granodiorite in the Biga peninsula (Fig. 3) forms a 20 km long and 3 - 4 km thick thrust sheet in an Alpide thrust stack. It is a leucocratic granodiorite consisting mainly of quartz, plagioclase and chloritized biotite, and is unconformably overlain by Upper Triassic arkosic sandstones. As the age of the granite is tectonically significant, zircons from a second sample from the (~amllk granodiorite were dated using the stepwise Pb-evaporation method. The details of the dating method have been given by Okay et al. (1996). Two zircon grains from the ~amllk granodiorite gave an Early Devonian age of 397.5 + 1.4 Ma (Fig. 7, Table 1), confirming the earlier less precise age of 399 +_ 13 Ma obtained by Okay et al. (1996). The relationship between the ~amlik Granodiorite and the other pre-Jurassic basement units of the Sakarya Zone are not known. However, the proximity of the highgrade Carboniferous metamorphic rocks of the Kazda~ and the essentially unmetamorphosed Devonian ~amllk granodiorite in NW Turkey (see Fig. 3) suggest major pre-Jurassic shortening between these two units.

Carboniferous deformation and metamorphism in the Sakarya Zone

The high-grade Variscan metamorphic basement of the Sakarya Zone is exposed in only a few areas throughout its 1500 km length. These include the Kazda~ and Uluda~ massifs in the west, the Devrekani Massif in the Central Pontides, and the Pulur Massif in the Eastern Pontides (Figs 3 and 4). These metamorphic regions are composed of gneiss, amphibolite and marble metamorphosed at amphibolite- to granulite-facies conditions, and in the Kazda~ and Pulur massifs there are also meta-ultramafic rocks within the sequence (Okay 1996; Okay et al. 1996; Duru et al. 2004; Topuz et al. 2004a). Isotopic age data exist only for the Pulur and Kazda~ massifs. Monazite Pb ages from a Pulur gneiss are late Early Carboniferous

396

A.I. OKAY ETAL.

Fig. 7. Histograms showing the distribution of radiogenic Pb isotope ratios derived from the evaporation of two zircon grains from the ~amhk granodiorite (a) and from a gneiss and an amphibolite of the Kazda~ Group (b) in the Sakarya Zone, NW Turkey.

(331-327 Ma, Namurian), considered as the age of high-grade metamorphism (Topuz et al. 2004a). The 315-310 Ma (Westphalian) N d - S m , R b - S r and A r - A r ages from the Pulur gneisses are regarded as cooling ages. Zircons from two gneiss samples from the Kazda~ Massif, dated by the stepwise Pb-evaporation method, gave an age of 308 • 16 Ma (Okay et al. 1996). To further refine the age of high-grade metamorphism in the Kazda~ Massif, we have dated a gneiss and an amphibolite from the Kazda~ Massif using the same method (Okay et al. 1996). Six zircon grains from the gneiss sample produced a relatively precise age of 319.2 ___ 1.5 Ma (early Late Carboniferous, latest Namurian), and one zircon grain from the amphibolite gave an age of 329 + 5 Ma (Fig. 7, Table 1). The isotopic data indicate high-grade metamorphism and associated deformation in the midCarboniferous (Namurian) in the Sakarya Zone. Permo-Carboniferous plutonism in the Sakarya Zone

Pre-Jurassic granitoids are common in the Sakarya Zone, although few are dated. The S6~tit granite in the western Sakarya Zone gave an A r - A r biotite plateau age of 290 • 5 Ma (CarboniferousPermian boundary, Okay et al. 2002) confirming earlier U - P b zircon and K - A r biotite ages (~o~ulu et al. 1965; ~o~ulu & Krummenacher 1967). K - A r biotite ages from the G6nen and Karacabey granites, east and west of Bandlrma, respectively, are

Table 1. Isotopic data from single-grain 2~ Lithology and sample number

Kazda~ gneiss K14{

Kazda~ amphibolite K4 ~amhk metagranite CL1

Grain 1 2 3 4 5 6 mean 1 t 1 / 2 t mean

/2~

in the range 286-298 Ma (Delaloye & Bing61 2000). These data indicate late orogenic acidic plutonism in the Sakarya Zone in the latest Carboniferous to early Permian period. The Variscan granites and high-grade metamorphic rocks in the Sakarya Zone were exhumed and unconformably overlain by the latest Carboniferous continental to shallow marine sedimentary rocks in the Eastern Pont• (Okay & Leven 1996; Okay & ~ahinttirk 1997; ~apklno~lu 2003) and in the Caucasus (e.g. Khain 1975).

Accretionary complexes in Anatolia: data on the spatial and temporal aspects of the Tethyan oceans There are widely differing views on the number, location, age span and name of the Tethyan oceans that existed during the Phanerozoic (e.g. Seng6r & Ydmaz 1981; Robertson & Dixon 1984; ~eng6r et al. 1984; Dercourt et al. 1986; Ricou 1994; Robertson et al. 1996; Stampfli et al. 2001, 2002). One way to approach this problem is through a biostratigraphic and isotopic study of the accretionary complexes. Because of their relatively low density the accretionary complexes have a wide preservation potential, and crop out widely in orogenic belts. The accretionary complexes may comprise three types of constituents: (1) pelagic sedimentary and basic magmatic rocks scraped at subduction

evaporation analyses of zircons from the basement of the Sakarya Zone, NW Turkey

Number of scans

2~176

Z~176

Mean value of 2~176 ratios

124 1l0 66 372 143 217

0.00210 0.00115 0.000255 0.000179 0.000115 0.000095

8.2 8.4 10.6 5.8 8.2 8.5

0.053178 + 102 0.052621 + 104 0.052407 • 130 0.052832 _+ 59 0.052792 • 90 0.052609 • 88

68 296 165

0.000071 0.000074 0.000336

7.7 12.0 12.1

0.053009 ___112 0.054640 _+ 37 0.054631 • 87

Errors are given at 95% confidence level and refer to the last digits.

Z~176 age (Ma) 336.4 • 4.4 312.5 + 4.5 303.2 • 5.7 321.6 • 2.5 319.9 • 3.9 312.0 _+ 3.8 319.2 • 1.5 329.2 • 4.8 397.6 • 1.5 397.2 +_ 3.6 397.5 • 1.4

OROGENS IN THE EASTERN MEDITTERANEAN zones from the downgoing oceanic crust; (2) greywacke and shale, which represent the trench infill; (3) blueschists and eclogites brought up along the subduction channel. The pelagic sedimentary rocks in the accretionary complexes provide an age range for the subducted ocean, whereas the greywackes and the isotopic age of the blueschists give an indication of the duration of subduction. The structural position of the accretionary complexes provides clues to the location of the associated oceans. Accretion ends by collision, or when the subduction zone is clogged by large oceanic or continental edifices, such as oceanic islands, oceanic plateaux, or isolated continental slivers (e.g. Cloos 1993). Termination of subduction in the Eastern Mediterranean south of Cyprus by the collision of the Eratosthenes Seamount provides a presentday example (e.g. Robertson 1998). The age of the accretionary complex can be defined as the age of subduction. At least four distinct accretionary complexes can be defined in the Balkan-Anatolian region. Karaburun-Chios

accretionary complex (Carboniferous)

Carboniferous accretionary complexes are found on the Karaburun peninsula and the adjacent island of Chios (Fig. 3; Stampfli et al. 1991, 2003). Both areas are situated in the Aegean on the northwestern margin of the Anatolide-Tauride Block immediately south of the Neotethyan Jzmir-Ankara suture. The complexes consist of strongly deformed siliciclastic turbidites, regarded as a Franciscan-type trench infill, which are unconformably overlain

397

by Lower Triassic basinal sedimentary and volcanic rocks (Robertson & Pickett 2000; Zanchi et al. 2003). The Lower Triassic pelagic sediments pass up into a typical Tauride carbonate platform of Triassic to Early Cretaceous age (Erdo~an et al. 1990). The intensely deformed and tectonically sliced and repeated turbidites comprise exotic limestone, radiolarian chert and volcanic blocks, up to kilometre scale, and Silurian-Carboniferous in age (Fig. 8; Kozur 1995). The age of the turbidite matrix is probably Early Carboniferous (Groves et al. 2003; Zanchi et al. 2003). Karakaya-Kiire accretionary complex ( T r i a s s i c - E a r l y Jurassic)

Triassic-Early Jurassic accretionary complexes are widely exposed in the Sakarya Zone below the Jurassic unconformity (Figs 3 and 4; Tekeli 1981; Ttiystiz 1990; Usta6mer & Robertson 1993, 1994; Pickett & Robertson 1996; Yflmaz et al. 1997; Okay 2000; Okay & G6nctio~lu 2004). In the western part of the Sakarya Zone they are attributed to the Karakaya Complex, and in the central Pontides to the Kiire Complex. Some Triassic palaeogeographical reconstructions show the Karakaya and Ktire accretionary complexes as belonging to different oceans separated by a continental sliver, attributed to the western part of the Istanbul Zone and to the northern parts of the Sakarya Zone (e.g. Usta6mer & Robertson 1993; Stampfli et al. 2001; Ziegler & Stampfli 2001). However, no coherent continental fragment can be defined

Fig. 8. A chronostratigraphic chart showing biostratigraphicand isotopic data from the Anatolian accretionary complexes. Data for the Karaburun complex are from Kozur (1995) and Groves et al. (2003); for the KarakayaKiire Complex from Kozur & Kaya (1994), Okay & Mosfler (1994), Kozur (1997), Okay & Moni6 (1997) and Okay et al. (2002); for the izmir-Ankara accretionary complexes from Bragin & Tekin (1996), Sherlock et al. (1999) and Tekin et al. (2002).

398

A.I. OKAYET AL.

between the outcrops of the Kiire and Karakaya complexes (Fig. 3). Furthermore, no Phanerozoic accretionary complex exists in the Istanbul Zone (Fig. 5). The Kiire and Karakaya complexes are similar in lithology, tectonostratigraphy and in structural position, but slightly differ in age, and will be treated together. The youngest palaeontological ages from the Karakaya Complex are latest Triassic (Leven & Okay 1996; Okay & Altmer 2004), whereas the age of the Kiire Complex extends to Early Jurassic (Kozur et aL 2000), and the complex is cut by granitoids of Mid-Jurassic age (Boztu~ et al. 1984; Yflmaz & Boztu~ 1986). Before the Cretaceous opening of the Black Sea, the Kiire Complex was contiguous with the Taurian Flysch of the Crimean Peninsula. The Karakaya-Ktire Complex consists of a lower metamorphic unit made up of a strongly deformed thrust stack of metabasitephyllite-marble with tectonic slices of ultramafic rock, broadly referred to as the Niliifer Unit. The depositional age of the Niltifer Unit, based on scarce conodonts in the marbles in NW Turkey, is Early to Mid-Triassic (Kaya & Mrstler 1992; Kozur et al. 2000). The geochemistry of the metabasites in the Niliifer Unit suggests a within-plate tectonic setting (Genq & Yalmaz 1995; Pickett & Robertson 1996, 2004; Genq 2004). The Niliifer Unit generally shows a high-pressure greenschist-facies metamorphism, although, in several localities in the Sakarya Zone, it also includes tectonic slices of blueschist and eclogite. The H P - L T metamorphic rocks in the Niliifer Unit are dated in the Bandlrma and Eski~ehir regions of NW Turkey (Fig. 3) as latest Triassic (205-203 Ma) using A r - A r method on phengites (Okay & Moni~ 1997; Okay et al. 2002). The structural setting and the lithological, metamorphic and geochemical features of the Niliifer Unit suggest an origin as an oceanic plateau or oceanic island, which was accreted to a Late Triassic active margin (Pickett & Robertson 1996, 2004; Okay 2000; Genq 2004). Recently, Topuz et al. (2004b) reported Early Permian (263-260Ma) R b - S r and A r - A r hornblende and muscovite ages from a metabasite-phyllite sequence from the Pulur region in the Eastern Pontides (Fig. 4). The metabasite-phyllite sequence, which is correlated with the Niliifer Unit, is tectonically overlain by the granulite-facies gneisses of mid-Carboniferous age (Okay 1996; Topuz et al. 2004b). If these isotopic data are confirmed then the subduction-accretion represented by the KarakayaKtire Complex will extend back to the Early Permian (Fig. 8). In the Sakarya Zone, the Niltifer Unit is overlain by Triassic to Lower Jurassic siliciclastic and volcanic sequences, which were strongly deformed, probably in a subduction zone setting, in the latest Triassic-earliest Jurassic (Okay 2000). In NW Turkey, the siliciclastic rocks comprise olistostromes with numerous Carboniferous and Permian shallow marine limestone blocks (Leven 1995; Leven & Okay 1996), and smaller numbers of Middle Carboniferous (Bashkirian), Permian and Triassic radiolarian chert and pelagic limestone exoticblocks (Fig. 8; Kozur & Kaya 1994; Okay & Mostler 1994; Kozur 1997; Kozur et aL 2000; Grnctio~lu et al. 2004). Olistostromes with the Carboniferous and Permian shallow marine limestone blocks form a belt, over 150km long and 5 - 1 0 k m wide, in NW Turkey immediately NW of the Izmir-Ankara suture (Fig. 3). The origin of the Permo-Carboniferous limestone blocks is controversial; the fauna in the blocks is interpreted either as Laurussian (Leven & Okay 1996) or as Gondwanan in origin (Altiner et al. 2000). Strandja accretionary complex (Late T r i a s s i c - E a r l y Jurassic)

The Triassic-Middle Jurassic epicontinental sediments of the Strandja Massif are tectonically overlain by a highly deformed volcano-sedimentary complex of siliciclastic turbidites, carbonates, mafic and acidic volcanic rocks of Early to Late Triassic

age (Chatalov 1980, 1988). Sengrr et al. (1984) and Ustarmer & Robertson (1993) interpreted the Strandja allochthons as an accretionary complex, although definite evidence for the oceanic origin of the Strandja allochthons (e.g. ultramafic rocks or deep-sea radiolarian cherts) is missing. Dismembered ophiolites of Late Jurassic age, including peridotite, gabbro and basalt, occur on the eastern margin of the Rhodope Massif (Fig. 3; Tsikouras & Hatzipanagiotou 1998), and are associated with a slightly metamorphosed epicontinental sequence of Triassic to Early Jurassic age (Kopp 1969). The ophiolites of this Circum-Rhodope Zone may represent the root zone of the Strandja allochthons. Although the Karakaya-Ktire and the Strandja accretionary complexes are similar in age, they differ markedly in their structural setting and lithology, and, as discussed below, are ascribed to different oceans. i z m i r - A n k a r a accretionary complex (Late Cretaceous)

Late Cretaceous accretionary complexes cover large regions in the Anatolide-Tauride Block south of the Izmir-Ankara-Erzincan suture (Figs 3 and 4; Okay 2000). They generally form tectonic imbricates sandwiched between the ophiolites above and the Anatolide-Tauride carbonate platform below. In many regions near the izmir-Ankara suture, such as north of Eski~ehir (Okay et al. 2002), east of Ankara (Koqyi~it 1991) and in the Tokat Massif (Bozkurt et al. 1997) the Izmir-Ankara accretionary complexes are imbricated with those of the Karakaya-Ktire Complex (Figs 3 and 4). The Late Cretaceous accretionary complexes consist mainly of basalt, radiolarian chert, pelagic shale and pelagic limestone, and in the old literature were often referred to as ophiolitic mrlange or coloured mrlange. In the north near the Izmir-Ankara suture, the accretionary complexes have undergone low-grade blueschistfacies metamorphism dated at c. 80 Ma (Sherlock et aL 1999). Palaeontological study of radiolarian cherts and pelagic limestones in these accretionary complexes has shown the presence of Triassic, Jurassic and Cretaceous rocks (Fig. 8; Bragin & Tekin 1996; Tekin et al. 2002). In contrast, no Palaeozoic pelagic sedimentary rocks were described in the Izmir-Ankara accretionary complexes, the oldest ones being Late Triassic (Late Carnian) radiolarian cherts (Tekin et al. 2002). Age and location o f the Tethyan oceans

Biostratigraphic and isotopic data from the accretionary complexes in Anatolia indicate the presence of several Tethyan oceans. A Tethyan ocean north of the Anatolide-Tauride Block, called the izmir-Ankara ocean, had a minimum age span from Mid-Late Triassic to the Cretaceous, suggesting an opening as early as the Early Triassic. This is in accord with the Triassic stratigraphy from the Karaburun Peninsula, which is indicative of rifting in the Early Triassic (Robertson & Pickett 2000). The widespread unconformity at the base of the Lower Triassic rocks in the northern margin of the Anatolide Tauride Block (e.g. Eren 2001; G6ncfio~lu et al. 2003) is probably also related to shoulder uplift before the rifting. The outcrop pattern of the Karakaya-Kiire Complex indicates an ocean in a similar position to the lzmir-Ankara ocean (e.g. north of the Anatolide-Tauride Block and south of the Istanbul Zone) but older (at least Mid-Carboniferous to Early Jurassic). The Karakaya-Kiire ocean must also have been located south of the Variscan basement of the Sakarya Zone, as the exotic blocks in the Karakaya-Kiire Complex were most probably derived from the south rather than the north (Okay 2000). The absence of continental fragments between the Karakaya-Kiire and Izmir-Ankara accretionary complexes (Figs 3 and 4) implies that the Karakaya-KiJre ocean corresponds to the main Palaeotethys (Fig. 9), rather than to a small back-arc basin as

OROGENS IN THE EASTERN MEDITTERANEAN

399

Fig. 9. Palaeogeographicalreconstructions of the Tethyan realm for the Late Carboniferous (a), Early Triassic (b), Late Triassic (c) and Late Jurassic (d), showingthe possible locations of the terranes and oceans discussed in the text. The general palaeogeographicalframework is taken from Stampfliet al. (2001).

shown in most models (e.g. ~engrr et aL 1984; Stampfli et aI. 2001). Biostratigraphic data from the Karaburun-Chios accretionary complexes suggest an ocean of Silurian to Carboniferous age, again situated north of the Anatolide-Tauride Block. The Karaburun-Chios accretionary complex may be related to the Variscan subduction (Zanchi et al. 2003), in which case it must have been displaced eastwards from its original position by strike-slip faulting (Fig. 9a). The Strandja allochthons indicate the presence of a Triassic to Early Jurassic ocean between the Strandja Massif in the north and the Rhodope-Serbo-Macedonian massifs in the south (Fig. 3). The Triassic stratigraphy of the Istanbul Zone indicates rifting in the Early Triassic. This ocean probably formed an eastern extension of the Hallstatt-Meliata ocean, described

farther west in the Eastem Alps and the Balkans (Kozur 1991; Channel & Kozur 1997). The relation between these Tethyan oceans and the surrounding continental terranes, which gave rise to the orogenic events, is discussed below.

Variscan orogeny in the Balkans and the Black Sea region The Late Carboniferous orogeny in the Pontides forms a link between the Variscan orogen in Central Europe and the Uralides of Eastern Europe. The Variscan orogeny comprises Carboniferous to Early Permian deformation, metamorphism and magmatism linked to the collision and amalgamation of Gondwana, Laurussia and the intervening terranes (e.g. Matte 2001; Wart 2002). The

400

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eastward extension of the Variscan orogen towards the Balkans and Anatolia is obscured by the strong overprint of the Alpide orogeny, or is concealed by the younger cover. The East European Craton is bordered in the south by a narrow tectonic belt, called the Scythian Platform, which is generally considered as a Late Palaeozoic (Early Carboniferous) orogen (e.g. Nikishin et al. 1998, 2001). The Palaeozoic stratigraphy in the Scythian Platform is concealed beneath the Mesozoic and younger strata, and there are only patchy data from a few boreholes, which indicate a thick Lower Devonian continental sandstone succession overlain by Middle Devonian to Lower Carboniferous shallow marine limestones. The overlying Vis6an-Namurian sequence consists of paralic and limnic deposits, and the Permian of red clastic rocks (Vaida & Seghedi 1997). During the Devonian and Carboniferous the Istanbul Zone and the Moesian Platform were adjacent to the Scythian Platform, and formed a south-facing passive continental margin (Figs 6 and 9a). The western part of the Istanbul Zone was the site of deep marine sedimentation in the Devonian and Carboniferous, in a continental slope setting, and hence was closer to the ocean compared with its eastern part and the Scythian-Moesian platforms (Fig. 6). The Late Carboniferous deformation in the Istanbul Zone is coeval with the high-grade metamorphism in the Sakarya Zone. It is plausible to relate the deformation and regional metamorphism to collision of the Laurussia margin with an ensialic arc represented by the basement rocks of the Sakarya Zone and possibly of the Strandja Zone. Absence of Palaeozoic magmatism in the MOIS terrane suggests southward subduction, which is compatible with the general north to NE vergence of Carboniferous deformation in the Istanbul Zone (G6rtir et al. 1997) and the northward migration of the coal deposition in the Moesian Platform (Tari et al. 1997). The ocean between the Laurussia margin and the Sakarya-Strandja microplate probably started to close by the Early Devonian, producing a magmatic arc represented by the ~amlkk Granite in the Sakarya Zone. The Late Carboniferous collision was followed by the latest Carboniferous-Early Permian plutonism in the core of the orogen in the Strandja and Sakarya zones, possibly linked to crustal thickening. The latest Carboniferous-Early Permian molasse deposition in the Eastern Pontides and the Caucasus marks the end of the Variscan orogeny in northern Turkey. The Intra-Pontide suture between the Istanbul and Sakarya zones probably links up with the Late Carboniferous Rheic suture in Central Europe (Fig. 9a; Ziegler & Stampfli 2001). The Variscan evolution of northern Turkey and the Balkans appears to be similar to that of NW Europe, with the Istanbul-Moesia-Scythian Block corresponding to Avalonia, and the Sakarya-Strandja zones to Armorica (Stampfli et al. 2002; Winchester & The PACE TMR Network Team 2002).

Early Triassic rifting and magmatism The Early Triassic is characterized by widespread tiffing and mafic magmatism in the Eastern Mediterranean region, possibly associated with mantle plumes (e.g. Dixon & Robertson 1999). The Istanbul Zone started to rift from the Sakarya Zone along the former Carboniferous suture, as shown by the deposition of earliest Triassic continental sandstones and conglomerates intercalated with basaltic flows (Fig. 9b). In the Mid-Triassic the Istanbul Zone became separated from the Sakarya Zone, as the rift turned into the Intra-Pontide-Meliata ocean. On the Gondwana side in the south, mafic magmatism was associated with the break-up of Permo-Carboniferous carbonate platforms, and the separation of the Anatolide-Tauride Block from Gondwana (Fig. 8b). Possibly a thin carbonate sliver, corresponding to the Cimmerian continent of ~eng6r et al. (1984), rifted away from the Anatolide-Tauride Block in the Early Triassic. Associated with this rifting, major intra-plate mafic magmatism occurred and an abnormally thick oceanic crust or oceanic plateau was created adjacent to the

passive continental margin. The northward drift of this narrow continental sliver is shown to close the Palaeozoic Tethys and open up the Mesozoic Tethys in the Triassic (Fig. 9b and c), although as discussed below there is no unequivocal evidence for the Cimmerian continent in the Pontides.

Cimmeride orogeny in the Pontides In Turkey deformation and metamorphism of latest Triassic to earliest Jurassic age is particularly marked in the Sakarya Zone. It is associated with the emplacement of large oceanic allochthons over the Variscan basement. In contrast, the Cimmeride deformation is weak, and the Cimmeride metamorphism is absent in the other Pontic zones, where this period is generally marked as an unconformity (Fig. 2). The cause of the Cimmeride orogeny in Anatolia was generally thought to be the collision and amalgamation of a Cimmerian continent with the Laurasian margin (e.g. ~eng6r 1984; ~eng6r et al. 1984). However, it has not been possible to define a Cimmerian continent in the field, which would have been readily recognized by its Gondwana-type stratigraphy, free of Variscan deformation and metamorphism. In many regions along the Izmir-Ankara suture the accretionary complexes of the Izmir-Ankara and Karakaya-Ktire oceans are tectonically intercalated with no evidence of an intervening continental fragment (Figs 3 and 4; Bozkurt et al. 1997; Okay et al. 2002). Apparently, the narrow Cimmerian continental sliver, responsible for the opening of the Izmir-Ankara ocean, was completely subducted, with only its Permo-Carboniferous limestone cover providing blocks to the accretionary complex. The Cimmeride orogeny in the Pontides was largely accretionary, caused by the collision and partial accretion of an oceanic plateau to the Laurasian margin during the latest Triassic (Fig. 9c; Okay 2000). This is compatible with the short duration of deformation and regional metamorphism observed in the Karakaya-Ktire Complex.

Late Jurassic Balkan orogeny Apart from the Strandja Massif, Late Jurassic deformation is strangely absent, or is marked by only a slight disconformity in the Pontic zones. The Late Jurassic was a period of opening of the Alpine Tethys in the west, where contractional deformation is also not reported. This leaves a relatively small space for the Balkan orogeny on the southern margin of Laurasia (Fig. 9d). The Balkan orogeny is possibly linked to the subduction of the Intra-Pontide-Meliata ocean between the Strandja and the Rhodope-Serbo-Macedonian massifs, and the ensuing collision (Fig. 9d). The north-vergent deformation in the Strandja Massif indicates a southward subduction under the Rhodope Massif, with the implication that the latest Jurassic-Early Cretaceous granitoids in the Serbo-Macedonian Massif were generated in a magmatic arc. The eastern part of the Intra-Pontide-Meliata ocean between the Sakarya and Istanbul zones did not close until the mid-Cretaceous, suggesting the existence of a transform fault between Pelagonia and the Sakarya Zone (Fig. 9d).

Conclusions The Pre-Alpide geological history of the Eastern MediterraneanBalkan region can be viewed as the episodic growth of Laurussia by the accretion of oceanic and continental terranes, interrupted by the opening of narrow back-arc basins on the southern margin of Laurussia. The continental terranes were invariably derived from Gondwana, and were accreted to Laurussia during the Late Ordovician-Early Silurian and Late Carboniferous,

OROGENS IN THE EASTERN MEDITTERANEAN

whereas a major accretion of oceanic crustal material occurred during the Late Triassic-Early Jurassic. Orogenic deformation associated with the Late OrdovicianEarly Silurian accretion of the Istanbul-Moesia-Scythian Platform (the MOIS terrane) is buried under young cover in the northern margins of the Black Sea. The Carboniferous accretion of the Strandja-Sakarya terrane to the Laurussian margin, along a south-dipping subduction zone, resulted in strong deformation, mid-Carboniferous metamorphism, and latest CarboniferousEarly Permian post-orogenic plutonism. The ensuing suture is probably an extension of the Rheic suture in Central Europe (Ziegler & Stampfli 2001). In contrast to these Palaeozoic continental collisions, a major accretion of oceanic crustal rocks occurred during the Late Triassic-Early Jurassic. The accretionary complexes in the Pontides comprise voluminous metabasic rocks with latest Triassic blueschist and eclogite ages. The Late Jurassic deformation and metamorphism, observed only in the Balkans, were the result of the closure of a narrow back-arc basin, the Meliata ocean between the Rhodope-SerboMacedonian and the Strandja massifs. In contrast to the polyorogenic history of the Pontides and the Balkans, the Anatolide-Tauride Block south of the I z m i r - A n k a r a suture was largely free of Palaeozoic-early Mesozoic deformations, except along its northwestern margin, where a Carboniferous accretionary complex has been recognized (Stampfli et al. 1991). Biostratigraphic and isotopic data from the Anatolian accretionary complexes and their structural position indicate the presence of three oceanic realms north of the Anatolide-Tauride Block during the Phanerozoic. Two of them correspond to the mid-Palaeozoic-Early Jurassic Palaeotethys, and Early Triassic-Tertiary Neotethys, respectively, both of which were subducted along the I z m i r - A n k a r a suture, which represents the main boundary between Laurussia and Gondwana. The third ocean, the Meliata-Intra-Pontide ocean, opened as a marginal back-arc basin on the Laurussian margin (e.g. Stampfli 2000). A major point from this review and that is also implicit in some recent studies (e.g. Dean et al. 2000) is the mobility of the small plates that make up Anatolia. The assumption of conjugate margins, common in the old Tethyan reconstructions (e.g. ~engrr & Ydmaz 1981; Robertson & Dixon 1984) is clearly not correct. Prior to the Tertiary, the Pontides and the Anatolide-Tauride Block never formed a single contiguous terrane, which is implicit in recent palaeogeographical reconstructions (e.g. Stampfli et al. 2001). Translation of continental terranes oblique to the rifted margin, and margin-parallel strike-slip faulting led to the juxtaposition of unrelated continental fragments. For example, in the Early Ordovician both the Anatolide-Tauride Block and the Istanbul Zone were probably located on the northern margin of Gondwana but separated by several thousand kilometres (see Dean et al. 2000). The sutures separating the terranes in the Eastern MediterraneanBalkan region were major zones of weaknesses, and were rejuvenated at various times. For example, the Intra-Pontide suture started as a Late Carboniferous suture, and later became the site of an Early Triassic rift, which developed into the MeliataIntra-Pontide ocean. This ocean closed in the Late Jurassic-Early Cretaceous, generating a second suture. In the Miocene the suture was reused by the North Anatolian Fault, which at present defines the northern margin of the Anatolia microplate. This study was partly funded by the Turkish Academy of Sciences. We thank J. Winchester, A. Saintot and L. Jolivet for constructive and helpful comments on the manuscript.

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OROGENS IN THE EASTERN MEDITTERANEAN

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Tectonic processes in the Southern and Middle Urals: an overview D. B R O W N 1, V. P U C H K O V 2, J. A L V A R E Z - M A R R O N 1, F. BEA 3, & A. P E R E Z - E S T A U N t

Xlnstitute of Earth Sciences 'Jaume Almera', CSIC, c/Llufs Sold i Sabarfs s/n, 08028 Barcelona, Spain (e-mail: dbrown @ija. csic. es) 2Ufimian Geoscience Center, Russian Academy of Sciences, ul. Karl Marx 16/2, Ufa 45000 Bashkiria, Russia 3Department of Mineralogy and Petrology, Fuentenueva Campus, University of Granada, 18002 Granada, Spain

Abstract: The tectonic evolution of the Uralide orogen began during the Late Palaeozoic as the continental margin of Baltica entered an east-dipping (today's coordinates) subduction zone beneath the Magnitogorsk and Tagil island arcs. The subsequent arc-continent collision resulted in the development and emplacement of an accretionarycomplex over the continental margin, the development and deformation of a foreland basin, and the extrusion of high-pressure rocks along the arc-continent suture. There is mounting evidence that, at about the same time as arc-continent collision was occurring along this margin of Baltica, eastward-directed subcontinental subduction of the Uralian oceanic crust was also taking place beneath the Kazakhstan plate. This subcontinental subduction is thought to have resulted in the formation of a continental volcanic arc. The final closure of the Uralian ocean basin and the start of collision between the Baltica and Kazakhstan plates occurred during the Late Carboniferous. This continent-continent collision resulted in development of the Late Carboniferous to Early Triassic western foreland fold and thrust belt and foreland basin of the Uralides. The foreland fold and thrust belt displays a large amount of basement involvement, extensive reactivation of pre-existing faults, and a small amount of shortening. At the same time, widespread strike-slip faulting accompanied by melt generation and granitoid emplacement took place in the interior part of the Uralides, leading to the transfer of material laterally along the strike of the orogen. The final crustal structure of the Uralides that resulted from the combination of all of these tectonic events is bivergent, with a crustal root reaching c. 53 km depth.

Extending for nearly 2500 km from near the Aral Sea in the south to the islands of Novaya Zemlya in the Arctic Ocean, the Uralide orogen of Russia marks the eastern boundary of the Early Palaeozoic continent Baltica and its collision zone with the Siberian and Kazakhstan plates during the Palaeozoic assembly of Pangaea. For descriptive purposes the Uralides have traditionally been divided into a number of longitudinal zones (Fig. la) that are largely based on the ages and palaeogeography of the dominant rocks within them (e.g. Ivanov et al. 1975; Khain 1985; Fershtater et al. 1988; Puchkov 1997). From west to east these zones are the Pre-Uralian zone, the West Uralian zone, the Central Uralian zone, the Magnitogorsk-Tagil zone, the East Uralian zone and the Trans-Uralian zone. Additionally, the Uralides have been divided geographically into the Southern, Middle, Northern, CisPolar and Polar Urals. The Pre-Uralian, West Uralian and Central Uralian zones contain syntectonic Late Carboniferous to Early Triassic sediments of the foreland basin, Palaeozoic platform and continental-slope rocks, and Archaean and Proterozoic rocks of the East European Craton (part of Baltica). These three zones were affected by Uralide deformation and make up the foreland thrust and fold belt (e.g. Kamaletdinov 1974; Brown et al. 1997b). The Magnitogorsk-Tagil zone consists of Silurian to Devonian intra-oceanic island arc volcanic rocks and overlying volcaniclastic sediments. The Magnitogorsk-Tagil zone is sutured to the former continental margin of Baltica along the Main Uralian fault. The East Uralian zone is composed predominantly of deformed and metamorphosed volcanic arc fragments with minor amounts of Precambrian and Palaeozoic rocks thought to represent continental crust (Puchkov 1997, 2000; Friberg et al. 2000b). The East Uralian zone was extensively intruded by Carboniferous and Permian granitoids (Fershtater et al. 1997; Bea et al. 1997, 2002), forming the 'main granite axis' of the Uralides. The East Uralian zone is juxtaposed against the Magnitogorsk-Tagil zone along the East M a g n i t o g o r s k - S e r o v - M a u k fault system. The Trans-Uralian zone is composed of Carboniferous volcano-plutonic complexes (Puchkov 1997, 2000). Ophiolitic material and high-pressure rocks have also been reported (Puchkov 2000). The contact between the East Uralian and Trans-Uralian zones is exposed only in the Southern Urals, where it is a serpentinite mrlange. Rocks that unequivocally belong to either the Kazakhstan or Siberia plates do not crop out in the Uralides.

It is generally accepted that the tectonic evolution of the Uralides (Hamilton 1970; Zonenshain et al. 1984, 1990; Puchkov 1997, 2000; Brown & Spadea 1999; Alvarez-Marron 2002; Bea et al. 2002) began with the development of intra-oceanic island arcs in the palaeo-Uralian ocean, which were then accreted to the margin of the East European Craton. Meanwhile, subcontinental subduction is thought to have been taking place along the margin of the Kazakhstan plate, forming Andean-type arcs. The Uralian orogeny began in the latest Carboniferous as the Uralian ocean basin closed and the Kazakhstan plate, followed by the Siberia plate, collided with Baltica. Continent-continent collision continued until the Early Triassic. With the exception of minor Triassic transtension, intra-plate volcanism, erosion and basin inversion during the development of the West Siberian Basin, the Uralide orogen has been preserved, relatively intact, since the Permian, providing an ideal place to study Palaeozoic orogenic processes. The aim of the paper is to summarize a number of the key tectonic processes that formed the Southern and Middle Urals (Fig. lb). It begins with the earliest recognizable event and progresses through time to the final crustal structure that is observable today. Emphasis is placed on two transects, which are focused around two deep seismic surveys, EUROPROBE's Seismic Reflection Profiling in the Urals (ESRU) survey in the Middle Urals and the multicomponent Urals Seismic Experiment and Integrated Studies (URSEIS) survey in the Southern Urals (Fig. lb).

Tectonic units and processes

Arc-continent collision (Mid-Devonian to Early Carboniferous) The Tagil and Magnitogorsk volcanic arcs developed during the Silurian-Devonian (Tagil) and the Early Devonian-Early Carboniferous (Magnitogorsk) in an intra-oceanic setting (Seravkin et al. 1992; Yazeva & Bochkarev 1996; Spadea et al. 1998, 2002; Brown & Spadea 1999; Herrington et al. 2002) and began to collide with the margin of Baltica in the late Mid-Devonian (Magnitogorsk) and the Early Carboniferous (Tagil) (Puchkov 1997; Brown & Spadea 1999). The Tagil arc, in the Middle Urals, is made up of Silurian andesitic basalts and Lower Devonian trachytes and volcaniclastic rocks, overlain by 2000 m of Lower

From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 407-419. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. (a) Map showing the zones of the Urals and its geographical divisions from north to south. The area discussed in this paper is indicated by the box. (b) Geological map of the Southern and part of the Middle Urals. The legend shows the disposition of the various tectonic units discussed in this paper. The locations of the cross-sections in Figures 2 and 5 are shown, as is the location of Figure 6 and the ESRU and URSEIS seismic profiles.

and Middle Devonian limestone that, in the east, is intercalated with calc-alkaline volcanic rocks (Antsigin et al. 1994; Yazeva & Bochkarev 1994). The Tagil arc has been deformed and folded into an open synformal structure (e.g. Bashta et al. 1990; Ayarza et al. 2000b) and has been metamorphosed to lower greenschist facies. By far the best preserved and exposed, and therefore the most studied of the Uralide arcs, is the Magnitogorsk arc in the Southern Urals. It is composed of Emsian boninite-bearing arc-tholeiites in the forearc region, followed by Emsian to Givetian arc-tholeiite to calc-alkaline volcanic rocks of the Irendyk volcanic front; all of which display a clear intra-oceanic island arc signature (Fig. 2a; Seravkin et al. 1992; Spadea et al. 1998, 2002; Brown & Spadea 1999; Herrington et al. 2002). These volcanic units form the basement on which up to 5000 m of Frasnian- to Famennian-age forearc basin volcaniclastic sediments were deposited (Fig. 2a; Maslov et al. 1993; Brown et al. 2001). Lower Carboniferous shallow-water carbonates and, locally, basalt-rhyolite volcanic rocks unconformably overlie the arc edifice. Locally, Lower Carboniferous granitoids intrude the arc. Deformation in the Magnitogorsk volcanic arc is low, with only minor open folding and thrusting (Brown et al. 2001). The metamorphic grade barely exceeds sea-floor metamorphism. In the Southern Urals, a well-preserved accretionary complex developed during the Magnitogorsk arc-continent collision (Figs lb and 2b) (e.g. Bastida et al. 1997; Brown et al. 1998;

Brown & Spadea 1999; Alvarez-Marron et al. 2000). The accretionary complex is composed of Silurian to Middle Devonian continental slope and platform sedimentary rocks (Suvanyak Complex) that were detached from the East European Craton, and were overthrust by c. 5 km of late Frasnian and Famennian syncollisional volcaniclastic turbidites (Zilair nappe) sourced predominantly from the accretionary complex to the east with minor input from the Magnitogorsk arc (e.g. Puchkov 1997; Brown et al. 1998; Brown & Spadea 1999; Alvarez-Marron et al. 2000; Willner et al. 2002) (Figs lb and 2b). These units are flanked to the east by eclogite- and blueschist-bearing gneisses of the Maksutovo Complex that record a peak metamorphic pressure and temperature of 20 _+ 4 kbar and 550 __ 50 ~ (Beane et al. 1995; Hetzel et al. 1998; Schulte & Blfimel 1999), and a peak metamorphic age of c. 380-370 Ma (Fig. 2c; Matte et al. 1993; Lennykh et al. 1995; Beane & Connelly 2000; Glodny et al. 2002). Recently, microdiamond aggregates have been described from the Maksutovo Complex, suggesting that even higher pressures were achieved than those recorded by the metamorphic mineral assemblages (Bostick et al. 2003). The highest structural level of the accretionary complex is the Sakmara Allochthon in the south and the Kraka lherzolite massif in the north. The accretionary complex is at present sutured to the Magnitogorsk arc along the east-dipping Main Uralian fault zone, a m~lange that contains several kilometre-scale ultramafic fragments, one of which records

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409

Fig. 2. (a) Geochemicaland isotope data for Magnitogorskextrusive rocks. Plots of Emsian age-corrected Nd and Sr isotope ratios for Baimak-Buribai, Irendyk and Karamalytashformationsshow depleted mantle sources and secondary radiogenic Sr enrichment.Th/Yb v. Ta/Yb plot shows mostly intraoceanic arc affinities.A stratigraphiccolumn for forearc basin stratigraphyis also shown. After Brown & Spadea (1999). (b) Upper crustal cross-sectionacross the Magnitogorskforearc and the accretionary complex showing the structural architectureof the arc-continent collision zone in the SouthernUrals (after Mvarez-Marron et al. 2000). The location of the section is shown in (a). (e) Radiometric age determinationsfrom the Mindyakand Maksutovo complexes and a P - T path of the lower unit of the Maksutovo Complex. Data are taken from the sources discussed in the text. The upper path is for a garnet-mica schist and the lower path for an eclogite. The open arrows indicate a generalized retrograde path.

metamorphism under mantle conditions (Savelieva & Nesbitt 1996; Savelieva et al. 1997, 2002; Scarrow et al. 1999). The geochemistry of the Magnitogorsk arc volcanic rocks (Spadea et al. 1998, 2002; Herrington et al. 2002), the structure of the accretionary complex and its forearc (Brown et al. 1998, 2001; Alvarez-Marron et al. 2000), the high-pressure rocks beneath and along the suture zone (e.g. Hetzel et al. 1998; Hetzel 1999; Beane & Connelly 2000; Brown et al. 2000), and the ophiolitic, mafic and ultramafic material (Savelieva et al. 1997, 2002; Scarrow et al. 1999) show that the Palaeozoic tectonic processes that went into its formation can be favourably compared with those in currently active settings such as the west Pacific (Fig. 3a; Puchkov 1997; Brown et al. 1998; Brown & Spadea 1999; Herrington et al. 2002; Spadea et al. 2002). For example, boninitic lavas found in the oldest arc volcanic units provide a geodynamic marker that records the initiation of intra-oceanic subduction and the early development of the arc (Spadea et al. 1998; Brown & Spadea 1999). High-pressure rocks along the backstop of the accretionary complex were in part derived from continental margin material (Hetzel 1999), and the Mid-Devonian age of the high-pressure metamorphism provides a constraint for determining the timing of the entry of the continental crust into the subduction zone (Brown et al. 1998). The pressure,

temperature and thermochronology of the Maksutovo Complex and other high-pressure rocks along the arc-continent suture provide evidence for the flux of material in the subduction zone channel during its evolution (Fig. 3b; Brown et al. 2000). The sediments overlying the volcanic arc record (near) surface processes such a growth folding (Brown et al. 1998, 2001; Alvarez-Marron et al. 2000). The widespread occurrence of debris flows within the Late Devonian Zilair formation is thought to represent seismic events (seismites), and may be related to the arrival of the full thickness of the continental crust at the subduction zone (Brown et al. 2001). The accretionary complex was subsequently reworked during the formation of the foreland fold and thrust belt (see below).

S u b d u c t i o n b e n e a t h the K a z a k h s t a n p l a t e (Late D e v o n i a n to Late Carboniferous)

To date, little is known about what happened along the margin of the Kazakhstan plate prior to or during its collision with Baltica, as no rocks that can be unequivocally assigned to its plate margin have been recognized in the Uralides. Nevertheless, some recent studies suggest the presence of a continental volcanic arc that may have developed on the active margin of the Kazakhstan

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Fig. 3. (a) The early convergent history in the Southern Urals is marked by the generation of boninite-bearing arc-tholeiites in the Magnitogorsk forearc (T1), followed by arc-tholeiite to calc-alkaline volcanism. With the entry of the East European Craton continental crust into the subduction zone, volcanism waned and stopped, and high-pressure metamorphism of its leading edge took place (T2). The arrival of the full thickness of the continental crust at the subduction zone is marked by increased sedimentation in the forearc basin and deposition of arc-derived volcaniclastic turbidites across the subducting slab (T3). These, together with offscraped continental material, the exhumed high-pressure rocks, and a lherzolite massif, formed an accretionary wedge. A broad m61ange zone containing ultramafic fragments separates the forearc basement from the accretionary wedge, and marks the damage zone that developed along the backstop region. From Brown & Spadea (1999). (b) T l: during the Early Devonian suprasubduction-zone material was subducted to upper mantle depths. T2: by the end of the Early Devonian, when the East European Craton appears at the subduction zone, steady-state intra-oceanic subduction was under way. Geotherms are from van den Beukel (1992). T3: with the entrance of the East European Craton into the subduction zone the thermal regime would have departed from steady state. Geotherms are from van den Beukel (1992) for a continental heat flow of 70 mW m -2. The dotted line indicates van den Beukel's continental crust, whereas we have chosen to show a thinned continental crust (dark grey). The lowest frame shows an enlargement of the area shown in the box in T3. When the downgoing slab had reached a depth of 50-70 km, the Proterozoic sediments with a quartz rheology were detached, interacted with the mantle wedge, and the exhumation history began.

plate. In particular, data f r o m granitoids in the East Uralian zone point in this direction (Bea et al. 2002). A n u m b e r o f Uralide granitoids formed in what is thought to be two subduction settings from the Late D e v o n i a n to Late Carboniferous (Fig. 4; B e a et al. 1997, 2002; Montero et al. 2000). The first subduction-related m a g m a t i s m occurred from about 370 M a to 350 Ma, and is found in the eastern sector o f the East Uralian zone. B e a et al. (2002) interpreted this phase of m a g m a t i s m to have b e e n related to an east-dipping subduction zone located to the east o f the accreted Magnitogorsk arc, and to have produced I-type granitoids such as those o f the C h e l y a b y n s k and the Chernorechensk batholiths. A n older continental c o m p o n e n t in these granitoids can be interpreted to be the result o f their formation on the continental margin o f the Kazakhstan continent (Bea et al. 2002). A second phase of subduction m a g m a t i s m occurred from about

335 M a to 315 Ma, and is found in the western part of the East Uralian zone, between 55~ and 58~ (Bea e t al. 2002). B e a et al. (2002) have interpreted this phase to have been related to a subduction zone located to the east of the accreted Tagil arc, and that dipped eastward underneath the older continental arc. This subduction event produced batholiths c o m p o s e d of I- and M-type granitoids with little, if any, continental component. Magmatic activity directly related to subduction e n d e d before the Permian. Friberg et al. (2000b) h a v e described mafic to felsic gneisses o f largely Silurian and D e v o n i a n age (note, h o w e v e r , that there are large errors on the age determinations) and v o l c a n o - s e d i m e n t a r y rocks in the East Uralian zone that have b e e n interpreted to represent a volcanic arc complex. It is into this arc c o m p l e x that the above-discussed granitoids intrude, suggesting that the gneisses m a y be a deep, m e t a m o r p h o s e d part of the arc. In the

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411

Fig. 4. (a) Schematic map of the Southern and Middle Urals, outlining the late orogenic strike-slip fault system and the location of subduction-type granitoids. (b) Continental-crust normalized REE plots of Early Carboniferous subduction-related granitoids (from Beaet al. 2002). These subduction granites are enriched in trace elements of continental affinity such as Rb, Ba, Th, U and Li, suggesting that the protolith was composed of oceanic materials plus a significant fraction of old crustal materials. (e) eNa(t) V. esr(t) of Uralide subduction granitoids (from Beaet al. 2002). Neither 87Sr/S6Sr(t) nor 143Nd/144Nd(t) values bear any relation to the age, but depend on the geographical longitude. The Early Carboniferous batholiths in the east, at Chelyabinsk and Chernorechensk, are composed of granitoids with significantly higher 87Sr/86Sr(t) but lower 143Nd/144Nd(t)than similar rocks of the Late Carboniferous batholiths in the west, which have identical (in some cases more primitive) 87Sr/86Sr(t) and only slightly lower 143Nd/144Nd(t) compared with oceanic plagiogranites.

Southern Urals, however, the East Uralian zone is primarily composed of amphibolite-facies metapelites that are thought to represent continental crust (e.g. Puchkov 1997, 2000). The presence of Early Carboniferous subduction-related granitoids in this zone may indicate that the continental crust was part of the Kazakhstan plate at some stage. Finally, the eastern parts of the URSEIS and ESRU seismic reflection profiles image west-dipping reflectivity throughout the crust of the Trans-Uralian zone which has been interpreted to possibly represent east-vergent structures related to imbrication along the margin of the Kazakhstan plate (Tryggvason et al. 2001; Brown et al. 2002). The f o r e l a n d f o l d a n d thrust belt (Late C a r b o n i f e r o u s to E a r l y Triassic)

The foreland fold and thrust belt of the Middle and Southern Urals (which includes the Pre-Uralian, West Uralian and Central Uralian zones) contains syntectonic Late Carboniferous to Early Triassic sediments of the foreland basin, Palaeozoic platform and slope sediments, the Archaean and Proterozoic basement of Baltica, and the a r c - c o n t i n e n t collision accretionary complex. The foreland fold and thrust belt developed from the Late Carboniferous to the Late P e r m i a n - E a r l y Triassic (Kamaletdinov 1974; Brown et al. 1997b; Puchkov 1997). The foreland fold and thrust belt

between c. 56~ and 59~ is a narrow, north-south-trending, west-verging basement-involved thrust stack measuring c. 5 0 75 km in width from the Main Uralian fault (the a r c - c o n t i n e n t suture) to the frontal folds (Fig. 1). In this area it is flanked to the east by the Precambrian-cored Kvarkush Anticline, and to the west by the foreland basin (Fig. 1). Balanced cross-sections and the amount of shortening have not been determined for this part of the orogen, and farther discussion of it is beyond the scope of this paper. By far the best studied area of the foreland fold and thrust belt is in the Southern Urals (from c. 56~ to 51~ where its architecture has often been compared with that of other thrust belts from around the world, especially that of the Appalachians (e.g. Kamaletdinov 1974; Kruse & McNutt 1988; Rodgers 1990). However, recent structural mapping and seismic reflection data have shown the southern Uralides to be different (see below) (Brown et al. 1997b, 1998, 1999; Perez-Estaun et al. 1997; Giese et al. 1999; Alvarez-Marron 2000; Alvarez-Marron et al. 2002). Between c. 56~ and 53~ the Southern Urals foreland fold and thrust belt is a c. 150 km wide, west-vergent thrust wedge made up of Precambrian basement in the Bashkirian Anticline, the accretionary complex, Palaeozoic platform and foreland basin sediments (Figs 1 and 5). Palaeozoic shortening in this part of the thrust belt is c. 20 km or less (Fig. 5a and b; Brown et al. 1996, 1997b, 1999; Perez-Estaun et al. 1997; Giese et al. 1999). South of 53~ the foreland fold and

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Fig. 5. (a) Balanced and restored cross-section across the central Bashkirian Anticline (after Perez-Estaun et al. 1997). The calculated shortening is c. 20 km. The location is shown in Figure 1. (b) Balanced and restored cross-section across the southern Bashkirian Anticline (after Brown et al. 1997b). The calculated shortening is c. 17 kin. The location is shown in Figure lb. (c) Surface slope v. basal dip angle relationships for critical wedges. Calculation of the surface slope of section 3 using the equation a = arctan(tan f i / H ~ ) - fi (where a is the surface slope, fl is the basal slope and H is shortening) yields an c~ value of 1.1 ~ for a taper (r -----ct + t ) of 4.1 ~ This value of r requires only a small amount of material to have been eroded from the frontal part of the belt, and is in agreement with fission-track data (Seward et al. 1997; 2002). It also yields realistic values for average strain (AR = (tan r/tan/3)) of 1.3 : 1, and would place the section within the subcritical field. (See Brown et al. (1997a) for further explanation).

TECTONICS IN SOUTHERN AND MIDDLE URALS

thrust belt is dominated by the Southern Urals accretionary complex (Brown et al. 1998, 2004; Alvarez-Marron et al. 2000; Fig. lb). Cross-sections by Brown et al. (2004) indicate a very different structural style from that of the thrust belt to the north, in the Bashkirian Anticline, although the amount of shortening in this part of the thrust belt also appears to be small. The Uralides foreland fold and thrust belt exhibits a number of features that differentiate it from other Palaeozoic thrust belts. For example, the amount of shortening is very small, with vertical displacement along faults nearly equal to horizontal displacement (Brown et al. 1997b; Perez-Estaun et al. 1997). Mechanically, the thrust belt may never have reached a critical taper, and developed as a subcritical wedge (Fig. 5c; Brown et al. 1997a). The amount of basement involvement is high, and in many cases thrusting appears to have been localized by reactivation of two sets of pre-existing structures in the basement (Brown et al. 1997b, 1999; Perez-Estaun et al. 1997; Giese et al. 1999). Reactivation of structures parallel to the developing Uralide structural grain resulted in the incorporation of crystalline thrust sheets into the thrust belt at an early stage in its development, whereas those at a high angle to the Uralide structural grain influenced the location and development of lateral structures that can explain along-strike structural changes (Fig. 6; Perez-Estaun et al. 1997; Brown et al. 1999). The small amount of shortening, together with the localization of thrusts along pre-existing structures, suggests that the basal detachment may also be controlled by a Precambrian feature within the basement, or is absent completely.

413

Late o r o g e n i c strike-slip f a u l t i n g a n d granitoid e m p l a c e m e n t (Late C a r b o n i f e r o u s to E a r l y Triassic)

The internal part of the Uralides is made up of a late orogenic strike-slip fault system (e.g. Echtler et al. 1997; Friberg et al. 2002; Hetzel & Glodny 2002) that extends for more than 700 km along the Uralides before it disappears beneath Mesozoic cover in the south and north (Fig. 7a). Throughout much of the Middle and Southern Urals this strike-slip fault system corresponds to the East Uralian zone, although the currently defined Main Uralian fault in the Middle Urals appears to be its western limit there (Ayarza et al. 2000a; Brown et al. 2002). Dating on one segment of the strike-slip fault system indicates a Late Permian to Early Triassic age (247-240 Ma) for the development of fault-related mylonites (Hetzel & Glodny 2002), and latest Carboniferous (305-291 Ma) ages for associated metamorphic rocks (Echtler et al. 1997; Eide et al. 1997). The late orogenic strike-slip fault system was extensively intruded by latest Carboniferous to Permian granitoids, first in the southern part (292-280 Ma) and then in the northern part (270-250 Ma; Fig. 7; Bea et al. 1997, 2002, 2006; Montero et al. 2000). In general, the granitoids were emplaced at a high level in the crust, at c. 12-15 km depth (Fershtater et al. 1997). These granitoids have a high SiO2 content, and are mildly peraluminous, with elevated Rb, Cs, Ba, Th and U contents (Fig. 7b), but with an unusually primitive Sr and Nd isotopic composition (Fig. 7c) (Bea et al. 1997 2002; Fershtater et al. 1997;

Fig. 6. (a) Geological map of the northern Bashkirian Anticline (location is shown in Fig. lb). (b) Simplified, balanced and restored cross-sections across the northwestern part of the Bashkirian Anticline (locations are shown in (a)) (after Brown et al. 1999). Comparing the map and the cross-sections, it should be noted how the Yurmatu anticline changes abruptly along strike into the Inzer syncline. Such a change is strongly indicative of a lateral structure (the Inzer lateral ramp in (a)). Also, the Karatau fault is an excellent example of a lateral structure across which displacement is transferred toward the foreland. (c) Schematic block diagram showing the relationships between hanging-wall structures and basement topography (after Brown et al. 1999). It should be noted that, because of problems in the projection, the Karatau fault and the Inzer lateral ramp have not been drawn in their true orientation relative to the transport direction; in reality they are somewhat oblique to the orientation shown.

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Fig. 7. (a) Schematic map of the Southern and Middle Urals outlining the late orogenic strike-slip fault system and the location of continental-type granitoids. (b) Continental-crust normalized trace element and REE plots of Permian collision-related granitoids. [3, gabbros; O, diorites; ~ , granodiorites; open crosses, granites. For Dzhabyk, additionally, crossed squares and open circles represent the Mochagi and Rodnichki quartz monzonites, respectively (from Bea et al. 2002). The average Permian granite of the Uralides has a trace element composition characteristic of continental granites, in which some trace element anomalies characteristic of arc magmas, although attenuated, are still recognizable. The only materials able to produce partial melts with this conjunction of mantle-like isotope and crust-like chemical composition are subduction-related rocks with a short crustal residence time of a few tens of million years. (c) eNd(t) V. esr(t) of Uralide continental granitoids (from Bea et al. 2002). The isotopic signature of the Permian continental-type granitoids is very primitive, with 87Sr/86Sr(t) and 143Nd/la4Nd(t) values that match the subduction granites. This feature excludes continental materials older than Silurian as a possible protolith.

Montero et al. 2000; Gerdes et al. 2002, pp. 3-19). Bea et aL (2002) interpreted this to have resulted from recycling of the older continental arc material that was deeply buried after the collision; they also interpreted Permian crustal melting to be the result of a combination of radiogenic heating of an overthickened sialic crust, from local underplating by mafic magmas, and from local accumulation of heat and fluids related to the oblique, crustal-scale strike-slip shear zones that finally assembled the Uralides. The existence of a late orogenic strike-slip fault system along the entire interior of the Uralides suggests that widespread mass transfer took place along the axis of the orogen during the late stages of its tectonic evolution. Estimates of displacement along some strands of this fault system range from a few tens of kilometres to more than 100 km (Ayarza et al. 2000a; Hetzel & Glodny 2002). The presence of high-grade metamorphic rocks near the surface at the time of granitoid generation suggests extensive exhumation of material from the lower crust and its emplacement into the upper crust. The widespread melting of deep crustal material and its subsequent emplacement in the upper crust is also indicative of mass transfer. Both these processes are also suggestive of heat transfer from the lower crust, as hot material in the form of granulites and melt ascends and is emplaced in the colder upper crust. The evolution of the melt emplacement from south to north is also suggestive either of heat transfer along strike in the

orogen, or differential heating from south to north. Much work is needed on the structure, kinematics, granitoids, and geochronology of this important strike-slip fault system before it is possible to fully understand its relevance to orogen-parallel mass and heat transfer during the late stages of the Uralian orogeny.

F i n a l c r u s t a l s t r u c t u r e ( L a t e T r i a s s i c to R e c e n t )

The ESRU (Juhlin et al. 1998; Fig. 8a), URSEIS (Berzin et al. 1996; Fig. 8b), and reprocessed Russian seismic reflection or refraction surveys provide significant new data for interpreting the crustal structure of the Uralides (Steer et al. 1995, 1998; Carbonell et al. 1996, 1998, 2000; Echtler et al. 1996; Knapp et al. 1996; Friberg et al. 2000a, 2002; Brown et al. 2002). In the Southern (URSEIS) and Middle (ESRU) Urals the East European Craton crust thickens eastward from c. 40 km to c. 48 km, and is imaged by subhorizontal to east-dipping reflectivity that can be related to its Palaeozoic and older evolution (Fig. 8). The suture zone between Baltica and the accreted terranes, the Main Uralian fault, is poorly imaged in the URSEIS section, but in the ESRU section it is imaged as a zone of east-dipping reflectivity that extends from the surface into the middle crust; it marks an abrupt change to weakly subhorizontal reflectivity in the Tagil

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415

Fig. 8. (a) Interpreted line drawings of the coherency filtered, depth-migrated ESRU data (after Brown et al. 2002). (See Fig. lb for location). The main suture zones that bind the tectonic units together have been interpreted to end at the Moho. Their exact location at depth cannot be unambiguously interpreted, and has therefore been shown as a zone in which they may possibly occur. The location of the UWARS wide-angle Moho is from Juhlin et al. (1998). (b) Interpreted line drawings of the coherency filtered, depth-migrated URSEIS vibroseis data (after Tryggvason et al. 2001). (See Fig. lb for location). The main suture zones that bind the tectonic units together have been interpreted to end at the Moho. Their exact location at depth cannot be unambiguously interpreted, and has therefore been shown as a zone in which they may possibly occur. The location of the URSEIS explosion-source reflection Moho (Steer et al. 1998) and the refraction Moho (Carbonell et al. 1998) are shown along with the Moho imaged in this dataset.

arc (e.g. Ayarza et al. 2000a). East of the Main Uralian fault, the Magnitogorsk (Southern Urals) and the Tagil (Middle Urals) volcanic arcs display moderate to weak upper crustal reflectivity, and diffuse middle to lower crustal reflectivity. The Moho beneath both arc complexes is poorly imaged in the reflection data, but based on refraction data is interpreted to be at 5 0 - 5 5 km depth (Fig. 8; Thouvenot e t al. 1995; Juhlin e t al. 1996; Carbonell e t al. 1998). East of the arc complexes, the wide zone of anastomosing strike-slip faulting and granitoids of the East Uralian zone is imaged in the seismic sections as clouds of diffuse reflectivity interspersed with, or cut by sharp, predominantly west-dipping reflections. In the Southern and Middle

Urals, west-dipping reflectivity of the Trans-Uralian zone extends from the middle crust into the lower crust, where it appears to merge with the Moho (EchOer e t al. 1996; Knapp e t al. 1996; Steer e t al. 1998; Friberg et al. 2000a, 2002; Brown et al. 2002). The URSEIS experiment imaged a number of sub-Moho reflections (Knapp e t al. 1996; Steer et al. 1998) that may represent deformation scars related to the Uralian orogeny. The overall seismic reflection pattern of the Uralide crust as imaged by the URSEIS and E S R U data is bivergent, perhaps representing the original collision-related crustal architecture (Fig. 9). With the exception of possible minor extensional features in the eastern part of the E S R U section (Friberg e t al. 2002), there

Fig. 9. Generalized crustal-scale structural cross-section of the Southern Urals along the URSEIS profile. The location of the Moho is from Carbonell et al. (1998) and Steer et al. (1998).

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D. BROWNETAL.

Fig. 10. Schematicplate model for the Southem Urals, outlining the key tectonic processes over time that went into the building of the Uralides.

is little evidence in the seismic reflection fabric for large-scale extensional collapse of the Uralides. The URSEIS and ESRU sections indicate that crustal thickness and Moho topography change somewhat between the Southern and Middle Urals, although the crustal root can be seen to extend along the western volcanic axis of the orogen. Recently, Diaconescu & Knapp (2002) argued that the formation of eclogite in the root zone may have led to an isostatically balanced system that ultimately preserved the Uralide structure. However, petrophysical modelling of the Uralide crust along the URSEIS transect indicates that the root zone is made up of mafic garnet granulite and not eclogite (Scarrow et al. 2002; Brown et al. 2003), so perhaps other as yet unidentified processes have been active. The Uralide orogen records a long and complex subduction-accretion history (e.g. arc-continent collision along the margin of Baltica, Andean-type subduction beneath Kazakhstan) prior to the final collision that gave it its final bivergent architecture. The complex late orogenic history, which involved extensive wrench faulting accompanied by widespread melt generation and granitoid emplacement in the interior of the orogen (see above), probably significantly overprinted and/or reworked much of the subduction- and accretionrelated tectonic fabric, giving this zone its varied and complex reflection seismic character. For example, orogen-parallel mass transport of material, as outlined above, may account for the subhorizontal reflectivity in the lower crust imaged in the ESRU seismic reflection profile (Koyi et al. 1999).

Conclusions The Uralide orogen of Russia was one of the main orogens built during the Palaeozoic assembly of Pangaea. Unlike the Variscide-Appalachian orogenic system, which was largely rifted apart by the opening of the Atlantic Ocean or extensively overprinted by post-orogenic processes, the Uralides have been preserved intact, providing an opportunity to study the tectonic processes that went into forming this Palaeozoic orogen. Clearly, subduction and accretion processes dominated during the Mid-Devonian to Early Carboniferous, as intra-oceanic island arcs collided with Baltica. During the same time period, we interpret that Andean-type continental arc(s) were forming on the margin of Kazaldastan (Fig. 10). The Southern Urals is of particular importance in the subduction and accretion history of the Uralides because it contains one of the best preserved examples of an arc-continent collision in any Palaeozoic orogen. The state of preservation and the level of exposure allow this arc-continent collision to be compared in detail with those that are currently active around the world, providing unprecedented insight into Palaeozoic tectonic processes. With the closure of the Uralian ocean, deformation began in the western Uralides foreland fold and thrust belt and, concomitantly, deposition of the foreland basin began (Fig. 10). The Uralides foreland fold and thrust belt is distinct from most other thrust belts, in particular in the amount of shortening, the amount of basement

TECTONICS IN SOUTHERN AND MIDDLE URALS

involvement, and the along-strike structural changes. W h y the shortening is so small is not clear. Perhaps the far-field stress induced by a highly oblique c o n t i n e n t - c o n t i n e n t collision was too small to imbricate the dense crust of the island arc systems that formed the margin at that time, and merely resulted in the reactivation of earlier structures in the basement. Whatever the reason, the small amount of shortening allows the relationship between the pre-existing basement structures and changes in structural style to be correlated. At the same time as the foreland fold and thrust belt was forming, the interior part of the orogen underwent extensive strike-slip faulting, metamorphism, melt generation and emplacement, and exhumation (Fig. 10). Finally, the bivergent crustal structure of the Southern and Middle Urals reflects the crustal stacking that occurred on both sides of the orogen during the subduction and accretion stage and during the c o n t i n e n t - c o n t i n e n t collision stage.

References ALVAREZ-MARRON, J. 2002. Tectonic processes during collisional orogenesis from comparison of the Southern Urals with the Central Variscides. In: BROWN, D., JUHLtN, C. & PUCHKOV,V. (eds) Mountain Building in the Uralides: Pangea to Present. Geophysical Monographs, American Geophysical Union, 132, 83-100. ALVAREZ-MARRON, J., BROWN, D., PEREZ-ESTAUN, A., PUCHKOV,V. & GOROZHANINA, Y. 2000. Accretionary complex structure and kinematics during Paleozoic arc-continent collision in the Southern Urals. Tectonophysics, 325, 175-191. ANTSmtN, N. I., ZOLOEV, K. K., KLUZINA, M. L., ETAL. 1994. Descriptions of the stratigraphic schemes of the Urals. Urals Stratigraphic Committee, Ekaterinburg [in Russian]. AYARZA, P., BROWN, D., ALVAREZ-MARRON, J. & JUHLIN, C. 2000a. Contrasting tectonic history of the arc-continent suture in the Southern and Middle Urals: implications for the evolution of the orogen. Journal of the Geological Society, London, 157, 1065 - 1076. AYARZA, P., JUHLIN, C., BROWN, D., ETAL. 2000b. Integrated geological and geophysical studies in the SG4 borehole area (Tagil volcanic arc--Middle Urals): location of seismic reflectors and source of the reflectivity. Journal of Geophysical Research, 105, 21333-21352. BASHTA, E. G., KUSHKOV, V. N., SHATORNAYA,L. N., GLUSHKOV,A. N. & SERGEEV, V. A. 1990. First results from the drilling and investigations of the Urals superdeep borehole (SG-4). Mineral Geology SSSR, 19-30 [in Russian]. BASTIDA, F., ALLER, J., PUCHKOV, V. N., JUHLIN, C. & OSLIANSKI, A. 1997. A cross section through the Zilair Unit (southern Urals). Tectonophysics, 276, 253-264. BEA, F., FERSHTATER, G., MONTERO, P., SMIRNOV, V. & ZIN'KOVA, E. 1997. Generation and evolution of subduction-related batholiths from the central Urals: constraints on the P - T history of the Uralian orogen. Tectonophysics, 276, 103-116. BEA, F., FERSHTATER, G. & MONTERO, P. 2002. Granitoids of the Urals: implications for the evolution of the orogen. In: BROWN, D., JUHLIN, C. & PUCHKOV,V. (eds) Mountain Building in the Uralides: Pangea to Present. Geophysical Monographs, American Geophysical Union, 132, 211-232. BEA, F., FERSHTATER,G. B., MONTERO, P., SMIRNOV,V. N. & MOLINA, J. F. 2005. Deformation-driven differentiation of granitic magma: the Stepninsk pluton of the Uralides, Russia. Lithos, 81, 209-233. BEANE, R. J. & CONNELLY, J. H. 2000. 4~ U - P b and Sm-Nd constraints on the timing of metamorphic events in the Maksyutov Complex, southern Urals, Ural Mountains. Journal of the Geological Society, London, 157, 811-822. BEANE, R. J., LIou, J. G., COLEMAN,R. G. & LEECH, M. L. 1995. Petrology and retrograde P - T path for eclogites of the Maksyutov Complex, southern Urals Mountains, Russia. Island Arc, 4, 254-266. BERZIN, R., ONCKEN, O., KNAPP, J. H., PEREZ-ESTAUN, A., HISMATULIN, T., YUNUSOV,N. t% LIPILIN, A. 1996. Orogenic evolution of the Ural Mountains: results from an integrated seismic experiment. Science, 274, 220-221.

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roots beneath mountain belts. Geophysical Journal International, 139, 556-562. KRUSE, S. & MCNUTT, M. 1988. Compensation of Paleozoic orogens: a comparison of the Urals to the Appalachians. Tectonophysics, 154, 1-17. LENNYKH, V. I., VALISER, P. M., BEANE, R., LEECH, M. & ERNST, W. G. 1995. Petrotectonic evolution of the Makysutov complex, southern Ural Mountains, Russia: implications for ultrahigh-pressure metamorphism. International Geology Review, 37, 584-600. MASLOV, V. A., CHERKASOV, V. L., TISCHCHENKO, V. T., SMIRNOVA, A. I., ARTYUSHKOVA,O. V. & PAVLOV,V. V. 1993. On the stratigraphy and correlation of the middle Paleozoic complexes of the main copper-pyritic areas of the Southern Urals. Ufimsky Nauchno Tsentr, Ufa (in Russian). MATTE, P., MALUSKI, H., CABY, R., NICOLAS, A., KEPEZHINSKAS,P. & SOBOLEV, S. 1993. Geodynamic model and 39Ar/4~ dating for the generation and emplacement of the high pressure (HP) metamorphic rocks in SW Urals. Comptes Rendus de l'Acaddmie de Sciences, Sdrie II, 317, 1667-1674. MONTERO, P., BEA, F., GERDES, A., ET AL. 2000. Single-zircon evaporation ages and Rb-Sr dating of four major Variscan batholiths of the Urals. A perspective on the timing of deformation and granite generation. Tectonophysics, 317, 93-108. PEREZ-ESTAUN, A., ALVAREZ-MARRON, J., BROWN, D., PUCHKOV, V., GOROZHANINA, Y. & BARYSHEV, V. 1997. Along-strike structural variations in the foreland thrust and fold belt of the southern Urals. Tectonophysics, 276, 265-280. PUCHKOV, V. N. 1997. Structure and geodynamics of the Uralian orogen. In: BURG, J.-P. & FORD, M. (eds) Orogeny through Time. Geological Society, London, Special Publications, 121, 201-236. PUCHKOV, V. N. 2000. Paleogeodynamics of the central and southern Urals. Ufa Pauria, Ufa (in Russian). RODGERS, J. 1990. Fold-and-thrust belts in sedimentary rocks. Part 1: Typical examples. American Journal of Science, 290, 321-359. SAVELIEVA, G. N. • NESBITT, R. W. 1996. A synthesis of the stratigraphic and tectonic setting of the Uralian ophiolites. Journal of the Geological Society, London, 153, 525-537. SAVELIEVA, G. N., SHARASKIN,A. Y., SAVELIEV, A. A., SPADEA, P. & GAGGERO, L. 1997. Ophiolites of the Southern Uralides adjacent to the East European continental margin. Tectonophysics, 276, 117-138. SAVELIEVA, G. N., SHARASKIN, A. Y., SAVELIEV, A. A., SPADEA, P., PERTSEV, A. N. & BABARINA, I. I. 2002. Ophiolites and zoned mafic-ultramafic massifs of the Urals: a comparative analysis and some tectonic implications. In: BROWN, D., JUHLIN, C. & PUCHKOV, V. (eds) Mountain Building in the Uralides: Pangea to Present. Geophysical Monographs, American Geophysical Union, 132, 135-154. SCARROW, J. H., SAVELIEVA,G. N., GLODNY, J., MONTERO,P., PERTSEV, A. N., CORTESOGNO,L. & GAGGERO, L. 1999. The Mindyak Palaeozoic lherzolite ophiolite, Southern Urals: geochemistry and geochronology. Ofioliti, 24, 239-246. SCARROW, J., AYALA, C., & KIMBELL, G. S. 2002. Insights into orogenesis: getting to the root of the continent-ocean-continent collision in the southern Urals, Russia. Journal of the Geological Society, London, 159, 659-672. SCHULTE, B. A. & BLOMEL, P. 1999. Prograde metamorphic reactions in the high-pressure Maksyutov Complex, Urals. Geologische Rundschau, 87, 561-576. SERAVKIN, I., KOSAREV, A. M. & SALIKHOV, D. N. 1992. Volcanism of the Southern Urals. Nauka, Moscow [in Russian]. SPADEA, P., KABANOVA,L. Y. & SCARROW,J. H. 1998. Petrology, geochemistry and geodynamic significance of Mid-Devonian boninitic rocks from the Baimak-Buribai area (Magnitogorsk Zone, southern Urals). Ofioliti, 23, 17-36. SPADEA, P., D'ANTONIO, M., KOSAREV, A., GOROZHANINA, Y. & BROWN, D. 2002. Arc-continent collision in the Southern Urals: petrogenetic aspects of the fore arc complex. In: BROWN,D., JUrILIN, C. & PUCHKOV,V. (eds) Mountain Building in the Uralides: Pangea to Present. Geophysical Monographs, American Geophysical Union, 132, 101-134.

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The Southern Urals: deep subduction, soft collision and weak erosion PHILIPPE M A T T E Laboratoire de Dynamique de la Lithosphbre, Universitd Montpellier 2, Place E. Bataillon, 34095 Montpellier, France (e-mail: matte @dstu. univ-montp2.fr)

Abstract: The Urals are a linear north-south-trending belt (2500 km in length, from Novaya Zemlya to the Aral Sea). Their apparent narrowness (100-150 kin) in the polar and cispolar parts is mainly due to the Siberian Meso-Cenozoic post-tectonic cover. In the southern, broadest part that crops out, the width of the Urals is close to 500 km. The Urals are very different from the other European Palaeozoic belts, the Caledonides and Variscides: despite subduction that built volcanic arcs in Silurian-Devonian times and pushed continental crust to great depth, where it underwent UHP metamorphism, the global shortening is relatively small, without great nappes, and the level of erosion very high, mainly east of the oceanic suture, where high-grade Uralian metamorphism is scarce. Another unusual feature is the preservation of an orogenic root in the centre of the orogen. These characteristics are due to the plate tectonic history that led to the orogeny: eastward subduction of the European passive margin stopped quickly after the Devonian because it was difficult for a so large a continent to sink further. Orogeny continued by westward subduction, east of the volcanic arc, closure of oceanic basins and accretion of small continental blocks and arcs without large underthrusting, and thus with little metamorphism or erosion (soft collision) but with large strike-slip motion. Preservation of the root is thought to be due to the high density of the central volcanic arc at depth (mantle and probably mafic granulites), which precluded strong uplift and erosion.

The Urals are a Late Palaeozoic fold belt, part of a much larger complex Palaeozoic belt that includes the Tien-Shan and the Kazakhstan, Altai and Mongolian belts, which were formed between the East European and the Siberian (Angara) cratons between 550 and 230 Ma. The Urals are an in situ chain, which was not deformed after Triassic time (except for very recent Quaternary uplift) and not dismembered by Mesozoic plate tectonics like the Scandinavian Caledonides or the West European Variscides. They are also generally less eroded and oceanic crust and volcanic arcs are very well preserved, only slightly affected by metamorphism. Thus the Urals are a key area to understand Palaeozoic plate tectonics. The Urals are now one of the best known Palaeozoic belts, as a result of the E U R O P R O B E programme, which brought together Russian and western Earth scientists and allowed the completion of two long deep seismic reflection profiles, which cross the entire belt in its southern (URSEIS) and middle part (ESRU).

Anatomy of the belt The Southern Urals, the widest part of the belt, are relatively well exposed between the Bashkirian and Kazakhstan plains (Fig. 1), with heights varying between 600 and 1500 m. Deformation affected the rocks in this region from Precambrian to Triassic times. The belt was eroded before the Jurassic and covered unconformably by horizontal continental deposits. The trend of folds, thrusts and strike-slip faults is roughly north-south. Although the most conspicuous structures are west-vergent nappes and thrusts, the belt is bivergent: east of the so-called Magnitogorsk syncline, folds and thrusts face east. All the large structures (anticlines, synclines, main faults) are relatively cylindrical and may be followed along strike for hundreds of kilometres, with a slight axial plunge of the folds to the south. The various zones from west to east (Fig. 2) have been described in detail by various researchers (Ivanov et al. 1975; Matte 1995, 1998, 2002; Puchkov 1997; Brown et al. 2002). A major oceanic suture, the Main Uralian Fault (MUF) separates the belt into a western part, consisting of material from the European continental passive margin with westward thrusts, and an eastern part, consisting of island arcs, oceanic sutures and microcontinents with eastward thrusts. This fan-like shape is clear from 50 ~ to 60~ and was well documented from the surface to the Moho by the URSEIS profile (Knapp et al. 1996) in the southern part and the ESRU

profile in the middle part (Juhlin et al. 1998; Brown et al. 2002), and is schematized in the sections shown in Figure 2.

Eastward oceanic, early subduction and building of volcanic arcs The Middle and Southern Urals are characterized by a wellpreserved mature island arc built during Ordovician and Devonian time as a result of calc-alkaline volcanism (basalts, andesites, tufts and cherts) and equivalent intrusive rocks (gabbros, diorites, norites and pyroxenites) (Puchkov 1997; Brown et al. 2001). In the southern section (Magnitogorsk-Irendick arc) only the upper part of the volcanic succession is present (andesites and overlying sediments). The lower oceanic part (gabbros, sheeted dyke complex, basalts) is better preserved along strike in the southernmost extremity of the Urals (see Zonenshain et al. 1990, for a detailed description). The age of this succession is Early to Mid-Devonian for the sheeted dyke complex and the tholeitic basalts (Shuldak complex), Mid- to Late Devonian for the island arc, and Late Devonian to Early Carboniferous for the overlying sediments. In the Middle Urals, the Tagil arc is characterized by a lineament of concentrically zoned d u n i t e - c l i n o p y r o x e n i t e - g a b b r o massifs of Alaskan type (Bea et al. 2002), the famous Platinumbearing Belt. Eastward subduction, east of the M a g n i t o g o r s k - T a g i l syncline during Late Devonian and Early Carboniferous time, beneath the Mugodzhar microcontinent may explain the geochemical polarity of the subduction-related granitoids of the East Uralian zone (Bea et al. 2002). The most striking tectonic feature of the Magnitogorsk synform is the very mild deformation (a spaced vertical fracture cleavage has been observed only locally). Over large areas, beds are subhorizontal and show neither internal deformation nor metamorphism.

U H P - L T Maksyutov Complex: deep eastward continental subduction and fast exhumation One of the peculiarities of the Southern Urals is the presence, west of the suture (Fig. 1), of a narrow ultrahigh-pressure-lowtemperature ( U H P - L T ) unit (600 ~ and 2 0 k b a r minimum),

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 421-426. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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the Maksyutov Complex, composed of eclogites, glaucophane micaschists and jadeite-bearing gneisses inferred to be metasediments of continental affinity (Dobretsov et al. 1996) sandwiched in an antiform, the Uraltau ridge, between low-grade to anchimetamorphic units. This UHP event was considered to be Precambrian ( 5 5 0 - 6 0 0 M a ) by Coleman et al. (1993) and Beane et al. (1995). Despite the first 4~ plateau ages around 380 Ma, obtained on phengite by Maluski (Matte et al. 1993), this age was considered as a retrogressive event much younger than the UHP metamorphism (Dobretsov et al. 1996). on phenAfter redating the Maksyutov UHP rocks by 4~ gite (Lennykh et al. 1995), U - P b on rutile (Beane & Connelly 2000) and S m - N d (Shatsky e t al. 1997), it became evident that 380 Ma was really the age of the UHP event. The Maksyutov Complex is thus, up to now, the best dated UHP metamorphism, determined by three methods and three different and independent teams. The similarity in age between minerals having different closure temperature, particularly between published S m / N d data (Matte et al. data (Shatsky et al. 1997) and 4~ 1993; Lennykh et al. 1995), suggests that the UHP Maksyutov Complex rose rapidly, with exhumation rates much higher than the 2.5 mm a -1 proposed by Leech & Stockli (2000) and probably around 1 cm a -1 or more (Chopin 2003). This is easily explicable by the physical analogue modelling of Chemenda et al. (1997; Fig. 3), followed by Boutelier et al. (2004), in which deep continental subduction and the buoyancy driven rise of a slice of continental crust dragged the subducted continental metasediments and parts of the overlying oceanic crust and mantle. A prerequisite of this model is the subduction of the forearc (at least its mantle part), which acted as a thermal screen, protecting the subducted crust from overheating by the asthenosphere. This explains the LT character of the metamorphism. Fig. 1. Tectonic sketch map of the Southern and Middle Urals, modified after Brown et al. (2002), with the location of sections A and B shown in Figure 2.

Fig. 2. Two interpretative simplified sections across the Southern (A) and Middle Urals (B) after the URSEIS and ESRU profiles, taking into account the studies by Friberg et al. (2002) and Kashubin et al. (2006) for section B. 1, inferred Archaean basement; 2, Riphean-Vendian sediments with or without pre-Palaeozoic deformation and metamorphism; 3, shallow-marine Lower Palaeozoic sediments and Permian molasse of the foredeep; 4, Zilair flysch, mainly Devonian; 5, Magnitogorsk-Tagil and Valerianov volcanic arcs; 6, oceanic crust and mantle; 7, .Kazakhstan Lower Palaeozoic series deformed before Devonian time. 8, MUF, Main Uralian Fault; TF, Troitsk fault.

SOUTHERN URALS

423

Fig. 3. Model for the exhumation of HP rocks in the SouthernUrals. +, continental crust; dashes ductile continentallower crust; black, oceanic crust; dotted pattern, sediments; chevron pattern, island arc; light grey, continental mantle; dark grey, oceanic mantle; ,,, HP metamorphicrocks. (a) Oceanic subductionand arc construction; (b) continental subduction; (c) forearc subductionand breaking of the continental crust; (d) upward extrusion of the continental slab; (e) continental lithospherebreak-off.

The collage of the eastern microcontinents and arcs: the result of westward subduction? From the previous observations it appears clearly that a great part of the Southern and Middle Urals evolution is the result of an eastward, first oceanic, then continental subduction of the East European margin beneath a Uralian ocean. Nevertheless, although the eastern half of the Urals is much less known, we suspect that westward subduction also occurred there. The first feature that favours this argument is the fan-like shape of the belt, visible on the two sections of the Middle and Southern Urals (Fig. 2). Another factor is the eastward flat thrusting of Silurian-Devonian ophiolitic m41anges: large (kilometre-scale) massifs of serpentinites, rooted at the eastern boundary of the Magnitogorsk-Tagil zone, lie

on top of shelf Devonian-Carboniferous sediments of the East Uralian zone, mainly in the Middle Urals, east of Sverdlovsk. This eastward obduction probably reflects a change in the sense of subduction at the end of the Devonian, as proposed by Zonenshain et al. (1984), Matte (1995) and Friberg et al. (2002). Such subduction reversal has also occurred in modern orogens such as Taiwan (Chemenda et aL 2001). Finally, if we consider the general westward dip of the thrusts and sutures in the Trans-Uralian Zone (Fig. 2), it is difficult to build M i d - L a t e Carboniferous volcanic arcs such as the Valerianov arc by eastward subduction. A westward subduction seems more likely to accrete against the western half of the Urals, the S a l d a - M u g o d z h a r microcontinent, the Valerianov arc and the Kazakhstan older Palaeozoic belt, as proposed by Seng6r et al. (1993).

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The late eastward subduction and Late Carboniferous-Permian Uralian orogeny The late eastward subduction is marked by the westward thrust tectonics, which formed the Bashkir and Kvarkush anticlinoria and the d6collements in the foredeep (Fig. 2). There is a time gap between the obduction stage in the Late Devonian, recorded by the deposit of the Frasnian Zilair chromite-bearing distal flysch, the emplacement of the Kraka oceanic nappe, the generation and exhumation of the UHP rocks, and the final Mid-Carboniferous-Permian thrust tectonics. The eastward subduction is also recorded by the sedimentation in the foredeep during the Late Carboniferous, and particularly the transition from molasse to the east to flysch and deep basinal shales and cherts to the west (Puchkov 1997). The total shortening is probably less than 100 km but certainly more than the 17 km proposed by Brown et al. (1996) on the basis of balanced sections, because those workers did not take into account the widespread

occurrence of dissolution cleavage in the Kvarkush and Bashkir anticlinoria.

Conclusion The Urals are characterized by a long subduction history (Fig. 4) that involved the following elements. (1) Eastward oceanic subduction (Ordovician-Devonian) led to the building of the Magnitogorsk-Tagil island arc. (2) Closure of the ocean and continuous eastward continental subduction to great depth of the continental crust of the East European continental passive margin led to the westward obduction of ophiolites and the generation and fast exhumation of UHP metamorphic rocks during the Mid-Devonian. This 'Oman' stage (Chemenda et al. 1996) probably did not produce a strong orogeny because the back-arc ocean was still open.

Fig. 4. Schematic illustration of a possible plate-tectonic evolution for the Southern Urals from Silurian to Late Carboniferous time: 1, continental mantle; 2, oceanic mantle; 3, continental crust; 4, arc crust; 5, oceanic crust; 6, Palaeozoic sediments in front of the suture.

SOUTHERN URALS

(3) Probably eastward subduction of the back-arc ocean beneath the Mugodzhar block during Latest Devonian to Early Carboniferous time, generated subduction-related granitoids (Bea et al. 2002). (4) Westward subduction of the back-arc ocean beneath the M a g n i t o g o r s k - T a g i l island arc led to the final closure of the eastern oceanic basins and generation of eastern arcs, and gave the fan-like aspect of the belt. Eastward thrusting and thickening of the crust in the Mugodzhar block may explain the generation of crustal granites such as the Dzhabyk and Murzinka batholiths (Bea et al. 2002). (5) Moderate eastward subduction of the East European continent produced westward thusting and the onset of the Uralian orogeny during Permian time. (6) Final orogen-parallel movements generated long vertical strike-slip faults such as the Sisert, East Magnitogorsk and Troitsk faults. The lack of granitoids and the scarcity of high-grade Uralian metamorphism west of the suture (MUF) indicate that crustal stacking was too weak to produce melting of the crust. As discussed by Scarrow et al. (2002) and Kashubin et al. (2006), the preservation of the root is due to the presence of dense rocks such as mafic granulites and mantle, which probably form the deepest parts of the M a g n i t o g o r s k - T a g i l island arc, as seen in the roots of obducted island arcs (Bard 1983), which can explain the very weak erosion of this part of the Urals. I thank D. Gee, who encouraged me to write this paper, for his comments and improvement of the English, F. Bea for his constructive remarks, and A. Delplanque for the final version of the drawings.

References BARD, J.-P. 1983. Metamorphism of an obducted island arc: example of the Kohistan sequence (Pakistan) in the Himalayan collided range. Earth and Planetary Science Letters, 65, 133-154. BEA, F., FERSHTATER, G. B. & MONTERO, P. 2002. Granitoids of the Uralides: implications for the evolution of the orogen. In: BROWN, D., JUHLIN, C. & PUCHKOV, V. (eds) Mountain Building of the Uralides, Pangea to the Present. Geophysical Monographs, American Geophysical Union, 132, 211-232. BEANE, R. J. & CONNELY, J. N. 2000. 4~ U - P b and Sn-Nd constraints on the timing of metamorphic events in the Maksyatov complex, Southern Ural Mountains. Journal of the Geological Society, London, 157, 811 - 822. BEANE, R. J., LIOU, J. G., COLEMAN, R. G. & LEECH, M. L. 1995. Petrology and retrograde P - T path for eclogites of the Maksyutov Complex, Southern Urals Mountains, Russia. The Island Arc, 4, 254-266. BOUTELIER,D., CHEMENDA,A. & JORAND,C. 2004. Continental subduction and exhumation of high-pressure rocks: insights from thermomechanical laboratory modelling. Earth and Planetary Science Letters, 222, 31-45. BROWN, D., ALVAREZ-MARRON,J. & PEREZ-ESTAUN, A. 1996. Architecture of the footwall to the Main Uralian fault, Southern Urals. EarthScience Reviews, 40, 209-216. BROWN, D., ALVAREZ-MARRON, J., PI~REZ-ESTAUN, A., PUCHKOV, V., GOROZHANINA, Y. & AYARZA, P. 2001. Structure and evolution of the Magnitogorsk forearc basin: identifying upper crustal processes during arc-continent collision in the southern Urals. Tectonics, 20, 364-375. BROWN, D., JUHLIN, C., TRYGGVASON, A., STEER, D., BECKOLMEN, M., RYBALKA, A. & BLIZNETSOV, M. 2002. The crustal architecture of the Southern and Middle Urals from the URSEIS, ESRU and Alpaev reflection seismic surveys. In: BROWN, D., JUHLIN, C. & PUCHKOV, V. (eds) Mountain Building of the Uralides, Pangea to the Present. Geophysical Monographs, American Geophysical Union, 132, 33-48. CHEMENDA, A., MATTAUER, M. & BOKUN, A. N. 1996. Continental subduction and a mechanism for the exhumation of high-pressure

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metamorphic rocks: new modelling and field data from Oman. Earth and Planetary Science Letters, 143, 173-182. CHEMENDA, A., MATTE, P. • SOKOLOV,V. 1997. A model of Palaeozoic obduction and exhumation of high-pressure/low-temperature rocks in the southern Urals. Tectonophysics, Urals Special Issue, 276, 217-227. CHEMENDA, A. I., YANG, R.-K., STEPHAN, J.-F., KONSTANTINOVSKAYA, E. A. & IVANOV,G. M. 2001. New results from physical modelling of arc-continent collision in Taiwan: evolutionary model. Tectonophysics, 333, 159-178. CHOPIN, C. 2003. Ultrahigh-pressure metamorphism: tracing continental crust into the mantle. Earth and Planetary Science Letters, 212, 1-14. COLEMAN, R. G., LIOU, J. G., ZANG, R. Y., DOBRETSOV, N., SHATSKY, V. 8,: LENNYKH,V. 1993. Tectonic setting of the UHPM Maksutov complex, Ural Mountains, Russia. EOS Transactions, American Geophysical Union, 74, 547. DOBRETSOV, N. L., SHATSKY, V. S., COLEMAN, R. G., ETAL. 1996. Tectonic setting and petrology of ultrahigh-pressure metamorphic rocks in the Maksyutov Complex, Ural Mountains, Russia. International Geology Review, 38, 136-140. FRIBERG, M., JUHLIN, C., BECRHOLMEN, M., PETROV, G. A. & GREEN, A. G. 2002. Palaeozoic tectonic evolution of the Middle Urals in the light of the ESRU seismic experiments. Journal of the Geological Society, London, 159, 295-306. IVANOV, S. N., PERFILIEV, A. S., EFIMOV, A. A., SMIRNOV, G. A., NECHEUKHIN, V. M. & FERSHTATER, G. M. 1975. Fundamental features in the structure and evolution of the Urals. American Journal of Science, 275, 107-130. JUHLIN, C., FRIBERG, M., ECHTLER, H., GREEN, A. G., ANSORGE, J., HISMATULIN, T. & RYBALKA, A. 1998. Crustal structure of the Middle Urals: results from the (ESRU) Europrobe Seismic Reflection Profiling in the Urals Experiments. Tectonics, 17, 710-725. KASHUBIN, S., JUHLIN, C., FRIBERG, M., ETAL. 2006. Crustal structure of the Middle Urals based on reflection seismic data. In: GEE, D. G. & STEPHENSON, R. A. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 427-442. KNAPP, J. H., STEER, D. N., BROWN, L. D., ET AL. 1996. A lithosphere-scale image of the southern Urals from explosion-source reflection profiling. Science, 274, 226-228. LEECH, M. L. & STOCKLI, D. F. 2000. The late exhumation history of the ultrahigh-pressure Maksyutov Complex, south Ural Mountains, from new apatite fission tracks. Tectonics, 19, 153-167. LENNYKH, V. I., VALIZER, P. M., BEANE, R., LEECH, M. & ERNST, W. G. 1995. Petrotectonic evolution of the Maksyutov Complex, Southern Urals, Russia: implications for ultrahigh-pressure metamorphism. International Geology Review, 37, 584-600. MATTE, P. 1995. Southern Uralides and Variscides: comparison of their anatomies and evolutions. Geologie en Mijnbouw, 74, 151-166. MATTE, P. 1998. Continental subduction and exhumation of HP rocks in Paleozoic belts: Uralides and Variscides. Geologiska FOreningens i Stockholm FCrhandlingar, 120, 209-222. MATTE,P. 2002. The Variscides, between the Appalachians and the Urals, similarities and differences between subduction-collision Paleozoic belts. In: MARTINEZ CATALAN,J. R., HATCHER, R. D. JR., ARENAS, R. & DIAZ GARCIA, F. (eds) Variscan-Appalachian Dynamics: the Building of the Late Paleozoic Basement. Geological Society of America, Special Papers, 364, 239-251. MATTE, P., MALUSKI, H., CABY, R., NICOLAS, A., KEPEZHINSKAS,P. & SOBOLEV, S. 1993. Geodynamic model and 39Ar/4~ dating for the generation and emplacement of the HP metamorphism in SW Urals. Comptes Rendus de l'Acaddmie des Sciences, 317, 1667-1674. PUCHKOV, V. N. 1997. Structure and geodynamics of the Uralian orogen. In: BURG, J.-P. & FORD, M. (eds) Orogeny through Time. Geological Society, London, Special Publications, 121, 201-236. SCARROW, J. H., AYALA, C. & KIMBELL, G. S. 2002. Insights into orogenesis: getting to the root of a continent-ocean-continent collision, Southern Urals, Russia. Journal of the Geological Society, London, 159, 659-671.

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ZONENSHAIN, L. P., KORINEVSKY, V. G., KAZMIN, V. G., PECHERSKY, D. M., KHAIN, V. V. • MATVEENKOV, V. V. 1984. Plate tectonic model of the Southern Urals development. Tectonophysics, 109, 95-135. ZONENSHAIN, L. P., KUZMIN, M. I. & NATAPOV, L. M. (eds) 1990. Geology of the USSR: a Plate-Tectonic Synthesis. American Geophysical Union, Geodynamics Series, 21.

Crustal structure of the Middle Urals based on seismic reflection data S. K A S H U B I N 1, C. J U H L I N 2, M. FRIBERG 2, A. R Y B A L K A l, G. PETROV 3, A. K A S H U B I N 2, M. B L I Z N E T S O V 1 & D. STEER 4

1Bazhenov Geophysical Expedition, Communotov Street 17, Zarechny, 624051, Russia 2Department of Earth Sciences, Uppsala University, Villaviigen 16, 75236 Uppsala, Sweden (e-mail: christopher.juhlin @ geo. uu. se) 3Urals Geological Survey Expedition, 55 Veinera, Ekaterinburg, Sverdlovsk District, 620014, Russia 4Department of Geology, University of Akron, Akron, OH, 44325-4101, USA

Abstract: EUROPROBE-related seismic reflection surveys in the Middle Urals, Russia (latitude 56-62 ~ since 1993 have led to an increased understanding of the crustal structure and tectonic evolution of this region. A 400 km long profile now extends from the foreland basin in the west well into the West Siberian Basin in the east. Bivergent structures characterize the upper crust of the Uralide orogen, whereas the middle and lower crust generally contain gently west-dipping reflections. A crustal root is imaged down to almost 60 km beneath the exposed Urals. Below the foreland and the West Siberian Basin the lower crustal reflectivity is pronounced and the Moho lies at a depth of 40-45 km. Below the foreland on the recently acquired Serebrianka-Beriozovka profile, two sets of late arriving (20-25 s) reflections are present. One set reflects from a zone in the mantle at about 60-70 km depth that strikes ENE and dips about 45 ~to the SSE. The other set may represent imbricated lower crust. Major events during the Palaeozoic tectonic evolution of the Middle Urals were: continental and oceanic rifting (Late Cambrian to Early Ordovician); development of a passive continental margin (Mid-Ordovician to Mid-Carboniferous); intra-oceanic subduction below the Tagil arc (Silurian to Devonian); east-dipping subduction of the Baltica plate (Silurian to Early Devonian); possible subduction reversal with formation of the Alapaevsk island arc and the Krasnoturjinsk-Petrokamensk active continental margin (Devonian to Early Carboniferous); active building of a mountain belt and intrusion of collision-related granitic plutons (Carboniferous to Permian).

Several multinational geoscientific studies, co-ordinated within the EUROPROBE programme, have been active in the Urals since 1992. The combined geological and geophysical URSEIS project in the Southern Urals (Berzin et al. 1996; Carbonell et al. 1996; Echtler et al. 1996; Knapp et al. 1996) includes a 465 km long seismic reflection profile across the orogen at latitude 54~ URSEIS was the first to reveal the bivergent character of the orogen, with east-dipping structures in the west and west-dipping ones in the east. The ESRU (Europrobe Seismic Reflection profiling in the Urals) project, presented in this paper, began in 1993 and conducted five seismic reflection experiments in the Middle Urals (Juhlin et al. 1995, 1996, 1997, 1998; Brown et al. 2002; Friberg et al. 2002) at latitude 58~ (Fig. 1), crossing the orogen close to the Urals deep drillhole, SG-4 (currently at c. 5.5 km depth). The composite ESRU seismic line is about 335 km long. The ESRU profile was recently extended to the west in three campaigns along the Serebrianka-Beriozovka (SB) profile (Fig. 1), parts of which are presented in this paper. The combined E S R U - S B profile is about 440 km long (Fig. 2), comparable in length with the URSEIS profile further south. In addition to the acquisition of these seismic reflection lines, EUROPROBE co-operation has resulted in joint interpretation of several Russian seismic datasets including the R l 1 4 and Rl15 profiles in the Southern Urals (Brown et al. 1998), the R17 profile in the northern Middle Urals (Juhlin et al. 1996), the Michailovsky profile in the southern Middle Urals, and the Chernoistoichinsk-Alapaevsk profiles (Steer et al. 1995) just south of the ESRU profile (Fig. 1). Although the tectonic history of the Uralides in the Middle Urals is similar to that of other Palaeozoic orogens, there are important differences between the present-day Uralides and these orogens. Foremost is the presence of a crustal root (Ryzhiy et al. 1992; Druzhinin et al. 1997), which appears to be absent in the Caledonide (Matthews & Cheadle 1986), Variscide (Meissner et al. 1987) and Appalachian (McBride & Nelson 1991) orogens. The fact that the Uralides lie within a tectonic plate (Eurasia) may be the reason why their collisional architecture is preserved at a low erosion level. This allows us to develop ideas about collisional orogenesis that are not possible for the Palaeozoic orogens listed above. Other important features of the Uralides

include: (1) relatively minor syn- or post-collisional collapse in the exposed part of the orogen, at least in the Southern Urals (Brown et al. 1998); (2) extremely well-preserved ophiolites and volcanic arc assemblages (Savelieva & Nesbitt 1996); (3) foreland and hinterland basins (including the Timan-Pechora and West Siberian basins) containing some of the largest hydrocarbon reserves in the world; (4) great mineral wealth, making the Uralides one of the most mineralized orogens, with exceptionally large ore deposits (Herrington et al. 2002); (5) peneplanation in the Jurassic followed by continuous uplift since the Tertiary (Borisevich 1992; Puchkov 1997; Seward et al. 2002). Earlier reported anomalously low heat flow along the axis of the Urals can be explained by palaeoclimatic effects and differences in the borehole depths at which the measurements were made in (Kukkonen et al. 1997). Although the Uralides crustal root is anomalous when compared with other Palaeozoic orogenic belts, roots are observed beneath both younger Alpine (e.g. Valasek et al. 1991) and older Palaeoproterozoic (e.g. BABEL Working Group 1990; Hajnal et al. 1996) orogens. By comparing data from the Uralides with these other orogens, it may be possible to determine which parameters govern the creation and the preservation of crustal roots. Deep seismic reflection profiling is an essential tool for an improved understanding of the Uralide crustal root and its relationship to continuing uplift. Whether the present-day crustal root is a relict of the Palaeozoic collision or is the result of this more recent tectonic activity is still an open question. In this paper, we summarize the results from the ESRU and other seismic reflection profiles and interpret the crustal-scale structure of the Middle Urals based upon them. To support our interpretations we draw upon information from geological mapping (see Friberg 2000, for details) and key isotopic studies. We also incorporate results from the other seismic reflection profiles in the Urals, gravity and magnetic data, and the topography.

Geological setting The Urals mark the boundary between Europe and Asia. Ivanov et al. (1975) divided the orogen into six geological subdivisions

From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 427-442. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig, 1. (a) Geological, (b) topographic, (e) magnetic, and (d) gravity maps of the Middle Urals. The Main Uralian Fault Zone and the western margin of the West Siberian Basin sediments are shown on the topographic, magnetic and gravity maps. Magnetic data are the total field magnetic anomaly from the National Geophysical Data Center (NGDC 1997). The data have been shifted 10 krn to the west so that the anomalies in the area of the ESRU profile coincide with more detailed Russian magnetic maps. The Bouguer gravity anomaly map is from Kaban (2001). Seismic profiles are shown as black lines on all maps. Seismic sections have been projected onto the straight light grey lines shown on the topographic, magnetic and gravity maps, with white dots marking every 20 km and red dots every 100 kin. EEP, East European Platform; WSB, West Siberian Basin.

CRUSTAL STRUCTURE OF THE MIDDLE URALS

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running parallel to the axis of the mountain belt. The three westernmost zones are of East European affinity (Fig. 1) and include the Urals Foredeep (a molasse basin deposited as Baltica platform sediments), the West Uralian Zone (part of the foreland foldand-thrust belt of Palaeozoic sediments), and the Central Uralian Zone (uplifted Precambrian rocks, consisting mainly of a Neoproterozoic succession in the Kvarkush Anticline in the Middle Urals). The three accreted complexes in the east are divided into the Tagil-Magnitogorsk Zone (Tagil Volcanic Arc Complex), the East Uralian Zone (consisting of arc sediments and metamorphic complexes), located to the east of the Tagil Volcanic Arc Complex, and the Trans-Uralian Zone (largely covered by West Siberian sediments). In the Middle Urals, the foreland fold-and-thrust belt is a westverging thrust stack that involves both the Proterozoic basement and the Palaeozoic cover. Thrusting during the Uralide orogeny also reactivated earlier, Proterozoic structures. The Palaeozoic sequences include continental and shallow marine deposits of Early Ordovician to Mid-Carboniferous age, overlain by molasse deposits of Late Carboniferous to Early Triassic age (Puchkov 1997). The accreted terranes further to the east consist mostly of wellpreserved Silurian to Carboniferous volcanic arc material with fragmented ophiolites, but there are also large areas of high-grade metamorphic rocks (Antsigin et al. 1994; Sobolev et al. 1964). Metamorphic terranes of the East Uralian Zone in the vicinity of the ESRU-SB profile are represented by the Salda Metamorphic Complex and the Murzinka-Adui Metamorphic Complex (MAMC), which consist of Mid- to Late Palaeozoic high-grade gneisses (Friberg et al. 2000) and granitoids (Bea et al. 1997; Friberg & Petrov 1998). Included in the East Uralian Zone is a mrlange of allochthonous arc volcanic and volcaniclastic rocks that partly cover the high-grade rocks (Friberg & Petrov 1998). Some of the metamorphic complexes have been interpreted as micro-continents composed of Proterozoic or older continental crust (Zonenshain et al. 1990; Puchkov 1997), but isotope age investigations indicate that they are, at least in part, the roots of Paleaozoic volcanic arcs (Friberg et al. 2000). The tectonic contact separating the accreted terranes from the East European continental margin is known as the Main Uralian Fault Zone and can be traced as an up to 20 km wide deformation zone along most of the orogen (Fig. 1). The structural character of this fault zone varies along the orogen. In the Southern Urals, it apparently only records the Late Palaeozoic collision (Brown et al. 1998), whereas, in the Middle and Northern Urals, it is partly overprinted by normal faulting (Juhlin et al. 1993; Echtler et al. 1997; Knapp et al. 1998; Friberg et al. 2002) and strike-slip faulting (e.g. Ayarza et al. 2000a; Brown et al. 2002; Hetzel & Glodny 2002). The central axis of the crustal root is generally located somewhat to the east of the Main Uralian Fault Zone and produces a 500 km wide long-wavelength gravity low across the orogen (Fig. 1). A distinct short-wavelength (c. 100 km wide) gravity high (McKenzie & Fairhead 1997; Kimbell et al. 2002) is also observed east of the Main Uralian Fault Zone (Fig. 1). This high is found somewhat to the east of the approximate location of the central axis of the crustal root and can be correlated with dense, probably oceanic rocks and peridotite-gabbro island arc intrusions, located at shallow crustal levels in the Tagil and Magnitogorsk accreted arcs (Petrov & Puchkov 1994; Scarrow et al. 2002; Kimbell et al. 2002; Brown et al. 2003). Evolutionary models of the Uralides (Hamilton 1970; Ivanov et al. 1975; Zonenshain et al. 1984, 1990) all infer that significant tectonic and magmatic activity in the area began in the Late Cambrian to Early Ordovician. In Neoproterozoic times there were few continental rifting events in the Middle and Southern Urals. Continental tiffing, continental break-up and sea-floor spreading during the Early Palaeozoic resulted in a passive continental margin along the eastern edge of the present-day East

European Craton (Puchkov 2002; Bogolepova & Gee 2004). Further east within the Uralian palaeo-ocean, island arcs were active during Silurian to Devonian times. In the Late Palaeozoic, the island arc terranes, and possibly some microcontinents, within the Uralian palaeo-ocean accreted to and were then thrust onto the East European Craton (Seng6r et al. 1993; Brown et al. 1998). Closure of the ocean basin and continent-continent collision, involving the surrounding Precambrian cratons to the west and east of the island arc assemblages, took place in Late Carboniferous to Permian times. Although convergence between the cratons probably continued into the Early Triassic, by Triassic times extension was an important factor in the development of the West Siberian Basin along the eastern flank of the orogen. The cratons and many of the accreted terranes in the east are covered by Mesozoic sediments, limiting the understanding of the eastern side of the orogen. The relative plate motions of the continents prior to orogeny are not well constrained; however, existing palaeogeographical reconstructions (Torsvik et al. 1992, 1995; Seng6r et al. 1993) infer oblique collision and clockwise rotation of Siberia relative to the East European Craton. New palaeomagnetic data from East Uralian terranes and the Uralian Palaeozoic continental margin (Sviazhina et aL 1999, 2003; Petrov 2003) suggest anticlockwise rotation of Baltica before the Devonian and clockwise rotation in the Late Palaeozoic.

Seismic reflection data acquisition Seismic reflection data along the ESRU profile were acquired during five field campaigns, beginning in 1993 (Table 1). The 60 km long (receiver station length) ESRU93 seismic reflection line was acquired first and was oriented in a SW-NE direction crossing the Main Uralian Fault Zone (Fig. 1). During acquisition of ESRU93, data were recorded by both Uppsala University and the Bazhenov Geophysical Expedition. Uppsala University recorded data along the entire line whereas the Bazhenov Geophysical Expedition recorded data along a 35 km long stationary spread in the middle part of the profile. In earlier publications only results from the Uppsala data have been presented. The ESRU93 part of the section presented in this paper now includes the Bazhenov recorded data. The ESRU95 survey consisted of two profiles, one running roughly east-west (ESRU95_we) and the other north-south (ESRU95_ns), crossing each other near the SG-4 deep borehole (Juhlin et al. 1997; Ayarza et al. 2000b). ESRU96 was the eastern extension of ESRU95_we and it imaged the eastern part of the Tagil Arc and most of the Salda Metamorphic Complex. ESRU98 overlapped the eastern end of ESRU96 within the Salda Metamorphic Complex and then crossed the Sisert Fault, the Murzinka-Adui Metamorphic Complex, and the Alapaevsk Arc before running in a northeasterly direction over Mesozoic strata of the West Siberian Basin (Fig. 1). A 2 km gap exists between ESRU98 and ESRU99 (Fig. 1) as a result of tape reading problems. ESRU99 continued to the east over Mesozoic strata of increasing thickness and ended where these strata reach a thickness of about 1 km. In 2001, about 50 km of seismic reflection data were acquired to the west of ESRU93 in the Uralian foreland along the SerebriankaBeriozovka profile (SB01). In 2002, an additional 42 km of profile (SB02) were acquired between the western end of ESRU93 and the SB01 profile. Finally, 73 km of profile (SB03) were acquired in 2003 as a western extension of the SB01 profile. Seismic reflection data were acquired along the Michailovsky profile in 1998 and 1999 (Fig. 1) with acquisition parameters similar to those for the ESRU-SB profiles (Table 1). The profile begins in the western part of the Uralian foreland and extends onto the foreland fold-and-thrust belt, ending about 5 km west of the Main Uralian Fault Zone. In addition to the ESRU-SB and Michailovsky profiles, data from the R17 profile (Juhlin et al. 1996), in the northern part of the area

CRUSTAL STRUCTUREOF THE MIDDLE URALS

431

(Fig. 1), and the Alapaevsk and Chernoistoichinsk profiles (Steer et al. 1995), just south of the ESRU profile (Fig. 1), are relevant

> o

for the large-scale structural interpretation of the Middle Urals. .~

x

Seismic reflection data processing

~T

Results from previous processing of the ESRU profiles have been presented in earlier publications (Juhlin et al. 1995, 1998; Brown et al. 2002; Friberg et al. 2002). Different software packages were used and different people did the actual processing, leading to a somewhat heterogeneous image when all the ESRU data were merged. To produce a more homogeneous section, all of the ESRU data have now been reprocessed using similar processing parameters (Table 2) with a focus on imaging the large-scale features of the crust. For this reason, dip moveout (DMO) has not been included in the processing flow on the sections presented in this paper, with the result that some of the upper crustal images differ from earlier published sections (e.g. compare with Juhlin et al. (1997) and Ayarza et al. (2000b)). Another difference between the present processing and earlier work is that the common depth point (CDP) stacking lines follow the stations, except for along SB02 and the Michailovsky profile, resulting in CDP lines whose azimuths vary greatly. In the earlier processing, piece-wise straight lines were chosen and the data were projected onto these straight lines. Here, we have chosen to project the resulting stacked sections onto three separate straight-line segments. To extract the most prominent reflections from the sections, automatic line drawings based on reflection coherency have been generated. Figure 2 shows a line drawing section of the entire E S R U - S B profile where the profile approximately follows the acquisition line. Figure 3 shows variable area plots of the straightline segments onto which the data represented in Figure 2 have been projected. The crooked nature of the acquisition geometry along the E S R U - S B profile (Fig. 1) has both advantages and disadvantages when processing and interpreting the seismic data. Some of the problems with the crooked-line geometry may be avoided by projecting the seismic data onto piece-wise straight CDP lines (Wu et al. 1995) and it was this strategy that was employed in the previous processing. However, tests on other datasets show that the image can also be degraded as a result of this projection and that a superior image may be obtained if the CDP line is chosen to follow the acquisition line as closely as possible. If the latter strategy is chosen, as was done along the ESRU and

x=

t'q t'q t'q

t'q t"q

Table 2. Generalized processing sequence for the ESR U - SB and Michailovsky profiles

s I

~x

X

c~ r

tt3 ~1-

~

~

Step

d

i

~

('-,I

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

Process Read SEGB-SEGY data Trace edits (automatic and manual) Correction for geometrical divergence Elevation statics: datum - 100 m Air blast attenuation Spectral whitening Bandpass filter Refraction and residual static corrections NMO correction DMO (only on Michailovsky) Trace equalization Trim statics: maximum shift 2 ms CMP stacking Trace equalization F-X deconvolution Display Trace equalization

432

S. KASHUBIN ETAL.

Fig. 3. Seismic data from the merged profile projected onto three straight lines as shown in Figure 1. The distance scale differs from that in Figure 2, as it now refers to distance along the straight-line segments. However, the distance scale is still referenced to the same zero point, the western limit of the Main Uralian Fault Zone.

SB01 profiles in this paper, then it is possible to orient, in three dimensions, those reflections that are present on those portions of the CDP line where the azimuth varies significantly. Based on this strategy, travel times to where major reflections or reflection zones are present or are expected to be present have been calculated for planes of arbitrary orientation in a constant velocity medium. Given that only 2D profile data are available, the assumption that the reflectors are planes in constant velocity media is reasonable. More complicated geometries can generally not be verified with the present dataset. Zaleski et al. (1997) used a similar strategy to calculate travel times of reflections from nearvertical dykes along crooked lines running nearly parallel to the strike of the dykes. A velocity of 6 km s-1 has been assumed for all crustal reflections. Increasing the velocity to 6.6 km s -~,

a reasonable average crustal velocity, increases the apparent dip only slightly. Figure 4 shows detailed images of portions of the E S R U - S B profile where travel times to some of the reflectors listed in Table 3 have been modelled as outlined above. Line drawings were also generated from the three sections shown in Figure 3. These line drawings were then migrated and merged into a single section (Fig. 5) where the distance scale is relative to distance from the western boundary of the Main Uralian Fault Zone. A 1D migration velocity based on results from the GRANIT refraction-DSS (deep seismic sounding) profile (Juhlin et al. 1996) was used, except for east of km 160 (Fig. 5), where the 1D velocity function was adjusted to take into account the low-velocity (1800 m s -1) Mesozoic sediments of the West Siberian Basin.

CRUSTAL STRUCTURE OF THE MIDDLE URALS

433

Fig. 4. Detailed variable area images of parts of the seismic section shown in Figure 2. Travel times of reflections shown on the sections are based on the orientations given in Table 3. (a) ST, Serebrianka Thrust; KA1, Kvarkush Anticline fault; MUTF, Main Uralian Thrust Fault; MUNF, Main Uralian Normal Fault; (b) D1, D2 and D3, reflections from the lower crust below the West Siberian Basin; (e) M1, reflection from within the mantle; LC, reflection from the lower crust. Vertical lines are traces with no data.

Data from the Michailovsky profile were processed in a manner similar to the earlier processing of the ESRU data (Friberg et al. 2002), including DMO. These data were projected directly onto straight lines prior to stacking. Only line drawing versions of the Michailovsky profile are shown in this paper (Fig. 6). No reprocessing of the Alapaevsk-Chernoistoichinsk profiles (Steer et al. 1995) has been carried out; the section used by Brown et al. (2002) is presented here (Fig. 6). Migrated line drawings from the Michailovsky, Alapaevsk-Chernoistoichinsk and E S R U - S B profiles have been plotted as fence diagrams on the magnetic map in Figure 7.

Crustal structure Based on reflectivity character, the E S R U - S B section can be divided into various zones (Fig. 5): sedimentary cover, upper crust, middle and lower crust, and upper mantle. Upper crustal

reflectivity within the exposed Uralides is generally characterized by distinct single reflections or sets of reflections, very often moderately to steeply dipping. Middle and lower crustal reflectivity is more diffuse, with 'clouds' of rather gently dipping scattered reflections. The most prominent upper crustal reflectors can be traced to depths of 1 0 - 1 5 km. Most of them coincide with known tectonic zones. However, some tectonic zones (e.g. the Prianichnikova shear zone; PSZ) were discovered and studied by the ESRU project. It is possible that the imaged tectonic zones below the exposed Uralides have been reactivated, resulting in better reflectivity compared with the nearly transparent upper crust below the sedimentary cover to the east. It is generally difficult to trace reflectivity from the middle and lower crust to the surface. Only one reflective zone appears to extend from the Moho into the upper crust (TF in Fig. 5). Furthermore, the reflectivity in the middle and lower crust is not confined to the terranes mapped at the surface. These observations suggest that the upper crust may, in part, have been decoupled from the

434

S. KASHUBINETAL.

Fig. 4. Continued

middle and lower crust. Although much of the observed reflectivity is probably related to processes that were active in the Palaeozoic and earlier, some of the reflections may represent relatively young features, possibly as young as Mesozoic.

Upper crustal features Foreland fold-and-thrust belt. Flat-lying reflections characterize the 5 - 1 0 km thick sedimentary rocks of the Uralian foreland west of km - 7 5 on the E S R U - S B section and km - 5 0 on the Michailovsky profile (Figs 6 and 7). They are mainly composed of Neoproterozoic and Early Palaeozoic rocks lying on top of relatively transparent Baltica crust. East of these locations, the sedimentary rocks become progressively more deformed as the Main Uralian

Table 3. Strike and dip of reflection zones Zone MUTF MUNF SM ST KA1 PSZ SIS D1 D2 D3 LC M1 NF2 X

Strike (degrees)

Dip (degrees)

Intersectssurface

- 13 - 13 178 - 25 - 25 160 190 145 145 145 120 70 - 5 140

45 60 60 4O 40 30 40 25 25 25 45 40 40 40

Y Y Y Y Y Y Y N N N N N Y Y

Fault Zone is approached. Equivalents to the Kvarkush Anticline and the Serebrianka Thrust on the E S R U - S B section are missing on the Michailovsky profile. Instead, a nappe, consisting of Ordovician to Devonian volcanic, shallow and deep marine rocks (Puchkov 2002), is present at about km - 2 0 to km -10. This nappe was overthrust onto the early Permian sedimentary rocks during the final stages of the collision. The seismic images indicate shortening of the East European basement along the E S R U - S B and Michailovsky sections. It is not possible from the images alone to determine how much of this shortening is Uralian and how much much is pre-Uralian. Within the Kvarkush Anticline, a Timanian thrust complex of Neoproterozoic metasedimentary rocks, dolerites and gabbro includes a glaucophane garnet-bearing allochthon (Beckholmen & Glodny 2004). This Timanian thrust assemblage composes the footwall to a Neoproterozoic juvenile island-arc complex, itself flanked to the NE by hinterland terranes, including both ophiolites and microcontinents. The Neoproterozoic ocean derived allochthons are absent in the Middle Urals; they may have been rifted off the eastern margin of Baltica during development of the Early Palaeozoic passive margin. It is likely that the Kvarkush Neoproterozoic thrust system was reactivated both in the extensional mode during the attenuation of Baltica's passive margin and subsequently during Uralian compression. Below c. 5 km in the Michailovsky profile, between km - 6 0 and km - 5 horst and graben structures covered by thick sedimentary sequences are present (Fig. 6). Extended Neoproterozoic sedimentary rocks are a likely candidate for these structures. Therefore, we interpret the structure along this section of the Michailovsky profile to have formed during the opening of the palaeo-Uralian ocean in the Mid-Ordovician. The Serebrianka Thrust (ST) projects to the surface close to the start of the ESRU93 (km - 30) profile near the boundary between Lower Vendian and Riphean rocks in the Kvarkush Anticline.

CRUSTAL S T R U C T U R E OF T H E M I D D L E URALS

435

~ r

~< ..= ~)

.L~~

r

r

[.-,~ 9

I

. ....,

f.,t3

~D

o~

o~

436

S. KASHUBIN ETAL.

Fig. 6. Migrated line drawings of the Michailovsky, Alapaevsk-Chernoistoichinsk and western part of the ESRU-SB profiles. Sections are plotted so that km 0 corresponds to the western limit of the Main Uralian Fault Zone as defined earlier. The flat-lying reflections on the western parts of the ESRU-SB and Michailovsky profiles represent Neoproterozoic to Permian sedimentary rocks.

Reflections from this zone can be traced to at least 5 s (Fig. 4a) and possibly to as deep as 1 0 - 1 1 s. The zone has been interpreted as extending down into the middle crust (Friberg et al. 2002) and acted as a detachment zone when the Tagil Arc accreted onto the East European Craton. Knapp et al. (1998) presented an alternative geometry for this zone, suggesting that it becomes subhorizontal at a depth of about 10 km and that it merges with the Main Uralian Fault Zone further east. We prefer the first interpretation, as its projection down into the crust corresponds to the westernmost limit of the strong reflectivity between depths of 20 and 30 km at km 0 to km 30 along the section (Fig. 5). A possible equivalent to the Serebrianka Thrust in the South Urals is the Zuratkul Fault (Brown et al. 1996). There, hanging-wall rocks that were metamorphosed and deformed in the Proterozoic are were thrust over similar, but less metamorphosed units. Brown et al. (1996) interpreted the Zuratkul Fault as a Proterozoic structure that was reactivated during the Uralide orogen; the same may be true for the Serebrianka Thrust. Based on the seismic image (Fig. 5) the hinge of the Kvarkush anticline (km - 45 to km 0) is located at km - 30. East of km - 2 0 the seismic image of the upper crust is poor. This may be due to the poor data quality in the upper crust between km - 2 0 and km 0 or structural complexity. However, surface mapped faults all dip to the east, suggesting that moderately to steeply east-dipping

structures are present. A reflection (KA 1) with similar strike and dip to the ST reflection is imaged from about 3 s at km 2.5 to at least 6 s at km 15 (Fig. 4a). This reflection projects to the surface at km - 1 1 , close to a surface mapped fault. This reflection was not as clearly imaged on previous processing and Friberg et al. (2002) did not focus at all on its interpretation. Knapp et al. (1998) suggested that this zone does not extend to the surface, but represents a pre-Uralian thrust that merges into a detachment zone at 10 km depth. However, the correlation between the projection of this reflection to the surface and a surface mapped fault and its orientation suggest that it is related to the same thrust system as the Serebrianka Thrust. M a i n Uralian Fault Zone. The pronounced reflectivity dipping at

4 5 - 6 0 ~ between km 0 and km 10 originates from the Main Uralian Fault Zone, as it has been defined in previous work (Juhlin et al. 1998; Friberg et al. 2002). This zone corresponds to the highly deformed margin of the East European continent and contains some tectonic fragments of oceanic rocks forming a m~lange. Strongly deformed sedimentary and basalt complexes of Neoproterozoic and Early Palaeozoic age are present in this m~lange between km 0 and km 10. The eastern boundary of the reflective zone marks the contact between the deformed former continental margin and the Tagil Arc. Controversy exists as to

CRUSTAL STRUCTURE OF THE MIDDLE URALS

437

Fig. 7. Perspectiveview of the Michailovsky,Alapaevsk-Chernoistoichinskand ESRU-SB profilesplotted onto the magnetic map. Reflectors have been projected to the surface as with the geometries definedin Table 3. Red reflectors can be traced to the surface on the seismic sections whereas blue ones cannot.

the definition of the Main Uralian Fault along the ESRU profile. Some workers (Knapp et al. 1998; Brown et al. 2002) considered the boundary to be well defined, separating the deformed East European continental margin from the Tagil Arc, and placed it at about km 10. About 40 km to the south along the Chernoistochinsk seismic profile (Fig. 1) the Main Uralian Fault was interpreted to correspond to the base of a highly reflective package (km 0 in Fig. 6) of similar appearance and geometry to that observed along the ESRU profile (Steer et al. 1995). It should be noted that Brown et al. (2002) placed the boundary 19 km further east along this profile. Linear magnetic anomaly highs along the deformed margin rocks on the 1:200 000 magnetic map (unpublished) suggest that mafic oceanic rocks may be present within these reflection packages. Therefore, we prefer to use the term Main Uralian Fault Zone as defined by Friberg et al. (2002) and do not mark the boundary between the East European continent and the oceanic rocks in the Tagil Arc by a distinct fault. Tagil arc. Apparent east- and west-dipping reflections characterize the upper crust of the Tagil Arc. Some of the west-dipping ones can be attributed to out-of-the-plane effects, but a number of them are related to east-directed thrusting that occurred during the accretion of the Tagil Arc. East-dipping reflections below km 25 have been penetrated by the c. 5.5 km deep SG-4 borehole and have been shown to be generated by fracture zones dipping at 35-45 ~ (Ayarza et al. 2000b). Further east,

the east-dipping reflections (Friberg et al. 2002).

represent

Palaeozoic

thrusts

S e r o v - M a u k zone. The Serov-Mauk Zone, at the latitude of the ESRU profile, is marked by a highly magnetic north- south trending serpentinite belt (Fig. 1). The zone is not well imaged on the seismic profile, but weak steeply dipping reflections (SM in Fig. 5, and SMF in fig. 10 of Juhlin et al. 1998) indicate that the zone dips 60 ~ to the west. High-grade rocks to the east of the Serov-Mauk Zone indicate that the serpentinite belt is bounded by a normal fault, and the seismic data have been interpreted accordingly (Juhlin et al. 1998; Friberg et al. 2002). It should be noted that the SerovMauk Zone is part of a larger strike-slip system and the reflectors that we have imaged may be local features. It is possible that on a larger scale the Serov-Mauk Zone is a major strike-slip fault system as suggested by Brown et al. (2002), and this possibility is also indicated in Figure 5. Other workers have interpreted the Serov-Mauk Zone at latitude 58~ to be both less steeply dipping to the west (Knapp et al. 1998), and more steeply dipping, but also to the west (Steer et al. 1995). Our interpretation of a 60 ~ west-dipping geometry at this latitude is consistent with earlier Russian interpretations of the Serov-Mauk Zone (i.e. Sokolov 1992), but differs from the east-dipping geometry of the zone in the South Urals (Sokolov 1988).

metamorphic complex. A major upper crustal feature on the seismic image is the prominent set of west-dipping reflections,

Salda

438

s. KASHUBINETAL.

the base of which projects to the surface near km 80 (PSZ in Fig. 5). This reflection package corresponds to the Prianitchnikova Shear Zone, which was discovered by the ESRU project. It separates granulites (mainly mafic) of island-arc affinity, in the west, from low-grade amphibolites to the east (Friberg & Petrov 1998). At about 10 km depth at km 65 there is a break in the reflection package. This break in the reflections together with the presence of more steeply dipping reflections indicate a normal fault associated with the shear zone (Juhlin et al. 1998), although the apparent geometry could be due to 3D effects. It should be noted that the Prianitchnikova Shear Zone has a non-linear geometry on the surface (fig. 2 of Juhlin et al. 1998). However, the synformal package of reflections above the PSZ, which are better imaged in the present processing, are consistent with a west-dipping normal fault (NF1 in Fig. 5). Sisertfault. At km 102, the profile crosses the Sisert Fault, known to have had a large component of strike-slip motion on it. The data quality just east of the fault is poor; however, there are few reflections that can be traced across it after migration (Fig. 5). The Sisert Fault appears to be near-vertical, but it is not clear how far it extends into the crust. The relatively continuous lower crustal reflectivity below it indicates that it does not penetrate to Moho depths as a near-vertical fault. metamorphic complex. At the latitude of the E S R U - S B profile the boundary between the metamorphic complexes to the west and the Alapaevsk arc complex to the east coincides with the Sisert Fault. However, to the south (Fig. 1) the Murzinka-Adui Metamorphic Complex (MAMC) is thrust onto the Alapaevsk Volcanic series along the Deevo Fault (Friberg et al. 2002). Ophiolite sequences and metamorphic rocks have been thrust eastwards onto Late Devonian volcanic rocks, Early Carboniferous marine strata, and Late Carboniferous molasse deposits. East of the Sisert Fault, along the E S R U - S B profile, metamorphic rocks are exposed. These are partly covered by Carboniferous limestones of the Alapaevsk Arc. Between km 112 and km 120, a Permian granite (Grevtsova et al. 1970) intrudes into the metamorphic rocks. A reflection dipping at about 40 ~ to the east (NF2 in Fig. 5) projects to the surface at km 120, near the boundary between this granite and the Alapaevsk rocks to the east. This reflection has been interpreted to represent a normal fault (Friberg et al. 2002), based on the lack of contact metamorphism in the Alapaevsk limestones (Sobolev et al. 1964). The strong reflectivity at about 10 krn depth at km 120 may originate from mafic intrusions associated with the granite. Murzinka-Adui

Alapaevsk arc. Relatively transparent crust is present below the

exposed Alapaevsk Arc (Fig. 5) except at about 1 - 2 s at km 140. These gently west-dipping reflections at 140 km could originate from a west-dipping thrust fault or represent layered structures within the Alapaevsk volcano-sedimentary strata. West Siberian Basin Mesozoic sediments cover the area east of km 147. However, the transparent upper crust continues to about km 190, suggesting that the volcanic and sedimentary units of the Alapaevsk Arc are present up to this point. Between about 3 and 8 s from km 200 to km 260 piece-wise strong reflections are present in the crust, suggesting a crust of a different nature from the Alapaevsk crust. Trans-Uralian zone. The boundary between the Alapaevsk Arc (part of the East Uralian Zone) and the Trans-Uralian Zone is covered by Mesozoic and Cenozoic sediments and exposed only in a few river valleys. Where exposed, it appears as a west-dipping blastomylonite zone that can be projected onto the E S R U - S B profile between km 180 and km 200. Brown et al. (2002) argued that the prominent magnetic high running in a north-south direction between 62~ and 63~ south of 58~ (Fig. 1), represents the boundary, based on observations in the Southem Urals. It is not

clear from the magnetic map how to extend this boundary up to the E S R U - S B profile, but the anomaly appears to intersect the profile between km 160 and km 200. The change to a significantly more reflective crust is approximately defined by a line starting at km 195 at the surface and extending to km 160 at 40 km depth (Fig. 5). Although the exact location of the boundary differs somewhat from that interpreted by Brown et al. (2002), we interpret the crust east of this line as Trans-Uralian and that to the west as East Uralian.

M i d d l e a n d l o w e r crustal f e a t u r e s

A reflective zone, about 12-15 km thick, characterizes the East European lower crust on the E S R U - S B section (Fig. 5). This reflective zone has been interpreted as East European Craton crystalline lower crust (Juhlin et al. 1995); it deepens towards the east (Fig. 5). The East European crust is also observed to thicken to the east further south along the Michailovsky profile (Fig. 6). Loading of the East European continental margin with the accreted Tagil Arc resulted in bending of the crust. The reflectivity extends to about km - 1 0 along the E S R U - S B profile (Fig. 5), suggesting that this is the eastern limit of intact East European crystalline crust. Further east, intense middle crustal reflectivity is present between km 0 and km 30. This reflectivity dips to the east and appears to be limited to the west by the Serebrianka Thrust (Fig. 5). The most likely explanation for this reflectivity is that it represents highly deformed rocks of the passive continental margin (Friberg et al. 2002). However, it could also represent East European lower crust that has been thrust upwards during the accretion process. Below the strong middle crustal reflectivity, west-dipping reflections are present beginning at km - 3 0 at 60 km depth and extending to km 100 at about 25 km depth (TF in Fig. 5). These west-dipping reflections were not imaged as clearly below the Tagil Arc on the earlier processing (i.e. Juhlin et al. 1998), which did not include the Bazhenov component of the ESRU93 data. They indicate the deepest part of the crustal root to lie under the Kvarkush Anticline, not the Tagil Arc. The west-dipping nature of these reflections suggests that they are related to arc material that was underthrust from the east, implying that the lowermost crust below the Kvarkush Anticline is Palaeozoic in age, not Precambrian. The reflective zone labelled TF in Figure 5 has earlier been referred to as the Trans-Uralian Thrust Zone (Juhlin et al. 1998) or the Deevo Fault (Friberg et al. 2002). We prefer not to attach a name to this reflective zone, as we cannot trace it to the surface. However, arc material appears to have been underthrust along this zone as far west as km - 3 0 (Fig. 5) and we consider it to be a major feature of the Middle Urals. East of km 50 to about km 150, the lower crust contains fairly weak subhorizontal to slightly west-dipping reflections (Fig. 5). Although the reflectivity is weak, the reflective Moho is reasonably well defined along this portion of the profile. East of km 150, the lower crust, as well as much of the entire crust, becomes more reflective. There are a number of clear reflective zones in the lower crust with a westerly dip component. Both cross-dip analysis (Friberg 2000) and crooked-line reflection modelling (Fig. 4b) indicate these reflective zones to strike N35W and dip about 25 ~ to the SW. These zones are distinct in the lower crust, but cannot be traced into the upper crust. East of km 150, the reflective Moho is not as well defined, but it generally appears to shallow towards the east.

M a n t l e reflections

Two clear reflections are present at times later than 20 s below the SB01 profile (Fig. 2). The more steeply dipping one can be

CRUSTAL STRUCTURE OF THE MIDDLE URALS

modelled as reflecting from a zone striking NNW and dipping 45 ~ to the SW close to the crust-mantle boundary (LC in Fig. 4c). It does not necessarily originate from a structure in the mantle. However, it is tempting to connect the deep west-dipping reflections below the Kvarkush anticline to this reflection. The less steeply dipping one, at about 23 s, is best modelled as reflecting from a structure striking WSW and dipping 45 ~ to the SSE (M1 in Fig. 4c) at depths of 6 5 - 7 0 km in the mantle.

Strike-slip f a u l t s y s t e m s

Major, longitudinal, transcurrent faulting has been inferred to dominate at least the later stages of the Uralian orogeny. Three candidate zones for major late- to post-orogenic strike-slip movement are crossed by the E S R U - S B profile. These are the Main Uralian Fault Zone, the Serov-Mauk Zone and the Sisert Fault (Fig. 5). Of these, the Sisert Fault is the most apparent on the seismic section (Fig. 5), as discussed above. Transparent upper crust below the Serov-Mauk Zone could be interpreted as the result of strike-slip movement; Brown et al. (2002) interpreted the Serov-Mauk zone to extend down to Moho depths based on some disruptions in the reflectivity in the middle and lower crust below the fault zone (Fig. 5). There is little evidence for nearvertical strike-slip motion along the Main Uralian Fault Zone in the seismic section (Fig. 5). However, minor strike-slip motion could have occurred along the existing east-dipping structures. Regardless of the amount of strike-slip motion that has taken place along these zones, movement out-of-the-plane of the profile needs to be considered in any interpretation.

The crustal r o o t

Scarrow et al. (2002) and Brown et al. (2003) recently presented integrated interpretations of seismic, gravity, magnetic and topographic data from the Southern Urals. In their interpretation the crustal root consists of mafic granulite material that is denser than the lower crustal material in the adjoining East European Craton. This dense root, together with mafic bodies in the upper crust, results in near-isostatic compensation at 70 km depth, in spite of the topographic variations at the surface. A key component in this model is that the East European lower crust does not extend very far to the east below the Main Uralian Fault. This is consistent with both recently proposed models based on interpretation of seismic refraction data (Stadtlander et al. 1999) and older models based on gravity data and topography (Kruse & McNutt 1988). However, this model contrasts sharply with those suggesting that East European crust extends far to the east below the Magnitogorsk-Tagil arcs (Berzin et al. 1996; Poupinet et al. 1997; Knapp et al. 1998; Drring & G6tze 1999; Diaconescu & Knapp 2002). A similar integrated interpretation has not been made for the Middle Urals, as a complete dataset is not yet available. However, Kimbell et al. (2002) concluded that magnetic East European crystalline crust cannot be present below the Tagil Arc. Instead, they interpreted magnetic Trans-Uralian crust to be present below the Tagil Arc. This interpretation implies that the Tagil Arc and the East Uralian Zone were thrust towards the east onto Trans-Uralian crust and the crustal root consists mainly of Trans-Uralian crust, not East European crust. Our interpretation differs somewhat from that of Kimbell et al. (2002) in that we extrapolate the boundary between the TransUralian crust and Alapaevsk Arc down at a much steeper angle based on the reflectivity contrast between the two interpreted crustal terranes (Fig. 5). Instead, we suggest that Alapaevsk crust comprises the main portion of the crustal root. Although the Alapaevsk crust includes significant volumes of limestone and appears to be relatively non-magnetic near the surface

439

(Kimbell et al. 2002), it is possible that it contains older basalts, andesites and their intrusive analogues at depth. It would then become more magnetic at depth, resulting in a magnetic field consistent with that along the ESRU profile (Fig. 2). Most researchers favour an interpretation where the present-day topography of the Ural mountains is a relatively recent feature (e.g. Borisevich 1992; Seward et al. 2002), but where the crustal root is a remnant of the Palaeozoic collision (e.g. Kruse & McNutt 1988; Thouvenot et al. 1995; Juhlin et al. 1998). Our crustal model is consistent with this interpretation, where the gently west-dipping reflectivity in the crustal root represents mafic Alapaevsk crust that was underthrust below the Tagil Arc, and possibly under the East European lower crust. Compression in the Tertiary resulted in uplift along pre-existing Palaeozoic, or earlier, structures in the foreland fold-and-thrust belt, producing the present-day Ural Mountains.

Implications for tectonic evolution Based on the ESRU seismic reflection data and geological mapping in the Middle Urals, the following tectonic evolution for the area was suggested by Friberg et al. (2002). Rifting of the East European continent began in the Early Ordovician and a passive margin developed along the eastern margin of Baltica, which lasted until the Carboniferous. By the Early Silurian, an intra-oceanic east-dipping subduction zone located east of the spreading centre was active, resulting in the development of the Tagil island-arc and back-arc complexes in the east (Petrokamensk-Salda and Alapaevsk terranes). In this preferred evolutionary model, Devonian magmatism was active in the eastern part of the Tagil Arc (Krasnoturjinsk volcanic belt), and in the Petrokamensk and Alapaevsk island arcs, as a result of a west-dipping subduction zone further east. At this time, granitic complexes intruded into the westernmost arcs as this composite terrane was sutured to the East European continent. In the Late Permian, the Siberian craton collided with this composite terrane, resulting in uplift. Subsequent to the collision, or possibly contemporaneously with it, north-south strike-slip faults were active in the hinterland, controlled mainly by the terrane boundaries (Ayarza et al. 2000a). Mesozoic extension resulted in minor stretching of the crust below the present-day western margin of the West Siberian Basin. Some of this extension may have been accommodated along SW-dipping structures in the lower crust below the basin, leaving many of the Palaeozoic collisional structures to the west intact, as well as the crustal root (Juhlin et al. 1996; Morozova et al. 1999). The additional seismic data presented here are consistent with the above scenario. However, some details of the history may be highlighted. As on the URSEIS profile to the south (Steer et aL 1998), the East European crust at 58~ is seen to thicken towards the east as the Main Uralian Fault Zone is approached (Fig. 5). Based on the geometry of the East European lower crustal reflectivity, we interpret this thickening to be due to loading by stacking of thrust sheets onto the lower crust. The gently west-dipping reflectivity in the crustal root indicates that it reaches a depth of nearly 60 km just west of the surface exposure of the Main Uralian Fault Zone. This is somewhat deeper than previous interpretations. The revised interpretation is due to the inclusion of the Bazhenov ESRU93 and the SB02 seismic data over the root zone, resulting in a better image of the deep reflectivity. The deeper estimate from the seismic reflection data presented here, compared with that from refraction-DSS data (Thouvenot et al. 1995; Juhlin et al. 1996), may be due to the better resolution of the former. Arc crust is interpreted to be the main component of the crustal root, not East European crust. We have earlier suggested that the lower crustal reflectivity below the Asian terranes may be related to the formation of the West Siberian Basin (Juhlin et al. 1998). However, this appears to be unlikely

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based on the geometries in the lower crust east o f k m 160 (Fig. 5). The SW-dipping structures in the lower crust east of km 160 are probably related to the accretion process that formed the crust in the later stages of the orogeny. These structures indicate that the accretion was from the present-day NE, not from the east. Their existence shows that lower crustal deformation related to the development of the West Siberian Basin must have been minor below the eastern part of the E S R U - S B profile.

Conclusions The seismic images presented in this paper are the result of more than 10 years of co-operation between Russian and western scientists. Our present interpretations will certainly be revised in the future, but the information that the seismic images contain will not change significantly. Based on these images, surface geology and other geoscientific data we have made the following interpretations. Most upper crustal reflective elements can be correlated with significant faults and geological boundaries at the surface. The major shear zones appear to be listric into a n d / o r truncated by zones of middle to lower crustal reflectivity. One major reflective zone (TF in Fig. 5) appears to extend from the lower crust into the upper crust. The west-dipping reflectivity at the base of the crust below this reflective zone, which has become more apparent with the present processing, suggests that the crustal root extends to depths of 5 5 - 6 0 km and is composed of Asian arc material. The hinterland east of the Main Uralian Fault Zone in the Middle Urals appears to be dominated by Palaeozoic island-arc terranes, in contrast to the South Urals where continental terranes are present further east. Except for under the Tagil Arc, the Moho under these accreted terranes is generally flat and continuous at 4 0 - 4 5 km depth; however, the character of the lower crustal reflectivity changes towards the east where more SW-dipping structures are present. We interpret these to represent collisional features in Trans-Uralian lower crust. Their presence implies that the amount of stretching below the westernmost part of the West Siberian Basin must have been minor. This is in contrast to earlier suggestions of major extension in the area (i.e. Juhlin et al. 1993, 1998; Knapp et al. 1998). We would like to thank the field teams who carried out these experiments, especially the BGE and Uppsala field leaders, T. Hismatulin (deceased), S. Nesterov and H. Palm. Presentation and discussion of Uralide geology and geophysics at EUROPROBE workshops laid the foundation for much of this work. We thank all our colleagues in the Urals Working Group for their input. C.J. thanks the Swedish Research Council (VR) for funding.

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SVIAZHINA, I., PUCHKOV, V. & PETROV, G. 2003. Paleomagnetism of Middle-Paleozoic series in Urals. In: KOROTEEV, V. A. & OSIPOVA, T. (eds) Evolution of intracontinental Mobile Belts: Tectonics, Magmatism, Metamorphism, Sedimentology, Useful Minerals. Institute of Geology and Geochemistry, Yekaterinburg, 62-64 [in Russian]. THOUVENOT, F., KASHUBIN, S. N., POUPINET, G., MAKOVSKIY, V. V., KASHUBINA, T. V., MATTE, P. & JENATTON, L. 1995. The root of the Urals: evidence from wide-angle reflection seismics. Tectonophyics, 250, 1-13. TORSVIK, T. H., SMETHURST, M. A., VANDERVOO, R., TRENCH, A., ABRAHAMSEN, N. & HALVORSEN, E. 1992. Baltica--a synopsis of Vendian-Permian paleomagnetic data and their paleotectonic implications. Earth-Science Reviews, 33, 133-152. TORSV1K, T. H., TAIT, J., MORALEV, V. M., MCKERROW, W. S., STURT, B. A. & ROBERTS, D. 1995. Ordovician paleogeography of Siberia and adjacent continents. Journal of the Geological Society, London, 152, 279-287.

VALASEK, P., MUELLER, S., FREI, W. & HOLLIGER, K. 1991. Results of NFP-20 seismic reflection profiling along the Alpine section of the European Geotraverse (EGT). Geophysical Journal International, 105, 85-102. Wu, J., MILKEREIT, B. & BOERNER, D. E. 1995. Seismic imaging of the enigmatic Sudbury structure. Journal of Geophysical Research, 100, 4117-4130. ZALESKI, E., EATON, D. W., MILKEREIT, B., ROBERTS, B., SALISBURY, M. & PETRIE, L. 1997. Seismic reflections from subvertical diabase dikes in an Archean terrane. Geology, 25, 707-710. ZONENSHA1N, L. P., KORINEVSKY, V. G., KAZMIN, V. G., PECHERSKY, D. M., KHAIN, V. V. & MATVEENKOV, V. V. 1984. Plate tectonic model of the South Urals development. Tectonophysics, 109, 95-135. ZONENSHAIN, L. P., KUZMIN, M. & NATAPOV, L. M. 1990. Uralian foldbelt. In: ZONENSHAIN, L. P., KUZMIN, M. & NATAPOV, L. M. (eds) Geology of the USSR: a Plate-Tectonic Synthesis. American Geophysical Union, Geodynamics Series, 21, 27-54.

U - P b Silurian age for a gabbro of the Platinum-bearing Belt of the Middle Urals (Russia): evidence for beginning of closure of the Uralian Ocean DELPHINE B O S C H 1, OLIVIER B R U G U I E R 2, A L E X A N D E R A. E F I M O V 3 & A R T H U R A. K R A S N O B A Y E V 3

tLaboratoire de Tectonophysique, Universiti de Montpellier II, UMR 5568-CNRS/UMII, Place E. Bataillon, 34095 Montpellier Cedex 05, France (e-mail: bosch@dstu, univ-montp2.fr) 2Service ICP-MS, ISTEEM, Universitd de Montpellier II, Place E. Bataillon, 34095 Montpellier Cedex 05, France 3Institute of Geology and Geochemistry, Pochtovyi Pereuloic 7, Ekaterinburg 620151, Russia

Abstract: The Platinum-bearing Belt of the Urals consists of a series of zoned ultramafic bodies obducted onto the passive continental margin of the East European Craton during the Palaeozoic. Conventional U-Pb and secondary ionization mass spectrometry U-Th-Pb analyses of single zircon grains from a pegmatitic gabbro of the Kumba massif provide Mid-Silurian ages of 425 + 3 Ma and 419 _+ 10 Ma, respectively, interpreted as dating crystallization of the gabbroic magma. This contrasts with the c. 360 Ma age of the neighbouring Kytlym massif and indicates that the Uralian Platinum-bearing Belt is best interpreted as remnants of an island-arc oceanic lithosphere that started forming in Mid-Silurian times and had a protracted lifetime of about 70 Ma. The Uralian Platinumbearing Belt is thus coeval with the Sakmara arc of the Southern Urals, and is of similar age to other ultramafic bodies such as the Kempersai and Mindyak massifs. Ophiolitic fragments, now preserved along the Main Uralian Fault, may thus represent the assemblage of contemporaneous oceanic and arc terranes brought together during the final stages of the evolution of the Uralide orogen. The age of the Kumba massif, in addition, indicates that the Uralian Ocean underwent contractional events as early as the Mid-Silurian, which suggests a possible link with the pivotal rotation of Baltica.

The convergence of the East European Craton (EEC) and the Siberian-Kazakhstan terrane assemblage that ultimately led to the formation of the Urals mountain belt at the end of the Palaeozoic was preceded by closing of the Uralian Ocean that separated the various continental masses. Remnants of oceanic lithosphere that became trapped along the suture between the colliding continents provide important time markers for the pre-collisional history of the orogen. Ophiolite complexes are numerous in the Urals (Savelieva & Nesbitt 1996) and crop out in the hanging wall of the Main Uralian Fault (MUF), the main suture zone that runs the entire length of the mountain belt (Matte 1995). These massifs appear as lens-shaped bodies mainly of lherzolitic and harzburgitic composition elongated parallel to the MUF (Savelieva et al. 1997). The continuity of the belt and similar structural position of the massifs (i.e. allochthonous position overlying either the Precambrian basement or Palaeozoic sediments of the EEC margin) suggest a common origin. However, whereas some massifs show a nearly complete petrological association typical of ophiolite sequences, others lack the upper gabbro, sheeted dyke and lava complexes, and crop out as zoned plutons of mafic-ultramafic rocks. Thus, it is unclear whether the c. 2000 km long oceanic terrane exposed in the Urals represents fragments of a single extensive dismembered ophiolite (where, in some cases, the upper lava complexes have been removed by tectonic processes) or whether some massifs are remnants of island-arc terranes derived from the Uralian Ocean and progressively telescoped with the EEC margin. Despite their importance in understanding the evolution and destruction of the Uralian Ocean and thus the earliest evolutionary stages of the Uralide Orogen, isotopic age information is sparse on these complexes. Most of our knowledge is from a few isotopic datings on Alaskantype zoned mafic-ultramafic complexes located in the Uralian Platinum-bearing Belt (UPB; Ivanov & Kaleganov 1993; Bea et al. 2001; Ronkin et al. 2003), and from typical ophiolitic massifs displaying well-developed plutonic and volcanic units, such as the Kempersai (Southern Urals) and Voykar massif (Polar Urals) (Edwards & Wasserburg 1985; Sharma et al. 1995; Melcher et al. 1999). In this study, we present conventional U - P b single-zircon analyses and U - T h - P b in situ secondary ionization mass spectrometry (SIMS) analyses for a pegmatitic

gabbro of the Kumba massif, which crops out in the Uralian Platinum-bearing Belt. The significance of these results is discussed in relation to the available isotopic data from other massifs in the Platinum-bearing Belt and from other occurrences along the Main Uralian Fault in an attempt to understand the evolution of the Uralian Ocean.

Geological setting The Uralian Platinum-bearing Belt (UPB) forms a giant structure exposed in the Central and Northern Urals (see Fig. l a). It is located east of the MUF where the suture sector widens and is characterized by the abundance of zoned mafic-ultramafic complexes. These massifs are similar to Alaskan-type complexes (Irvine 1967) and show the typical association dunite-clinopyroxenite-gabbro (DCG units of Fershtater et al. 1997). The UPB consists of a series of 13 rounded to elongated massifs forming a discontinuous but linear elongated zone, which crops out for a distance of about 900 km (Fig. la). The mafic-ultramafic bodies are associated with rocks of the Tagil Zone, which consists of Early Ordovician to Early Silurian volcanic and volcano-sedimentary units, tectonically associated with shelf-facies and continental-slope deposits. They commonly developed a thermal aureole in the volcanosedimentary sequences (Savelieva & Nesbitt 1996). The massifs are mainly composed of gabbros (olivine gabbros, gabbro-norites, clinopyroxene-hornblende gabbros), and also of clinopyroxenites, wehrlites, dunites, pegmatitic gabbros and plagiogranites (Fig. lb) with a progressive magma differentiation from dunite through wehrlite- clinopyroxenite to gabbro. Sharp compositional changes of rock associations suggest that intrusions were pulsating (Savelieva et al. 1999). In the study region, field relationships indicate a complex allochthonous association with at least two main nappe systems. Structurally lower units consist of Late Ordovician to Early Silurian deposits of palaeo-oceanic affinities (greenstone basalts, metacherts, psammitic and pelitic tufts, along with tectonized relics of possibly Ordovician ophiolites), whereas the structurally upper system is composed of a stack of tectonic slices. These are composed of volcano-sedimentary and volcaniclastic terrigenous units of Late Ordovician to Early Silurian age in contact with and apparently

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 443-448. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

443

444

D. BOSCH E T A L .

Fig. 1. Simplified geological map of the Uralide orogen (a) and a detailed map of part of the Platinum-bearing Belt close to the sample locality (b). Square with cross indicates sample location. The simplified map is modified after Friberg et al. (2000); (b) is from Efimov et al. (1993).

intruded by the mafic-ultramafic bodies of the UPB (Savelieva & Nesbitt 1996).

Sample description The Kumba massif is located in the centre of the UPB, and is one of the smallest of the 13 massifs of the belt. The general shape of the Kumba massif is a 30~ dipping oval form. This massif consists of the typical association of dunite and clinopyroxenite enclosed by gabbros and gabbro-norites (Fig. lb). In the northern part, it mainly consists of olivine-anorthite gabbro with rare pyroxenites, the central part contains gabbro-norites, and the western part comprises dunites, pyroxenites and some metamorphic products. The study sample is an amphibole pegmatitic gabbro enclosed in gabbro-norites. The mineralogical assemblage consists of plagioclase, clinopyroxene, orthopyroxene, hornblende, spinel, ilmenite and magnetite. The sample is fresh without evidence of alteration. Zircon appears as a scarce accessory mineral in this rock and more than 50 kg of rock fragments have been crushed to perform this study.

Analytical procedures The minerals were prepared for U - P b analyses according to standard techniques (e.g. Bosch e t al. 1996) where only the best quality zircons (non-magnetic and free of visible inclusions and fractures) were selected for analysis. For thermal ionization mass spectromery (TIMS) analyses, single grains were weighed on a Cahn electrobalance and then processed according to Bosch & Bruguier (1998). Total Pb and U blanks over the period of the analyses were < 15 pg and < 5 pg ( + 5 0 % ) , respectively. The

isotopic composition of radiogenic Pb was determined by subtracting first the blank and then the remainder, assuming a common Pb composition at the time of initial crystallization as determined from the model of Stacey & Kramers (1975). Isotopic measurements were carried using on a VG Sector mass spectrometer at the University of Montpellier II, with a Daly detector. SIMS analyses were carried out on the same zircon population as was analysed by TIMS. The grains, together with chips of standard zircon, were mounted in epoxy resin and polished to approximately half their thickness to expose internal structure. U T h - P b analyses were performed with an elliptical spot size of about 25 p~m • 35 I~m on the CAMECA IMS 1270 ion microprobe at the CRPG Nancy (France), following the technique outlined by Deloule e t al. (2001). Isotopic ratios were measured with a primary O 2 beam of 10 nA at a mass resolution of c. 5000, at which no significant interferences on the Pb, U and Th isotopes were detected. Oxygen flooding was used to enhance sensitivity. Under these operating conditions, the sensitivity for Pb isotopes was c. 20 cps per ppm per nA of primary beam. P b / U ratios were normalized using quadratic working curves, to values measured on the G91500 standard zircon (Wiedenbeck e t al. 1995). Common Pb was corrected using measured 2~ and a composition taken from the model of Stacey & Kramers (1975). Corrected isotopic ratios, regression lines and intercepts were calculated according to Ludwig (1987). Decay constants are those recommended by the IUGS Subcommission on Geochronology (Steiger & Jager 1977).

Results Zircons occur as a uniform population of large ( > 100 ~zm) colourless translucent grains. They show euhedral shapes with sharp

SILURIAN AGE FOR PT-BEARING BELT, MIDDLE URALS

angle terminations, often interrupted by irregular boundaries (see left upper termination of grain B in Fig. 2), which we interpret as reflecting late magmatic crystallization in a small melt volume adjacent to already crystallized minerals. SEM images of polished surfaces (Fig. 2a and b) show that the grains exhibit almost structureless internal domains, although irregular, broad-banded zoning can be recognized, which is often observed in magmatic zircons from mafic lithologies (e.g. Rubatto 2002). All grains have high

445

U concentrations ranging from c. 200 to 2000 ppm (see Tables 1 and 2). These U contents are surprising for zircons from gabbroic rocks, which usually have U concentrations in the range 5 0 400 ppm (Sch/irer et al. 1986; Pedersen & Dunning 1997), but concentrations up to 6000 ppm have been reported (Dunning & Pedersen 1988). This is likely to reflect a large degree of differentiation before zircon crystallization, consistent with the pegmatitic nature of the rock studied and with crystallization of the zircon in the latest stages of the magmatic evolution. Four of the six grains analysed by TIMS were air abraded (Krogh 1982) to minimize Pb losses. On a concordia diagram (Fig. 3a), the abraded grains cluster close to concordia and, with the two unabraded grains, yield a discordia line with upper and lower intercept ages of 424.9 +_ 2 . 7 M a and - 2 6 + 8 M a (MSWD = 0.72). The unabraded grains (Zr5 and Zr6) are discordant, which indicates Pb loss. The near-zero lower intercept indicates recent, surface-correlated, Pb loss rather than an ancient, metamorphic disturbance of the U - P b system of these zircons. The high degree of concordance of the abraded zircons and the magmatic morphology of the grains indicate that the 424.9 _+ 2.7 Ma intercept age corresponds to the age of zircon growth during crystallization of the pegmatitic magma. Sixteen SIMS analyses were performed on 14 grains, which displayed T h / U ratios typical of magmatic zircons ( T h / U > 0.1; after Williams & Claesson 1987). Overall, the SIMS analyses yielded a pattern consistent with the TIMS results (Fig. 3b). Although all analysed grains are concordant within error margins, spot analyses yield a tendency to scatter along the discordia between 390 and 430 Ma, suggesting that the grains have undergone various degrees of Pb loss. The best age information is thus given by a discordia line where the upper intercept with the concordia is 419 + 10 Ma. This age is slightly younger, but within error, of the more precise TIMS age, which is interpreted as our best estimation of the age of the Kumba gabbro.

Discussion and conclusion

Fig. 2. Back-scattered SEM images of zircons from the Kumba gabbro (Platinum-bearing Belt, Central Urals).

The Wenlock U - P b zircon age of 424.9 _+ 2.7 Ma obtained for the Kumba gabbro provides a new time constraint for the formation of the zoned mafic-ultramafic massifs that crop out in the UPB, which has implications for the early evolution of the Uralide orogen. Available geochronological data for mafic and ultramafic bodies of the UPB are sparse and limited to K - A r ages of minerals from gabbros of the Kachkanar massif (Ivanov & Kaleganov 1993) and to recently published P b - P b evaporation and SIMS U - P b zircon ages from a dunite of the Kytlym massif (Bea et al. 2001). The latter study resulted in the discovery of inherited zircons, but no firm age was proposed for emplacement of the massif. Excluding very old grains (Proterozoic and Archaean in age), the data yielded

Table 1. Conventional single-zircon U-Pb isotopic data (Kumba pegmatitic gabbro, Platinum-bearing Belt, Central Urals) Sample number

1, ab 2, ab 3, ab 4, ab 5 6

Weight (rag)

0.001 0.001 0.002 0.002 0.005 0.007

Concentrations

Corrected ratios

U (ppm)

Pb (ppm)

2~176 measured

2010 543 475 302 507 596

326 84 71 43 61 60

2384 666 1399 791 7511 12240

2~176

0.277 0.215 0.179 0.130 0.203 0.217

2~176 ( +_20-) 0.0554 • 0.0551 • 0.0551 + 0.0552 • 0.0555 • 0.0556 •

p

Z~ ( • 20-) 2 6 7 3 1 1

0.0700• 0.0687• 0.0693• 0.0675• 0.0548• 0.0458•

2o6pb/238U

2~ ( • 20-) 2 3 2 2 1 1

0.534 • 0.522 • 0.527 • 0.514 • 0.419 • 0.351 •

Apparent ages (Ma)

(• 2 6 7 3 1 1

0.76 0.48 0.57 0.60 0.94 0.96

436 __+3 428 • 4 432 +__3 421 +__2 344 __ 1 288 + 1

2ovpb/Z06pb (• 427 415 417 421 431 436

_____07 + 23 + 27 • 11 + 02 _____02

Grains were selected from non-magnetic separates at full magnetic field in a Frantz magnetic separator, ab, grains that have been air abraded following Krogh (1982). Isotopic ratios have been corrected for fractionation, blank (204:206:207:208 = 1:18.31:15.59:37.88), and initial common Pb after the growth curve of Stacey & Kramers (1975). Errors are 20- and refer to the last digits.

D. BOSCH ET AL.

446

Table 2. SIMS in situ U-Th-Pb isotopic data (Kumba pegmatitic gabbro, Platinum-bearing Belt, Central Urals) Spot number

U (ppm)

1-1 1-2 1-3 2 3 4 5 6 7 8 9 10 11 12 13 14

Th (ppm)

486 499 525 507 469 626 650 897 1518 527 259 378 718 758 264 551

Pb* (ppm)

189 194 213 294 273 349 239 469 1374 294 125 156 412 272 134 300

Th/U

29 28 30 29 25 34 36 51 86 30 14 21 40 42 15 31

2~ 2~

0.39 0.39 0.41 0.58 0.58 0.56 0.37 0.52 0.91 0.56 0.48 0.41 0.57 0.36 0.51 0.54

2~ 2~

0.0002 0.0002 0.0002 0.0002 0.0001 0.0001 0.0001 0.0001 0.0001 0.0005 0.0003 0.0003 0.0003 0.0001 0.0003 0.0002

2~ 2~

0.122 0.124 0.129 0.185 0.182 0.176 0.115 0.167 0.292 0.179 0.155 0.131 0.180 0.113 0.161 0.172

4- (lcr)

2~

-I- (1o-)

235U

0.056 0.056 0.055 0.055 0.056 0.055 0.055 0.055 0.055 0.056 0.054 0.055 0.054 0.056 0.054 0.055

0.009 0.009 0.008 0.008 0.008 0.008 0.011 0.006 0.003 0.015 0.025 0.013 0.012 0.006 0.014 0.009

0.530 0.508 0.514 0.502 0.483 0.485 0.491 0.499 0.502 0.511 0.477 0.479 0.482 0.497 0.483 0.493

2~

4- (1 or)

p

Apparent age (Ma)

238U

0.012 0.009 0.011 0.011 0.009 0.010 0.011 0.010 0.011 0.014 0.017 0.013 0.011 0.010 0.013 0.011

0.0692 0.0659 0.0675 0.0663 0.0626 0.0637 0.0642 0.0657 0.0660 0.0668 0.0636 0.0635 0.0642 0.0649 0.0644 0.0648

0.0015 0.0011 0.0014 0.0013 0.0011 0.0011 0.0012 0.0013 0.0014 0.0015 0.0016 0.0015 0.0012 0.0012 0.0015 0.0013

0.92 0.87 0.93 0.93 0.92 0.91 0.87 0.96 0.99 0.83 0.71 0.88 0.85 0.94 0.85 0.91

2~ 2o6pb

2~ 238U

4- (lcr)

433 448 423 407 449 424 428 417 418 434 387 400 391 434 384 418

431 411 421 414 392 398 401 410 412 417 397 397 401 406 403 405

9 7 8 8 7 7 7 8 8 9 10 9 7 7 9 8

*Radiogenic. Errors are given at 1or. 0.08

Kumba

(a) .. 4~r..>~j 6~"

Gabbro

(Platinum-bearing Belt)

TIMSda_a

0.07-

380 ~/,'- "

/~/-"'" 424.9_+2.7M a

0.06 -

~ 0.05-

#s

"/

(MSWD

=0.7)

300~/./'""""

J

/J

. ../~=t=9 0.04

#6 Ma

0.34

2~

I

I

I

,

0.38

0.42

0.46

0.50

0.54

0.58

0.076 Kumba

0.072 ~,,

Gabbro

/V

(Platinum-bearing Belt)

_SIMSdata

0.068

y

419+10 Ma

//(,~/ 3 v ~ ' ~'""" "

(MSWD = 0.7)

p,"

0.060

0.3

/ /

l

0.064

0.056

(b)

0~/ / ,'/' /,,/' /

37 ,

~ / /

0.4

;

-105•

Ma

=

0.5

,

207pb/235U i

0.6

Fig. 3. U - P b concordia diagram for zircon grain analyses of the Kumba gabbro. (a) TIMS single-grain analyses. Error ellipses are 2or. (b) SIMS U - T h - P b analyses. Error ellipses are lcr.

a continuum of ages from 435 4- 18 Ma to 315 + 13 Ma (lo-). On the basis of age clustering at 350-370 Ma, Bea et al. (2001) suggested emplacement and crystallization of the massif at 360 Ma, broadly coeval with the high-pressure metamorphism in the Urals, dated at 370-390 Ma (Matte et al. 1993; Glodny et al. 2002). The inferred age of the Kytlym massif, located only c. 50 km south of the Kumba massif, differs by 60-70 Ma, although some values at c. 400-430 Ma reported by Bea et al. (2001) are in good agreement with our proposed age. The observed discrepancies may be explained in two ways, which have different geodynamical implications. (1) The Kumba and Kytlym massifs are contemporaneous and the continuum of ages observed for the Kytlym dunite zircons by Bea et al. (2001) reflects varying degrees of recrystallization and associated lead loss of a zircon population that crystallized at c. 425 Ma. This is consistent with some of the SEM images provided by Bea et al., which show oscillatory zoned areas interrupted by unzoned patchy domains. Because the zircon lower intercept for the Kumba massif is nearly zero, recrystallization and Pb losses in the Kytlym massif should be related to a local phenomenon. Late gabbroic or granitic intrusions associated with locally developed high-temperature plastic deformation have been observed in the Kytlym massif (Savelieva et al. 1999) and may have driven recrystallization and associated lead loss in the 425 Ma zircons. However, this would result in an episodic disturbance of the zircon U - P b system instead of the continuum of ages observed by Bea et al. (2001). This continuum of ages, on the contrary, requires multiple resetting of the U - P b zircon ages in the Kytlym massif during a protracted time interval. The occurrence of metasomatized ultramafic rocks around the massifs (Savelieva et al. 1999) and particularly around the Kytlym massif, where the so-called kytlymite developed at the expense of basement country rocks, indicates fluid flows associated with metasomatic episodes. Leaching of radiogenic lead and/or low-temperature recrystallization or dissolution-reprecipitation processes through fluid flow under experimental (Geisler et al. 2002) or natural (H6gdahl et al. 2001) conditions can, episodically, strongly disturb the U - P b systematics of zircon. However, this requires that the zircons in the Kytlym dunite have responded nonuniformly to a common history, a possibility that could be explained by proposing that some zircons may have been armoured against Pb loss by incorporation in other grains, such as large olivine crystals.

SILURIAN AGE FOR PT-BEARING BELT, MIDDLE URALS

(2) The age difference between the two massifs is evidence that the UPB contains an association of mafic-ultramafic rocks that formed over a long time interval ranging from the Mid-Silurian to Mid-Devonian (425-360 Ma). The geochemical signatures of mafic and ultramafic rocks of the Kytlym massif show island arc-like composition, interpreted as evidence for a suprasubduction-zone setting (BeG et al. 2001). The geology of the surrounding Late Ordovician to Early Silurian volcanosedimentary and volcaniclastic units of the Tagil Zone is consistent with this view. Combining the ages of the two massifs would therefore reflect a long-lived island-arc history of about 6 0 70 Ma. Alternatively, the two contrasting ages could also indicate a more complex history in which the c. 425 Ma Kumba massif represents remnants of a Mid-Silurian oceanic lithosphere on top of which a Late Devonian-early Tournaisian arc, represented by the Kytlym massif, subsequently developed. Although further studies are needed, all the massifs in the UPB have similar characteristics and the latter hypothesis is thus considered unlikely. Existing models for the Palaeozoic tectonic development of the Uralides involve break-up of the EEC margin and initiation of the opening of the Uralian Ocean during Ordovician times, at c. 4 8 0 - 5 0 0 Ma (Zonenshain et al. 1984; Fershtater et al. 1997). The subsequent evolution of this oceanic domain is characterized by the building of several volcanic island arcs within or peripheral to the Uralian Ocean. Several independent lines of evidence concur with a Silurian age for the UPB. The schematic distribution of the Uralian magmatism in space and time given by Fershtater et al. (1997), in particular for rocks of the dunite-clinopyroxenite-gabbro series, is in very good agreement with the Wenlock age of the Kumba massif. Those researchers indicated a Late Ordovician-Early Silurian age for rocks of the suture sector, east of the MUF, including the UPB. Lastly, K - A r ages of amphiboles and phlogopites from clinopyroxenites and gabbros of the Kachkanar massif, south of the Kytlym massif, range from 420 to 430 Ma, and a mean amphibole age of 4 2 3 _ 3 Ma was proposed for the gabbroic rocks (Ivanov & Kaleganov 1993). Further north, the Chistop massif yielded a S m - N d isochron age of 419 +_ 12 Ma (Ronkin et al. 2003). All these ages are similar to the U - P b zircon age of the Kumba massif. Given the similarities between all the massifs of the UPB and the suprasubduction-zone setting of the Kytlym massif, we currently favour a model in which the UPB represents remnants of an island-arc oceanic lithosphere that existed in Mid-Silurian times, or earlier, by subduction of the Uralian oceanic domain. In Early Silurian times, anticlockwise rotation of Baltica (Torsvik et al. 1996; Torsvik & Rehnstrom 2001; Roberts 2003) may have been accommodated by subduction of the Uralian oceanic domain and development of an island arc. The Mid-Silurian UPB could have formed in such a geodynamic setting, contemporaneously with the Sakmara arc of the Southern Urals (Zonenshain et al. 1984). The 425 Ma U - P b zircon age of the Kumba massif would thus represent a minimum age for early contraction processes and for onset of the closure of the Uralian oceanic domain. According to model (1), the UPB island arc may have been short lived and would thus represent closure of a discrete marginal basin of the Uralian Ocean. Conversely, a continuous arc history of the UPB implies the occurrence of a major arc whose lifetime (from 425 to 360 Ma), broadly corresponds to the time interval between the opening of the Uralian Ocean (480-500 Ma) and the beginning of its destruction (425 Ma). The currently available dataset of ages does not exclude either of these models. Nevertheless, the occurrence of gabbroic rocks (known as tilaites) dated at 340 ___ 22 Ma (Pushkarev et al. 2003) within the Kytlym massif, with a low 87Sr/S6Sr initial ratio (<0.705), suggests that rocks of the UPB were still within an oceanic environment at that time. We thus consider model (2) more likely, and suggest that the UPB represents a long-lived island arc (and its associated marginal basin) that has been preserved from the sequential subduction-obduction cycles and

447

particularly from the Late Silurian-Early Devonian westward obduction events, during which the MUF already acted as a major thrust (Matte 1995; Echtler et al. 1997). Available geochronological information from other ophiolitic fragments that crop out along the MUF allows a direct comparison with massifs from the UPB. A S m - N d age for the Kempersai massif in the SW Urals is 397 _+ 20 Ma (Edwards & Wasserburg 1985), consistent with the 394-427 Ma age range yielded by S m - N d mineral isochrons from various lithologies in the same massif and with a 420 _+ 10 Ma Pb/Pb apparent age on one zircon fraction (Melcher et al. 1999). The Voykar massif in the northeastern Urals yielded a less precise younger, but identical within error, S m - N d age of 387 + 34 Ma (Sharma et al. 1995). Lastly, although its age is not precisely known, the Mindyak ophiolite in the Southern Urals formed before 414 +_ 4 Ma, the age of peak metamorphism affecting this massif (Scarrow et al. 1999). All these dates are identical to or within the limits of error of the ages available for the UPB and, given the large errors on most dates, no significant age discrepancies can be detected for the various oceanic fragments. In the light of the present dataset, this would suggest that the fragments of oceanic lithosphere preserved in the hanging wall of the MUF may represent a single ophiolite, tectonically dismembered during collision of the continental masses, rather than an assemblage of various oceanic and arc terranes brought together subsequently during the final stages of the evolution of the Uralide orogen. As the closure of the Uralian Ocean began as early as the Mid-Silurian, it may thus show similarities to the closure of the Iapetus Ocean in the sense that the earliest stages of contraction of the Uralian Ocean may be the counterpart of the rotation of Baltica that preceded the main Scandian continent collision with Laurentia. We thank H. Austrheim, F. BeG and J. Glodny for detailed constructive reviews that improved the manuscript. We thank the ISTEEM institution for financial support during the 1996 field trip in the Urals. F. Boudier, P. Matte, A. Nicolas, A. Pertsev and G. Savelieva are thanked for scientific interaction and fruitful discussions on the Urals geology.

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SMIRNOV, V. N. & BEA, F. 1997. Uralian magmatism: an overview. Tectonophysics, 276, 87-102.

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Isotope Geology, Salvador-Bahia, Brazil. Journal of Conference Abstracts, 762-765. RUBATTO, D. 2002. Zircon trace element geochemistry: distribution coefficients and the link between U - P b ages and metamorphism. Chemical Geology, 184, 123-138. SAVELIEVA, G. N. & NESBITT, R. W. 1996. A synthesis of the stratigraphic and tectonic setting of the Uralian ophiolites. Journal of the Geological Society, London, 153, 525-537. SAVELIEVA, G. N., SHARASKIN, A. Y., SAVELIEV, A. A., SPADEA, P. & GAGGERO, L. 1997. Ophiolites of the southern Uralides adjacent to the East European continental margin. Tectonophysics, 276, 117-137. SAVELIEVA, G. N., PERTSEV, A. N., ASTRAKHANTSEV,O. V., DEN1SOVA, E. A., BOUDIER, F., BOSCH, D. & PUCHKOVA,A. V. 1999. Structure and dynamic history of the Kytlym Pluton, North Urals. Geotectonics, 33, 119-141. SCARROW, J. H., SAVELIEVA,G. N., GLODNY, J., MONTERO, P., PERTSEV, A. N., CORTESOGNO, L. & GAGGERO, L. 1999. The Mindyak Palaeozoic lherzolite ophiolite, Southern Urals: geochemistry and geochronology. Ofioliti, 24, 239-246. SCHARER, U., KROGH, T. E. & GOWER, C. F. 1986. Age and evolution of the Grenville Province in eastern Labrador from U - P b systematics in accessory minerals. Contributions to Mineralogy and Petrology, 94, 438-451. SHARMA, M., WASSERBURG,G. J., PAPANASTASSIOU,D. A., QUICK, J. E., SHARKOV, E. V. & LAZ'KO, E. E. 1995. High 143Nd/144Nd in extremely depleted mantle rocks. Earth and Planetary Science Letters, 135, 101-114. STACEY, J. S. & KRAMERS, J. D. 1975. Approximation of terrestrial lead isotope evolution by a two-stage model. Earth and Planetary Science Letters, 6, 15-25. STEIGER, R. H. & J~GER, E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. TORSVII~, T. H. & REHNSTgOM, E. F. 2001. Cambrian palaeomagnetic data from Baltica: implications for true polar wander and Cambrian palaeogeography. Journal of the Geological Society, London, 158, 321-329. TORSVlK, T. H., SMETHURST, M., MEERT, J., ET AL. 1996. Continental break-up and collision in the Neo-proterozoic and Palaeozoic--a tale of Baltica and Laurentia. Earth-Science Reviews, 40, 229-258. WIEDENBECK, M., ALLt~, P., CORFU, F., ETAL. 1995. Three natural zircon standards for U - T h - P b , Lu-Hf, trace element and REE analyses. Geostandards Newsletter, 19, 1-23. WILLIAMS, I. S. & CLAESSON, S. 1987. Isotopic evidence for the Precambrian provenance and Caledonian metamorphism of high grade paragneiss from the Seve Nappes, Scandinavian Caledonides. Contributions to Mineralogy and Petrology, 97, 205-217. ZONENSHAIN, L. P., KORINEVSKY, V. G., KAZMIN, V. G., PECHERSKY, D. M., KHA1N, V. V. & MATEVEENKOV, V. V. 1984. Plate tectonic model of the south Urals development. Tectonophysics, 109, 95-135.

The Vendian-Early Palaeozoic sedimentary basins of the East European Craton SAULIUS S L I A U P A 1, P A V E L FOKIN 2, J U R G A LAZAUSKIENI~ 1 & R A N D E L L A. S T E P H E N S O N 3

1Department of Earth Sciences, Vilnius University, Vilnius, Lithuania (e-mail: sliaupa @get. lt) 2Department of Regional Geology and Earth History, Moscow State University, Vorobievy Gory, 119899, Moscow, Russia 3Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV, Amsterdam, Netherlands

Abstract: Vendian-Early Palaeozoic sedimentation on the East European Craton (EEC) was confined to the cratonic margins with limited intracratonic subsidence. Generally, there are two geodynamic stages involved: in stage 1, basins formed in response to continental break-up processes; in stage 2, basins formed in response to the reassembly of continental lithosphere fragments and associated continental accretionary processes. The establishment of the Peri-Tornquist passive margin was a polyphase process that commenced in the Early Vendian in the SW and ended in the Cambrian in the NW. Neoproterozoic rifting along the Scandinavian margin was essentially a long process (250 Ma), whereas the fragmentation of the continent along the earlier Timanian orogen in the east and establishment of Peri-Uralian basins took place in a rather short time span. A similar scenario is suggested for the basement of the Peri-Caspian Basin. The rift-to-drift transition is expressed differently on the various EEC margins and this could be a reflection of the relative strengths of the underlying lithosphere. The change to a convergent margin setting is recorded in Peri-Tornquist basins in the Late Ordovician, climaxing with high rates of subsidence during Late Silurian time. Subsidence rates on the cratonic margins were governed by the emplacement of orogenic loads. Where there was a short time span between Stages 1 and 2, continuing thermal subsidence from the former was superimposed onto the flexural subsidence of the latter, such as on the Dnestr margin. Other processes, such as dynamic loading related to mantle flow, are implied for the anomalously broad Stage 2 Baltic Basin. Peri-Uralian basins developed as passive continental margin basins throughout the Early Palaeozoic. Stress regime changes generated at the craton margins are reflected in the structure and subsidence patterns recorded in the intracratonic Moscow Basin.

The continental crust of the East European Craton (EEC) was formed during Archaean-Mid-Proterozoic times (see Bogdanova & Gorbatschev 2006). V e n d i a n - E a r l y Palaeozoic sediments are widely distributed within the supracrustal successions of the East European Platform and form a number of distinct basin depocentres, mostly opening towards the margins of the craton, giving the appearance, at least, of having some affinity with processes occurring on the margins of the craton (Fig. 1). The mechanisms responsible for the formation of these basins, inferred from basin fill (lithostratigraphic correlation and subsidence analysis; see Figs 2 - 4 ) and the tectonic style of intrabasinal structures, form the subject of this paper. These are related, in terms of their occurrence in time and space, to the inferred plate-tectonic processes affecting the EEC during this time. In general, this involves the break-up of Rodinia (Powell et al. 1993), with oceans developing around the EEC and passive continental margin basins forming on the peripheries of the craton followed by the closure of these oceans and the encroachment of other continental fragments and concomitant 'foreland' basin development during Early Palaeozoic (Caledonian) crustal accretion. A renewal of tectonic subsidence occurred in a number of Early Palaeozoic basins in the (post-Caledonian) Devonian. A concomitant shift of the main locus of sedimentation from the craton margins to the craton interior during the Early Palaeozoic implies a significant change in the geodynamic mechanisms involved. The Late Palaeozoic evolution of these basins is described and discussed further in a companion paper by Stephenson et al. (2006). The V e n d i a n - E a r l y Palaeozoic basins near the margins of the EEC are categorized in this paper according to the present-day geography (Fig. 1). The western (Peri-Tornquist) marginal basin system comprises the Dnestr Basin in the south (west Ukraine, Romania and Moldova) and the B a l t i c - P o d l a s i e Basin to the north (east Poland, Baltic region). The separation into distinct Podlasie and Baltic basins is the result of a Late Palaeozoic uplift event. The eastern (Uralian) marginal basin system comprises the Peri-Uralian-Pechora Basin in its central and northern segments and the Peri-Caspian Basin in its southern segment. The northern (Scandinavian) margin is treated as a single entity

as is the main intracratonic basin system, referred to as the Moscow Basin. Little can be said about V e n d i a n - E a r l y Palaeozoic basin development on the southern margin of the EEC (east of the Dnestr Basin and west of the Peri-Caspian Basin) given the degree of Late Palaeozoic and younger extensional and compressional tectonic overprinting and deformational events, including those derived from Cimmerian and Alpine orogenic processes (see Saintot et al. 2006).

Stage 1: break-up of Rodinia and the formation of passive margin and extensional peri- and intracratonic basins Peri- Uralian basins The Mezen and Central Peri-Uralian basins (Fig. 1) overlie Palaeoproterozoic and older basement inferred to have long, narrow rifted depressions of presumed Riphean age (Fig. 5). These are thought to be the earliest expression of regional extensional processes that eventually led to the breaking apart of the Rodinia supercontinent in this area (e.g. Lobkovsky et al. 1996) and formation of the Timanian Ocean in Riphean and Early Vendian times (Nikishin et al. 1996; Puchkov 2002). This (presumably small) ocean basin is thought to have closed shortly thereafter, resulting in the Timan orogenic belt during the Late Vendian and probably the earliest Cambrian (Wilner et al. 2001; Roberts & Siedlecka 2002; Gee & Pease 2004; Gee et al. 2006). The tuffaceous-siliciclastic Vendian succession exceeds 2 k m in thickness (Fig. 2) and was mainly sourced from the east (Grazhdankin 2004). A thick succession of Upper Vendian siliciclastic deposits in the M e z e n - C e n t r a l Peri-Uralian area is interpreted as a foreland basin. The width of the foreland is in the range of 3 5 0 - 5 0 0 km, and if this basin is purely flexural in origin, then an extremely strong cratonic lithosphere is implied. More probably, there remains a residual component of thermal subsidence in the underlying, peri-cratonic, extensional basins superimposed upon the eventual flexural foreland subsidence (e.g. Grazhdankin 2004).

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 449-462. 0435-4052106/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Locations of Vendian-Early Palaeozoic basins of the East European Craton (EEC) discussed in this paper, Palaeorifls are indicated by dashed lines with names in white circles: L, Ladoga; M, Mezen; P, Pachelma;PK, Pechora-Kolva; R, Roslyatino;V, Valday; Vo, Volhyn. Locations of modelled wells (Fig. 4) are also indicated (numbered black dots). Other abbreviations: CD, Central Dobrogea; LB, Lysog6ryblock; LS, Lviv slope; MP, Moldova platform comprising Dnestr marginalbasins; MM, Malopolska massif; PD, Pre-Dobrogea depression; PB, Podlasie basin; RKFH, RingkCbing-Fynhigh; SE, South Emba.

The Peri-Uralian EEC margin was uplifted during most of the Cambrian period (see Figs 2 and 3) with the subsequent Lower Palaeozoic succession lying with distinct angular unconformity on the Vendian (Bogolepova & Gee 2004). Subsidence was re-established in Late Cambrian-Early Ordovician times and expanded cratonwards during Ordovician and Silurian times. The Pechora and Barents Shelf basins (see Stephenson et al. 2006) are also thought to have been established at this time, overlying crust that had been accreted to the EEC during the Timan orogeny (e.g. Puchkov 1997; Gee & Pease 2004). In general, the Lower Palaeozoic succession can be divided into a lower part and an upper part. The former, composed of siliciclastic deposits with volcanic rocks, accumulated during Late Cambrian-Early Ordovicima times and is overlain by the Middle-Upper Ordovician and Silurian platform carbonates passing to shales in the east (Dembovsky et al. 1990; Bogolepova & Gee 2004). The Upper Cambrian-Lower Ordovician succession was deformed and eroded during the Uralian orogeny but restorations indicate an originally 0 . 3 - 6 km thick succession controlled by normal faulting (Zhemchugova et al. 2001). The thickness of the Middle-Upper Ordovician succession ranges from 100-200 m to > 1 km in the east. The Early Silurian is marked by a new transgression, which resulted in

deposition of limestones and dolomites up to 1.5 km thick, overlain by a regressive Upper Silurian-Lower Devonian succession of 2.5 km thickness. These stages are attributed respectively to the rift and drift stages. Based on petrographic studies in the Southern Urals, oceanic crust was formed by the end of Arenig (MasIov et al. 1997), with the opening of the ocean propagating to the north throughout the late Early and Mid-Ordovician (Savelieva et al. 1997). The rift-to-drift transition is marked by the lower Middle Ordovician unconformity in the Pechora Basin (Fig. 3). On the other hand, this also correlates with a Pre-Scandian unconformity affecting the Baltic and Moscow basins. It is notable that the Barentsia and Pechora basins overlying the Timanide basement are much wider ( > 5 0 0 km) than the craton marginal basins in the south. This may be a response of the younger, weaker, lithosphere to tectonic extension. Tectonic subsidence curves for these two basins are shown in Figure 4 (wells 6-9). Rapid tectonic subsidence during the Early Silurian (20 m Ma-1) was followed by deceleration during the Late Silurian (5 m Ma-1), describing an apparently concave shape, suggesting an episode of continental separation along the earlier Timanian orogenic belt. Stratigraphic data from the Southern Urals suggest some localized rifting even as early as the Early Cambrian (Maslov et al. 1997).

EARLY PALAEOZOICBASINS OF THE EEC

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Fig. 2. EEC sedimentary depocentres (isopach thicknesses shown in metres) and lithofacies for the Late Vendian, earliest Cambrian, Cambrian, Ordovician, and Late Ordovician (Ashgill)- Silurian.

Two subsidence events occurring in the latest CambrianOrdovician and Silurian, recognized in the Pechora Basin, may be interpreted to reflect two distinct extensional pulses (IsmailZadeh et al. 1997). Alternatively, the phase of rapid Silurian subsidence can be attributed to subduction loading as intra-oceanic subduction is thought to have started during the latter part of Ordovician or Early Silurian time (Puchkov 1997).

Peri-Caspian Basin

Most of the Neoproterozoic(?) and Lower Palaeozoic succession of the Peri-Caspian Basin is buried by up to 12 km of younger sediments (Volozh et al. 2003; see Stephenson et al. 2006). A few hundred metres of Cambrian-Silurian strata have been drilled on the northern margin of the basin (in the Orenburg area). Lower-Middle Ordovician siltstones, shales and carbonates

(the last abundant in the upper part of the succession) are overlain by Middle Ordovician-Lower Silurian carbonates that pass into clastic sediments of Late Silurian-earliest Devonian age. According to published models based on seismo-stratigraphy and seismic velocities, an up to 4 km thick Riphean succession of terrigenous and carbonaceous rocks is inferred to overlie basement of Early Precambrian age. Vendian-Ordovician siliciclastic rocks are interpreted to be 2 km thick, as are the Upper Ordovician-Silurian carbonates (Lobkovsky et al. 1996; Volozh et al. 2003). However, the presence of a Riphean basin, in view of the indirect nature of available data, cannot be considered as very certain (see Stovba et al. 1996; Stephenson et al. 2006). Given the lack of any even poorly constrained stratigraphy, little can be said on the basis of tectonic subsidence curves (see Brunet et al. 1999). Lobkovsky et al.'s (1996) synthetic Mid-CambrianSilurian subsidence curve has a concave shape, with rapid

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Fig. 3. Lithostratigraphicalcorrelation chart of the EEC Early Palaeozoic basins. Legend as in Figure 2.

subsidence during mid-Cambrian to mid-Ordovician times followed by a reduction in subsidence rate during Late Ordovician-Silurian times (Fig. 4). Volozh et al. (2003), in contrast, published a convex-shaped subsidence curve indicating an increase in subsidence rate during the Late Ordovician and Silurian. Making any reconstruction even more difficult is that there are only scarce data available on the Neoproterozoic-Early Palaeozoic evolution of the southern margin of the EEC in the area between the Peri-Caspian depression and Dnestr slope (see Saintot et al. 2006). Middle Cambrian and Silurian sediments reported from the Kislovodsk area in the Caucasus (Markovski 1958) tend to imply a platformal setting for sedimentation on the southern margin of EEC at this time, a view also adopted by Saintot et al. (2006). Accordingly, given the general lack of solid evidence otherwise, the tectonic evolution of the Peri-Caspian Basin is often considered to parallel that of the Peri-Uralian margin of the EEC, that is, a passive margin followed by a Timanian orogen foredeep during Late Riphean-Early Vendian times, with the remnants of the orogenic belt presumably now underlying the North Caspian-Aktyubinsk uplift in the south of the basin (Volozh et al. 2003). In such a model, the subsequent break-up of Rodinia resulted in the formation of an east-west elongated Peri-Caspian depression during Late Cambrian-Ordovician times. Extensional faults cutting Early Palaeozoic and older rocks are thought to have been inverted prior to the Mid-Devonian (e.g. Ural-Orenburg 'zone of elevations'; Nevolin 1993). In the east, the basin was linked to the Uralian rifted system, whereas the southern flank remained uplifted as the North CaspianAktyubinsk block that separated the basin from the South Emba intercontinental rift to the south. As such, the main basin

depocentre was established above the EEC-Timanian suture. The depocentre widened to the south during Late OrdovicianSilurian times, explained either as a thermal sag or as the result of back-arc extension behind a southern ocean (e.g. Volozh et al. 2003), these two scenarios corresponding to the two alternative subsidence patterns mentioned above. In any case, rapid, active rift subsidence was re-established by Mid- to Late Devonian times (Brunet et al. 1999; see Stephenson et al. 2006). Peri- Tornquist basins

The Peri-Tornquist basins were established during Early to Late Vendian times along some 2000 km of the western margin of the EEC (Fig. 1). They are about 200-300 km wide, except for the Baltic Basin, which is 650 km wide. Lithofacies trends for the Early Palaeozoic are similar all along the craton margin (Figs 2 and 3). In general, two segments can be distinguished, the Baltic-Podlasie depocentre and Dnestr slope (Fig. 1). Lower Vendian clastic and volcanic rocks up to 500 m thick are reported from the Dnestr slope. Igneous activity started at 6 0 0 - 5 8 0 Ma and terminated at about 551 Ma (Compston et al. 1995). The Lower Vendian units are overlain by Upper Vendian arkosic sandstones and mudstones of 200-300 m thick. In the southernmost part, Upper Vendian Histra turbidites up to 5 km thick were deposited in Central and South Dobrogea (Oaie 1998); they grade into the 2 km thick Avdarma Series in the Pre-Dobrogean depression and to 350 m thick mudstones and sandstones of the Moldova platform. The Podlasie and Baltic regions remained uplifted until Late Vendian arkosic sedimentation started in narrow basins located along the future Baltica continental margin (Aren 1988).

EARLY PALAEOZOIC BASINS OF THE EEC

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Baltic Basin was uplifted. Subsidence curves for the Late Vendian-Cambrian Baltic Basin are generally concave, whereas subsidence was accelerating on the Lviv slope at this time (Fig. 4). Clastic sedimentation was replaced by shale-carbonate deposition in the Peri-Tornquist basins during the Ordovician (Laskovas 2000), with thicknesses of 5 0 - 2 5 0 m (Fig. 2) reflecting a general deceleration of subsidence rates. Carbonates were deposited on the eastern peripheries of basins and shales predominate in the west, the sedimentary architecture pointing to clastic starvation of the marine basin. The Moldova platform remained uplifted throughout the most of the Ordovician. Thus, continental break-up associated with volcanic activity took place initially in the southernmost part of the PeriTornquist zone during the Early Vendian and propagated to the NW in the Late Vendian. The Baltic-Podlasie platform shows general deceleration of subsidence during latest Vendian-Cambrian to Mid-Ordovician times (Fig. 4) so that by the Mid-Ordovician a common tectonic regime had been established all along the Peri-Tornquist margin of the EEC, comprising the passive continental margin of Baltica with the Tornquist Sea to the west (Nikishin et al. 1996; Sliaupa et al. 1997; Poprawa et al. 1999), as seen in Figure 2. The Lysogrry, Malopolska (Narkiewicz 2002) and Lezhaysk blocks of southern Poland and western Ukraine have similar subsidence trends and lithological patterns to those of the craton margin. The proximity of the Lysogrry and Malopolska massifs to Baltica is also suggested by palaeontological data (Cocks & Torsvik 2005).

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Fig. 4. Early Palaeozoic tectonic subsidencecurves for basins of the EEC. The tectonic subsidencecurve of the Peri-CaspianBasin is adopted after Lobkovsky et al. (1996) and that of the Scandinavianmargin after Greiling et al. (1999). The high subsidencerate stage of the western margin of Baltica is highlightedin grey. Well locations are shown in Figure 1.

Sedimentation occurred on the entire western margin of the EEC during the Cambrian, beginning with the Blue Clay succession, which attains a thickness of 400 m on the Lviv slope (Garetsky 1981, 1990; Zinovenko 1986). The Blue Clays are overlain by Cambrian quartz sandstones, siltstones and shales increasing in thickness westwards. No deposition is recorded on the Moldova platform at this time (Fig. 2). The thickness of the Cambrian succession exceeds 1 km on the west of the Lviv slope, apparently comprising a deep-water flysch succession, attributed by Vishniakov et al. (1982, 1997) to an early 'Caledonian' tectonic phase. Dipping Cambrian strata are reported to be unconformably overlain by flat-lying Ordovician-Silurian shales, a geometry that also could imply fault-block rotation in an extensional environment. The Baltic-Podlasie basin has an up to 500 m thick Cambrian succession; sandstones predominate in the east and shales in the west (Jankauskas & Lendzion 1994). A drastic rearrangement of the sedimentation pattern occurred during mid-Mid-Late Cambrian times when the central part of the

Most of the Scandinavian platform basin along the EEC western margin was deformed and eroded during the Caledonian orogeny or eroded during succeeding stages of Fennoscandian shield uplift. The restored platform facies extends for 2 0 0 270 km (up to 400 km) from the foreland into the hinterland of the Caledonian orogen (Roberts & Gee 1985). Protracted Riphean-Early Vendian sedimentation and volcanic activity were confined to rifted basins (Kumpulainen & Nystuen 1985; Kumpulainen 1986; Svenningsen 1994, 1995; Bingen et al. 1998; Greiling et al. 1999). This was followed by platform sedimentation of Late Vendian arkosic clastic deposits, transgressively overlain by the Cambrian sequence (Bergstrom 1980; Bergstrom & Gee 1985). The lowermost Cambrian units comprise shallowwater quartzose siltstones and sandstones up to 150 m thick that pass into an upper Lower Cambrian siltstone-sandstone succession, a few tens of metres thick, and then to the deeper-water Middle-Upper Cambrian Alum shales of even less thickness, indicating a progressing sediment starvation of the basin. The Ordovician succession, of several hundred metres, consists of carbonates passing westwards into shales and greywackes that are more than 1 km thick (Gee et al. 1974; Gee 1975). Shallow-marine limestone environments expanded cratonwards during the Late Ordovician and Early Silurian (Gee & Wolff 1981). Continental break-up along the Scandinavian margin was a long-lived process (from 850-800 Ma to 610-600 Ma), leading to the accumulation of a very thick synrift sedimentary package (Soper 1994). The drift stage (after opening of the Iapetus Ocean) was marked by a cratonward expansion of sedimentation from the Late Vendian-Cambrian to the Ordovician. The subsidence curves (Fig. 4) have concave shapes typical for passive continental margins. Several events of increased subsidence are documented in the earliest Cambrian, late Early-early MidCambrian (Greiling et al. 1999), and Mid-Ordovician. The first event is probably related to an increase in tectonic extension. The latter two events took place in a background of a narrowing Iapetus Ocean. Accordingly, the Mid-Cambrian subsidence event (518-505 Ma) has been attributed to subduction loading

454

S. SLIAUPAETAL.

Fig. 5. Distributionof Middle and Upper Riphean and Lower Vendian (including volcanic rocks) sedimentarysuccessions on the EEC. Neoproterozoic rifted and oceanic basins along the EEC margins are also indicated.

(Gee 1987), as indicated by 505-480 Ma eclogites (e.g. Cocks & Torsvik 2002). The Mid-Ordovician pulse of increased subsidence may evince a Pre-Scandian orogenic event.

subsidence occurred in the Roslyatino palaeorift. In general, the tectonic subsidence curves show a deceleration during the Late Vendian-Early Cambrian and the Late Cambrian-Ordovician (Fig. 4).

M o s c o w Basin and other intracratonic basins

The epeirogenic Moscow Basin was established in the Late Vendian and earliest Cambrian by deposition of 900 m (1200 m in the east) and 350 m thick mudstones and clays (Kirsanov 1968; Nalivkin & Jakobson 1985), as shown in Figure 3. After an Early to early Mid-Cambrian break in sedimentation that is associated with some structural inversion, 200-250 m of late M i d - L a t e Cambrian sandstones and 3 5 0 m of Ordovician carbonates were deposited. During the Vendian the Moscow Basin was linked to the Peri-Timan Basin, whereas it was connected to the Baltic Basin during the Ordovician, periodically forming part of a seaway linking the Baltic Basin in the west to the Peri-Uralian Basin in the east (Sablukov 1984; Kheraskova et al. 2005). Some 5 0 - 7 0 m of erosion took place before the Late Ordovician in the Roslyatino palaeorift (Fig. 1). Early Silurian (Llandovery) dolomites, 8 0 - 2 9 0 m thick, occur only in the central part of the Moscow Basin. The Silurian succession was partially eroded prior to the Devonian, most significantly (200-300 m) on the inverted northern flank of the Roslyatino palaeorift. The Moscow Basin superimposes the Volhyn-Orsha and Central Russian palaeorift systems that mark the juxtaposition of the Fennoscandia and Volgo-Uralian segments of the EEC (see Bogdanova et al. 2006). There is a general correlation between maximum thickness of the Upper Vendian-Lower Palaeozoic deposits and the underlying older rift basins (Fig. 6). Maximum

Stage 2: the transition from passive to convergent margin basin settings Peri- Tornquist basins

The Peri-Tornquist basins show drastic changes reflected in a systematic increase of the subsidence rate since the Late Ordovician in the NW and the late Llandovery (Early Silurian) in the SW; subsidence climaxed during the Pfidolf (latest Silurian) along the entire Peri-Tornquist margin of the EEC (Fig. 4). The Ordovician-Early Silurian starved basin gave way to compensated sedimentation during the late Silurian. Accumulation of carbonates was confined to the shallow eastern periphery of the basins whereas graptolitic shales accumulated in the west. The thickness of the Llandovery succession is in the range of 1 0 - 4 0 m , the Wenlock sequence is 5 0 - 2 5 0 m thick, the Ludlow reaches 1 km in thickness, and the Pf'fdolf is up to 1.5 km thick, but originally was thicker (Vejbaek et al. 1994). The > 1 km thick Tiver Group mudstones, attributed to the Late Silurian or earliest Devonian, are distributed to the SW of the carbonate platform in the Pre-Dobrogean depression. The latest Ordovician-Silurian succession of the westernmost part of the Baltic Basin contains sandy interbeds (Jaworowski 2000); they are reported also from the western Silurian facies of the Dnestr slope (Cegelniuk 1989) implying that sediment was derived from land masses situated in the west. Palaeontological, isotope

EARLY PALAEOZOIC BASINS OF THE EEC

455

(Vishniakov et al. 1997). Compressional deformations are mapped in the Lysogbry and Malopolska blocks (Pozaryski 1990). The Baltic Basin was subjected to widespread reverse faulting during latest Silurian-earliest Devonian times (Fig. 7), in a stress regime related to the Scandinavian Caledonides (Sliaupa 1999). The Late Ordovician-Silurian subsidence of the PodlasieBaltic Basin was governed by the flexural bending of the EEC margin induced by docking of the Eastern Avalonia microcontinent (Sliaupa et al. 1997; Poprawa et al. 1999; Lazauskiene et al. 2000), whereas the mechanism responsible for the high-rate subsidence of the Dnestr slope and amalgamated Malopolska and Lysog6ry blocks is not clear. The Lviv and Moldova slopes display basement-controlled subsidence patterns. The Early Precambrian basement of the EEC dips at a low angle (0-500 m within a zone of 150 km width), whereas flexural bending increases sharply over the adjacent Cadomian crust (e.g. Rava-Russian zone) where the thickness of the Upper Silurian succession increases from 500 m to > 1000 m within a few tens of kilometres (Bondartshuk 1974) (Fig. 8). The accreted Cadomian crust (adjacent to the older Precambrian craton) remained uplifted during most of the Silurian and then was severely flexed down at the end of Silurian (beginning of Devonian?) whereas the cratonic crust in the east displayed accelerating subsidence throughout the Silurian. This suggests that the foreland basin architecture is at least in part controlled by structural decoupling or basement rheology, perhaps with the uplift explicable in terms of a narrow forebulge developed on the weaker Cadomian crust. The eastern periphery of the forebulge is well recorded by lithofacies indicating shallowing in the west of the Moldova carbonate platform (Fig. 8). Later, as convergence progressed, the Cadomian basement would have become involved in a more rapid subsidence as the forebulge was transferred to the eastern periphery of the Moldova platform.

Scandinavian basin

Fig. 6. (a) Earliest Cambrian, (b) Ordovician, and (e) Ordovician-Silurian isopachs for the Baltic and Moscow basins (not restored). Underlying Neoproterozoic basins are indicated in grey.

and geochemical studies indicate sources of Gondwanan affinity in the west (Vecoli & Samuelsson 2001; Sliaupa 2002). The Ordovician-Silurian subsidence trends of the Malopolska and Lysog6ry blocks are similar to those of the Peri-Tornquist basins (Poprawa et al. 1997; Narkiewicz 2002), displaying subdued sedimentation (100-230 m) of carbonates and shales during the Ordovician (with the specific Tremadoc sandy deposition analogous to the marginal basins) and increased subsidence rates in the Early Silurian (shales of a few hundred metres thickness) culminating with more than 2 km of Late Silurian shales and greywackes, the latter increasing in abundance to the west. The establishment of the convergence regime is indicated by contractional structures identified in the westernmost part of the Baltic Basin (Beier et al. 2000). Low-grade metamorphism is dated at 440 Ma in the Ringkcbing-Fyn High (Frost et al. 1981). Folding of Early Devonian platform sediments overlain by undeformed deposits of Mid-Late Devonian age is reported from the westernmost Lviv slope (Rava-Russian area)

The Scandinavian passive margin was open to the gradually narrowing Iapetus Ocean throughout the Ordovician. Convergent processes modified the generally decelerating subsidence rate of the craton margin. A distinct increase in subsidence rate, as mentioned above, that correlates with an enhanced influx of sandy lithologies reflecting the Pre-Scandian orogenic event (Dunlap & Fossen 1998) is recorded during the Mid-Ordovician (Fig. 4). The main Scandian orogeny took place during the Silurian and the earliest Devonian (e.g. Cocks et al. 1997; Cocks & Torsvik 2002). As a result of erosion, the Caledonian front is exposed considerably to the west of its original position and the foreland basin was apparently obliterated (Gee 1975). Its absence has been explained by invoking a 'flat toe' geometry for the orogenic wedge (Garfunkel & Greiling 1998), although other available evidence seems to favour the previous existence of a deep foreland basin. The high thermal maturity of sediments preserved to the east of the deformation front in combination with fission-track data imply 2.5-6 km depth of foredeep sediments (Middleton et al. 1996; Larson et al. 1999). Another argument supporting the presence of a flexural depression in front of the Caledonides is the apparent forebulge that separated the Scandinavian and Baltic basins during Silurian-earliest Devonian times (Fig. 2). This forebulge would have migrated to the SE as suggested by mapped lithofacies variations (Baarli 1990; Sliaupa 2003). Preserved platform sediments (e.g. in the Oslo area) show that the subsidence of the craton margin significantly increased from the Early Silurian time (latest Llandovery-Wenlock) and was associated with deposition of Old Red Sandstone facies, indicative of a foreland molasse basin.

456

S. SLIAUPAETAL.

Fig. 7. Seismic profile showing Late Vendian-Cambrian extensional structures and latest Silurian-earliest Devonian transpressionalfaults in the central Baltic Basin, western Lithuania (location of section is shown in Fig. 10). PR1, crystallinebasement; Cm, Cambrian;O, Ordovician;S, Silurian;D, Devonian;T1, Lower Triassic.

Peri-Uralian and Peri-Caspian basins

The eastern margin of the craton, comprising the Peri-Uralian and Peri-Caspian basins, was open to the adjacent ocean during the Ordovician (see Saintot et al. 2006). The geometry of the Peri-Caspian Basin changed during the Silurian-Early Devonian. Its southern flank, formed by the North Caspian-Aktyubinsk uplift, was flexed down while the basin became less marked in the north. It was suggested by Volozh et al. (2003) that the passive margin was transformed into a convergent one at the end of Silurian, based on inversional tectonics reported for earliest Devonian sediments of the Peri-Caspian Basin and replacement of Ordovician-Early Silurian carbonate sedimentation by terrigenous deposits of Late Silurian-Early Devonian age (Fig. 3). Accordingly, the subsidence of the southern flank of the basin might have been related to flexural bending. This mechanism alone, however, is not sufficient to explain subsidence across the 600 km wide foreland, which probably was driven by thermal relaxation inherited from the Early Palaeozoic. The onset of intra-oceanic subduction in the Uralian Ocean, as mentioned above, was suggested by Puchkov (1997) to have occurred during latest Ordovician to Silurian times but the available data neither strongly support nor conflict with this view. In the north, the Pechora-Kolva palaeorift (Fig. 1) was subject to tectonic compression and structural inversion that involved Lochkovian (earliest Devonian) and older sediments. Erosion is estimated to be as great as 800-1200 m. The thickness distribution and dextral sense of strike-slip movements suggest a transpressional regime related to N W - S E compression, implying the possibility of a Caledonian source for the far-field stresses.

Moscow Basin

Only Early Silurian sediments are preserved in the Moscow Basin, indicating a cessation of subsidence, which evolved later to uplift.

This seems to correlate directly with the initiation and progression of collision processes along the Scandinavian and Peri-Tornquist margins.

Further discussion Baltic Basin: anomalously wide

The break-up of the Rodinian supercontinent and establishment of the Baltican passive continental margin basins took place during Late Vendian-Ordovician times but occurred at different times and with different styles on the various margins of the EEC. The Scandinavian passive margin formed after a long Neoproterozoic rifting stage that apparently lasted as long as 2 0 0 - 2 5 0 M a . The Peri-Tornquist margin showed propagation of break-up from the SE to the NW, begun in the Early Vendian and completed in the earliest Ordovician (c. 150 Ma in duration). The Early Palaeozoic rifting stage on the eastern (Peri-Uralian) margin was considerably shorter, lasting from the (Mid-) Late Cambrian to Mid-Ordovician, and sedimentation was confined to fault-controlled basins. Adjacent cratonic areas, perhaps characterized by weaker lithosphere than elsewhere, were also sometimes involved in subsidence that was accompanied by faulting. A particularly long-wavelength (650 km) example of this occurs in the Baltic Basin (Fig. 2) segment of the Tornquist continental margin basin system that was established during the Late Vendian-Cambrian. The Baltic Basin comprises two distinct components. The westernmost part consisted of a 160 km wide zone, flexed down by the thermomechanical and sediment loads of the adjacent rifted margin (like the remainder of the Tornquist margin), where the Cambrian succession is as thick as 250600 m. The 500 km wide pericratonic part of the Baltic Basin, however, is anomalous with respect to the remainder of the Tornquist margin. Its development was explained in terms of a weaker

EARLY PALAEOZOIC BASINS OF THE EEC

457

Fig. 8, Thickness and lithofacies distribution of sediments in the Early (left) and Late (right) Silurian for the Lviv slope and Moldova platform. The cratonic crust shows moderate subsidence whereas the Cadomian domain indicates a marked increase in flexural bending. than normal lithosphere that subsided for a few hundred metres under moderate extension but warped up under in-plane compression (Sliaupa & Ershov 2000; Sliaupa 2002). Numerous Vendian-Cambrian low-amplitude normal faults in the region (Fig. 7) provide ample proof of the extensional regime responsible for the Baltic Basin during what is referred to here as its Stage 1 development (Brangulis & Kanevs 2002; Sliaupa 2003). Lithofacies patterns indicate the absence of a western sedimentary provenance as early as the late Early Cambrian and a rift-drift (break-up) unconformity of mid-Late Cambrian age is seen in its central part. Seismic data from the Rtigen area, however, indicate that extensional faulting persisted until the earliest Ordovician (Shlfiter et al. 1997). Continued subsidence during the Ordovician was generally dominated by a system of west-east- and WSW-ENE-trending sub-basins (Fig. 2) that are aligned with Precambrian structural trends (e.g. Liepaja-Saldus zone, Middle Lithuanian Suture), suggesting a role for basement weaknesses in controlling Stage 1 basin development. During its Stage 2 (convergence) development in the Silurian, the Baltic Basin, as in Stage 1 (extension), also consisted of two prominent sub-basins (Fig. 9). Orogenic loading (e.g. the overriding Eastern Avalonian terranes) can explain the subsidence of the 180 km wide high-gradient westernmost part of the basin (Fig. 2) but other processes have to be invoked to account for the subsidence of the rest of basin. One possible explanation is dynamic loading induced by viscous mantle corner flow above an obliquely subducting slab (Lazauskiene et al. 2000). This idea was also developed by Daradich et al. (2002). A compressional tectonic regime is clearly indicated by the establishment of a reverse fault system in the basin at this time (Fig. 10). There is also the suggestion of forebulge development between the Scandinavian foredeep and the Baltic Basin, with its cratonward migration as the Scandinavian orogen advanced evident in mapped facies trends. The collapse of the forebulge is marked by a marked influx of the sandy clastic deposits into the Baltic Basin during the late Early Devonian (e.g. Sliaupa 2003).

D n e s t r area." a n o m a l o u s c o n v e r g e n c e signature

A divergent aspect of the two-stage approach taken for the Early Palaeozoic basins of the EEC, representing a simple continental break-up-reassembly (Stage 1-Stage 2) cycle, is displayed by the Peri-Tornquist basins. They show different subsidence trends during Late Vendian-Cambrian times although similar lithologies are deposited. Subsidence rates suggest that the Dnestr slope was involved in convergent processes whereas rifting was active in the Podlasie-Baltic area. A thick succession of Late Vendian-Cambrian turbidites accumulated on the Dnestr slope and a northwestward propagation trend is discernible. Convergence may have taken place in the Pre-Dobrogean area during the Late Vendian; it ceased in the earliest Cambrian, as evidenced by deposition of 90 m thick shales of earliest Cambrian age followed by Cambrian-Ordovician non-deposition in the Moldova platform. After some delay, flexural bending of the Lviv slope took place in Cambrian-earliest Ordovician times, preceded by Vendian rift and drift stages. An orogenic event is corroborated by zircon ages from Malopolska and Upper Silesia of 6 1 1 - 5 2 7 M a (Zelazniewicz et al. 2001), suggesting the docking of Cadomian terranes to the EEC, such as the Lezhajsk (west Lviv area) and Malopolska (Poland) massifs, at this time. Non-penetrative deformation of pre-Ordovician sediments is observed in the Malopolska massif (Pozaryski 1977, 1990; Winchester et al. 2002) and on the westernmost Lviv slope (Vishniakov et al. 1997). The proximity of Upper Silesia and the Malopolska massif to Baltica at this time is also suggested by palaeontological data (Orlowski 1975; Fortey & Cocks 2003).

Intracratonic basins

Tectonic changes along the craton margin affected intracratonic areas. The largest of the intracratonic epeirogenic basins is the Moscow Basin, established in the Late Vendian above a Neoproterozoic rift system (Lobkovsky et al. 1996; Nikishin et al. 1996; Kostyuchenko et al. 1999) contemporaneously with

458

S. SLIAUPA ETAL.

Fig. 9. Isopachs for Early (a) and Late (b) Silurian (restored, in metres) mad lithofacies of the Baltic Basin. the initiation of continental break-up along the western margins of the craton. Post-rift widening of the basin is a typical feature, generally explained by processes such as thermal contraction and increasing lithosphere strength (e.g. England 1983; Sawyer 1985; Ziegler 1994; Bertotti et al. 1997). Mantle fertilization was recently suggested as a factor in the development of the Moscow Basin (Artemieva 2003). However, none of these models are consistent with the subsidence history of the Moscow Basin, with its main subsidence phase being established long after synrift activity ceased in the underlying rift zones. In contrast, this occurred immediately after the termination of synrift processes in other Early Vendian rifts of the EEC. After

cessation of igneous activity the lithosphere of rifts became too strong to be deformed, leading to redistribution of the extension-induced strain, as it is documented, for example, in the East African rift system (Nyblade & Brazier 2002). The coring rifts are confined to the junction zone between the Early Proterozoic Fennoscandian block and older Sarmatian and Volga-Uralian blocks (Bogdanova et al. 1996). Most of the Late Vendian-earliest Silurian Moscow Basin is situated within the younger Fennsocandian segment. The reactivation of the pre-weakened zones had only a partial effect reflected in maximum subsidence rates confined to those zones. The major factor seems to be the heterogeneity of the lithosphere strength, as the underlying rifts mark suture zones between crustal segments having very different characteristics that may lead to differentiated vertical movements when subjected to horizontal tectonic forces (Artyushkov et al. 2000; Sliaupa & Ershov 2000).

Summary and conclusions: lithosphere plate boundary processes and sedimentary basin formation on the EEC in the Vendian-Early Palaeozoic

Fig. 10. Late Caledonian faults and isopachs (metres) for Early Devonian (Lochkovian) time the Baltic Basin. P shows the location of seismic profile shown in Figure 7. Arrows indicate the presumed source and direction of horizontal compression during Late Silurian-Early Devonian times.

Vendian-Early Palaeozoic sedimentation on the EEC was confined to the cratonic margins with limited intracratonic subsidence. Generally, there are two geodynamic stages involved: in stage 1, basins on cratonic lithosphere formed in response to continental break-up processes (in this case Rodinia); in stage 2, basins on cratonic lithosphere formed in response to the reassembly of continental lithosphere fragments and associated continental accretionary processes (in this case, the early assembly of Pangaea). A general summary as applied to the Vendian-Early Palaeozoic sedimentary record of the EEC is shown in Figure 11. Neoproterozoic sedimentation was restricted to long narrow rift zones that perhaps related to the ancient sutures of major cratonic crustal segments. Some of these progressed into oceanic basins. The establishment of the Peri-Tornquist passive margin was a polyphase process. Break-up of the lithosphere took place first in the SE during the Early Vendian (Dobrogea) and mid-Vendian

EARLY PALAEOZOICBASINS OF THE EEC (Lviv margin). A second phase is recorded in the NW with rifting taking place along the Baltic-Podlasie margin during latest Vendian-mid-Mid-Cambrian times. This eventually led to continental break-up, recorded by a marked rearrangement of sedimentation patterns in the Baltic Basin in the mid-Cambrian. Even as continental separation was progressing to the NW, some convergent processes evidently took place along the Dnestr margin in the Late Vendian (Pre-Dobrogea) and Cambrian-earliest Ordovician (Lviv slope), possibly related to an early docking phase of Cadomian blocks onto the thinned margin of the EEC. The 2000 km long passive margin was open to the Tornquist Sea throughout the Ordovician. Compensated Cambrian sedimentation gave way to clastic starvation during the Ordovician. The Scandinavian passive margin was established after what was apparently a very long (250 Ma) period of rifting. The drift stage is marked by the onlapping geometry of a Late VendianCambrian terrigenous platform succession that shows sediment

459

starvation during the late Cambrian. Increased subsidence rates during the earliest Cambrian are related to extension of the craton margin whereas mid-Cambrian and mid-Ordovician accelerations in subsidence reflect the establishment of a convergent, Pre-Scandian tectonic regime (Stage 2). This is also recorded in the Baltic and Moscow basins by basin regression and structural inversion. The progressive narrowing of the Iapetus Ocean led to a change from an uncompensated to a compensated sedimentation regime on the craton margin during the Ordovician. Continental break-up in the east took place along the earlier Timanian orogenic belt during Late Cambrian-Early Ordovician times, later than it did in the west. A similar scenario is suggested for the basement of the Peri-Caspian Basin. The rifting stage, compared with the western margins, was much shorter. The drift stage is recorded by platformal sediments of Mid-OrdovicianSilurian age overlying Late Cambrian-Early Ordovician basins that are inferred to be fault-controlled. Saintot et al. (2006) have

Fig. 11. Tectonic setting and major driving mechanismsof Late Vendian-Early Palaeozoic sedimentarybasins of the EEC for differenttime slices (a) Late Vendianearliest Cambrian; (b) earliestmid-Cambrian;(c) Late Cambrian;(d) Ordovician; (e) Late Ordovician-Silurian.

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suggested that there may not have been actual continental break-up on the eastern margin of the EEC at this time but rather that these basins developed on collapsed, subsided crust of the late Neoproterozoic orogens (Timanian or Baikalian). The rift-to-drift transition is expressed differently on the various EEC margins and this could be a reflection of the relative strengths of the underlying lithosphere. The drift-stage sediments of the Scandinavian and Peri-Uralian margins show onlap geometries with respect to rift-stage basins, explicable in terms of an increasing lithosphere rigidity through time. The Baltic Basin, in contrast, displays a clear break-up unconformity and basin narrowing that could be the response of a weaker lithosphere to establishment of the convergent (Stage 2) tectonic regime. The first indication of plate convergence processes along the Peri-Tornquist margin is the occurrence of sandy lithologies associated with increased subsidence in the western part of the Baltic Basin during the Late Ordovician. The whole Peri-Tornquist margin displays accelerated subsidence during the Silurian. The approach of western land masses is reflected by the gradual change from starved basin sedimentation to a compensated deposition regime associated with the occurrence of sandy clastic lithologies in the westernmost parts of these basins. In the NW the EEC margin was flexed down by the Eastern Avalonian terrane(s), whereas the flexural load in the SE remains unclear. A possible explanation is a shift of various Cadomian terranes, previously amalgamated to the EEC margin, as Eastern Avalonia collided with Laurentia, as discussed by Winchester et al. (2006). Such a model could also explain the coincidence of these processes with the Scandian orogeny along the Scandinavian margin. In general, subsidence rates on the cratonic margins during Stage 2 development were governed by the orogenic loads. Where there was a short time span between Stages 1 and 2, continuing thermal subsidence from the former was superimposed on the flexural subsidence of the latter, such as on the Dnestr margin. Other processes, such as dynamic loading related to mantle flow, need to be invoked for the anomalously broad Stage 2 Baltic Basin. The Peri-Uralian basins developed as passive continental margin basins throughout the Early Palaeozoic. However, some changes in the subsidence trends may be considered as evidence of the establishment of a plate convergence regime in the Uralian Ocean during the Late Silurian-Early Devonian. Similarly, subduction-related subsidence superimposed upon subsidence driven by other mechanisms, such as thermal relaxation, cannot be ruled out in the Peri-Caspian Basin during the Silurian-Early Devonian. Stress regime changes generated at the craton margins are reflected in the structure and subsidence patterns recorded in the Moscow Basin. These include mid-Cambrian uplift following earliest Cambrian extension along the Scandinavian margin and re-establishment of the subsidence regime in the Late Cambrian contemporaneously with the initiation of the basin development along the Peri-Uralian margin. The Caradocian inversion event recognized in the Moscow Basin correlates with a Pre-Scandian orogenic event. It implies far-field stress transmission into the craton interiors from its margins. The Scandian orogeny, at its climax along the Scandinavian margin, induced a dense system of reverse faults in the Baltic Basin, at a distance of >1000 km. 'Hard' coupling between Laurentia and Baltica is implied, in contrast to the 'soft' docking of Eastern Avalonia to the Baltica. The former is also indicated by inversional tectonic features recognized in the lowermost Devonian succession of the Pechora Basin. This study was performed within the framework of the European Science Foundation EUROPROBE programme. It is a part of the project K-104 financed by the LithuanianScience Foundationand AB Geonafta (S.S.).

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Late Palaeozoic intra- and pericratonic basins on the East European Craton and its margins R. A. S T E P H E N S O N 1, T. Y E G O R O V A 2, M.-F. B R U N E T 3, S. S T O V B A 4, M. W I L S O N 5, V. S T A R O S T E N K O 2, A. S A I N T O T 1'6 & N. K U S Z N I R 7

1Netherlands Centre for Integrated Solid Earth Sciences, Faculty of Life and Earth Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands (e-mail: randell, stephenson @f alw. vu. nl ) 2Institute of Geophysics, National Academy of Sciences of Ukraine, Kyiv, Ukraine 3Laboratoire de Tectonique, Universiti Pierre et Marie Curie, 4 place Jussieu, 75252 Paris cedex 05, France 4Naukanaftogaz, Naftogaz of Ukraine, Uritckogo 45, 03035 Kyiv, Ukraine 5Institute of Geophysics and Tectonics, School of Earth and Environment, Leeds University, Leeds LS2 9JT, UK 6present address: Geological Survey of Norway (NGU), Leiv Eirikssons vei 39, N-7491 Trondheim, Norway 7Department of Earth and Ocean Sciences, University of Liverpool, Liverpool L69 3GP, UK

Abstract: The (Mid-) Late Devonian to Early Carboniferous was a time of widespread rifting on the East European Craton (EEC) and its margins. The most prominent basin among these and, accordingly, the best documented is the Dniepr-Donets Basin (DDB) in Ukraine and southern Russia. The DDB is associated with voluminous rift-related magmatism and broad basement uplift. Two other large, extensional, basin systems developed along the margins of the EEC at the same time: the East Barents Basin (EEB) and its onshore prolongation the Timan-Pechora Basin (TPB), and the Peri-Caspian Basin (PCB). Rifting, associated magmatism, and possible domal basement uplift are also reported elsewhere within the EEC, suggesting a common, 'active', rifting process, involving a cluster of thermal instabilities (or generalized thermal instability) at the base of the lithosphere beneath widely separated parts of the EEC by Mid-Late Devonian times. The DDB is an intracratonic rift basin, cutting across the Archaean-Palaeoproterozoic structural grain of its basement and, as such, differs from the EBB-TPB and PCB, which are pericratonic rift basins developed on reworked and juvenile crystalline basement accreted to the EEC during the Neoproterozoic. The DDB opened into a deep basin, possibly having oceanic lithospheric affinity, to the SE, in the area where it adjoins the southern PCB, suggesting the possibility that rifting led to (limited?) continental break-up in this area at this time. Post-rift compressional tectonic reactivations and basin inversion in the DDB, leading to the formation of its prominent Donbas Foldbelt segment, are related to Tethyan events (Cimmerian and Alpine orogenies) occurring on the nearby southern margin of the EEC. Post-rift compressional inversions in the PCB and TPB, which lie closer to the Urals margin of the EEC, are related to Uralian tectonics.

The Late Palaeozoic, in particular the Late Devonian, was an important time for extensional basin development on the East European Craton (EEC) and along its margins (see Fig. 1). The most prominent basin among these, and the one that received the most attention by E U R O P R O B E (Stephenson 1996), is the D n i e p r - D o n e t s Basin (DDB). This is an intracratonic rift basin with well-defined syn- and post-rift sedimentary successions within the Archaean-Palaeoproterozoic Sarmatian segment of the EEC. Two other large, extensional, basin systems developed along the margins of the EEC during the Late Palaeozoic: the East Barents Basin (EBB), mainly below sea level at present, with its onshore prolongation the T i m a n - P e c h o r a Basin (TPB), and the Peri-Caspian Basin (PCB), also in part below present-day sea level (northern Caspian Sea). Late Devonian intracratonic tiffing, associated magmatism, and possible domal basement uplift are also reported elsewhere within the EEC, on the Kola Peninsula (Kontozero Graben) and in the Vyatka Rift (Fig. 1). Whereas extensional tectonics and basin formation characterized the whole of the EEC during much of the Late Palaeozoic, the margins of the European continent had been or were shortly to be strongly affected by orogenesis during this time (Fig. 1). The main aim of this paper is to make a critical reassessment of what is actually known about Late Palaeozoic basin development on the EEC and to judge this in terms of lithospheric processes that may or may not be linked to plate boundary (convergence a n d / o r divergence) tectonic events taking place at about the same time. The more complete knowledge of the DDB is used as a point of departure in discussing these issues as they pertain to the other Late Palaeozoic basins of the EEC.

Overview of major Late Palaeozoic rift basins: architecture, magmatism and crustal structure

Dniepr-Donets Basin (DDB) The D n i e p r - D o n e t s Basin (DDB) is located in the southeastern part of the EEC along a N W - S E - t r e n d i n g axis between the present-day Ukrainian Shield and Voronezh Massif (Figs 1 and 2). It is part of the same rift basin system as the shallower Pripyat Trough to the NW (mainly in Belarus) and the inverted Donbas Basin (Donbas Foldbelt (DF), straddling the U k r a i n e Russia border; Fig. 3) and its prolongation to the SE, the presentday Karpinsky Swell (Fig. 2). The sedimentary succession of the DDB can be readily subdivided into pre-, syn-, and post-rift series, corresponding to pre-late Frasnian ( D 2 _ 3 ) , late F r a s n i a n Famennian (D3), and post-Devonian units, respectively (Fig. 4). The sedimentary thickness increases southeastwards to more than 20 km in the DF. Rifting may have begun slightly earlier in the SE, propagating northwestwards (see Stephenson et al. 2001; McCann et al. 2003). The post-rift succession is well developed, displaying evidence of multiple extensional reactivations as well as compressional tectonic events. A lack of stratigraphy of suitable age in the DF (as a result of subsequent uplift, deformation, and erosion affecting it) and, therefore, an absence of diagnostic structural relationships led to uncertainty regarding the timing and nature of post-rift tectonic events controlling its development (see Fig. 3). However, by comparing the exposed and drilled geology of the DF with seismic images from the adjacent Donets segment of the DDB (e.g. Fig. 4d), Stovba & Stephenson (1999) demonstrated that the main pre-inversion events affecting the DF

From: GEE, D. G. & STZPI-IENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 463-479. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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Fig. 1. Late Devonian palaeogeography of the EEC and surrounding areas (modified from Ziegler 1988). Pink, 'shield', areas; brown, inactive orogenic belts; grey, active orogenic belts; light blue, sedimentary platforms; green, ocean basins. Rift basins (dashed lines, labels italicized): BrS, Barents Shelf; DDB, Dniepr-Donets (Rift) Basin; KG. Kontozero Graben (Kola Peninsula); PCB, Peri-Caspian Basin; PKR, Pechora-Kolva Rift; PT, Pripyat Trough; VR, Vyatka Rift. Other abbreviations: EEP, East European Platform; UkS, Ukrainian Shield; VM, Voronezh Massif. Also shown, with boxes labelled accordingly, are the approximate locations of the maps shown in Figures 2 (DDB), 6 (PCB), and 8 (EBB-TPB). do not significantly differ from those of the DDB. This has been confirmed by subsequent structural studies (Saintot et al. 2003a,b) and by DOBREflection deep seismic profiling (Storba et al. 2005). The oldest sediments in the DDB are of Eifelian to MidFrasnian age, the so-called 'undersalt', pre-rift sediments. These were deposited in platformal terrestrial and shallow marine

environments and comprise sandstones, siltstones, clays and carbonates. These pre-rift Devonian sediments correlate with equivalent Devonian sequences of the East European Platform (Eisenverg 1988). They are characterized by homogeneous lithofacies, have an average thickness of 300-400 m, and include a series of stratigraphic gaps, the most significant being between the Eifelian and Givetian and between the Givetian and Frasnian. Thickness variations of this pre-rift succession are independent of the modern basement relief, although it is observed only locally on the rift shoulders. They were probably deposited over a much wider area, but were eroded during synrift uplift of the rift flanks. Similarly, they are locally absent atop intrabasinal structural highs developed during rifting. The marine Mid-Devonian sediments are not recorded in the southern DF area, where Eifelian to early Frasnian sediments are continental clastic deposits, transported northwards, with a few lacustrine carbonate intercalations, mainly deposited in a fluvial or delta-plain setting (McCann et al. 2003). Basal conglomerates rest unconformably on weathered Precambrian basement and are reportedly associated with fissural basaltic extrusive rocks already in the Eifelian (McCann et al. 2003). There is no evidence for the presence of a coaxial, but narrower, pre-Devonian, perhaps Riphean-aged graben underlying the DDB, as reported in much of the older literature (e.g. Chekunov et al. 1992). This was based on deep seismic sounding (DSS) velocity models, but is not observed on seismic reflection profiles recorded up to 12 s two-way travel time (TWT) (Stovba et al. 1996). No strata older than Mid-Devonian have been encountered in any of the numerous boreholes that penetrate basement beneath the Palaeozoic sediments of the DDB (Chirvinskaya & Sollogub 1980; Eisenverg 1988). Rather, the Devonian-Carboniferous succession revealed by the reflection data is much thicker than inferred from the earlier velocity models and occupies those parts of these models thought previously to represent Riphean strata (see Stovba et al. 1996). Therefore, tectonic models suggesting that a precursor Riphean rift basin was reactivated during the Devonian are no longer viable. Although modified by post-rift tectonic and especially salt movements (Stovba & Stephenson 2003), the basic architecture

Fig. 2. Tectonic map of the southern EEC, showing the extent of the Late Devonian Pripyat-DDB-DF rift basin. The dashed-line box indicates the location of the map shown in Figure 3. The locations of cross-sections shown in Figure 4 are also shown (red lines).

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Fig. 3. Cenozoic subcrop map of the Donbas Foldbelt area, showing the location of the DOBREflection profile (continuous line), shown in Figures 4e and 5. PreC, Precambrian crystalline basement; other stratigraphic labels are as in Figure 4.

of the DDB, seen in Figure 4, was developed during its Late Devonian rifting stage. High and laterally variable synrift subsidence rates, accompanied by the development of grabens and half-grabens, resulted in a wide range of local depositional environments and considerable palaeogeographic heterogeneity both in space and time, intense volcanism, and multidirectional tectonic movements (Stovba et al. 1996). Synrift deposits in the DDB reach a maximum thickness of about 4 km (e.g. Ulmishek et al. 1994; Stovba et al. 1996). They overlie the pre-rift sequence discordantly and are absent on some interior fault blocks, as a result of reduced sedimentation and subsequent erosion. Much of the lower part of the synrift sequence consists of Frasnian salt, called the 'lower salt', that alternates with clastic deposits and carbonates rocks in a complex laterally variable pattern. The depositional thickness of the Frasnian series is at least 1000 m and reaches a maximum (up to 2 km) in the axial zone of the southeastern part of the DDB. The upper part of the synrift series consists of a thinner Famennian 'upper salt' that thickens in the northwestern part of the DDB. In the southern DF, McCann et al. (2003) described the formation of half-grabens along major normal faults, filled with fluviatile continental clastic deposits and some lacustrine limestones. Short-lived but very frequent subaerial fissural extrusions (making up about two-thirds of the sequence) are always preceded by clastic input showing relief formation and erosion. Synrift volcanic and intrusive rocks, consisting of a variety of alkali basalts and their differentiates and associated pyroclastic deposits, occur in two main series of late Frasnian and late Famennian age in the DDB, attaining thicknesses of more than 2000 m (e.g. Wilson & Lyashkevich 1996). Additionally, there was widespread intrusion of tholeiitic basalt dykes, sills and stocks, which cross-cut formations ranging from Frasnian to late Famennian in age. The synrift phase s e n s u stricto terminated by the end of the Devonian and, in general, the Carboniferous and younger post-rift sedimentary fill of the DDB has the configuration of a broad syncline centred on the rift axis, overlapping the rift shoulders, and increasing in thickness towards the SE (Fig. 4). Seismic profiles published by Stovba et al. (1996) clearly demonstrate, however, that the DDB was affected during its Permo-Carboniferous

evolution by a series of post-rift extensional reactivations, generally synchronous with salt movements (Stovba & Stephenson 2003), but tectonic in origin; these occur at the end of the early Vis6an, during the mid-Serpukhovian, and during latest Carboniferous-earliest Early Permian times. Evidence of these events is visible in the regional cross-sections (Fig. 4); they have been comprehensively documented by Stovba & Stephenson (1999). The intensity of each of the Permo-Carboniferous extensional events increases in the DDB southeastwards towards the DF, where the late early Vis6an rift reactivation is clearly in evidence in the field, with uplifted and clearly rotated blocks along active normal faults, and associated magmatic activity (McCann et al. 2003). Saintot et al. (2003a) inferred a N N E - S S W extension in the DF that clearly affected the Early Carboniferous and older succession and, also, a younger transtensional stress regime thought to correspond to the latest Carboniferous-earliest Early Permian event recognized in the DDB by Stovba et al. (1996). Additional evidence of Early Permian extensional deformation along the northern margin of the DF, documented widely but generally not in published literature, was presented and discussed by Stovba & Stephenson (1999). Elsewhere in the DF, sediments of Late Cretaceous age directly overlie block-faulted and rotated Devonian and Carboniferous strata. The lack of a Permian-Early Cretaceous sedimentary record prevents a definite interpretation of the age of these faults; however, it is likely that the faulting and block rotation seen along the southwestern margin of the DF are part of the widespread phase of Early Permian extension (transtension) seen throughout the DDB (Stovba & Stephenson 1999). The Carboniferous succession is represented by continental deposits in the northwestern part of the DDB (e.g. Ulmishek et al. 1994; Dvorjanin et al. 1996; Izart et al. 1996). Elsewhere in the DDB it is characterized by continuous rhythmic sedimentation and comprises mainly siliciclastic rocks (with some clastic-carbonate sequences) deposited in shallow marine and lagoonal environments. There is little variation in the position of the basin depocentre. Only in the axial part of the southeastern DDB, where the Lower Carboniferous sequence includes marine carbonates, did the depth of deposition exceed 200 m. Exposed and drilled Lower Carboniferous sediments in the DF are mainly

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Fig. 4. Structural cross-sections through the DDB (a-d) and DF ((e); from DOBREflection), based on depth-converted versions of interpreted regional seismic reflection profiles (from Stovba et al. 1996, 2005; Stovba & Stephenson 1999). (For locations, see Fig. 2.) Light blue, salt bodies, brown, Devonian sediments, shades of grey, Carboniferous Ar-PR1, Archaean-Palaeoproterozoic; D2_3, Middle Devonian-Upper Devonian; C, Carboniferous; C1, Lower Carboniferous, (t, Tournaisian, vl, lower Visean, v2, upper Visean, s, Serpukhovian); Ca, Middle Carboniferous (Ukrainian-Russian usage: b, Bashkirian; m, Moscovian); C3, Upper Carboniferous (Ukrainian-Russian usage: e.g. Kasimovian and Gzelian); P1, Lower Permian, (as, Asselian; s, Sakmarian); Mz, Mesozoic; T, Triassic; J, Jurassic; K, Cretaceous; K2, Upper Cretaceous; Kz, Cenozoic; Pg, Palaeogene.

LATE PALAEOZOICBASINS ON THE EEC marine limestones overlain by sandy-clay deposits interbedded with thin coal and limestone beds. The (uppermost Famennian?-) Lower Carboniferous limestones were probably deposited in a very quiet shallow-water inner platform with occasional terrigeneous input. The uppermost part of this succession where it is exposed in the DF shows subaerial karstification, suggesting its emergence at the time of late early Vis~an extensional reactivation. The overlying silica-rich unit shows numerous synsedimentary deformational features such as normal faults and slumps (McCann et al. 2003). Middle and Upper Carboniferous successions are exposed throughout most of the DF and consist mainly of arenaceous-argillaceous rocks interbedded with coal and limestone. With the exception of coal beds and sandy-clay continental intercalations, most were deposited in a shallow-marine environment. Carboniferous sediments in the DDB reach thicknesses of 11 km, with the maximum depth of their base at about 15 km (Stovba et al. 1996). The present-day total thickness of Carboniferous sediments in the DF area is about 20 km based on the DOBREflection profile acquired as part of the EUROPROBE programme (Figs 4e and 5). The lowermost Lower Permian sediments are represented by monotonous sand-shale series containing rare interbeds of limestones and coals that, similar to the Upper Carboniferous units, reflect coastal-continental facies. Asselian sediments consist of five to seven layers of rock salt, separated by clastic deposits and carbonates, and also include numerous beds of gypsum, anhydrite and dolomite. The thickness of the salt layers, and per cent volume, increases upward in the section (Eisenverg 1988). The Sakmarian part of the series consists of a single salt layer probably representing redeposited Devonian salt dissolved from diapirs piercing the depositional surface in the Early Permian (Stovba & Stephenson 2003). In the southern pre-shoulder zone of the DDB the Lower Permian sequence abruptly decreases in thickness and pinches out as a result of a decrease in depositional thickness as well as subsequent erosion. In contrast, its thickness decrease towards the northern shoulder of the basin is far more gradual. There are no sediments of Early Permian age preserved within the DF, although Upper Carboniferous and Lower Permian sediments are documented beneath the eastern extension of the northern margin of the DF. A general absence of Upper Carboniferous and Lower Permian sediments in the northwesternmost part of the DDB can be explained by a decrease in the rate of post-rift subsidence within a platform-wide regime of relative sea-level fall. Elsewhere within the DDB, the basin margins, particularly the southern one, were exposed during Early Permian times whereas the axial part of the basin continued to subside (see Fig. 4). Uplift of the southern margin of the DDB was very shortlived, lasting no more than 2 - 3 Ma between the late Asselian and early Sakmarian (Stovba et al. 1996). Extensive erosion occurred, with progressively older sediments subcropping beneath the erosion surface in the direction of the Ukrainian Shield; by implication, considerable erosion of the Ukrainian Shield may also have occurred. Locally more than 2 km of Upper and Middle Carboniferous sediments were eroded at this time and during an ensuing dormant phase, which lasted until the Triassic. The widespread regional Permian unconformity observed throughout the DDB is, therefore, interpreted to be the result of the Early Permian event followed by a relative sea-level lowstand during the later Permian. Sedimentation resumed in the DDB in the Triassic, a time of tectonic quiescence, rising sea levels, and the resumption or continuation of post-rift subsidence. Most of the Mesozoic succession, comprising both marine and continental sediments, occurs throughout the area, overlying the rift axis as well as its flanks. Exceptions are the Upper Triassic and Lower Jurassic units, which occur only in the southeastern part of the DDB, and the Upper Cretaceous marls and chalks, which were eroded from large parts of the southern flank. The Upper Cretaceous succession

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was, as a whole, characterized by Chirvinskaya & Sollogub (1980) as 'close to' platform type, although subsidence coincident with the Devonian rift axis exceeds that of the marginal zones (see Fig. 4). It is up to 2000 m thick in the central part of the DDB, with maximum thicknesses of the Triassic, Jurassic and Cretaceous units being 900, 700 and 1000 m, respectively. In the vicinity of the DF, no marginal facies or developments are observed near the erosional edges of the Mesozoic successions. It is, therefore, likely that the entire area of the DF underwent post-rift subsidence during the Mesozoic and that, depending on relative sea-level variations, Mesozoic successions were deposited within its confines, but were later eroded. There is little evidence of post-rift magmatic activity in the DDB; however, this is not the case for the DF, where igneous rocks of Early Carboniferous, Early Permian and Mesozoic ages have been reported. A summary of the available geochronological data has been given by Alexandre et al. (2004). A widespread angular unconformity in the DDB developed at the end of Cretaceous-beginning of Palaeogene (Kabyshev et al. 1998). The magnitude of inferred relative uplift increases towards the Ukrainian Shield and, as during the Early Permian, its maximum occurred in the area bordering the DF. In this area, Upper Cretaceous, Jurassic and Triassic sediments were eroded (see Fig. 4). In the axial part of the southeastern DDB, there are local folds, domes and salt diapirs defining linear trends, which correspond to the trends of the main folds of the DF as seen on the Cenozoic subcrop map (Fig. 3). Within the DF itself, structural relationships determining the age of formation of folds, thrust and reverse faults can be observed only near its margins, where Lower Permian, Mesozoic and Cenozoic sediments are preserved. Stovba & Stephenson (1999) reported that no single geological section could be found in the published literature showing tightly constrained, structurally defined pre-Triassic folding or reverse faulting in the DF. In contrast, where Cretaceous sediments are present, for example along the northern margin of the DF, reverse faults and/or folds younger than the Cretaceous sediments and exposed at the surface are relatively common. Reverse faulting of Late Cretaceous age is also evident on the southern margin of the DF (Stovba & Stephenson 1999). Saintot et al. (2003a,b) determined that the palaeostress field associated with compressional structures observed in the Cretaceous sediments on the margins of the DF is identical to that recorded by the outcropping Carboniferous sediments. Thus it can be concluded that the inversion of the DDB and formation of the DF occurred mainly in the Late Cretaceous (see Stovba & Stephenson 1999; Stephenson et al. 2001). The DOBREflection profile (Fig. 5) shows that the shortening of the DF occurred at the crustal scale as a 'mega-pop-up', which involved a major detachment fault through the entire crust and an associated back-thrust (Maystrenko et al. 2003; Stovba et al. 2005). The Cenozoic section of the DDB unconformably overlies Upper Cretaceous and older series and reaches a maximum thickness of 500 m in the NW DDB (Eisenverg 1988). The Palaeogene sequence consists mainly of sands, clays and marls, and the Neogene sequence mainly of sands with clayey interbeds. Deep seismic sounding (DSS; e.g. Chekunov et al. 1992; Ilchenko 1996) and more recent wide-angle reflection-refraction (WARR) seismic studies (DOBREfraction'99 Working Group 2003) show that the amount of crustal thinning beneath the DDB increases to the SE, concurrently with increasing sedimentary thickness (see Stephenson et al. 2001). The most recent profile is DOBRE (DOBREfraction'99 Working Group 2003), crossing the inverted DF segment of the DDB (Fig. 5). The sedimentary basin itself is well-defined, overlying a main crustal layer that thins significantly beneath the main sedimentary depocentre. In turn, a high-velocity lower crustal layer thickens significantly in the same part of the profile. The shape of the sedimentary basin is asymmetric, with the steepest crystalline basement surface on the southwestern margin of the basin, whereas the asymmetry

468

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The Peri-Caspian Basin (PCB; also sometimes referred to as the Precaspian, Pricaspian or North Caspian Basin) is situated on the southeastern margin of the East European (Russian) platform and extends into the northern part of the Caspian Sea. It runs 900 km east-west and 600kin north-south, bordered on the east by the Ural Mountains and to the SE and SW by crustal terranes that have an uncertain relationship with the EEC (see Saintot et al. 2006). The sedimentary succession of the PCB is about 20 km thick (Figs 6 and 7). There is a prominent Lower Permian salt layer some 4-4.5 km thick with its base at a depth of 7-9.5 km in the central part of the basin. Overlying sediments comprise the 'postsalt' layer, which is up to 7 km thick, and range in age from Late Permian to Quaternary. What lies below (called the 'sub-salt layer', which is up to about 9 - 1 0 km thick) is characterized primarily on the basis of seismic data (e.g. Volozh 1991). The conventional view holds that four seismo-geological successions can be recognized in the 'sub-salt' section: Riphean, Lower Palaeozoic, Devonian-Lower Carboniferous and Middle Carboniferous-Lower Permian, separated by hiatuses seen as erosional unconformities in marginal areas. Even though a huge volume of seismic (reflection and refraction) data exists, the age of the sub-salt sediments in the centre of the basin is controversial because seismic correlation from the margins to the deep basin is rather uncertain, given complications arising from seismic facies, steep slopes and interruption of key seismic markers. This contrasts with the view of many authors that the reference horizons can be traced (robustly) through the entire depression (e.g. Lobkovsky et al. 1996). Thus, the oldest sediments in the basin could be as old as Riphean or as young as Devonian (see Brunet et al. 1999; Volozh et al. 2003a). Riphean ages are based on seismic velocities and fabrics, and on strata thought to be Riphean (but undated) encountered in wells on the northern and northeastern margins of the basin (e.g. Soloviev et al. 1989). It is worth noting that the now rejected (based on modern regional seismic reflection profiling (Stovba et al. 1996)) postulate of a thick Riphean sequence deep in the DDB was based on very similar arguments. According to Zonenshain et al. (1990) and others, all these sediments are of Devonian age (as in the DDB) and, as such, they could overlie Devonian-aged oceanic crust. The regional interpretation shown in Figure 7 is based on the conventional interpretation that Devonian sediments are underlain by Neoproterozoic and Early Palaeozoic successions (e.g. Volozh et al. 2003a). What can be stated with certainty is that Vendian strata occur above the Precambrian basement on the margins of the PCB, on the Russian Platform to the north and west, and on the eastern and southeastern margins, for example, in the South Emba region (Fig. 6). A sedimentary hiatus occurred in the Early Palaeozoic; probably no Cambrian series exists within most of the basin; and Ordovician-Silurian strata are limited in extent, although up to 1000 m thick in pericratonic troughs such as the South Emba (e.g. Brunet et al. 1999). Thick, more terrigenous deposits with marginal carbonate reef complexes form the Upper DevonianCarboniferous succession. The basin was more or less filled (probably accompanied by a sea-level drop) by Early Permian times and became isolated from the open sea by structural highs developed especially on its south and southeastern margins, after which the thick salt layer was deposited. The initial thickness of this layer (prior to reconfiguration by diapirism) is estimated to be about 4.5 kin, deposited mainly during the Kungurian (Early Permian). Clastic rocks of this age are present in the eastern part of the basin, sourced from the eroding Urals Mountains and

LATE PALAEOZOIC BASINS ON THE EEC

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Fig. 6. Main tectonic units of the PCB (modified from Volozh 1991; Volozh et al. 2003) showing surrounding marginal uplifts. The light green area in the central PCB indicates the extent of the high-velocity layer at the base of the crust (eclogites?). Also shown are the locations of the Aralsor and Khobda positive gravity anomalies ('AA' and 'KA', respectively), the location of the schematic cross-section shown in Figure 7 (with position C marked by an X for reference), and the 12 and 20 km depth contours to top basement.

showing basinward progradation. Post-depositional movement has resulted in the development of about 1800 salt structures in the PCB (Volozh et al. 2003b), of various types, some related to hydrocarbon production. Permian and younger sediments in the PCB were deposited in shallow-water or continental conditions but, because of the dominance of salt movement during this time in producing local, intradiapiric depocentres, little can be said about post-Permian tectonic controls on basin subsidence. The nature of unconformities in the PCB and the timing of fault activity are poorly described in the literature and subject to some inconsistency (e.g. Brunet et al. 1999). The Frasnian (beginning in

the Givetian?) has been reported by some workers as a time of active rifting in the PCB and the F a m e n n i a n - T o u r n a i s i a n as a time of relatively stable subsidence with the formation of a topographic depression not compensated by sediments (Nikishin et al. 1996; Volozh et al. 1999). However, it is extremely difficult to document this from the existing literature because stratigraphic boundaries in the PCB itself are poorly resolved and defined because different authors use different interpretations for the same seismic horizons. Soloviev et al. (1989) reported that the Riphean U z e n - S a k m a r a graben was reactivated in mid-Devonian (Eifelian) times and that tectonic movements occurred in the

Fig. 7. Simplified sketch (after Brunet et al. 1999) of a north-south basin-crustal cross-section of the PCB (from an unpublished interpretation and compilation of the seismic line Zhambay-Uralsk by Yu.A. Volozh, V.I. Kozlov & Yu.G. Yurov); location is shown in Figure 6 (with position C marked for reference). Seismic refraction velocities are indicated in the crust and the high-velocity layer. The presence of the Riphean and Lower Palaeozoic sedimentary layers in the deep basin is based on the interpretation of seismic velocities and is not confirmed by drilling.

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Late Devonian (Frasnian), accompanied by rapid subsidence towards the PCB. The Frasnian lies unconformably on Middle Devonian deposits in some wells on the eastern margin of the PCB (Akhmetshina et al. 1993). Abrupt changes in DevonianCarboniferous sequence thickness, limited by faults, as well as facies heterogeneity, are observed on the Astrakhan Dome (Brodsky et al. 1994), which is a structural high in the southwestern part of the PCB adjacent to the Karpinsky Swell (Fig. 6). Active faulting from the Late Devonian to the Mid-Carboniferous has been reported by Kirukhin et al. (1983) and Brodsky et al. (1994) in this area. A number of authors (e.g. Lobkovsky et al. 1996; Brunet et al. 1999) have reported tectonic subsidence curves for the PCB, but these are not well-constrained given the degree of actual borehole penetration and uncertainties of stratigraphic identification from seismic data (see Brunet et al. 1999). Backstripping of wells on the margins of the basin does not illuminate the Late Palaeozoic, and backstripping of synthetic stratigraphic columns from the basin centre are dependent on the authors' choice regarding the interpretation of key seismic horizons. However, independent of these considerations, a robust tectonic event did occur in the evolution of the basin during the Devonian, with other events dependent upon age interpretations and assumptions regarding paleobathymetry. There exists very little solid reference to volcanogenic deposits in the PCB. Volozh (1991) reported late Riphean continental volcanoclastic deposits in the exterior zone of the northwestern region, Ordovician-Silurian marine to continental volcano-sedimentary deposits in the southeastern region (Primugodzhar, South Emba) and in the early Kungurian in the SE (Primugodzhar, Koltyk-Zamstan). Shein et al. (1989) showed on a north-south cross-section volcanogenic rocks of mid-Devonian and Carboniferous (Vis6an-Bashkirian) age in the centre of the basin and of Triassic age in the south. According to Kostyuchenko et al. (2004), the presence of pyroclastic rocks of variable composition indicates volcanic activity during the Early and Mid-Carboniferous (Kalashnikov 1974; Vishnevskaya & Sedaeva 2000). Kostyuchenko et al. (2004) also reported the interpretation of Brodsky et al. (2000), from seismic data (reflection and refraction) across the Astrakhan Dome (Fig. 6), of the presence of a large basic extrusive magmatic body (about 40 km wide) lying below a depth of about 12 km, down to about 24 kin. The age of sediments overlying this proposed volcanic unit, and therefore helping to date it, is subject to exactly the same uncertainty, discussed above, relating to the PCB in general; that is, it is dependent upon the acceptance or not of the unverifiable deep seismic stratigraphy. The inferred volcanic body is overlain by sediments older than the Devonian succession according to the conventional view (e.g. Volozh et al. 2003a) and, therefore, of Ordovician-Silurian age. The alternative view, that the whole of the sedimentary succession is Devonian (e.g. Zonenshain et al. 1990), would allow this inferred magmatic body to be of Devonian age. The crystalline basement beneath the central PCB is thin, being only some 10-12 km thick, characterized by the absence of an upper crustal velocity layer (velocities <6.6 km s-J; Fig. 7). Its origin is still controversial; it could be either thinned continental or (see Zonenshain et al. 1990) oceanic crust. At the base of the crust, an 8 - 1 0 km thick layer of velocity 8.0- 8.1 km s- 1 is recognized by seismic and gravity observations; this has been interpreted by some authors as being eclogitic in composition (Volozh 1991; Volozh et al. 2003a). Kostyuchenko et al. (1999) considered this layer to be part of the upper mantle, underlying strongly thinned continental crust (probably affected by intrusion of some mantle material). Indeed, upper mantle velocities of 8.1 km s -1 beneath a rift basin, the subsidence of which is easily explicable without the invocation of upper mantle phase changes, is not particularly unusual (cf. Christensen & Mooney 1995). In this regard, it is also worth noting that the subsidence produced by the postulated eclogite body beneath the PCB is usually ascribed to events earlier than the Late Palaeozoic

(e.g. Volozh et al. 2003a). Volozh et al. in fact actually developed the idea that the crust underlying the PCB comprised an accretionary belt of Neoproterozoic age, interpreting the eclogite body as a lens emplaced during 'Baikalian Proto-Urals' subductionorogenesis (see also Saintot et al. 2006). East Barents Sea- Timan-Pechora Basin (EBB- TPB)

The East Barents Sea is that part of the Barents Sea lying contiguously with the Russian mainland and is separated from the West Barents Sea by the north-trending Central Barents (Sea) Uplift. The sedimentary basins of the latter were formed by rifting during Mid-Carboniferous times and form a prolongation of an Arctic-North Atlantic rift zone between Greenland and Norway (Ziegler 1988; Dor6 1991; Gudlaugsson et al. 1998). The Sverdrup Basin in Arctic Canada may have formed part of the same rift basin system (Embry 1989; Stephenson et al. 1987). The relationship between these (Arctic-North Atlantic-West Barents) basins and those of the East Barents Sea is not considered further here. The eastern Barents Shelf (Fig. 1) is dominated by the Late Palaeozoic-Triassic age East Barents Basin (EBB), trending NNE, with a length of 1500 km from south to north and a width of 3 0 0 - 6 0 0 k m (Gramberg 1988; Dor~ 1991; Verba et al. 1992). It comprises two sub-basins (northern and southern) separated by an east-west axial uplift (Fig. 8). The deeper, faultcontrolled, parts of these basins are interpreted as synrift successions of either Devonian (Zonenshain et al. 1990; Johansen et al. 1993) or Carboniferous-Permian (Verba et al. 1992) age. The Russian onshore equivalent of the EBB is the TimanPechora Basin (TPB), as shown in Figure 8. It lies on the most northeastern part of the EEC, bounded on its SW margin by the Timan Range (see below). To the east and NE the TPB merges with, or is overlain by, the foreland basin of the Northern Polar Urals Foldbelt. The basin reaches 400 km in width along the Barents shore with a maximum depth to basement of up to 10 km onshore (except in the Urals foreland where sediment thickness is greater). The EBB and TPB are sometimes considered as separate entities but here they are treated together, as a single Late Palaeozoic basin system, the EBB-TPB. An immediate issue that arises in so doing has to do with the affinity of the crust underlying the basins. There has been some disagreement about this in the past, with models that included Archaean and Palaeoproterozoic massifs (being part of what Zonenshain et al. (1990) referred to as Barentsia) embedded within younger terranes (e.g. Khain 1977), although there is now a general consensus that the basement of both the TPB and EBB mainly comprises crust accreted to the EEC during the Neoproterozoic Timanide Orogen (e.g. Gee & Pease 2004, and references therein). The suture between crust accreted at this time and the older, pre-existing margin of Baltica (see Gee et al. 2006; Kostyuchenko et al. 2006), lies somewhere in the vicinity of the subsequently formed Pechora-Kolva Trough (Fig. 8), sometimes referred to as the Pechora Rift. Potential field data suggest that basement structures are continuous between the onshore TPB and the offshore EBB (e.g. Dedeev & Zaporozhtseva 1985). According to seismic data, the thickness of the sedimentary succession of the EBB exceeds 15-18km, comprising Upper Silurian-Devonian to Triassic-Lower Jurassic units that are covered by Pliocene-Quaternary sediments. Figure 9a shows a schematic geological interpretation based on seismic data. Identification of the deeper horizons, as in the PCB, is ambiguous. The total sedimentary thickness decreases into the TPB, to a maximum of 7 - 8 lon (Fig. 9b), where the oldest sediments are thought to be terrigenous Cambrian-Ordovician (Ismail-Zadeh et al. 1997) or Ordovician-Silurian (Belyakov 1994; Lobkovsky et al. 1996; Bogolepova & Gee 2004) units. Vendian-aged rocks are included in the profile shown in Figure 9b, taken from Spiridonov et al. (1999). The Vendian-Early Palaeozoic units are considered to

LATE PALAEOZOICBASINS ON THE EEC

and Kanin Peninsula, and drilled in a number of wells, occurred in the TPB in the Late Devonian (Khain 1977; Grachev et al. 1994). dating of dolerite sills from the Kolva superdeep well (on the Kolva Swell), intruded into Early Devonian sediments, indicates an Early Carboniferous (Visran-Tournaisian boundary) emplacement age (Wilson et al. 1999). Numerous intrusions are reported in the Lower-Middle Devonian, Permian-Triassic and Jurassic successions of the EBB (Komarnitckii & Shipilov 1991; Shipilov & Tarasov 1998), in part identified on the basis of seismic sections (Bogolepov et al. 1991; Sakoulina e t al. 2003). Similarly, the Triassic succession in the TPB is also thought to contain a significant component of basic magmatic rocks, inferred mainly on the basis of seismic data. According to Nikishin et al. (2003), this magmatism is related to a distinct Permo-Triassic rifting event affecting the TPB. Zonenshain et al. (1990) considered the magmatism to be (Early) Triassic age and related to flood-basalt extrusion and rifting in adjacent parts of Eurasia. The Devonian grabens within the TPB are reported to be partly inverted during the Carboniferous-Early Permian (Nikishin et al. 1996; Fokin et al. 2001; lasting through the Triassic according to Lobkovsky et al. 1996), leading to the development of the Pechora-Kozva and Kolva swells (Fig. 8b), which by the end of the Triassic were eroded into basement uplifts, as seen today, according to Khain (1977). The compressional structures appear to have controlled the distribution and evolution of Permian reefs and bioherms (Belyakov 1994). Models of DSS data collected in the 1970s and 1980s across the southern EBB indicate a Moho depth of 40 km on the continental shelf transition to the TPB, with sediments being some 7 - 8 km thick (although with relatively low velocities in the underlying crustal units; Verba et al. 1986; see also Davydova et al. 1985; Morozova et al. 1995). Under the main, NNE-trending depocentre of the EBB, the Moho shallows to about 35 km and the thickness of the crustal layer (beneath up to 18 km of sediments) to less than 20 km. However, much of the 'crustal' layer is interpreted to comprise deeply buried metamorphosed sediments, so the thickness of sediments associated with Late Palaeozoic rifting could be greater than 18 kin. The crustal layer itself is characterized by elevated velocities, said to be characteristic of basalt (Verba et al. 1986; Tulina et al. 1988), and the absence of a more felsic upper crust (given the proposed interpretation of the metamorphosed sediment layer) below the deep part of the basin. Neprochnov et al. (2000), based on an analysis of all available crustal seismic data for the EBB at the time, concluded that the crust beneath the basin has (thinned) continental affinity, having both a felsic as well as a mafic layer (though the former may be absent beneath the axial parts of the rift). According to Aplonov (1992), mainly based on potential field data, oceanic crust underlies the EBB, presumed to be of Devonian age, being a remnant of either Iapetus (Barents Sea shelf and northern part of Kara Sea) or the Uralian Ocean (southern part of the Kara Sea). Nikishin et al. (1996) also considered that this area is underlain by oceanic crust (of unstated age) or by very thin continental crust. More recently, there has been considerable new seismic profiling, including the acquisition of deep CDP data in the EBB. New seismic studies were carried out in the area between the Kola coastline, Frans-Jozef Land and Novaya Zemlya along two profiles (Mitrofanov & Sharov 1998a,b; Sakoulina et al. 2000, 2003; Verba et al. 2005). In general, these studies have confirmed the earlier inferences: a very thin to absent upper crustal layer beneath the deepest part of the basin and, correspondingly, high seismic velocities just below the basin, the sedimentary infill of which is estimated to be more than 14-16 km in thickness in the deepest part of the basin (Sakoulina et al. 2000, 2003; Verba et al. 2005). Sakoulina et al. (2003; see also Morozova et al. 1995) explained the absence of an upper felsic layer in the crust of both the South and North Barents basins, at least in part, by mafic magmatism affecting the crust during Palaeozoic rifting.

4~

Fig. 8. CombinedEBB-TPB depth to basement map (contours in km; from Bogolepov et al. (1991) for the EBB and Spiridonov et al. (1999) for the TPB), showing the main regional structures (red lines, from Bogatski et al. 1996) and locations of geologicalcross-sections shown in Figure 9.

be platformal-type deposits preserved beneath Middle-Upper Devonian synrift sediments, analogous to the pre-rift sediments of the DDB. This information is based on numerous boreholes as well as a considerable body of seismic reflection data recorded to 7 - 8 s (e.g. Spiridonov et al. 1999). According to O'Leary et al. (2004), however, there is evidence for mild extension driving Ordovician- Silurian subsidence. The main Devonian rift structure in the TPB is the PechoraKolva Trough (Figs 8 and 9b), typically filled with marine argillaceous-carbonate sediments having a maximum thickness of about 1.5 km. Post-rift subsidence ensued, culminating in the accumulation of evaporites in the Early Permian (Kungurian, as in the PCB and DDB), on what was at that time the shelf of a Uralian palaeo-ocean that subsequently closed during Uralian orogenesis. The depositional environment changed dramatically in the mid-Permian (Khain 1977), when carbonate-evaporite successions were replaced by thick siliciclastic deposits represented, in the TPB, by red sandstones, conglomerates and argillites sourced from the developing northern Urals mountains (Khain 1977; Zonenshain et al. 1990; see also O'Leary et al. 2004). MidJurassic and younger sediments lie unconformably on older strata and were deposited in a quiescent tectonic environment. They do not reach any significant thickness (rarely greater than 1 km). Intense, rift-related, basaltic volcanism, represented by both lavas and numerous dolerite sills exposed on the Timan Range

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Fig. 9. (a) Geological model from seismic and well data along line II in the EBB (from Bogatski et al. 1996) with inferred intrabasinal volcanic rocks marked with a F symbol and Precambrian (crystalline) basement with ' + '; (b) geological model from seismic and well data along line 15PS in the TPB (from Spiridonov et al. 1999), showing well locations used in constructing the profile. The profile locations are shown in Figure 8. Stratigraphic labels are as in previous figures.

The crustal structure beneath the TPB has been studied along a number of seismic profiles acquired during the past two decades. There are various models and interpretations of these seismic data (Egorkin 1991; Kozlenko et al. 1994; Morozova et al. 2000; Pavlenkova 2000) but, in general, there are no significant differences between them. The velocity structure of the TPB crust is different from that of the adjoining EEC, related to its younger time of accretion during the Timanian orogeny (see Gee & Pease 2004, and references therein), having generally lower velocities and only two main (crustal) layers. There is some suggestion of Moho shallowing and concomitant crustal thinning beneath the main rift axis, but no obvious evidence of a high-velocity body commonly associated with magmatic processes in the lower crust accompanying tiffing (e.g. Kostyuchenko 1994). Such a layer does, however, appear towards the Urals foreland basin (Fore-Ural Trough), where it accompanies an abruptly deepening Moho (to 50 km), similar to the crustal structure of the Urals themselves (Egorkin 1991; Pavlenkova 2000). The more recent Beloe More (White Sea)-Vorkuta profile (Kostyuchenko 2005; see also Kostyuchenko et al. 2006) supports a continental affinity of the crust below the TPB, comprising juvenile Neoproterozoic oceanic crustal fragments and

microcontinental fragments (accreted during the Timanian orogeny). In the central part of the basin, approximately in the region of the Pechora-Kolva Trough, a local thickening of the crust (up to 46 km), accompanied by a thick (c. 20 km) body in the lower crust with increased seismic velocities, has been inferred. Kostyuchenko (2005; see Kostyuchenko e t al. 2006) considered this thickening to be the result of the Neoproterozoic collisional processes; however, it could also represent Palaeozoic magmatic underplating.

Discussion The Devonian was a time of a major change in plate configuration, leading up to the amalgamation of the Laurussian and Gondwanan supercontinents as the single world-continent Pangaea by the Permian. Laurussia (Fig. 1) was largely surrounded by subduction zones (see Saintot et al. 2006), the presence of which has been argued to be the driving mechanism of extension leading to formation of basins such as the DDB ('back-arc extension'; Ziegler 1990; see also Stephenson e t al. 2001; Saintot et al. 2006). However, two counter-observations can be made in this

LATE PALAEOZOICBASINS ON THE EEC regard: (1) extensional rifting, accompanied by magmatism, occurred throughout the European (EEC) part of Laurussia (for example, in the TPB and elsewhere, as discussed above) but the orientation of these rift zones is strongly divergent from that of the DDB; (2) convergent zones also surrounded the Laurentian (Canadian Craton; CC) part of Laurussia but Late Palaeozoic intra- or pericratonic extension is not notable for the CC. Minor, rift-like extensional provinces and basin development may have occurred in the Devonian in the Peace River Arch area (pericratonic, oblique to the margin of the CC; see Stephenson et al. 1989) and in Hudson Bay (intracratonic, reportedly E a r l y mid-Devonian, not Late Devonian; Hanne et al. 2004) and were not accompanied by significant magmatism. Nevertheless, it is easy to surmise that what drove rifting within the EEC was indeed related to plate boundary processes (remembering that the base of the lithosphere is also a plate boundary), even if a direct link to subduction is not explicit. The DDB differs from the PCB and the E B B - T P B in that it is truly intracratonic, formed in Archaean-Paleoproterozoic lithosphere as opposed to younger Neoproterozoic basement in a pericratonic setting. What follows is a discussion of tift-forming mechanisms (and subsequent deformation) relevant to the Late Palaeozoic rifts of the EEC, as described above, and the implications of these for the tectonic setting of the EEC during this time, using the DDB' s role as a point of reference given its better constrained crustal structure and basin architecture.

W a s there a m a n t l e p l u m e b e n e a t h the E E C

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alone, assuming anhydrous mantle and a normal asthenosphere potential temperature (McKenzie & Bickle 1988). To generate partial melting by extension-driven decompression at such low levels of extension, either volatiles such as water and carbon dioxide are required in the lower lithosphere (or asthenosphere) or the mantle must be anomalously hot. The latter suggests the presence of a mantle plume. The occurrence of significant volcanic activity during the earliest synrift stage in the Frasnian (Wilson & Lyashkevich 1996), when stretching factors were less than 1.02 for the Pripyat Trough (Kusznir et al. 1996a), suggest that lithospheric extension is not the most important trigger for magmatism. The major and trace element and S r - N d isotope characteristics of the magmas emplaced within the PDD rift also provide some supporting evidence for involvement of a mantle plume in their petrogenesis (Wilson & Lyashkevich 1996). The alkalineultramafic magmas appear to have been generated at considerable depths in the mantle ( > 150-200 km), whereas the alkali basalts may have segregated close to the base of the lithosphere (c. 100km). The extensive tholeiitic dolerite dykes and sills intruded into the basin fill suggest high degrees of partial melting towards the end of the tiffing period. Consequently, it seems likely that the (Pripyat-)DDB magmas were generated not by extension-driven decompression melting, but rather by decompression melting within a rising mantle plume, which also largely drove contemporaneous rifting and basement uplift. In this respect, it should be noted that extension, volcanism, and (reportedly) basement uplift also occurred during the Late Devonian within the Kola and Vyatka regions of the EEC (Fig. 10). The

in the L a t e D e v o n i a n ?

The DDB appears to be characterized by more intensive synrift magmatism than either the E B B - T P B or the PCB (although it is more difficult to assess for the latter two) and is associated with a regional basement high. It is bounded by two Precambrian basement massifs, the Ukrainian Shield to the south and the Voronezh Massif to the north, together defining a domal structure, some 1000km in diameter, transected by the rift. As mentioned above, the Ukrainian Shield and the Voronezh Massif were partially covered by marine middle Devonian sediments prior to the onset of rifting. Analysis of stratigraphic data suggests that uplift of the Voronezh Massif probably commenced in the midFrasnian, reaching a maximum at the Frasnian-Famennian boundary (Alexeev et al. 1996). The time of most rapid rifling coincides with the most active period of volcanism documented by Wilson & Lyashkevich (1996). Kusznir et al. (1996b) proposed, based on unconformities related to footwall uplift on the rift-bounding fault systems, that c. 300 m of regional uplift occurred during the rifting phase in the NW DDB. Subsequently, the same 300 m was recovered, according to the model, as a regional downwarp during the postrift stage, to provide sufficient sediment accommodation space in the Carboniferous. Kusznir et al. (1996b) suggested that the inferred transient regional vertical crustal motions were generated by the dynamic (i.e. not thermal) effects of a mantle plume. This uplift event commenced no earlier than mid-Frasnian, peaked in late Famennian and was largely dissipated by Visran-Serpukhovian times (mid-Carboniferous). Thus, peak uplift also corresponds to the time of most rapid rifting and most active volcanism. The sheer volume of magma generated and its geochemical signature, in the context of the relatively low lithospheric stretching factors inferred from subsidence modelling (Kusznir et al. 1996a,b; van Wees et al. 1996; Starostenko et al. 1999) and crustal structure (e.g. Ilchenko 1996), are in themselves indicative of mantle plume activity, supporting earlier hypotheses (e.g. Gavtish 1989; Chekunov 1994; Wilson & Lyashkevich 1996). Stretching factors in the Dniepr segment of the DDB (and adjoining Pripyat Trough) lie in the range 1.1-1.3, insufficient to generate decompression melting driven by passive extension

Fig. 10. Intra- and pericratonic Late Devonianrifling (dashed lines) and volcanism (• on the EEC and its margins, modified(somewhat simplified)after Nikishin et al. (1996). The veracity of reported extensional faulting and magmatism,as shown, has not been checked. Rifting, as shown, for the PCB is inferred only (as discussedin this paper); inferred volcanism on the Astrakhan Dome (southern PCB), reportedly of Early Palaeozoic age, has been added (as discussed in this paper). Basement doming around the DDB (not shown) is clear (comprisingthe Ukrainian Shield and Voronezh Massif; see Fig. 1), and Nikishin et al. (1996) referred to similar large-scale doming around the Kontozero Graben (KG) and Vyatka Rift (VR). The shaded zone represents basement proposed to have been accreted to the EEC during the Neoproterozoic; this is fairly well established beneath the TPB (Gee & Pease 2004) and is more hypothetical elsewhere (as discussed in this paper; see also (Seghedi et al. 2005; Saintot et al. 2006). EBB, East Barents Basin.

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mid-Late Devonian magmatism of the Kola-Arkhangelsk region, occurring between 380 and 360 Ma, constitutes one of the world's largest alkaline igneous provinces and coincided with subsidence of the NE-SW-trending Kontozero Graben (Kramm et al. 1993). Rifting and magmatism within the Vyatka Rift commenced during the late mid-Devonian (Givetian, 381-377 Ma) and terminated before the end of the Devonian (Nikishin et al. 1996). Magmatism and associated basement uplift may therefore have been widespread on the EEC in the mid-Late Devonian and, if these were caused by thermal processes beneath the lithosphere, then it follows that there was not a single, isolated, plume head but rather a cluster of thermal instabilities impinging on the base of the lithosphere beneath widely separated parts of the EEC.

C o n t i n e n t a l b r e a k - u p in the L a t e P a l a e o z o i c ?

Shatsky (1964) considered the DDB as a type example of an aulacogen. Although the DDB rift is clearly a 'failed rift', in the sense that it did not itself lead to continental break-up and ocean crust formation, there remains disagreement on whether it represents the failed third arm of a triple-rift system, the other arms of which formed a passive continental margin (e.g. Saintot et al. 2003a; see also Saintot et al. 2006). Subsequent to Shatsky's work and until quite recently, some workers (e.g. Chekunov et al. 1992; Volozh et al. 1999) have preferred an alternative hypothesis, regarding the DDB to be part of a very long, intracontinental rift system, including the Karpinsky Swell (on the northern margin of the northern Fore-Caucasus domain; Figs 2 and 6), that crosses the Caspian Sea and continues into western Asia. Such a model appears at least in part to be untenable within a plate-tectonic framework. In their plate-tectonic synthesis of the geology of the former Soviet Union, however, Zonenshain et al. (1990) represented the DDB as the failed arm of a Late Devonian rift system that led to the development of subsequently closed, small ocean basins along the southern margin of the EEC (see also Saintot et al. 2003a; Kostyuchenko et al. 2004). In recent palaeogeographical and palaeotectonic reconstructions, Nikishin et al. (1996) appeared to agree with this view, although it was not explicitly expressed as such. A critical issue in this respect would is the relationship of the PCB to the DDB and, thus, the tectonic origins and age of the oldest sediments in the former. As mentioned above, some workers (e.g. Volozh et al. 2003a) have argued strongly and without equivocation that these are undoubtedly at least as old as Early Palaeozoic, based primarily on the presence of sediments of this age on the northern, pericratonic platformal margin of the basin. However, other workers have suggested that definitive evidence of the age of these very deeply buried and inaccessible sediments simply does not exist and that a tectonic scenario in which the basin is mainly a consequence of Late Devonian rifting (contemporaneously with the DDB) fully and more satisfactorily fits the available observations (e.g. Brunet et al. 1999). At best, these conflicting views remain an open question. It seems likely that intracratonic rifting in the DDB was related to asthenospheric processes, possibly including upwelling of thermally anomalous mantle at the base of the lithosphere, as argued above. The thickness of the fill of the DDB increases significantly to the SE and, accordingly, so does the concomitant degree of lithospheric stretching based on subsidence modelling studies (van Wees et al. 1996; Stovba et al. 2003). Further, van Wees et al. (1996) found that lithospheric thinning became increasingly dominated by subcrustal (thermal) thinning rather than crustal (mechanical) thinning in this direction, to explain subsidence, implying an increase to the SE in the DDB in the importance of sub-plate, asthenospheric processes for driving rifting. Deep seismic sounding (Chekunov et al. 1992; Ilchenko 1996) and more recent WARR seismic studies (DOBREfraction'99 Working Group 2003) have shown that the amount of crystalline

crustal thinning beneath the DDB also increases to the SE, concurrently with increasing sedimentary thickness (see Stephenson et al. 2001). The Moho depth in the DF, where there is about 20 km of Late Palaeozoic and younger sediments, is more or less constant ( 4 0 k m _ 2 km; DOBREfraction'99 Working Group 2003; see Fig. 5). Taking into account that the lower crustal high-velocity layer found in this area has presumably incorporated (and been thickened by) mantle-derived magmas during the rifting process, the 'stretching factor' beneath the DF is as much as 2.25-2.5 (DOBREfraction'99 Working Group 2003). This degree of crustal thinning is close to that required for continental break-up. Although the intensity of rifting increases in the DDB to the SE, in the direction of the supposed triple junction, the intensity of extrusive magmatism seems not to do so. However, the presence of a lower crustal rift pillow, thicker beneath the DF (as noted above; DOBREfraction'99 Working Group 2003) than further NW in the DDB (e.g. Ilchenko 1996), does suggest that intrusive magmatism in general increases to the SE. Potential field anomalies confirm that mafic rocks have intruded the crust throughout the DDB. Magnetic anomalies in the Dniepr segment, indicative of sources within the sedimentary layer as well as in the crystalline basement, correlate well with igneous bodies known to lie within the sedimentary sequence (Orlyuk & Pashkevich 1994; Stephenson et al. 2001). Gravity modelling has long suggested the presence of a high-density lower crustal body along the rift axis interpreted as an 'axial dyke' related to rifting (Starostenko et al. 1986). Yegorova et al. (1999, 2004) showed that this feature extends the entire length of the DDB, intensifying to the SE through the DF to its culmination in the Karpinsky Swell, where it makes a 120 ~ bend to the NE parallel to the recognizable margin of the EEC. This culmination, beneath the PCB, may be the locus of a rift-rift-rift triple junction (Fig. 1 la). The apparent lack of magmatic activity associated with the Devonian tectonics of the PCB argues against its being part of a triple junction leading to continental break-up (i.e. Zonenshain et al. 1990). However, as mentioned above, Kostyuchenko et al. (2004) have identified, from seismic reflection data, a large volcanic edifice in the area of the Astrakhan Dome, in the southern PCB. Kostyuchenko et al. considered the age of this edifice to be older than Devonian (Early Palaeozoic) but this is based on the same inferred chronostratigraphy for the PCB that leads to ambiguity in dating tectonic events in general. Indeed, the sediments overlying the Astrakhan body, assumed to be Ordovician and Silurian by Kostyuchenko et al. (2004), and therefore constraining the age of the volcanic edifice to Ordovician or older, must in any case be considered Devonian in the context of the Zonenshain et al. (1990) model of continental break-up. Thus, there is potentially a large volcanic massif associated with Late Palaeozoic rifting near the junction of the southeastern extremity of the DDB and the southern margin of the main PCB rift axis, in an area close to where crustal thinning beneath the DDB was extreme, as argued above. The crystalline crust beneath the centre of the PCB, based on the profile shown in Figure 7, is also rather thin (if the so-called 8.1 km s-1 eclogite layer is simply taken as riftzone upper mantle), although still probably too thick (and high velocity) for oceanic crust. More modern, robust crustal seismic data need to be collected to provide better constraints. Figure 1 lb shows the postulated triple junction involving the DDB and PCB rift axes with 120 ~ separation superimposed on a map of the crystalline crustal thickness of the North Caucasus area (seismically defined sedimentary layer removed) constructed by Kostyuchenko et al. (2004). It is interesting to note that the (south-trending) third arm corresponds to a distinct change of crustal structure (associated with the Stavropol High) north of the Caucasus and to a change of trend of the Greater Caucasus orogenic belt itself. Kostyuchenko et al. (2004) recognized this pattern and suggested that there were indeed two distinct crustal affinities sutured in the North Caucasus Scythian Platform.

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Fig. 11. (a) Residual gravity of DDB-PCB area (gravity effect of sedimentary layers removed; from Yegorova et al. 2004) with hypothesized triple-junction rift axes (see text for discussion). The location of the DOBREfraction profile, crossing the DF, is shown for reference. (b) Thickness of the crystalline crustal layer (thickness of overlying sedimentary layer subtracted from Moho depth; from Kostyuchenko et al. 2004) with rift axes from (a) superimposed.

Although they preferred a model in which this suture was older than the Late Palaeozoic, the age of the sediments overlying crystalline basement in this area show only that the suturing has to be older than the Mesozoic. Further discussion of the (palaeotectonic) implications of a Late Palaeozoic plate boundary in this area is beyond the scope of this paper (see Saintot et al. 2006, for additional information) except to say that it would be not be inconsistent with widespread rifting and related ocean basin development on the southern margin of Europe at this time that could be generally associated with the development of the Palaeotethys Ocean (and possibly the intra-Variscan Rheno-Hercynian Ocean; see Franke 2006).

Basin setting: controls on rifting a n d s u b s e q u e n t inversion

Of the major Late Palaeozoic rift basin systems of Precambrian eastern Europe, the DDB is clearly intracratonic, with rift

structures cutting obliquely across the Archaean-Palaeoproterozoic fabric of the underlying Sarmatian segment of the EEC. In contrast, the E B B - T C B and the PCB possess the important common characteristic of having (apparently) formed on pericratonic crust accreted to the EEC in Neoproterozoic-earliest Palaeozoic times (e.g. Timanide-Baikalian orogenies). The lithosphere underlying these basins, on the margins of the craton, appears to have been weak, and, during the Palaeozoic, was highly susceptible to tiffing (Fig. 10; see also Saintot et al. 2006; Sliaupa et al. 2006). Subsidence and sedimentary basin development appear to have occurred on this lithosphere soon after the time of its accretion, from at least the Early Ordovician, in what Saintot et al. (2006) referred to as the 'renewed passive margin' of the EEC. The processes responsible for Neoproterozoic accretion apparently led to the rapid collapse of orogenic topography and to moderately thin crust that subsided throughout most of the Palaeozoic. Thus, by the time that the Devonian and younger extensional tectonic regime established itself throughout the EEC, driven (apparently) by some sort of thermal instability

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affecting much of the eastern (European) part of the Laurussian plate (as argued above; see also Saintot et al. 2006), much of the concomitant rifting favoured the weak Neoproterozoic subsided margins, leading to the development of the Late Palaeozoic E B B - T P B and PCB rift systems. That a common process was involved for both the pericratonic PCB as well as the intracratonic DDB is strongly suggested by their proximity and similar style at this time (whether or not a model of continental break-up and ocean basin development is accepted). The link between the DDB and the E B B - T P B is via the Late Palaeozoic magmatism that characterizes both and, indirectly, via the PCB, which has similar crustal affinity to the E B B - T C B . Much of the 'renewed passive margin' of the EEC became involved in the Uralian orogeny not long after the major Late Devonian-Early Carboniferous rifting event (see Brown et aL 2006). In this regard, both the E B B - T P B and PCB probably represent the more pericratonic parts of the passive margin basin system, or perhaps have been preserved as such simply because of the geometry of the Neoproterozoic orogenic belt vis-h-vis the subsequent Uralide belt. Their proximity to the Urals, and their fundamentally different crustal setting from the intracratonic DDB, explains how their postrift tectonics (foreland basin development and compressional deformation) have been dominated by Uralian events; in contrast, the DDB, with its closer proximity to the southern margin of the EEC, has been more affected by post-Palaeozoic Cimmerian and Alpine events occurring in the adjacent Tethyan belt.

Summary and conclusions The ( M i d - ) Late Devonian to Early Carboniferous was a time of widespread rifting on the EEC and its margins in general. Some degree of diachronism within and between individual rift systems is evident. Early Carboniferous (especially Visran) extensional activity can be considered as a rejuvenation phase, after a short period of tectonic quiescence (as clearly demonstrated in the DF and DDB), that contributed to the formation of very deep basins with more than 20 km of sediments such as in the D D B - D F and PCB. Magmatism accompanied the earliest synrift stages of these basins and relationships between rifting, basin architecture, and emplacement of igneous rocks indicate an 'active' rifting process. The widespread nature of these 'active' phenomena suggests that there was not a single, isolated, plume head but rather a cluster of thermal instabilities (or generalized thermal instability) at the base of the lithosphere beneath widely separated parts of the EEC. The Devonian-Carboniferous rift basins can be classified as intracratonic or pericratonic according to the nature of their crystalline basement. The ( P r i p y a t - ) D D B - D F system is clearly intracratonic, cutting across the apparent structural grain of the Archaean-Palaeoproterozoic basement of the EEC. The PCB and E B B - T P B rift systems are mostly pericratonic, developed on reworked and, at least partly, juvenile crystalline basement accreted to the EEC during the Neoproterozoic. The pericratonic PCB is closely related to the intracratonic D D B - D F , and the Kontozero Graben on the Kola Peninsula is an intracratonic rift segment associated with the pericratonic E E B - T C B rift system. A number of observations and implicit arguments suggest that the DDB opened into a deep basin, possibly having oceanic lithospheric affinity, to the SE, in the area where it adjoins the southern PCB. As such, the Late Devonian rift axes of the DDB and PCB may comprise two arms of a r i f t - r i f t - r i f t triple junction that was possibly the locus of continental break-up at this time. Post-rift compressional tectonic reactivations ('basin inversion') are related to Uralian events in the PCB and TPB, which lie close to the Urals margin of the EEC, and to (Mesozoic) Tethyan events (Cimmerian and Alpine orogenies) in the D D B - D F .

The authors thank EUROPROBE for bringing them together and providing a stimulating framework in which to learn from each other. The senior co-author thanks other contributors to the European Lithosphere Dynamics memoir for their patience, and D. Gee for his vision, his warmth, and his support. P. Ziegler provided much support, friendship, and good humour during our initial work on the DDB.

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The evolution of the southern margin of Eastern Europe (Eastern European and Scythian platforms) from the latest Precambrian-Early Palaeozoic to the Early Cretaceous A L I N E S A I N T O T 1'2, R A N D E L L A. S T E P H E N S O N 1, SERGIY S T O V B A 3, M A R I E - F R A N ~ O I S E B R U N E T 4, T A M A R A Y E G O R O V A 5 & V I T A L Y S T A R O S T E N K O 5

1Netherlands Centre for Integrated Solid Earth Sciences, Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands 2present address: NGU, Leiv Eirikssons vei 39, N-7491 Trondheim, Norway (e-mail: [email protected]) 3Naukanaftogaz, Uritckogo 45, Kyiv 03055 Ukraine 4UMR 7072 Tectonique CNRS-UPMC, case 129, University Pierre et Marie Curie, 4 place Jussieu, 75252 Paris cedex 05, France 5Institute of Geophysics, National Academy of Sciences of Ukraine, Palladin av., 32, 252680 Kyiv-142 Ukraine

Abstract: The southern part of the Eastern European continental landmass consists mainly of a thick platform of Vendian and younger

sediments overlying Precambrian basement, referred to as the East European and Scythian platforms (EEP and SP). Some specificgeological features, such as the Late Devonian Pripyat-Dniepr-Donets rift basin, the Karpinsky Swell, the Permo(?)-Triassictroughs of the SP, and the deformed belt running from Dobrogea to Crimea and the Greater Caucasus, in which rocks as old as Palaeozoic crop out, form a record of the geodynamic processes affecting this part of the European lithosphere. Hard constraints on the Palaeozoic history of the SP are very sparse. The conventional view has been that the SP is a Late Palaeozoic orogenic belt. However, it is shown that the few available data are also consistent with an alternative interpretation in which it is the thinned margin of the Precambrian continent, reworked by Late Palaeozoic-Early Mesozoic rifting events. The geodynamic setting of the margin is classically reported as one of active convergence throughout the Late Palaeozoic and Early Mesozoic, with subduction of the Palaeotethys Ocean beneath Europe. Actually, there are no direct observations constraining the polarity of Palaeotethys subduction in this area although indirect evidence is not inconsistent with the conventional model. In such a case, the sedimentary-tectonic record of the SP suggests that convergence during the Permo-Triassic(?) and certainly during the Early and Mid-Jurassic was oblique. An Eo-Cimmerian (Late Triassic-Early Jurassic) event is widespread and implies a tectonic compressional regime with systematic inversion of most sedimentary basins. There is also a widespread unconformity at the end of the Mid-Jurassic and in the Late Jurassic. These can be interpreted as indicators of compressional tectonics; however, nowhere is there evidence of intense shortening or other orogenic processes. A revised tectonic model is proposed for the area but, given the degree of uncertainty characterizing the geology of this area, it is best considered as a basis for further discussion.

The southern part of the Eastern European continental landmass consists mainly of a thick platform of Vendian and younger sediments overlying Precambrian basement. In part, this is referred to as the East European Platform (EEP) where the sedimentary successions are overlying crystalline crust of the East European Craton (EEC). Fringing the EEP to the south, in southwestern Ukraine, Crimea, and in the North Caucasus, is a related physiographic platform called the Scythian Platform (SP). The sedimentary successions overlying the SP are generally thicker than those of the EEP and contain more post-Palaeozoic units. For this reason the crust underlying the SP has traditionally been thought to be younger than that of the EEP. Understanding the relationship between the crusts of these two platforms is a major focus of this paper. It is obviously necessary to know the tectonic setting of the southern margin of the EEC throughout geological history to understand what processes (tectonic and magmatic styles; accretionary processes) have been responsible for forming the continental lithosphere of the southern EEP and SP. It is generally considered that an active convergent system (Pacific-type a n d / o r Andean-type) lay to the south of the EEP from Late Palaeozoic to Neogene times (Zonenshain et al. 1990; Dercourt et al. 1993, 2000; Nikishin et al. 1998b, 2001). However, there are radical differences of opinion about the processes responsible for the tectonic development of the area, leading to two end-member conceptual models. One of these envisages continuous accretion of terranes accompanied by closure of back-arc and marginal seas a n d / o r wider oceans (i.e. Rheic, Tethys or Palaeotethys-Neotethys

oceans) and successive southward jumping of subduction zones. In this case, crustal fragments (or domains) are sutured and these suture zones remain essentially stable thereafter (e.g. Dercourt et al. 1993, 2000; Stampfli & Borel 2002). The other model envisages the repeated opening and closure of back-arc basin(s) behind a long-lived subduction zone, with partial or total inversion of related sedimentary basins. In this case, suture zones between continental domains are unstable, being reopened and reclosed episodically, forming a series of (sub-)oceanic basins that subsequently were inverted (Nikishin et al. 1998b, 2001). In the framework of international programmes (e.g. INTAS, EUROPROBE, DOBRE, Tethys and Peri-Tethys), important new studies on the geological evolution of southeastern Europe have been accomplished (e.g. Aubouin et al. 1986; Dercourt et al. 1993, 2000; Stephenson et al. 1996, 1999; Perez-Estaun et al. 1997; Crasquin & Barrier 1998, 2000; Ziegler et al. 2001; Bergerat et al. 2002; Brown et al. 2002; Brunet & Cloetingh 2003; Stephenson 2004). Most of these have dealt with specific tectonic targets within the EEP and SP and only very few (e.g. Nikishin et al. 1996, 1998a,b, 2001) have attempted to present a synthesis of the whole East European plate and its complete evolution through geological time. These papers have presented a series of palaeotectonic-palaeogeographical models that have been extremely useful in providing a general outline of the geological evolution of this area. However, given an acknowledged general absence of hard observational data and modern analytical results, it is difficult to assess how robust certain features of these models are. In this paper, an attempt is

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European LithosphereDynamics.

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made to document what is known with some certainty and, more importantly, what is known to be uncertain and, on this basis, propose a revised schematic model for the Palaeozoic to Mid-Mesozoic evolution of this area as a framework for further study. The key tectonic units of the southern part of Eastern Europe that are described and utilized to constrain the geological evolution of the southern EEP and SP include the Late Devonian Dniepr-Donets rift basin (DDB)-Donbas Foldbelt (DF), the Karpinsky Swell (KS), Peri-Caspian Basin (PCB), the Permo(?)Triassic troughs of the SP; the deformed belt running from North Dobrogea (ND) to Crimea and the Greater Caucasus (GC), and, to a lesser extent, the Pontides in northern Turkey and the Transcaucasus (TC) domain south of the GC (Fig. 1). These key tectonic units are described in turn, for various time slices beginning in the Neoproterozoic, in the framework of the following kinds of questions. (1) What are the features related to rift- and subduction-related tectonics, as well as to other convergent and divergent plate movements? (2) Where are the zones of suturing and basin closures and/or inversions, and how did they reopen and when? (3) What is/are the process(es) governing the observed changes of tectonic regime (extension v. compression) through time behind the convergent margin? For each time slice, the order in which the various geological units are discussed depends on their relative tectonic importance at that time and a judgement on the reliability of the available data for each.

The structural units and the present-day state of the lithosphere of the southern margin of Eastern Europe The southern margin of the EEC is poorly investigated compared with its western margin along the Trans-European Suture Zone (TESZ; see Pharaoh et al. 2006). There are very few available subsurface data characterizing the Palaeozoic and Mesozoic successions in the study area. Palaeozoic and Mesozoic outcrops do occur in the DF, which is an inverted part of the Late Palaeozoic intracratonic DDB rift, lying within the EEP, and to the south along the ND-Crimea-GC(-Pontides) deformed belt. Numerous drill-holes of various depths have reached Mesozoic, and occasionally Palaeozoic, units in Crimea, the KS, the SP, and the Fore-Caucasus region. The large-scale structures of southern Eastern Europe, presumed to be mostly Phanerozoic in age, buried under a thick pile of platform sediments, can be indirectly observed from geophysical data and several authors have published maps of the main crustal units underlying the area (e.g. Kostyuchenko et al. 2004; Yegorova et al. 2004). A number of recent papers have dealt with the state of the lithosphere of southeastern Europe (Yegorova & Starostenko, 2002a,b; Artemieva 2003; Kostyuchenko et al. 2004; Yegorova et al. 2004; Artemieva et al. 2006). The western and southern boundaries of the southern EEP lithosphere are evident in the mantle component residual gravity anomaly field (Yegorova & Starostenko 2002b). A strong positive anomaly runs along the southern margin of the SP, from the South Caspian Basin to and along the GC belt and along its western boundary, adjacent to the TESZ. The sources of these anomalies are considered to be high-density upper mantle domains marking the Alpine accretionary-collisional belt at the southern margin of the SP (Yegorova & Starostenko 2002a,b). The EEP and SP lithosphere is thicker than the Phanerozoic lithosphere of Western Europe (Yegorova et al. 2004). However, according to Artemieva & Mooney (2001), the 1300 ~ isotherm (base of the 'thermal' lithosphere) is shallower beneath the southern EEP and, in particular, the SP than under the Russian Platform (central EEC), being 1 2 0 - 1 4 0 k m and 180-210km, respectively (cf. Artemieva et al. 2006). The latter depth range corresponds to typical Early Proterozoic lithosphere whereas the

former is closer to that of Phanerozoic Western Europe. According to Artemieva (2003), Late Palaeozoic rifting within the EEC (e.g. DDB; see below and Stephenson et al. 2006) has contributed to significant rejuvenation of the entire lithospheric column in this area, possibly being a factor explaining the continuing subsidence of the southern part of the SP since the Late Palaeozoic. Like the lithosphere, the crust of the southern EEP and SP is also thicker than that of Phanerozoic Western Europe (Yegorova & Starostenko 2002a,b). In particular, the crust of the SP on the northern Crimean peninsula is also thicker, being 40-45 km thick (e.g. Stephenson et al. 2004), although it is thinner than that of the adjacent Ukrainian Shield (UkS; Fig. 1). A similar pattern is discernible for the SP of the North Caucasus area (Fig. 2), taking into account crustal thinning associated with Mesozoic extension evident in the area east of the Azov Sea.

Latest Precambrian to Early Palaeozoic evolution of the southern Eastern European Platform (EEP) and Scythian Platform (SP) The SP is a key zone in understanding the evolution of the southern margin of Eastern Europe. Because its basement is buried beneath between 5 and 12 km of Palaeozoic to Quaternary sedimentary cover, there is a significant lack of data on the nature, structure, and evolution of the SP (see Stephenson et al. 2004). It is a zone located between domains that are (relatively) better understood, the KS and the southern edge of the Ukrainian Shield to the north and the N D - C r i m e a - G C corridor to the south (Fig. 1), because they are better described by geophysical data or because rocks recording their geological history are exposed or lying at shallow depths. The SP is classically considered to be a wide Variscan belt between the EEC and the Alpine-Cimmerian folded belts on its southern border, referred to as the 'Scythian orogen' by Milanovsky (1987), Zonenshain et al. (1990), and others. Active orogenesis was supposed to have occurred from Early Carboniferous to Permian times (Nikishin et al. 1996, 2001; Stampfli & Borel 2002), with a platform stage of development in the Mesozoic (Muratov 1979). As such, it has usually been considered to be the link between the Variscan orogenic system of western and central Europe and the Uralian belt at the eastern edge of the EEP. According to this model, the southernmost EEP is the passive margin of Baltica from the Cambrian until collision and docking of SP crust in the Late Palaeozoic. However, there exists no irrevocable evidence that the basement crust of the SP comprises a Late Palaeozoic (Variscan) accretionary orogenic belt (Stephenson et al. 2004). Indeed, if the basement of the SP was consolidated during the Late Palaeozoic, it and overlying preEarly Carboniferous sedimentary strata should display intense and penetrative deformation (and concomitant metamorphism) of this age but there is no observed evidence for this at all. What is reported as Variscan metamorphism in the SP is weak, not higher than greenschist facies, with sedimentary layering intact (Khain 1977; Adamia et al. 1981). It is most easily interpreted in terms of simple sedimentary burial. Furthermore, as mentioned above, the geophysical structure of SP lithosphere and crust has greater affinity with the EEC than with the Variscan belt of Western Europe. Accordingly, the nature and evolution of the SP needs to be placed in the context of latest Precambrian-Early Palaeozoic (preVariscan) tectonic events affecting the EEC. Indeed, many previous authors have considered the SP as a passive margin of the EEC reworked by Late Proterozoic and younger tectonism (see Kruglov & Cypko 1988; Gerasimov 1994; Yudin 1995; Milanovsky 1996; Shnyukov et al. 1997; Stephenson 2004; Stephenson et al. 2004).

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Fig. 1. Southeastern Europe and adjacent areas showing the locations of tectonic units mentioned in the text. AM, Azov Massif; AD, Astrakhan Dome; Ca, Carpathian Belt; Cr, Crimean Mountains; DDB, Dniepr-Donets Basin; DF, Donbas Foldbelt; Dz, Dzirnla Massif; GC, Greater Caucasus; IZ, Istanbul Zone; KB, Kayasula Basin; KS, Karpinsky Swell; LC, Lesser Caucasus; Mg, Mangyshlak; Mo, Moesia; MB, Mozdok Basin; MT, Manysh Trough; MV, Mineralnie Vody Dome; NCA, North Crimea-Azov Basin; ND, North Dobrogea; NF, Novo-Fedorovsk Basin; NP, Nakhichevan Platform; NU, Novoselskoe Uplift; PCB, Peri-Caspian Basin; PT, Pripyat Trough; RH, Rostov High; SCB, South Caspian Basin; SH, Stavropol High; SNU, Simferopol and Novotsaritsyn Uplifts; SZ, Sakarya Zone; TC, Transcaucasus; UkS, Ukrainian Shield; VM, Voronezh Massif. The background is a digital topographic map from a global topography 2 min database, illumination from N135 and N270. The dashed line is the boundary of the Scythian Platform with the East European Craton as mapped by Nikishin et al. (1998b).

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A. SAINTOTETAL. references therein). Moreover, plate-tectonic reconstructions (based on palaeomagnetic data of Torsvik & Rehnstr6m 2001) show that the Baikalian deformation of the Baltican margin did not involve collision between Baltica and any other large continent (such as Siberia), implying that a wide oceanic system was still open to the east (present-day position) when the accretionary Baikalian belt formed. Baikalian deformation is observed along the present-day eastern margin of Baltica with the Vendian Proto-Urals (described in more detail by Puchkov 1997) and along the Timan-Pechora Belt NE of Baltica (Olovyanishnikov et al. 1997; Roberts 2000). It signifies amalgamation or accretion of (subduction-related) arc or oceanic complexes and microcontinental blocks on the margin of Baltica (exemplified by the Beloretzk terrane accreted along the Southern Proto-Urals, Glasmacher et al. 1999; or by the Pechora-Barentsia terranes to form the Timan-Pechora foldbelt, Nikishin et al. 1996; Torsvik & Rehnstr6m 2001). It follows that it is difficult to distinguish clearly between Baikalian and Cadomian-derived terranes. N e o p r o t e r o z o i c geology o f the Scythian Plate (SP) and the N o r t h D o b r o g e a ( N D ) - C r i m e a - G r e a t e r

Fig. 2. (a) Crystalline crustal thickness and (b) Moho depth of the southeastern EEP-SP-North Caucasus area (in kilometres; from Kostyuchenko et al. 2004). AM, Azov Massif; AD, Astrakhan Dome; DF, Donbas Foldbelt; PCB, Peri-Caspian Basin; RH, Rostov High; SH, Stavropol High.

N e o p r o t e r o z o i c orogens: B a i k a l i a n v. C a d o m i a n

Three terms are used to define roughly contemporaneous accretionary orogens along continental margins during Neoproterozoic time: Baikalian, Pan-African and Cadomian (see Ziegler 1990). The first is related to processes at the edge of Baltica (EEC), and the others are confined to the Gondwanan margin. It should be noted that competing Neoproterozoic palaeotectonic reconstructions, in this context, are irrelevant. The Cadomian terranes are remnants of a Late Proterozoic subduction-related belt on a Gondwana active margin (Fig. 3) that subsequently migrated (in the Palaeozoic) to be accreted onto the margins of Baltica (and Laurentia), with Avalonia being an example (see Dalziel 1997; Scarrow et al. 2001; Winchester et al. 2006). In the context of the Neoproterozoic reconstruction of Dalziel (1997) shown in Figure 3, there is very little tectonic distinction between what is a Cadomian terrane (adjacent to what becomes Gondwana) and a Baikalian terrane (adjacent to what becomes Baltica). The Baikalian orogenic system, along the (present-day) eastern margin of Baltica, is simply a prolongation of the Cadomian arc (~eng6r et al. 1993; Torsvik et al. 1996; Dalziel 1997; Scarrow et al. 2001; Torsvik & Rehnstr6m 2001; and

C a u c a s u s (GC)

The geological unit north of the ND (the Pre-Dobrogean Depression; Fig. 4) is generally referred to as part of the SP (e.g. Sandulescu 1990); however, its basement could be as old as Archaean (Patrulius & Iordan 1974). Granites, diorites and gabbros are reported that apparently have yielded Late Riphean and Vendian K - A r ages (790 Ma and 640-620 Ma, respectively; Belov et al. 1987), perhaps part of a widespread magmatic event throughout the EEC around this time related to Riphean intracratonic rifting (Bogdanova et al. 1996; Sliaupa et al. 2006). The crystalline basement is overlain by what appears to be passive margin Vendian and younger sediments, thickening to the SW. Obviously, what is called the SP north of the ND is a remnant the latest Precambrian-Early Palaeozoic passive margin of the EEC. The palaeography and crustal affinity of the Moesian Platform, lying south of the ND (Fig. 4), is a matter of much debate (e.g. Pharaoh et al. 2006; Winchester et al. 2006). Its northern part (referred to as Central Dobrogea; Fig. 4) and the ND both display Late Precambrian greenschist basement (Seghedi 1998, and references therein). In contrast, Moesian basement cropping out in the Carpathian belt, along strike with the southern Dobrogean segment of the Moesian Platform (location shown in Fig. 4), is derived from the Cadomian arc (e.g. developed along the Gondwanan margin), with a basement of higher grade metamorphism (e.g. Haydoutov & Yanev 1997; Seghedi 1998, and references therein; Crowley et al. 2000; Seghedi et al. 2000, 2003a). It follows that the northern part of Moesia (Central and Northern Dobrogea) probably formed a part of the Late Proterozoic Baltican passive margin, whereas the southern part of Moesia consists of a subsequently accreted Gondwana-derived terrane (cf. Winchester et al. 2006). As such, the CapidavaOvidiu Fault Zone (Fig. 4) would be the corresponding suture, as supported by recent seismic refraction data (Hauser et al. 2001). The Istanbul Zone of the Western Pontides in Turkey (IZ, Fig. 1) is classically associated with Moesia (Okay et al. 1994, 2006; Usta6mer 1999; Chen et al. 2002) as a Cadomian entity later accreted to Europe (see Winchester et al. 2006). For the reasons stated above, perhaps this should refer only to the southern segment of Moesia, south of the Capidava-Ovidiu Fault Zone (or thereabouts) and not to all of Moesia nor including the whole of the SP as suggested by Okay et al. (2006). The basement of the Crimean SP is reported to be primarily Baikalian (Khain 1994). It was reached by drill-holes in Crimea (Simferopol and Novotsaritsyn uplifts; Fig. 1) where it comprises Late Proterozoic greenschists (Muratov 1969; Muratov et al. 1984; Khain 1985), resembling those found north

PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE

485

Fig. 3. Plate-tectonic reconstruction for the Late Precambrian, 545 Ma (from Scarrow et al. 2001, after Dalziel 1997), showing the prolongation of the Cadomian arc to the present-day eastern margin of Baltica. EA, East Avalonia; WA, West Avalonia; S, Siberia (for additional details see Scarrow et al. 2001).

of ND (Khain 1985). According to Chekunov (1994) and Milanovsky (1996), these 'North Crimean' Precambrian highs developed during Riphean rifting on this part of the SP, thus implying that crustal consolidation occurred as early as the Mesoproterozoic. In this case, the observed greenschist metamorphism would be of Baikalian age (Milanovsky 1996). Ophiolites have been reported from the same area of Crimea (Gerasimov 1994).

Fig. 4. Tectonic zonation and the main faults in Dobrogea and northern Moesia, NW of the Black Sea (mainly Romania). #, location of inferred Moesian basement (Crowley et al. 2000; Seghedi et al. 2000, 2003a), as discussed in the text.

However, this is based upon nothing more than a drillcore reaching talc-bearing shales and serpentinites, thought to be of Late Precambrian or Early Palaeozoic age, and not actually on the observation of an ophiolite-type section. The protolith of these serpentinites is clearly a mafic to ultramafic magmatic rock, but could be related to Riphean (or other) rifting events (as suggested for the Pre-Dobrogean basement). There are no potential field data or other data to support the existence of an ophiolitic suture zone. The SP basement of the central north GC is uplifted as the Stavropol High (SH, Fig. 1) and exposed on the Mineralnie Vody Dome (MV, Fig. 1) as well as further to the south in the Main Range of the GC. The basement is heterogeneous in terms of crustal thickness and Moho depth, with the Stavropol High having a crustal thickness of 40 km and thin sedimentary cover (Fig. 2a and b). Precambrian and Early Cambrian metamorphism and magmatism is observed (Letavin 1980; Milanovsky 1987) and was considered to belong to the Baikalian orogeny by Khain (1975, 1994). In the northern GC belt itself, isotopic dating ( R b - S r ) gives a Neoproterozoic age (790Ma) for metamorphism (Belov 1981); Khain & Leonov (1998) reported 6 0 0 - 7 0 0 Ma mantle or subduction-type granitoids. The correlation of this crust to the basement of the Stavropol High is made using potential field data (in particular, magnetic anomalies; see Kostyuchenko e t al. 2004). Some new geochemical data suggest that the lithospheric mantle beneath the Kayasula Basin and Manysh Trough (KB and MT, Fig. 1; south of the KS, in the northern part of the SP) is similar to the lithospheric mantle of the DF (southeastern DDB) since a Late Proterozoic event at 650 + 50 Ma. This age comes from Nd ages on Late Devonian alkaline basalts from the DF as well

486

A. SAINTOT ETAL.

as from Triassic calc-alkaline basalts from the Kayasula Basin, and is interpreted to correspond to the age of youngest metasomatism of mantle sources during Baikalian subduction (F. Chalot-Prat, pers. comm.). Little can be said about the crustal affinity of the PCB basement. It is generally considered to be Precambrian, although Zonenshain et al. (1990) hypothesized that it comprised Devonian oceanic crust (cf. Stephenson et al. 2006). The basement of the Astrakhan Dome, a structural high in the southern PCB (Fig. 1), is consistently interpreted as Precambrian in age from numerous integrated seismic and potential field studies (e.g. Kostyuchenko et al. 2004; Yegorova et al. 2004) although little can be said on whether it is Neoproterozoic or older. It can also be mentioned that the generalized velocity structure of the crust in the eastern S P - K S southern PCB corridor (cf. Baranova & Pavlenkova 2003; Kostyuchenko et al. 2004) appears, on the basis of the synthesis by Artemieva (2001), to have somewhat more affinity with Precambrian crust than with Phanerozoic crust. A 10 km thick anomalously high-velocity lower crustal body, well known from regional seismic studies, lies beneath the PCB (e.g. Brunet et al. 1999; Volozh et al. 2003, and references therein). Volozh et al. (2003), having rejected earlier models in which it was related to rifting processes (Artyushkov 1993; Lobkovsky et al. 1993, 1996a,b), have interpreted it as an eclogitic lens emplaced during Baikalian Proto-Urals subduction or orogenesis, thus implying that the basement of the PCB is Neoproterozoic. The basement of the TC, cropping out in the Dzirula Massif (Dz, Fig. 1) and lying south of the GC belt, is characterized by highgrade metamorphism and significant magmatism, and was referred to by Zakariadze et al. (2001) as a Neoproterozoic arc accretion complex. Zakariadze et al. (1998) reported S m - N d ages ranging from 810 + 100 Ma to 657 + 78 Ma. It is either a Gondwana-or Cadomian-derived terrane, as shown in recent plate reconstructions (see Golonka 2000; Stampfli & Borel 2002), or a Baikalian terrane as described in the classical Russian literature (Khain 1985, 1994; Milanovsky 1987, 1996). Thus, its precise Late Precambrian affinity is at best speculative. The similarities between the Late Proterozoic of the GC (described above) and the TC imply that the latter could be Baikalian (part of Baltica in the Neoproterozoic). Alternatively, as a Cadomian terrane, it could be the eastern prolongation of a Moesia-Istanbul Zone lithospheric block, implying the presence of a Palaeozoic suture zone somewhere between the GC and the TC.

The r e n e w e d p a s s i v e m a r g i n o f the E a s t E u r o p e a n Craton ( E E C ) in the Early P a l a e o z o i c

On the western margin of Baltica, from Late Ordovician to Silurian times, the Iapetus and Tornquist oceans closed and Laurentia-Baltica, along with various Cadomian terranes including Avalonia, were amalgamated (see Cocks & Torsvik 2002, 2006; Fortey & Cocks 2003, and references therein). By this time, the Neoproterozoic Baikalian orogenic belt fringing eastern Baltica subsided and developed into a passive margin (Fig. 5). It is not clear whether this involved the rifting away of a continental mass and the formation of the Palaeo-Urals Ocean at this time (Zonenshain et al. 1990; Maslov et al. 1997) or whether this simply involved a post-orogenic collapse of the Baikalian belt adjacent to an already existing Palaeo-Urals Ocean (Matte et al. 1993; Glasmacher et al. 1999). In any case, subsidence is recorded on the margins of the PCB from earliest Ordovician times (Volozh et al. 2003) and, given that the Neoproterozoic orogenic belt probably extended across much of the southern margin of the EEC (as discussed above), it seems likely that this was also a passive margin at this time. This would correspond to much of what is now the eastern (Fore-Caucasus) SP.

Fig. 5. Schematicillustration of the Ordoviciantectonic setting discussed in the text: the renewed passive margin of the EEC on an extinct fringing Baikalian orogenic belt and active rifting on the southern margin of the subsided ScythianPlatform. Abbreviationsfor structural units are as in Figures 1 and 2. Indeed, the velocity structure of the SP south of the KS is not incompatible with that of a thinned continental margin (e.g. Baranova & Pavlenkova 2003; Yegorova et al. 2004). The boundary between the SP and the EEC is poorly known in terms of its geometry as well as in terms of its tectonic origins (see Stephenson et al. 2004). Traditionally, it has been defined on the basis of potential field data north of Crimea along with sparse data from a few wells that reach basement in the area. Various authors have considered the boundary to be the remnant of a Late Palaeozoic north-dipping S P - E E C suture (Kruglov & Cypko 1988), a south-dipping suture (Gerasimov 1994; Yudin 1995; Nikishin et al. 1996, 2001), or a set of normal faults or a single steep and deep normal fault with 3 - 5 km of displacement as young as Late Triassic, overprinting the pre-existing collisional suture (Kruglov & Cypko 1988). However, as discussed above, it seems unlikely that there is a Late Palaeozoic suture in this area. The SP basement contains metamorphic rocks, mainly MidOrdovician to Early Devonian greenschists (470-410 Ma; Belov 1981), which, in this context, are interpreted to represent the deeply buried sediments of renewed passive margin development in the Early Palaeozoic. The Cadomian basement of the Istanbul Zone (IZ, Fig. 1; Western Pontides) is overlain by an Early Ordovician (starting with continental basement-derived clastic rocks; Dean et al. 1997, 2000), to Devonian fossiliferous sedimentary succession (the 'Istanbul Palaeozoic' of ~engtr et al. 1980, 1984; ~engtr 1984; G6rtir et al. 1997). These deposits were initially assumed to be the record of an Atlantic-type passive margin of Eurasia (~engtr & Yalmaz 1981; Robertson & Dixon 1984; ~engtr et al. 1984; Ustatmer & Robertson 1993; Okay et al. 1996). In contradiction to this, the Early Ordovician (Arenig) endemic fauna content of the Istanbul Zone apparently shows an Avalonian affinity (Dean et al. 2000) that places the Istanbul Zone on the other margin of the Early Palaeozoic Ocean (Cocks & Torsvik 2002). The basement of the Macin Zone (northern Moesia in ND, Fig. 4) is overlain by Silurian-Early Devonian sediments also considered by some to have Gondwanan (Avalonian) affinity (Iordan 1988). Some of these issues have been discussed further by Winchester et al. (2006). In the central GC, where Precambrian and Palaeozoic rocks crop out, the Cambrian succession is missing or consists only of continental conglomerates. This was the time of erosion of the Baikalian orogenic belt (Khain 1975). Along the so-called Peredovoy or Fore Range of the GC (Fig. 6), highly metamorphozed Ordovician(?) mafic to ultramafic rocks are observed. Khain (1975) interpreted these as being the remnants of magmatism associated with a renewed stage of rifting in this area (Fig. 7a).

PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE

40o

44~

\ ..~~1

Scythian Platform High

9

48~ j (_. Caspian Sea

440

',er,

Black

Main thrusts of the present-day Caucasus belt

Leucasu~ I

I

400

Fig. 6. The Precambrian and Palaeozoic units of the Greater Caucasus area, compiled from Khain & Milanovsky (1963), Kotansky (1978), Dotduyev (1987), Philip et al. (1989), and Ruppel & McNutt (1990).

Late Devonian to mid-Carboniferous evolution of the southern East European Craton (EEC) and its margins North Dobrogea (ND) and Moesia

There are only a few reliable observations constraining the Late Devonian evolution of ND. The northernmost structure of ND (in the vicinity of Tulcea; Fig. 4) comprises a folded band of Ordovician to Devonian cherts and turbidites. This unit is considered to be part of an accretionary complex indicating the presence of a suture zone between Moesia and the SP (e.g. Seghedi 2001). Some I-type calc-alkaline granitoid plutons found in the ND basement are thought to be Devonian in age and are interpreted as being subduction related. The Palaeozoic history of ND is thought to have been one of an Early Palaeozoic passive margin evolving into a (?)Late Devonian-Early Carboniferous Andean-type system that continued subducting until collision in the Late Carboniferous-Early Permian. Following this, a Late Carboniferous-Early Permian thermal event determined in ND was interpreted by Crowley et al. (2000) as the signature of the collision between Moesia and the SP. However, the inferred accretionary prism contains rocks that are not younger than Devonian and there are no indications of a long-lived active margin throughout the whole of the Carboniferous, such as subduction-related magmatism. Indeed, as discussed above, the northern part of Moesia was probably not far from its present position vis-~t-vis the EEC from some time before the Devonian (Fig. 8). Winchester et al. (2006) have argued that Moesia was accreted to the EEC during the (eastern) Caledonian orogeny in the Early Palaeozoic (but considered Moesia as a single unit). There is simply no convincing record in ND of a complete orogenic cycle between Moesia and the EEC; that is, one that includes the development of an ocean basin and its passive margins, and subsequent closure by subduction and continental collision. It cannot be ruled out that ND was part of the widespread rift system that developed on the southern margin of the EEC during Devonian times. It can also be noted that the Devonian and younger sedimentary succession on Moesia, as for the Istanbul Zone, displays a European faunal affinity (e.g. Dean et al. 1997, 2000).

487

remnants of a Devonian oceanic basin) overthrust FamennianTournaisian limestones (Khain 1979). Rocks of late Visran age unconformably overlie the tightly folded Devonian and older Early Palaeozoic units of the Fore Range of the GC (Khain 1975, 1979; Zonenshain et al. 1990). Adamia et al. (1981), Belov (1981), Belov et al. (1990), Zonenshain et al. (1990) and Milanovsky (1991) all mentioned Visran folding and thrusting in the northern GC. This zone corresponds to a suture between the SP and a microcontinent called Maker (from the Makerska series of the Main Range of the GC; Fig. 9) as defined by Zonenshain et al. (1990). A metamorphic event prior to the Early Devonian (prior to 400 Ma), 400 Ma gneisses, and metamorphism of ophiolites at 4 6 0 - 3 7 0 Ma have been reported for Maker rocks by Somin (1998). Late Devonian granitoids (370 Ma) also occur (Belov 1981; Zonenshain et al. 1990). The Maker terrane was interpreted by Zonenshain et al. (1990) as a Gondwanan fragment, with the ophiolitic suture representing the closure of a wide ocean basin, that is, 'proto-Tethys' or a branch of the Iapetus-Tornquist oceanic system. However, there is no obvious record of long-lived subduction in the Silurian sedimentary succession on either side of the proposed suture. The Silurian sediments, which are phyllitic, could have been deposited on an Ordovician passive margin of the SP and GC (Fig. 7a). Furthermore, Adamia et al. (1981) reported petrological and geochemical analyses from the mafic and ultramafic rocks that suggest that they originated in a small oceanic back-arc basin. The enigmatic Dizi shale-slate (or flysch) series, 7 km thick, lies on the southern limb of the GC (Fig. 6). It contains Mid-Devonian to Triassic fauna (Khain 1975; Adamia et al. 1981; Kazmin 1991) and contains Devonian magmatic rocks (Khain 1994). According to Adamia et al. (1986) and Kazmin et al. (1986), the Dizi series corresponds to an accretionary prism. Nevertheless, according to the very scarce available data, there are neither high-pressure metamorphic rocks nor (ultra)mafic fragments such as are commonly incorporated in an accretionary complex. This series could represent therefore a deep marine rift(?) basin developed on a detached segment of the Baikalian margin of the EEC (GC-TC) and subsequently filled by turbidites series (Figs 7b and 8a,b; Dizi and TC positions A). The Dizi series does not display any Visran deformation (i.e. Fig. 7c) such as in the northern GC (Khain 1975; Zonenshain et al. 1990). Folding in the Dizi is Late Triassic or even Jurassic in age, related to Cimmerian events (Adamia & Lordkipanidze 1989; Kazmin 1991). As such, the Dizi series and its basement could be exotic (with respect to the GC during Visran time), being a 'protected' domain somewhere else on the SP margin (Fig. 8a; Dizi position B), corresponding to a subsiding basin on the margin of the SP (Fig. 8b) and later laterally displaced into its position south of the Main Range of the GC (discussed further below). In the above, only the interpretation of the TC as a Baikalian terrane (e.g. Khain 1994; Milanovsky 1996) has been considered (e.g. Figs 7 and 8a; TC position A). If the TC is a Cadomian terrane (as discussed above) its arrival adjacent to the SP would have been somewhat later but still prior to the Carboniferous. In this case, it would have affinity with Moesia and the Istanbul Zone (Fig. 8a and b; TC position B). Crimean Plain (Scythian Platform; SP)

Greater Caucasus (GC) and Transcaucasus (TC)

Oceanic crust was reportedly developed in the area of the presentday Fore Range of the GC in Devonian times (Khain 1975; Adamia et al. 1981; Figs 7b and 8a). Presumably related intrusions occur in the Main Range of the GC. Highly dismembered and metamorphosed mafic to ultramafic rocks (interpreted as ophiolitic

Drill-holes have reached low-grade metamorphic DevonianEarly Carboniferous deep-water shales and volcanic rocks of the subsequently uplifted Novoselskoe Uplift (Fig. 1; Letavin 1987) of the Crimean part of the SP (Fig. 1). This series has been interpreted as the remnant of a volcanic arc and associated accretionary complex, with the main orogenic event ascribed to the Visdan, to fit with what was described in the GC (Gerasimov 1994;

488

A. SAINTOT ETAL.

Fig. 7. Evolution of the Greater Caucasus (GC) belt without taking into account potential strike-slip movement along the belt (i.e. the Dizi basin is lying between the Main Range (Maker) of the GC and the Transcaucasus). (a) Ordovician: thinning of Baikalian belt; (b) Late Devonian: oceanic spreading in the Fore Range of the GC, development of the Dizi basin and widespread rift-related magmatism; (c) prior to Serpukhovian: subduction of Palaeotethys and the oceanic basin of the Fore Range is closed; (d) Late Carboniferous-Early Permian: dextral translation between the Gondwana and Eurasia (see text for further discussion); (e) Late Permian-Early Triassic: rift development in a right-lateral strike-slip setting of the margin. It should be noted that (as mentioned in the text) the Dizi series (b-d) does not show any record of tectonism from Devonian to Triassic times and that eclogites of the Fore Range-Peredovoy (c-e) formed at about 300 Ma, shortly before their exhumation. (a)-(c) are based mainly on Khain (1975) and Adamia et al. (1981).

N i k i s h i n et al. 2001). At best, this is very p o o r l y c o n s t r a i n e d ( S t e p h e n s o n et al. 2004). W h a t can be said with s o m e certainty is that there is no e v i d e n c e of Late P a l a e o z o i c h i g h - g r a d e metam o r p h i s m or p e n e t r a t i v e d e f o r m a t i o n as m i g h t be e x p e c t e d in an a c c r e t i o n a r y or o r o g e n i c complex. I n d e e d , d e e p - w a t e r s e d i m e n tation and v o l c a n i s m are also consistent w i t h a rift setting.

Pericratonic rifting (Dniepr-Donets Basin (DDB) and Peri-Caspian Basin (PCB)) T h e first i m p o r t a n t P h a n e r o z o i c tectonic events affecting the interior of the E E C o c c u r r e d in Late D e v o n i a n times with a

PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE

Fig. 8. (a) Schematic representation of the Late Devonian rifting process on the southern edge of the EEC with alternative positions of Transcaucasus (TC-A, as part of Moesia-Istanbul Zone, or TC-B, not) and the Dizi series (Dizi-A or Dizi-B). (b) Cross-sections (location shown in (a)), with positions A for the Dizi series and the TC on I and position B for the TC on II. The downward arrows on cross-section II indicate inferred subsidence rates along the passive margin. Extensional trends in (a) are the minimum stress axes o3 determined by Saintot et al. (2003a). Abbreviations for other structural units are as in Figure 1.

widespread tiffing event (Stephenson et al. 2006). The main tectonic feature formed at this time being the W N W - E S E trending intracratonic Pripyat T r o u g h - D D B rift system and its present-day inverted segment, the DF (Fig. 1). Numerous new geophysical and geological studies were carried out in the framework of the EUROPROBE 'Georift' project (e.g. Stephenson et al. 1993, 1996, 1999; Stephenson 2004, and references therein; McCann et al. 2003; Saintot et al. 2003a,b; Alexandre et al. 2004). The evolution of the DDB and the PCB is clearly linked (see Yegorova et al. 2004; Stephenson et al. 2006), although there is no explicit observable evidence of Late Devonian rifting in the PCB, except for a rapid increase of subsidence rate at that time (Brunet et al. 1999). From the Famennian until the Early Carboniferous, some of this increase of subsidence in the PCB may be ascribed to a sedimentary influx from the east, where the southern Urals and its Emba branch may have been emerging (Volozh et al. 1991). The DDB and DF were (extensionally) reactivated in the late early Visran (Khain 1985; Stovba & Stephenson 1999; McCann et al. 2003; Saintot et al. 2003a) with associated extrusive rocks, as well as during the Serpukhovian (Stovba & Stephenson 1999) and in the Early Permian (Stovba & Stephenson 1999; Saintot et al. 2003a,b). Some authors (e.g. Volozh et al. 1999) have regarded the DDB as part of a very long, singular intracontinental rift system, including the KS, that crosses the Caspian Sea to the Mangyshlak,

489

Fig. 9. Schematic representation of the evolution of the southern EEC margin from (a) Late Carboniferous subduction normal to the margin to (b) Late Carboniferous-Early Permian transcurrent or strike-slip margin without significant subduction and subsequent development of wrenchingrifting and uplift in Permo-Triassic times. Abbreviations for structural units are as in Figure 1.

north of the TP (Mg, Fig. 1). Late Devonian rifting of the KS is not documented by normal faulting and associated sedimentary architecture (prerift, synrift and post-rift layers) and magmatism, such as in the DDB and DF. A recent reinterpretation of a DSS profile between the DF and the KS was made by Baranova & Yegorova (2004) and Yegorova et al. (2004), and it is clear that at least the western KS is a continuation of the intracratonic DDB rift system.

Late D e v o n i a n to m i d - C a r b o n i f e r o u s g e o d y n a m i c setting

The observations and, in some cases, interpretations presented and briefly discussed above regarding the Late Devonian to midCarboniferous evolution of the southern EEC and its margins are schematically summarized in Figure 8. This was a time of widespread intra- and pericratonic rifting of the EEC itself, recorded most prominently in the DDB, the Timan-Pechora and Eastern Barent Sea basins, and probably in the PCB (see Stephenson et al. 2006). It seems possible, based on the observations described above, that this widespread event also affected areas on the subsequently tectonically overprinted southern margin of the EEC such as Crimea, the GC area and perhaps also the ND area. Artemieva (2003; see Artemieva et al. 2006)

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A. SAINTOTETAL.

also pointed out that the entire southern part of the Eastern European (mantle) lithosphere appears to have been rejuvenated by a process such as rifting. The driving mechanism of Late Palaeozoic rifting on the EEC and, by implication, throughout its southern marginal area, as proposed here, has been discussed in the companion paper by Stephenson et al. (2006). A number of observations and modelling inferences, including geochemistry of magmatic rocks (Wilson & Lyaskevich 1996), subsidence modelling (e.g. Kusznir et al. 1996a,b; van Wees et al. 1996; Starostenko et al. 1999), and structural relationships between sediments, faults and magmatic bodies in the DF (McCann et al. 2003), all favour an 'active' rifting process. This implies that mantle processes such as the impingement of a mantle plume and, therefore, excess temperatures at the base of the EEC lithosphere were involved in the initiation of rifting. Although the concept of a Late Devonian mantle plume was also proposed in the older literature (e.g. Chekunov 1994), many authors have also suggested that rifting in the EEC (namely, the DDB) was somehow facilitated by its 'back-arc' position behind the conventionally postulated active convergencesubduction belt from the Urals to Dobrogea at this time (e.g. Ziegler 1989; Zonenshain et al. 1990; ~engrr et al. 1993; Nikishin et al. 1996; Wilson et al. 1999). This kind of geodynamic setting was also used by some authors as a point of departure in modelling the Late Palaeozoic subsidence of the EEP as a whole (e.g. Mitrovica et al. 1996; Ismail-Zadeh 1998; Artemieva 2003). Both these geodynamic scenarios (mantle plume grid mantle flow behind a vigorous subduction system) involve the presence of thermal perturbations in the mantle beneath the lithosphere. In this respect, they are similar except that a mantle plume sensu stricto implies a more discrete thermal source at a deeper level. That Late Devonian magmatism is so widespread on the EEC and its margins therefore argues against a conventional plume model that should be associated with one well-defined 'hotspot'. As regards a 'back-arc' scenario, this is also problematic. Recent studies on the Urals suggest that Devonian subduction was beneath the (Magnitogorsk) intra-oceanic volcanic arc, therefore away from and not towards the EEC (Puchkov 1997; Brown et al. 1998, 2002; Brown & Spadea 1999; see also Brown et al. 2006). This is in contradiction to what is shown, for example, by Nikishin et al. (1996), which formed the basis of the Mitrovica et al. (1996) model of Devonian platform subsidence. Moreover, there is also little compelling argument that the southern margin of the EEC was the hanging wall of an active subduction zone at this time (e.g. the northern margin of Palaeotethys), despite the fact that it has very often been adopted as such in plate kinematic models (e.g. Adamia et al. 1981; Gamkrelidze 1986; Ziegler 1990; Zonenshain et al. 1990; Stampfli et al. 2001; Stampfli & Borel 2002; von Raumer et al. 2003). Based on the information reviewed above, the closest a subduction zone would be at this time is south of the TC (Adamia et al. 1981; Gamkrelidze 1986) and, according to the integrated model shown in Figures 7 and 8, this is not the case. At best, some of the volcanism of this age in the area of the GC (Dizi Basin) could be explained in such a way, but it seems unlikely to be responsible for widespread peri- and intracratonic magmatism elsewhere. What can be said is that by the Late Devonian some time has passed since the Caledonian amalgamation of Laurentia, Baltica and Avalonia (see Pharaoh et al. 2006), and that the eastern (proto-Urals) as well as at least parts of the southern (GC, Moesia-ND?) margins of the previously Baltican segment of the newly formed Laurussian continent are compressional plate boundaries. The thermal state of the underlying asthenosphere was probably in a state of flux in such a scenario. The possible imposition of a mantle plume (or broader, plume-like thermal perturbation from deeper in the mantle) would only serve to destabilize the situation further and perhaps this is sufficient, plume or not, to explain the character and distribution of magmatism on the EEC and its margins at this time.

Mid-Late Carboniferous to Permo-Triassic evolution of the southern margin of the East European Craton (EEC) Table 1 summarizes the main Permo-Triassic characteristics of the main zones of interest, based on Dercourt et al. (2000) with additional references. A more detailed description, which includes the Carboniferous, follows. North Dobrogea (ND) and Moesia

In southern ND-northern Moesia (Fig. 4) Late Carboniferous sedimentation was dominated by marine facies (with Lower Carboniferous radiolarian cherts, Haydoutov & Yanev 1997) and, going northwards, it becomes paralic to continental. A presumably Carboniferous-Early Permian thick continental succession of silici- and volcanoclastic rocks was deposited on the Macin Zone of Dobrogea (Fig. 4). Shallow, I-type calc-alkaline granitoid plutons intrude the base of the continental sequence as well as Middle Carboniferous and Lower Permian units; they might be of Early Permian age according to A r - A r ages (Seghedi et al. 1999). (In this case, it may be that I-type calc-alkaline granitoid plutons reported as Devonian in age, mentioned above, are also of Permian age.) Monazites and A r - A r dating both evidence a Late Carboniferous-Early Permian thermal event in ND (Seghedi et aL 1999, 2003b; Crowley et al. 2000). Rifting with bimodal basalt-rhyolitic volcanism developed in ND in Permo-Triassic times (Seghedi 2001, and references therein). This bimodal volcanism preceded the widespread extrusion of mid-ocean ridge basalt (MORB)-type-basalts (Niculitel units) during the Anisian that marked the final stages of rifting and possibly (sub)oceanic crust formation. This would correspond to the Meliata-Karakaya ocean of Stampfli et al. (2001). Structural analyses show that rifting occurred in a wide right-lateral shear zone, indicating that the (minor) displacement of Moesia was to the NW during Permo-Triassic rifting (as inferred by Banks & Robinson 1997). The Late Palaeozoic tectonic evolution of the Pre-Dobrogean Depression north of ND (the southwesternmost SP) was similar, with Permo-Triassic rift basin formation and igneous activity (Seghedi et al. 2003c). The associated faults are thought to be steeply dipping, based on correlation between wells but without seismic control. Pre-Dobrogean greenschists may be Early Permian in age, resulting from either deep burial and/or enhanced heat flow at this time. The P o n t i d e s

The Carboniferous succession of the eastern part of the Istanbul Zone (IZ, Fig. 1) of the Pontides (Zonguldak area) consists of an alternation of conglomerates, sandstones, coal, claystones and siltstones-limnic coal basins (similar to north Moesia-Dobrogea; Okay et al. 2006). In the western Istanbul Zone the sedimentary facies was evidently deeper, with radiolarian cherts and siliclastic turbidites similar to those of south Moesia (Okay et al. 2006). Early Late Permian granites near Istanbul are reported to have been intruded into the Carboniferous succession (Yllmaz 1977). Weak contractional deformation of Late Carboniferous age may be a record of the inferred Late Carboniferous-Early Permian collision in the Balkans on the southern margin of Moesia (Haydoutov & Yanev 1997). The Triassic continental clastic succession of the Istanbul Zone includes basaltic flows. As in the upper Palaeozoic succession, a change of facies occurs towards the east (Okay et al. 2006). Carboniferous granites and Carboniferous-Permian continental formations are reported in the Eastern Pontides (Okay & Sahinttirk 1997). According to Robinson et aI. (1995), the continental clastic deposits of the Eastern Pontides (Fig. 1) are Late Carboniferous to

PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE

491

Table 1. Evolution of some key zones of the southern margin o f the E E C in Permo-Triassic times

Urals and Uralian foredeep

East European Platform

Peri-Caspian Basin (PCB)

The DDB and Donbas

Scythian PlatformFore-Caucasus Scythian PlatformKarpinsky Swell

Early Permian, Artinskian (Vai & Izart 2000b)

Mid-Permian, Wordian (Gaetani 2000a)

Compression and migration of the foredeep depocentre to the west, affecting the PCB to the south Regression; shallow marine lagoon to evaporitic; weak subsidence, mainly compensated by sedimentation Persistence of deep marine turbiditic sedimentation; supply of sediments from the Urals transported through its foredeep

Final stage of the Uralian orogenesis; foredeep still evolving with intense subsidence

Shallow marine (in SW Russian part) and evaporitic to continental (towards the NW); (trans)tensional reactivation and uplift of the SE Donbas 1, volcanism2 Early Permian volcanic rocks2

Early Triassic, Olenekian (Gaetani 2000b)

Mainly exposed

Mostly emerged; continental clastic deposits towards the margins

Kungurian salt deposits followed by renewed sedimentation (still supplied by the Urals); sedimentation still accommodated by the PCB Continental clastic deposits

Dramatic increasing thickness of marine sediments towards the KS

Kayasula and Mozdok rift basins 4 with volcanism;5 with en echelon dextral distribution (transtensional basins5)

Shallow-water limestones and turbidites

Different blocks (uplifts, volcanism); rift basins 3

Highly variable settings; emergent versus highly subsiding blocks; (trans)tensional3

Greater Caucasus

Shallow marine to continental

Mid- and Late Permian very thick continental deposits (with some rhyodacitic pebbles) in grabens fed by the erosion of uplifted areas

Deep marine turbidites in the Dizi series (the only indicator for a wide seaway, south of the transtensional belt)

Marine limestones

Shallow-water limestones (occurring only in pebbles within the Late Triassic-Early Jurassic Tauric flysch) Synrift activity with volcanism in ND

Moesia-North Dobrogea (ND)

Pontides

Shallow marine to continental, and in ND rift activity

Late Triassic, Late Norian (Gaetani 2000d)

PCB still actively subsiding; gateway to the sea between SH and UkS; sediment supplied by incipient inversion along KS and Mangyshlak

Continental clastic rocks and areas of emergence, erosion

Tectonic inversion1

Turan Platform

Transcaucasus Crimea

Mid-Triassic, Early Ladinian (Gaetani 2000c)

Stavropol High emerged

Mostly emerged

End of mid-Triassic: possible onset of inversion of the KS; Kayasula and Mozdok basins filled by shallow-water epicontinental sediments Mostly emerged? (no Mid-Triassic sediments preserved); continental alluvial deposits along the Mangyshlak Development of rift basins in transtensional setting and subsequent inversion of basin; remnant of the back-arc deep basin as in the Dizi series

Tectonic inversion of the KS

Emerged Marine shales and shallow-water limestones on the Crimean Plain

Emerged Remnant of the deep basin (Tauric series)

Synrift activity with volcanism in ND; mainly alluvial sedimentation in Moesia Eastern and Central Ponfides emerged; conglomerates(?) in the Istanbul Zone

In ND: oceanic-like crust development in the basin (Niculitel units)

Basin inversion in northern Caucasus and fluvial conglomerates; Dizi series with an increasing turbiditic infill from the southern active convergent margin Emerged Shallow-water limestones in Simferopol area (north of Crimean Mountains); increasing turbiditic infill In ND: increasing turbidific infill

Remnant of the deep basin (Kfire series); south of Devrekani unit is the northward subducfing PalaeoTethys(6)

Evolution is according to relevant sections of the Peri-Tethys Atlas (Dercourt et al. 2000) and references therein. Other sources: 1Stovba & Stephenson (1999); 2Alexandre et al. (2004); 3Thomas et al. (1999); 4Nikishin et al. (1998a, b); 5Tikhomirov et al. (2004); 6Ustarmer & Robertson (1999).

492

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Triassic in age and were deposited in half-grabens developed in an extensional setting. The Sakarya Zone (SZ, Fig. 1) is a large Pontide domain including the Triassic subduction-accretion Karakaya complex (Pickett et al. 1992; Ustarmer & Robertson 1993, 1994, 1997, 1999; Altiner et al. 2000; Goncuoglu et al. 2000; Kozur et al. 2000; Stampfli 2000; Robertson et al. 2004; Stephenson et al. 2004; Yllmaz & Yllmaz 2004). The exact geology of the various units of the Sakarya Zone is controversial (see Okay et al. 2006), but the Karakaya complex, which forms the basement of the Sakarya Zone in general, is widely considered to be the product of the subduction of Palaeotethys (e.g. Stephenson et al. 2004). Crimea

There is considerable uncertainty concerning the formation and evolution of latest Palaeozoic-Early Mesozoic sedimentary basins in Crimea. The basement of the Simferopol and Novotsaritsyn uplifts (SNU, Fig. 1) is unconformably overlain by a thin Permian sedimentary layer (Zonenshain et al. 1990). In southern Crimea, Upper Carboniferous to Permian shallow marine limestones are found as olistoliths in the overlying Upper Triassic-Lower Jurassic unit. On the Crimean plain, Lower Triassic coarse clastic deposits and Triassic limestones occur. The oldest tectonostratigraphic unit of the Crimean Mountains is the Tauric series, a highly deformed clastic complex of Late Triassic-Early Jurassic age with turbiditic (flysch-like) deep-water to shallow-water mudstones to siltstones - sandstones (Byzova 1980, 1981). The North Crimean Azov Basin (NCA, Fig. 1) formed along the border of the Ukrainian Shield, probably being initiated in the Permian and almost certainly in existence by Triassic times (Nikishin et al. 2002; Stephenson et al. 2004). As mentioned above, the boundary between the Ukrainian Shield (EEC) and the SP was characterized by normal faulting throughout the Late Triassic (Kruglov & Cypko 1988).

The Greater Caucasus ( G C ) - T r a n s c a u c a s u s (TC)

Somin (1998) dated diorites and gneisses along the Main Range of the GC at 320Ma (Serpukhovian-Bashkirian boundary) and Zonenshain et al. (1990) mentioned Devonian and Lower Carboniferous sediments of the eastern SP that are deformed and intruded by Serpukhovian-Bashkirian rocks. Perchuk & Philippot (1997) reported that the age of massive eclogites found within a metamorphic complex in the Fore Range of the GC is BashkirianMoscovian (c. 310 Ma) and that their exhumation had already occurred by the Kazimovian (c. 300 Ma). These rocks belong to the same ophiolitic complex as discussed above, ascribed to Devonian ocean crust development but argued to have been obducted already by the Visran (Khain 1975, 1979; Adamia et al. 1981; Belov 1981; Belov et al. 1990; Zonenshain et al. 1990; Milanovsky 1991). Further analyses are thus needed to settle the position and nature of the mafic and ultramafic rocks of the Palaeozoic GC, although this particular discrepancy does not substantially affect the general tectonic model presented above (Fig. 7). The Upper Carboniferous succession of the GC is shallow marine along the Main Range and paralic to continental to the north, with alternations of conglomerates, sandstones, claystones and siltstones-limnic coal basins (Fore Range and northward). Subaerial calc-alkaline Late Carboniferous andesites, rhyolites and basalts, and Early Permian andesites, dacites and trachytes accompany these sediments (Khain 1975; Adamia et al. 1981; Zonenshain et al. 1990). Late Carboniferous sedimentation in the TC is clearly similar to that of the GC north of the Main Range (Khain 1975). (It is noted

that this sedimentation also resembles that described in Dobrogea and in the Zonguldak area of the Istanbul Zone). The sediments are intruded by I-type granitoids (Zonenshain et al. 1990) that are Late Carboniferous (300 Ma) and, according to Zakariadze et al. (2001), show a mantle signature. They are similar to the 300280 Ma granitic massifs of the GC (Milanovsky 1991; Somin 1998; Zakariadze et al. 1998, 2001). Mid-Permian continental graben-like basins developed in the area of the GC (Gaetani 2000a). Early and Mid-Triassic rifting (Nikishin et al. 1998a) accompanied by the extrusion of tholeiites (Shalimov et al. 1995) has also been reported in the GC, but no magmatic or sedimentary rocks exist in the area of the TC, which was uplifted and denuded during Late Permian (Adamia & Lordkipanidze 1989) and Triassic (Zonenshain et aL 1990) times. The deep marine flysch sedimentation in the Dizi series was uninterrupted from the Late Palaeozoic (Devonian) to the Triassic (Adamia et al. 1981).

The Peri-Caspian Basin (PCB) and the Turan Platform (TP)

Late Carboniferous deep marine turbiditic and shallow platform sedimentation occurred in the PCB. In the Early Permian, this trend of sedimentation persisted at an increasing rate in a deepening basin. Sediment supply came from the newly formed Uralian orogenic belt through its foredeep (Vai & Izart 2000b). Evaporites were deposited in Kungurian times. Early Triassic sedimentation was shallow marine (Nikishin et al. 1996). The Emba Basin (Fig. 1) is characterized by a bathyal to abyssal sedimentation from Kasimovian until Artinskian times, when evaporitic facies are recognized (Vai 2003). The basement of the TP crops out in some belts east of the Caspian Sea (Tuarkyr, Mangyshlak; Fig. 1). Garzanti & Gaetani (2002) have shown that the TP basement probably comprises an accretionary belt formed on an active Indonesian-type margin and that it was already in the process of consolidation during the Late Palaeozoic. Although many authors have correlated the TP to the SP, it seems unlikely that the SP is the lateral equivalent of the TP (given that Late Palaeozoic amalgamation is nowhere recorded for the SP basement). Late Carboniferous to Early Permian sedimentation in some areas of the TP shows deep marine facies (Vai 2003). Permo-Triassic sediments and volcanic rocks of the TP (intermediate to felsic and dated in the range 275-205 Ma) unconformably overlie Lower and Middle Palaeozoic rocks. A Permo-Triassic flysch-like series occurs in the Mangyshlak area (see Gaetani et al. 1998; Thomas et al. 1999; Gaetani 2000a; Garzanti & Gaetani 2002). Thomas et al. (1999) suggested that Late Permian to Triassic basins trending roughly north-south developed in a broad transtensional right lateral wrench zone. The D n i e p r - D o n e t s Basin (DDB), the Karpinksy Swell (KS) and the eastern Scythian Platform (SP)

The Late Carboniferous succession on the EEP is shallow marine and was deposited in a stable to slightly extensional tectonic setting (Vai & Izart 2000a,b). A global sea-level regression beginning in the Late Carboniferous dominated the Permian. Coastal siliciclastic sedimentation replaced the Late Carboniferous carbonate platform on the EEP (Vai & Izart 2000a,b) and the Early Triassic was mostly continental (Nikishin et al. 1996). The regional nature of the EEP Permian unconformity is demonstrated, for example, in the central part of the DDB where it does not display angularity. The small angularity seen on the northern margin of the DDB between the Triassic and underlying units

PALAEOZOIC-EARLY MESOZOICSOUTHERNEAST EUROPE reflects a relative sea-level fall in combination with differential post-rift subsidence rates from the basin margin to its axis in a typical 'thermal sag' basin (Stephenson et al. 2001; Stovba & Stephenson 2003). The Early Permian unconformity is expressed not only as a long wavelength phenomenon in the DDB-DF but also as an abrupt, rapidly developed, monoclinal uplift of significant amplitude (some kilometres) affecting especially the southern margin of the basin (Stovba & Stephenson 1999). Shatalov (1986), and more recently Alexandre et al. (2004), have identified an important Early Permian magmatic episode (286 -+-_11 Ma or 285-270Ma, respectively) in the DF and the adjacent AM (Fig. 1). Vitrinite and fission-track modelling also imply a thermal event in the DF of this age (Spiegel et al. 2004). Furthermore, the tectonic setting in the DDB-DF in latest CarboniferousEarly Permian times was one of dextral transtension (Stovba & Stephenson 1999; Saintot et al. 2003a). Thus, the Early Permian was a time of rift reactivation in the DDB accompanied by magmatism, including lava flows, domes and shallow dykes (McCann et al. 2003). Early Triassic volcanism (245-250 Ma) is recorded in the DF as a minor event (Alexandre et al. 2004). The recent studies of the DF and DDB mentioned above demonstrate that the tectonic setting and controls on sedimentation within the EEC during the Permian are not dominantly compressional, and therefore are not related to orogenesis in the SP area as proposed by Nikishin et al. (1996). A key issue in this regard is the history of the Main Anticline of the DF. This structure clearly was initiated during latest Carboniferous-Early Permian times, based on stratigraphic relationships, and this was often cited as proof of compressional tectonism at this time. However, it is now known with some certainty that salt movements, intimately related to the transtensional tectonics observed at this time (Stovba & Stephenson 2003), are responsible for the initial PermoCarboniferous development of the Main Anticline of the DF (Stovba & Stephenson 1999, 2003; Saintot et al. 2003b). Recent seismic and geological data reveal that the compressional events in the DDB-DF were in Late Triassic and (most importantly) Late Cretaceous times (Stovba & Stephenson 1999; Saintot et al. 2003a,b), and in the western KS in the Late Triassic (Sobornov 1995). The eastern SP is in general covered by middle to Upper Carboniferous clastic deposits with Upper Carboniferous and Permian subaerial andesites, rhyolites, basalts (Belov 1981; Khain 1985; Belov et al. 1990; Zonenshain et al. 1990; Milanovsky 1991) that are, in turn, buried by a thick overlying PermoTriassic and younger succession. The Devonian-Carboniferous units display low-grade metamorphism (to greenschist facies at most, probably related to burial) intruded by granites that are Serpukhovian-Bashkirian (Zonenshain et al. 1990) or Early Permian in age, as exemplified in the Manysh Trough (MT, Fig. 1) area (see Tikhomirov et al. 2004), just south of the central KS. Extrusive magmatism of the Fore-Caucasus has been recently dated as being Early Permian (285-270 Ma; Alexandre et al. 2004). A pre-Triassic uplift of as much as several kilometres was reported in Nikishin et al. (2001) in the Pre-Caucasus (south SP) region, comparable with the regional Permian uplift of the southern part of the EEP, including the southeastern part of the DDB (see Stovba & Stephenson 1999). Typically, an orogenic event has been proposed for the Permian in the eastern SP (Zonenshain et al. 1990; Nikishin et al. 1996, 2001). However (and keeping in mind that the observational dataset is rather limited), none of the observations noted above are particularly compatible with the occurrence of a tectonic event of such significance as an orogeny. The main evidence for an orogeny in this area seems to be that the Permian succession is unconformably overlain by Triassic marine deposits and calc-alkaline volcanic rocks. This model is based on angular unconformities between Permian and Triassic (and if they are absent, Jurassic) strata. However, this angular unconformity is not observed everywhere and, where it is, appears to be related to rifting processes. In Late Permian

493

times, the Stavropol High (SH, Fig. 1) had emerged and the sedimentation was shallow marine in remnant basins. The angular unconformity between Permian and Triassic rocks in the Stavropol High is similar to that in the Manysh Trough-Kayasula Basin (MT and KB, Fig. 1), which developed in a rifting environment, based on the nature of extrusive magmatism and sedimentation at this time (Tikhomirov et al. 2004). Subaerial to submarine basaltic volcanism and shallow marine sedimentation characterizes the Manysh Trough (southern margin of the KS) during the Early and Mid-Triassic (Nazarevich et al. 1986; Nikishin et al. 1996, 1998b, 2001; Tikhomirov et al. 2004). No angular unconformities are found in the basin depocentres. Volcanism also occurred in the Late Triassic, dominated by subaerial explosive rhyolites in Norian-Rhetian emerged areas of the Manysh Trough and Kayasula Basin (Tikhomirov et al. 2004). The geodynamic background of Triassic volcanism, in general, was extensional and related to rifting processes (Nazarevich et al. 1986; Nikishin et al. 1996, 1998a,b, 2001; Tikhomirov et al. 2004). Gaetani (2000b,c) mentioned that the en echelon distribution of Early Triassic basins on the SP, such as the Manysh Trough, and the Kayasula and Mozdok Basins (MT, KB and MB, Fig. 1), occurred in a general dextral transcurrent setting.

L i t h o s p h e r i c p l a t e d y n a m i c s in the latest Palaeozoic- Triassic

The Late Palaeozoic-Early Triassic was a time of significant global plate-tectonic reorganization. Pangaea was formed in the Late Carboniferous (e.g. Ziegler 1989, 1990; Ziegler & Stampfli 2001; Stampfli & Borel 2002), involving the collisions of Laurussia with Gondwana (Variscan Appalachian orogen) and Laurussia with Kazakhstan-Siberia (Uralian orogen); the Central Asian Orogen was also active at this time (Ziegler 1989, 1990; Kazmin & Natapov 2000). There is a vast literature dealing with the process of assembly of Pangaea followed by break-up and the opening of the Permo-Triassic (Neo)Tethys (e.g. Dercourt et al. 1986, 1993, 2000; Ziegler 1989, 1990; Stampfli & Pillevuit 1993; Nikishin & Ziegler 1999; Golonka 2000; Golonka & Ford 2000; Nikishin et al. 2002; Stampfli & Borel 2002; Vai 2003). In this section a brief review of general plate kinematic models and their constraints, as they pertain to southeastern Europe at this time, is made, followed by an assessment of how the information presented above fits or does not fit these models, including the implications of this. Late Carboniferous-Early Permian continental coal-beating rift-wrench basins are widespread in Western Europe (e.g. Oslo Rift, Polish Trough, North East German Basin, Lodeve and Autun Basins in France). Many of these basins, but not all of them, developed upon Variscan basement. They are linked to a regional wrench fault system under a transtensional stress regime (e.g. Ziegler 1990), thought to be related to dextral megashearing between Gondwana and Laurussia (Arthaud & Matte 1977). According to the Pangaea B configuration of Vai (2003), there was a lateral displacement of 800 km, developed continuously from Late Carboniferous to Triassic times (Vai & Izart 2000a; Vai 2003). At the eastern limit of this transform zone, the Permo-Triassic Tethys (Neotethys) began to form (see Dercourt et al. 1993, 2000; Vai & Izart 2000b). With respect to the immediate borders of the EEC itself (and earlier accreted marginal belts), continental collision occurred on the EEC eastern border, in the Urals, in the Late Carboniferous (Zonenshain et al. 1990; Nikishin et al. 1996; Puchkov 1997; Brown et al. 1998, 1999, 2002, 2006; Brown & Spadea 1999). The eastern part of the EEC began to evolve as a foredeep basin, which in part developed as a well-expressed foreland fold-and-thrust belt during the Permian (Perez-Estaun 1997; Puchkov et al. 1997). To the SW, a Balkanide terrane is thought

494

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to have been accreted to south Moesia in the Late Carboniferous (Haydoutov & Yanev 1997; Yanev 2000; Fig. 9b). Subduction of Palaeotethys is generally assumed to have occurred beneath the southern margin of the EEC marginal belts throughout the Carboniferous (e.g. Mossakovsky 1975; Kazmin 1991), evolving into a Pacific-type margin with continental graben and back-arc basin formation in the Late Carboniferous and Permian (Usta6mer & Robertson 1994; Nikishin et al. 2001), with subduction culminating during the Triassic (Nikishin et al. 2001; Stampfli et al. 2001; Ziegler & Stampfli 2001). The Karakaya complex (Sakarya Zone, Western Pontides) is thought to be a remnant of the accretionary prism of the Palaeotethys subduction zone (Okay & Mostler 1994; Ustarmer & Robertson 1994). Indeed, an Early Permian (Artinskian) seaway ran from Moesia to the Crimea and Caucasus, opening towards the Turan Sea (Vai & Izart 2000b), and Kazmin (1991) interpreted the almost contiguous Mid-Carboniferous to Late Permian-Triassic volcanic belt along the GC (see Khain 1979; Adamia et al. 1986) to be subduction related. He mentioned, however, that this was based only on the calc-alkalinity of the volcanic rocks. Carboniferous sediments are generally reported as shallow water and continental (Nikishin et al. 1996), and contain volcanic and volcano-clastic rocks. There are strong similarities between Late Carboniferous facies seen in northern Moesia-Dobrogea, Zonguldak (eastern Istanbul Zone), the Eastern Pontides, the Fore Range of the GC, and the TC. The Late Palaeozoic limestones, such as the pebbles in the Tauric series of Crimea, were deposited in a shallow back-arc basin behind the subduction trench of Palaeotethys (cf. Gaetani 2000a). Remnants of Permo(?)-Triassic back-arc(?) basins with ophiolitic fragments are found in ND (the Nalbant flysch), in Crimea (the Tauric flysch; Robinson & Kerusov 1997) and with ophiolitic fragments in the Central Pontides (the Ktire series and the Akgrl flysch; Ustarmer & Robertson 1994). In such a back-arc basin model, it follows that continental fragments such as the Istanbul and Devrekani units (Western and Central Pontides, respectively) became detached from the Eurasian margin (Ustarmer & Robertson 1994) at this time. Extensional basins also developed on the SP and the Turan Platform, in a fight-lateral transtensional setting according to Gaetani (2000a). Palaeomagnetic data reported by Lemaire et al. (1998) suggest a southeastern displacement of the Turan and Scythian platforms from the Permian to the Ladinian (Mid-Triassic). It can be concluded that the present study area was characterized by a transtensional right-lateral wrench system, similar to that of Central-Western Europe but whether or not this was, strictly speaking, a back-arc setting cannot be said with certainty. Robinson et al. (1995) in fact considered that the widespread Middle Triassic limestone cover on the whole of the Pontides is indicative of an EEP-SP passive margin and that subduction of Palaeotethys therefore had not begun prior to the Late Triassic (with the Eo-Cimmerian Late Triassic inversion of basins being associated with the onset of subduction). If this were the case, and no subduction beneath the EEP-SP occurred during the Late Palaeozoic, then the extensional basins of ND, the SP and the GC developed as a prolongation of the Permo-Carboniferous rift system of Western Europe along a broad dextral shear zone between Laurussia and Gondwana. Similar kinds of magmatic activity occurred during this time (Early Permian (285-270 Ma) and Late Triassic (230-200 Ma), with a minor Early Triassic (245-250Ma) event) in three regions representing different tectonic contexts: platform (SP), ancient shield (Azov Massif, AM, Fig. 1) and rift (DF). This implies the presence of a long-lived or intermittently active mantle plume according to Alexandre et al. (2004), assuming that the source of magmatism was the same for the whole region, that it was deep (in the mantle, and therefore independent of the tectonic context), and that magmatism was active over a fairly long period of time. Although Nikishin et al. (2002) made

a similar suggestion (short-lived Permo-Triassic mantle plumes in Eastern Europe), no isotopic or geochemical data are available to support the mantle source of magmatism as plume related. Indeed, in consideration of a diffuse wrench margin as described above, it can simply be related to the transtensional tectonic setting (with oblique subduction, if any, increasing the probability of magmatism in the extensional basins). Three major lithospheric units are generally inferred for this area during the Early Permian: the Palaeotethys oceanic domain and continental Laurussia dextrally displaced relative to continental Gondwana (e.g. Dercourt et al. 1993, 2000). Although Robertson et al. (2004) suggested the possibility of a short-lived period of subduction beneath Gondwana in the Late Carboniferous, it is generally considered that Palaeotethys was being subducted beneath the EEP-SP margin. As such, it would have been affected by the rotational movement of Gondwana, either as part of the same plate (with a Gondwanan passive margin) or with an incipient Neotethys in an intervening position. In this case, a geometric argument can be made that subduction became more oblique with time, with more and more dextral strike-slip movement being imparted to the EEP-SP margin (see Fig. 9). The broader driving mechanisms responsible for global plate reconfigurations at this time (including Pangaea break-up) occurred at the scale of mantle convection systems (e.g. Ziegler 1993; Nikishin & Ziegler 1999; Ziegler & Stampfli 2001; Ziegler et al. 2001; Nikishin et al. 2002). An outstanding issue regarding the southernmost part of the EEP-SP during Early Permian times is its inferred widespread and significant uplift, over and above a global drop in sea level (affecting the southern margin of the DDB-DF, for example; see Stephenson et al. 2001). In the context of the plate boundary model discussed above, in which subduction beneath the EEPSP becomes increasingly oblique throughout the Carboniferous, ultimately evolving into a dominantly transcurrent plate boundary in the Permian, one mechanism that could be responsible for Early Permian uplift is slab break-off. Enough time had elapsed (some 50-60 Ma during the Carboniferous) for the subducted slab to have reached the 660 km discontinuity, even with a convergence rate as low as 1-2 cm a -~. Shearing of the slab within the oblique subduction setting could have led to its detachment at this time, triggering not only the Early Permian uplift of the overriding plate but also heating of the lithosphere and widespread magmatism.

Late Triassic to Late Jurassic-Early Cretaceous evolution and Cimmerian tectonics Following the development of Permo(?)-Triassic rift basins, the southern EEC margin remained unstable behind the obliquely subducting Palaeotethys. The tectonic record is fairly abundant and various events can be described. The Greater Caucasus (GC)

Gaetani et al. (2005) described an Eo-Cimmerian orogenic event in the GC as follows. Pulses of Early and Mid-Triassic tectonic activity are indicated by intercalations of clastic wedges, the most important being a Lower Triassic conglomerate that contains pebbles of exhumed serpentinites, which was intercalated with the marine carbonate succession followed by late Anisian folding of these units. They were then unconformably overlain by Ladinian conglomerates and subsequent marine sediments. Gaetani et al. (2005) thus reassigned the Eo-Cimmerian event from Early to Mid-Triassic times. Nevertheless, the uplift and erosion of metamorphic rocks and the input of clastic wedges can also occur in a rifting environment. Evidence for a later compressional event is provided by outcrops of the Upper Triassic

PALAEOZOIC-EARLY MESOZOICSOUTHERNEAST EUROPE Khodz Formation. Its lower part shows highly folded limestones (similar to the Anisian limestones, displaying chevron folding) and its upper part is formed by conglomerates or/and debris flows that rework the underlying limestones. This suggests that the inversion of the GC basin occurred throughout latest Triassic times (Saintot et al. 2006). The Early and Mid-Jurassic corresponds to a renewed period of rift basin formation, with associated extrusive magmatism, in the GC region (Saintot et al. 2004, 2006). The SinemurianPliensbachian sediments of the western GC are represented by deltaic(?) coarse sandstones north of the belt, by deep marine mudstones-sandstones in the central part, and by mudstonessandstones with Pliensbachian rhyolitic tufts and basalts in the south (Panov & Gushin 1987; Lordkipanidze et al. 1989). In Toarcian times, from north to south, shelf to deep marine mudstones-sandstones were deposited without any record of volcanic activity in the western part (extrusive rocks have been reported in the eastern GC by Lordkipanidze et al. 1989). In AalenianBajocian times, lateral facies variations in sedimentation are important (from continental to deep marine; Saintot et al. 2004). Bimodal (from mantle and crustal sources) rhyolitic and basaltic magmatism in subaerial as well as shallow marine environments also occurs at this time, with the peak of volcanic activity in the Bajocian (Saintot et al. 2004). The volcanic edifices form structural highs, having been developed on elevated blocks within the rift system and/or having formed a segmented and elevated volcanic chain within the basin system. The Bajocian volcanic rocks can thus be considered to belong to the synrift succession of the GC basin. Mid-Jurassic magmatic activity (180 Ma) has also been reported in the Fore Caucasus area of the SP (Alexandre et al. 2004) and could be linked to the widespread rift-related volcanism in the GC. However, the widespread Bajocian volcanism could be also the record of the enlargement of the subduction-related volcanic arc of the TC, encompassing the whole of the GC-Crimea basinal structure (probably caused by the shallowing of the subducted slab; e.g. Saintot et al. 2006). Brittle tectonic structures in the Lower to Middle Jurassic series of the GC record a transtensional regime characterized by a nearly east-west trend of extension (Srbrier et al. 1997; Saintot et al. 2004). In the northern GC, the Early Jurassic extension more probably trends NE-SW, as recorded by NW-SE-trending normal faults and east-west right-lateral strike-slip faults with Pliensbachian synsedimentary normal faulting and with N W - S E and north-south dykes (Srbrier et al. 1997). Under such a stress regime, the present-day major WNW-ESE fault zones of the GC could have originally been transtensional right-lateral faults. Sedimentary facies variations are also compatible with W N W ESE intra-basinal structures (structural highs and depocentres). It follows that these structures, prior to their inversion as thrusts during Cenozoic shortening, may have controlled the development of the rift system as a series of pull-apart basins, thus implying oblique relative plate movement on the southern margin of the EEP at this time (see S~brier et al. 1997). Compressional structures (thrusts and folds) in the early MidJurassic units of the GC were mainly formed during Cenozoic NE-SW to NNE-SSW shortening (Saintot 2000; Saintot & Angelier et al. 2002). Beds often dip steeply and folds are close to tight, becoming tighter towards the south. (There is no intense Cenozoic deformation affecting Callovian and younger units of the northern limb of the GC.) No older compressional structures of significance exist in the Early to Mid-Jurassic series, where a Bathonian Mid-Cimmerian phase of compression has been reported, although locally there is evidence of a mild compressional pre-Callovian Mid-Cimmerian tectonic event (see Saintot et al. 2006). On the northern limb of the GC, the succession from Jurassic to Cretaceous is fairly continuous, displaying no spectacular gap that could be correlated with a Late Jurassic Neo-Cimmerian compressional event. Where Upper Jurassic limestones overlie, with a

495

strong angular unconformity, Middle Jurassic units, they are considered to be allochthonous Alpine structural units (S. Korsakov, pers. comm.). This seems to have been observed as early as the early 20th century by Renngarten (1929), who wrote: 'The plastic Middle Jurassic slates formed close folds and the thick Upper Jurassic limestone unit, more rigid, slid above them ... At other sites, we can observe the gradual conformity from the slates to the limestones.' Such behaviour resembles descriptions of the Alpine Tertiary phase of folding and thrusting, during which rocks deformed differently according to their competence. Khain (1994) also emphasized that the Jurassic Cimmerian deformations are confined to Dobrogea and Crimea, whereas the GC is mainly an Alpine edifice. Zonenshain et al. (1990) described the GC as a deep basin from Early Jurassic time that closed in the Tertiary. In the TC, dacites and rhyolites erupted during the Pliensbachian (Adamia et al. 1981). Bajocian volcanism is also widespread. The Jurassic depositional environment was shallow water to continental. Like the GC, the TC in the Early Jurassic could have been in a rift setting (intra-arc?) although it is classically interpreted as the locus of an active subduction volcanic belt. Certainly by the Mid-Jurassic, sedimentological and structural evidence clearly supports an arc environment for volcanism just south of the TC (Adamia et al. 1981, 1986; Gamkrelidze 1986; Panov 2004). A review of the evolution of this region, with a focus on the nearby South Caspian Basin (SCB, Fig. 1) and its margins, has been given by Brunet et al. (2003). In all surrounding areas the Eo-Cimmerian compressional event can be observed. The timing of rifting of the SCB is Early(?)-Mid-Jurassic, with (inferred) oceanic crust accretion during the Mid- to Late Jurassic. The SCB might therefore be considered as the continuation of the GC basin. However, no present data support a strike-slip tectonic setting for the development of the SCB as described above for the GC; nor do any present data rule this out.

C r i m e a ( T a b l e 2)

Hettangian rift inversion has been reported in Crimea (Mileyev et al. 1996; Nikishin et al. 1998b) leading to the development of folding and south-vergent thrusting in the Tauric (Late TriassicEarly Jurassic in age) flysch-like unit (Byzova 1980, 1981; Khain 1994). A Sinemurian to Toarcian complex is represented by deep-water clastic turbidites along with shallow-water conglomerates with minor volcanic rocks (pillow basalts, lava flows, dykes). Aalenian to Early Bajocian units are mostly sandy. The Early Mesozoic units (sediments and volcanic rocks) are strongly affected by strike-slip faults (A. Saintot, field observation) although there is no age constraint for this faulting. A key unit is the Late Bajocian to Bathonian volcanic complex with shallow to deeper-water volcanoclastic rocks and Bathonian shales and siltstones. According to Voznesensky et al. (1998), the volcanic activity occurred throughout the Callovian (as recorded in the tholeiitic to calc-alkaline extrusive Karadag complex). Local unconformities related to block rotation then appear from time to time from the Early Jurassic to the Callovian. These are typically interpreted as indications of compressive deformation (at Aalenian-Bajocian and Bathonian-Callovian times; Zonenshain et al. 1990; Nikishin et al. 2001). However, ambiguities exist. For example, the senior co-author has observed a nearly undisturbed (Middle Jurassic?) lava flow upon a slump unit of Lower Jurassic flysch that has also been interpreted as a unit folded during tectonic shortening. Such a depositional architecture is also explicable in the context of rifting and it was proposed that, following Early Jurassic compressive deformation and inversion of the 'Tauric flysch-like' basin, Crimea was affected by transtensional deformation and (as in the GC) the development of pull-apart basins separated by structural highs (Saintot et al. 2005a). The NW-SE Simferopol-Alushta Fault Zone (SAFZ) may have played an

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PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE important role during the evolution of the Jurassic basin in Crimea. It seems to control the locations of magmatic bodies and, furthermore, the Callovian unconformity disappears to the east of it (Zonenshain et al. 1990). There are no observations to preclude normal displacement along this fault and, indeed, the presence of magmatic bodies along the SAFZ supports the interpretation that it formed in an extensional regime. The obliquity of the N W - S E trend of the SAFZ relative to the southern margin of the EEP also suggests that the Early and Mid-Jurassic basin in this area developed as a pull-apart basin (analogously to the Early to Mid-Jurassic pull-apart basins in the GC). Early to Mid-Jurassic rifting terminated with the deposition of shallow- to deeperwater carbonates, reefs and conglomerates. In detail and highlighting the lateral facies variations during the Late Jurassic, a carbonate platform developed in the western part of Crimea whereas, going eastwards, conglomerates laterally passing to flysch were deposited in a much deeper environment. The pebbles of the thick Kimmeridgian conglomerates of Central Crimea probably are derived from the Pontides. As the conglomerates unconformably overlie the oldest strata, Zonenshain et al. (1990) considered that the Kimmeridgian was a time of continental collision between the EEP and a southern continental block (namely, the Pontides). That this unconformity disappears to the east was taken as evidence that no collision occurred to the east. However, the zone of facies transition in the Late Jurassic corresponds to the SAFZ, which, as described above, separated two distinct palaeo-environments (probably inherited from Mid-Jurassic tiffing tectonics: a structural high to the west and a basin depocentre to the east). Recent observations (McCann, pers. comm.) suggest that the Kimmeridgian conglomerates are in fact sourced from within the basin itself, not from an external emerged unit. In any case, the Upper Jurassic succession as a whole is clearly not intensely deformed. Furthermore, the Upper Jurassic platform unit is considered to be allochthonous according to Mileyev et al. (1996, 1997), emplaced during the Berriasian. In this case, the structures of the Crimean mountains resemble those of the GC except for the age of the thrusting that emplaced the allochthonous units, which is believed to be Berriasian in Crimea but Alpine (Tertiary) in the GC.

Pontides

In the southern western Pontides, the Karakaya complex is a southvergent imbricated complex of Triassic- Lower Jurassic flysch, magmatic units and radiolarites, with olistostromes of Carboniferous and Permian limestones. The Nilfifer unit of the Karakaya complex is newly identified as an oceanic seamount or seamounts on an oceanic plateau of the Palaeotethys oceanic plate. Its deformation and metamorphism dates from the Late Triassic, when it was accreted to the Eurasian margin (Okay 2000; Can Gent 2004), producing the Eo-Cimmerian compressional event. In the Central Pontides are the Upper Triassic-Lower Jurassic Kfire series and the Akg61 Formation, with turbidites and silts with Late Triassic ophiolitic fragments (Usta6mer 1993; Usta/Smer & Robertson 1993, 1994, 1997). The Kfire deep basin closed in Callovian times with a propagation of north-vergent thrusts to the north (Usta6mer & Robertson 1994). The Ktire series is similar to the Tauric unit of Crimea and the Akg61 Formation is equivalent to the Eskiordia Formation of Crimea, highlighting the probable proximity of this part of the Pontides to Crimea from Triassic to Early Jurassic times. The olistostromes of Carboniferous and Permian limestones are derived from a platform, the location of which is not known. The present-day Ktire unit may be the remnant of a back-arc basin to the north of the Palaeotethys subduction zone (Usta6mer & Robertson 1994, 1997, 1999). Robinson et al. (1995) reviewed the early Mesozoic history of the Eastern Pontides. Late Palaeozoic-Triassic grabens were

497

inverted in the Late Triassic (in pre-Sinemurian times). Lateral facies variations are common in the Lower Jurassic strata, and volcanic rocks are present in the eastern part (and extend into the TC area, as mentioned above). An Aalenian compressive event has been reported, followed by the widespread development of a midJurassic volcanic belt (subduction arc-related). The associated sediments were deposited in paralic to deep marine environments (as in the GC). Upper Jurassic limestones were deposited in the Pontides as early as Callovian and conformably overlie the Bajocian-Bathonian volcanogenic succession. These events all occurred prior to the opening of the present-day Black Sea, generally considered to have been in Late MesozoicEarly Cenozoic times (see Stephenson et al. 2004). However, it is clear that considerable evidence exists for a significant Jurassic crustal thinning and rifting event in the same area (cf. Spadini et al. 1997).

North Dobrogea (ND)

In ND, basin inversion is typically reported to have begun in Late Triassic (Carnian) times (e.g. Lupu et al. 1987), with the migration to the east of the depocentre of siliciclastic turbiditic deposition (Alba Formation), into the Tulcea Zone, which is inferred to indicate uplift of the Macin Zone as a source area. The same scenario was proposed to continue into the Mid-Jurassic (BajocianBathonian) with deposition of the Nalbant clastic turbidites. Sedimentation in ND at this time is in basins separated by structural highs reported to have formed in a transpressional stress regime (Gradinaru 1988). That transpressional deformation was continuous from the Late Triassic to the Mid-Jurassic (60 Ma) seems unlikely, as basin inversion events typically occur as short-lived pulses (e.g. de Lugt et al. 2003). What is clear is that there are compressional structures of Mid-Jurassic (mid-Cimmerian) age (Saintot et al. 2005b). Late Jurassic platform carbonates transgressed the deformed strata of the Tulcea Zone from the SE. According to Seghedi (2001), therefore, basin inversion had ceased by this time, although Banks (1997) assumed that compression reached its maximum in Kimmeridgian times. This is inconsistent with observed Kimmeridgian transtensional reactivation of the Peceneaga-Camena Fault Zone (PCFZ) on the SW margin of ND (Fig. 7), which controlled the extrusion of basalts at this time (Seghedi 2001). In the Early Cretaceous, the Late Jurassic carbonate platform was apparently uplifted and destroyed as a result of renewed compressive movements. Hippolyte et al. (1996) proposed reverse dextral movement along the PCFZ under a NNE compression during Late Jurassic-Early Cretaceous times. Seghedi (2001) pointed out that several key issues pertaining to the development of ND remain unknown. These include the amounts of extension and subsequent shortening, whether or not an Anisian (early Mid-Triassic) ocean formed, and the lateral displacement through ND of Moesia relative to the EEC.

Turan and Scythian platforms (TP and S P ) - D o n b a s Foldbelt ( D F ) - K a r p i n s k y Swell (KS)

An important compressional event in this area occurred at the end of the Triassic (Sobornov (1995) and Gaetani (2000d) for the KS; Popov (1963), Konashov (1980), Stovba & Stephenson (1999) and Saintot et al. (2003a) for the DF). Inversion of basins within the TP also occurred during Late Triassic times (Thomas et al. 1999). The accretion of the TP occurred at this time on the Kazakhstan plate along the Mangyshlak belt (Mg, Fig. 1) (Garzanti & Gaetani 2002), where the Permo-Triassic flysch-like series is strongly deformed and overlain by an underformed Lower Jurassic succession

498

A. SAINTOT ETAL.

(Gaetani et al. 1998). According to Zonenshain et al. (1990), these formed in a continental rise environment on the margins of a basin lying between the EEC (the Ustyurt Massif) and the Karabugaz Massif (of the TP). Permo-Triassic sec!iments and volcanic rocks of the TP are in turn intruded b y a-~skite granite for which latest Triassic (207 Ma) and late Early Jurassic (180Ma) dates have been reported (Belov 1981). The Mashad suture, lying south of the TP, is considered to be the Palaeotethys suture (Garzanti & Gaetani 2002). Prior to the beginning of the Jurassic, magmatism on the SP and D F - A M occurred more or less simultaneously but thereafter did not (Alexandre et al. 2004). One anomalous magmatic event in this respect occurred in the early Late Jurassic (150 Ma) in the DF and in M i d - L a t e Jurassic time (156 ___ 11 Ma) in the AM (Shatalov 1986, and references therein). Given that magmatism of this age is not reported elsewhere in the southern EEP area (with the possible exception of the Karadag volcanic complex of the east Crimea Mountains, which Voznesensky et al. (1998) proposed to be Callovian) it is probably local in nature, related to some 'internal' process in the evolution of the DF, rather than signifying a regional lithosphere-scale process. Thus, the interior of the southern EEP during Jurassic times evolved in a stable tectonic environment, with no major structures forming. Tectonic activity was restricted to the N D - C r i m e a - P o n t i d e s - G C ( - S C B ) corridor, where extensional basin formation predominated.

Terrane reconstructions

Lemaire et al. (1998) defined the TP and SP as two separate lithospheric blocks. A slight right-lateral and SE-directed displacement (relative to Laurussia) occurred in Permian times, roughly along the KS and the Mangyshlak belt (and thus, as mentioned above, in agreement with the development of pull-apart rift basins along the margin of the EEC at this time). Thereafter, according to the palaeomagnetic data of Lemaire et al. (1998), both terranes moved northwards in Carnian to Norian times to be welded finally in their present location by the end of the Triassic. According to Adamia et al. (1981), the subduction of Palaeotethys ended in Late Triassic time with the collision of Iran with the southern EEC and contiguous units. Kazmin (1991) stated that this occurred prior to the Norian. In general, according to most global plate reconstructions, the 'Cimmerian' blocks of Iran collided with Eurasia at the end of the Triassic, resulting in the development of the Elburz belt (Fig. 1; e.g. Baud & Stampfli 1989; Stampfli & Pillevuit 1993; Saidi et al. 1997). Palaeotethys subduction was still continuing to the east and west of this collisional front. The widespread Late Triassic (Eo-Cimmerian) inversion of basins on the southern EEC margin can therefore be linked to the accretion of Iran to Eurasia. Dercourt et al. (2000) considered the Dizi series as the remnant of a Permo-Triassic basin, based on the stratigraphic studies of Somin & Belov (1967), developed as part of the K t i r e - C r i m e a Svanetia Palaeotethys back-arc basin. This interpretation is not consistent with the presence of volcano-terrigeneous sedimentary layers of Devonian-Early Carboniferous age in the Dizi series as reported by Kazmin (1991) and Adamia et al. (1981). However, allowing that it may have initially developed in a 'protected' location on the subsiding margin of the EEC, as discussed above, its infill could have been metamorphosed during the eventual accretion of the TP, displaced by the 'indentation' of Iran (Fig. 10a). If the Dizi series represents an accretionary prism, its development would have ceased at this time, thus recording a pre-Sinemurian (Khain 1975; Adamia et al. 1981) or prePliensbachian metamorphism (Kazmin 1991). A similar scenario can also be envisaged for the TC, which may have been located, for example, to the east of the Fore-Range of the GC, and then displaced westward at the same time. Thereafter (from the Late Triassic), until the Tertiary, the subduction zone was south of

Fig. 10. Schematicrepresentation of (a) the Early Mesozoic Eo-Cimmerian widespread compressionalevent and (b) the Early to Mid-Jurassic development of pull-apart basins and the lateral displacementof Transcaucasus. Abbreviationsfor structural units are as in Figure 1.

TC, north of the Lesser Caucasus (Gamkrelidze 1986) or between the TC and Nakhichevan Platform (NP, Fig. 1; Adamia et al. 1981). Early to Mid-Jurassic rift basin development in the GC is thought to have occurred in a back-arc setting (Nikishin et al. 2001), in which case the 'driving' subduction zone lay south of the Pontides, south of TC, and south of the accreted Iranian blocks. In this case, the widespread Bajocian volcanic rocks of the GC and present Crimean Mountains are s e n s o stricto not remnants of a volcanic arc related to this subduction zone (e.g. Nikishin et al. 1998a, b), although there are some geochemical data (Page et al. 1998) that suggest this setting for Crimea. Rather, in view of the sedimentological and structural data in Crimea and GC as described above, the Bajocian volcanic rocks might have occurred in a synrift setting (i.e. synchronous with normal faulting) during the formation of the Crimea-(proto-East Black Sea)-GC and SCB basin system (Fig. 10b). What can be said with some certainty is that, from Early Jurassic to Aalenian or Bajocian times, these basins developed in a transtensional (presumably back-arc) regime as pull-aparts on W N W - E S E right-lateral fault zones (e.g. Saintot et al. 2006).

Conclusions: the main outstanding issues It is not the intention here to repeat by summarizing the general palaeotectonic history of the study area, which is, in any case, as has been demonstrated, very poorly constrained and involves considerable controversy. Rather, the main issues confronting advances to a better understanding of this area and the lithospheric processes that have governed its evolution are highlighted.

PALAEOZOIC-EARLY MESOZOIC SOUTHERN EAST EUROPE

(1) On the basis of what can be considered reliable, it seems likely that the crust of the Scythian Platform (SP), at least in part, as well as crustal units underlying the PCB and present-day Greater Caucasus (and possibly even parts or all of Moesia and TC) was accreted to the EEC in the Neoproterozoic-earliest Palaeozoic, more or less contemporaneously with the Timanide and Baikalian orogens recognized to the east and north of Baltica. This would seem to provide the 'missing link' in explaining the orogenic-accretionary history of the SP, long assigned to a Variscan (Hercynian) age for which there seems to be no real evidence and which seems to be in contradiction to other indications of tectonic setting in Late Palaeozoic times. There are obvious implications for plate-tectonic reconstructions at this time, with the location and passivity of this margin of Neoproterozoic Baltica as traditionally defined being in need of significant revision. (2) In general, the dominant tectonic style of the southern EEC and its margins was extensional or transtensional from the Devonian to the end of the Jurassic (and thereafter, including the Cretaceous development of the Black Sea and associated basins, although this is beyond the scope of this paper), punctuated in time and space by compressional events, the most important of which is the Late Triassic(-Early Jurassic) Eo-Cimmerian event, probably marking the (partial?) closure of an oceanic domain, possibly the Palaeotethys. A milder compressional (basin inversion) event occurred in Mid-Jurassic times (midCimmerian) throughout the area behind a renewed subduction zone, and is probably related to a change of geometry of the subducting slab. (3) There is widespread evidence of Late Palaeozoic (Late Devonian-Early Carboniferous) rifting and rift-related magmatism not only in an intracratonic EEC setting (e.g. Dniepr-Donets Basin) but also on what would have been the Neoproterozoic-Early Palaeozoic accreted, pericratonic (passive margin), crust. There is no strong evidence that the southern margin of the E E C - S P at this time included a north-vergent subduction zone. Thus, the distribution of the rift system and the tectonic setting in which rifting processes have developed are placed in an entirely new context. (4) Subduction-related processes are locally in evidence during the Carboniferous, but by the latest Carboniferous-Early Permian, wrenching and basin formation in what was a regional transtensional tectonic environment dominate. This is seen not only on the margin but also in the interior of the EEC. As such, the regional tectonic style at this time is similar to what is widely observed in Central-Western (Variscan) Europe, where it is sometimes regarded as evidence of post-orogenic gravitational collapse of the Variscan orogeny. As it is far more widespread than the Variscan orogenic belt, affecting even the interior of the EEC, other plate-scale processes are clearly involved. (5) In general, the changes in the tectonic record from Early to latest Carboniferous(-Permian) imply the possibility that the plate boundary along the southern margin of the EEC evolved from one of normal or slightly oblique convergence to one of increasing obliquity, possibly better characterized as a transcurrent plate boundary from the Early Permian. Regional uplift throughout the southern EEC and its margin in the Permian may be related to this apparent change in plate boundary type. One possible mechanism could be the steepening and eventual detachment of a subducted slab. Triassic basins on the EEC margin thereafter developed as pull-apart (back-arc?) basins. (6) Following the Late Triassic (Early Jurassic) Eo-Cimmerian interruption of extensional tectonics, rifting and basin formation was reactivated and a (North Dobrogean?)-Crimea-Greater Caucasus-South Caspian basin system developed from E a r l y Mid- to Late Jurassic times. Unlike the earlier (?Latest Carboniferous- Permian-Triassic) phase of basin development, this occurred in what was more likely a true back-arc setting (although with convergence still being highly oblique and the basins, accordingly,

499

having a pull-apart geometry), with arc volcanic rocks well developed in the Pontides and TC regions. The magnitude and direction of lateral terrane displacement involved in the configuration of the margin during this time is difficult to assess with the present dataset. (7) The long-lasting period of extensional tectonic style on the southern margin of the EEC from the Late Palaeozoic through much of the Mesozoic can be thought of in terms of three 'simple' stages: (a) the opening of Palaeotethys(?) in Late Devonian-Early Carboniferous times; (b) the further development, subduction and then closure of Palaeotethys(?) thereafter, until the Late Triassic; (c) the migration of active subduction followed by consumption of Neotethys during the Jurassic (and throughout the Cretaceous until closure in Late Cretaceous-Tertiary times). This paper is the Netherlands Research School of Sedimentary Geology Contribution 2005.05.01. The research on the Late Palaeozoic and Early Mesozoic evolution of the southern EEP is supported by the Netherlands Research Foundation (NWO) with a grant to the first two authors. Part of the research was funded by the MEBE programme and, in the past, by the Peri-Tethys Programme. We would like to warmly thank participants and all who helped to carry out field campaigns in the Donbas Foldbelt (1999 and 2000), in Crimea (2002 and 2004), and in the Greater Caucasus (2003) and especially, for fruitful discussions, F. Chalot-Prat, E. Baraboshkin, P. Fokin, A. Kitchka, S. Korsakov, S. Kostyuchenko, T. McCann, A. Nikishin, C. Pascal, V. Privalov, A. Seghedi, P. Tikhomirov and S. Vincent. The authors are grateful to J. Golonka for the review of the manuscript.

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The Timanide, Caledonide and Uralide orogens in the Eurasian high Arctic, and relationships to the palaeo-continents Laurentia, Baltica and Siberia D. G. GEE, O. K. B O G O L E P O V A & H. L O R E N Z

Department of Earth Sciences, Uppsala University, Villaviigen 16, SE-752 36, Uppsala, Sweden (e-mail: david.gee@ geo.uu.se)

Abstract: Recent studies of structure, stratigraphy and isotope geochronology on Svalbard and East Greenland have provided a foundation for reconstructing the Laurentian margin of the Arctic segment of the North Atlantic Caledonides. The axial zone of the high Arctic, Barentsian Caledonides has been inferred to trend northwards through the Barents Shelf to the northern edge of the Eurasian margin between Kvitcya (easternmost Svalbard) and western Franz Josef Land, based on analysis of drill-cores that sampled the preCarboniferous basement beneath Alexandra Island. The deformation front of the Barentsian Caledonides has been inferred to trend northeastwards between Franz Josef Land and Novaya Zemlya. The North Kara Terrane, reaching from Severnaya Zemlya (SZ) and northernmost Taimyr in the east to northern Novaya Zemlya in the west, comprises the northernmost foreland to the Barentsian orogen. Four lines of independent evidence are presented here demonstrating that the North Kara Terrane is a direct northerly continuation of the Timanide domain, the latter composing the Neoproterozoic accreted margin of Baltica in the Timan-Pechora-Urals region. These lines of evidence, all from October Revolution Island (SZ), include: (1) a westerly source for Old Red Sandstones successions, with 'Caledonian' fish fauna and detrital muscovites yielding Ar/Ar ages of c. 450 Ma; (2) Ordovician igneous rocks containing c. 550 Ma xenocrysts; (3) Cambrian turbidites with c. 545 Ma detrital muscovites; (4) Cambro-Silurian fauna with many species shared with Baltica. In addition, the Neoproterozoic turbidites of northern Taimyr have been previously reported to contain c. 560 Ma zircon populations, a signature that has been recently found in similar lithologies from Bol'shevik Island (SZ). All these late Vendian ages are characteristic of the Timanide Orogen of the Timan-Pechora-Novaya Zemlya region and, together, indicate that the North Kara Terrane was not an independent 'plate' or 'microcontinent' in the Palaeozoic, as previously proposed, but an essential part of southernmost (Ordovician coordinates) Baltica. Comparability of the evolution of the Timanian margin of the North Kara Terrane with the contemporaneous Baikalian evolution of adjacent Taimyr, together with the lack of evidence of Palaeozoic oceanic rocks and Uralian collisional, high-pressure metamorphic assemblages in Taimyr, suggests that the palaeo-continents Siberia and Baltica were never separated by a major ocean in the high Arctic.

The Palaeozoic geology of the Eurasian high Arctic, from Svalbard in the west to Severnaya Zemlya in the east (Fig. 1), has long been a matter of dispute and speculation. Most of the area is covered by sea, ice and a C e n o z o i c - M e s o z o i c blanket of sediments. The track of the Caledonide orogen northwards, from the Scandes of northernmost Norway, across the Barents Shelf to Svalbard, its discordant relationship to the NW-trending Timanide orogen of northwestern Russia, the extension of the latter to the north into N o v a y a Zemlya, and its truncation in the east by the Uralian orogenic front have been widely treated in the literature; the relationship of all these old orogens to the Palaeozoic to early Mesozoic deformation of Taimyr and Severnaya Zemlya has remained enigmatic. For many decades, parts of Svalbard have been known to have Laurentian affinities (Harland 1997, and references therein; Gee & Teben'kov 2004). The Timanides (Gee & Pease 2004) have been accepted as an essential late Neoproterozoic component of the eastern (today's coordinates) margin of Baltica, flanking the Archaean and Palaeoproterozoic core (the East European Craton) of this Early Palaeozoic continent. The southern and central domains of Taimyr have been shown to have unambiguous Siberian affinities (Sobolevskaya et al. 1995; Tesakov et al. 1995). Also, between Baltica and Siberia, a relatively small continental region, centred on Severnaya Zemlya and including northern Taimyr, referred to as the Kara 'block' or 'plate' or 'microcontinent' (called here the North Kara Terrane), has been thought by most workers to be an independent unit, perhaps part of a larger continental assemblage, Arctida, that once m a y have dominated the northern polar regions (Zonenshain & Natapov 1987; Zonenshain et al. 1991). Related to these questions about the possible prolongation of these Palaeozoic orogens into the Eurasian high Arctic shelf and their interrelationships, has been the speculation concerning their continuity even further northwards and westwards prior to the opening of the Arctic Basin (Fig. 1). The compelling evidence

of sea-floor spreading in the Eurasia Basin, with separation of a microcontinent, Lomonosova, from the Eurasian margin in the Early Cenozoic (Kristoffersen 2000; Jokat et al. 2003) to form the Lomonosov Ridge, implies that the latter is probably underlain by Mesozoic, Palaeozoic and Precambrian complexes comparable with those of northernmost Eurasia. This conclusion can thus be used to test hypotheses for the opening of the Amerasia Basin, a subject that will be left for another occasion. W e focus here on B a l t i c a - N o r t h K a r a - S i b e r i a relationships.

The Barentsian Caledonides Recent work on Svalbard, summarized by Gee & T e b e n ' k o v (2004), and on eastern Greenland by Higgins & Leslie (2000) and Higgins et al. (2004), has shown that most of Svalbard' s Caledonian terranes are direct northerly continuations of the Caledonides of eastern Greenland (Fig. 2). In particular, structural and stratigraphical similarities of the two eastern complexes on Svalbard, the West Ny Friesland Terrane (Witt-Nilsson et al. 1998) and the Nordaustlandet Terrane (Gee et al. 1995), with the N E Greenland and Central East Greenland thrust complexes, respectively, extend in time from the late Palaeoproterozoic to the Early Palaeozoic. They leave little doubt that eastern Svalbard's Proterozoic and Cambro-Ordovician bedrock composed an essential part of the western flank of the North A t l a n t i c - A r c t i c Caledonides during the Caledonian orogeny (Gee 2005a). The latest studies of Nordaustlandet (Teben'kov et al. 2002; Johansson et al. 2004, 2005) have shown that metamorphic grade and intensity of deformation increase towards the east. High-grade complexes, with widespread migmatization, assumed by the early Svalbard explorers, from Nordenskjold (1866) and Sandford (1956) to Krasil'shikov (1973) and Harland (1997), to be ancient cratonic basement, have proved to be Caledonian h i g h - T - l o w - P terranes with widespread zircon crystallization at

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 507-520. 0435-4052/06/$15.00

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Fig. 1. Tectonic elements of the western Eurasian Arctic in the Early Cenozoic, during initial opening of the Eurasian Basin and Norwegian-Greenland Sea (modified from Gee 2004). The North Kara Terrane is marked by a red dotted line, which follows the hypothetical suture between central and northern Taimyr (e.g. Inger et al. 1999) and crosses the Barents-Kara Shelf (Cocks & Torsvik 2002) to the Eurasian shelf-edge.

c. 440 Ma. Successions on Nordaustlandet range in age from late Mesoproterozoic metasediments, intruded by c. 950 Ma granites (Johansson et al. 2004), to unconformably overlying Neoproterozoic successions (Gee & Teben'kov, 1996; Sandelin et al. 2001; Teben'kov et al. 2002), which pass up, via Vendian tillites, into shelf sandstones and carbonates of Early Palaeozoic age. Caledonian migmatization has been identified as far east as KvitCya (Fig. 2) and the Caledonian suture(s) must be located even further to the east (Gee 2004). The early assumption that these Nordaustlandet-KvitCya highgrade complexes were ancient ('Urberg', Archaean, etc.) led previous researchers to conclude that the northern part of the Barents Shelf was underlain by cratonic lithosphere, part of the so-called Barentsian (Arkhangel'sky & Shatsky 1933; Stille 1958; Ziegler et al. 1977; Sharov 2000) or Barents (Egiazarov et al. 1972; Aam 1975; Siedlecka 1975; Harland 1997) Craton. This concept has persisted (Verba & Sakoulina 2001), partly because of the lack of reliable geochronological evidence. The Barents Craton has been assumed to occupy much of the basement of the northern Barents Shelf and extend far south into the Pechora Shelf and beneath the Pechora Basin (Stille 1958). Only at one location (Fig. 1), beneath the westernmost island (Alexandra) in the Franz Josef Land Archipelago, has pre-Palaeozoic basement been sampled. Here, the Soviet drilling project, Nagorskaya-1971,

passed through Lower Carboniferous basal sandstones and an unconformity at a depth of c. 1900 m and then sampled tightly folded turbidites and dark phyllites, metamorphosed at low greenschist facies. Vendian acritarchs have been reported from these metasediments (Smirnova in Dibner 1998). Detrital muscovites yielded c. 610 Ma ages by the A r / A r method (Kaplan et al. 2001) and zircon populations are dominated by 'Grenvillian' ionmicroprobe U / P b ages (Pease et al. 2001a). The A r / A r data were interpreted by Kaplan et al. (2001) to be related to the low greenschist-facies metamorphism, and by Pease et al. (2001a) to the age of the source rocks. This evidence favours affinity to the Timanide complexes of the southern Barents Shelf; also and that the deformation and metamorphism is of Caledonian age. A K / Ar whole-rock age on a phyllite of c. 360 Ma has been inferred to support this interpretation (Dibner 1998). Whether or not the Alexandra Island basement turbidites are correctly interpreted to have been subject to Caledonian deformation (late Timanian deformation is also possible, but less likely on the basis of the published information), it is clear that the main Caledonian suture(s) defining the axial zone of the Barentsian Caledonides must lie to the west of Franz Josef Land between this archipelago and Kvit0ya. The inference that the Caledonian deformation front lies to the east of Franz Josef Land is supported by other lines of evidence, presented below.

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Fig. 2. ArcticCaledonides and Laurentia-Baltica relationships in the early Mesozoic (from Gee & Teben'kov 2004)

Timanide-Baltiea relationships From VarangerhalvCya, in northernmost Norway, where the NWtrending Timanide successions and structures are abruptly truncated by the thrust front of the Caledonide orogen (Roberts & Olovyanishnikov 2004), the Timanide orogen (Gee & Pease 2004) continues nearly 2000 km southeastwards to the foreland fold-and-thrust belt of the northern Urals (Gee & Pease 2004, fig. 1). Within the Urals, both to the north via the Polar Urals, Pai Khoi and Vaigach to Novaya Zemlya, and to the south for 1000 km to westernmost Kazakhstan, the Timanian complexes can be recognized below a Late Cambrian to Early Ordovician unconformity. Rifted and passive margin Palaeozoic shelf and continental slope successions (Puchkov 1997; Bogolepova & Gee 2004) overlie these Timanian rocks, which together compose a vast region of Late Neoproterozoic accretion along Baltica's eastern margin. At the time of the main Caledonian orogeny (Scandian) in Scandinavia, starting in the late Llandovery and continuing through the Silurian and Early Devonian (Gee 1975), the passive margin successions of eastern Baltica (presentday coordinates) were little influenced by tectonic activity; in the Uralian Ocean further east, subduction (probably west-dipping) was generating volcanic-arc complexes and back-arc ophiolites (Perez-Estaun et al. 1997; Brown et al. 2002). The effect of this intra-oceanic subduction on Baltica's northeastern passive margin may have been the local Devonian extension and mafic magmatism in the Pechora Basin (Nikishin et al. 1996). The basal Palaeozoic unconformity of eastern Baltica (Bogolepova & Gee 2004) oversteps a wide variety of Timanide complexes, from the classical Riphean to Vendian successions of the Bashkirian Anticlinorium (Maslov 2004) to the blue schists of the Kvarkush Anticlinorium (Beckholmen & Glodny 2004), the Neo-Mesoproterozoic and perhaps older complexes of

the Sub-Arctic Urals (Belyakova 1988; Pystin 1994) and the ophiolites of the Polar Urals (Dushin 1997). Vendian magmatism, from c. 550 Ma (Gee et al. 2000) to 580 Ma (Remizov & Pease 2004) and c. 610Ma (Larionov et al. 2004), is widespread within the Timanide complexes as far as central Novaya Zemlya (Lopatin et al. 2001; Korago et al. 2004). The main characteristics of the Timanide orogen (Gee & Pease 2004; Gee 2005b) within the Timan-Pechora type area (Fig. 3) are known from extensive deep drilling of the Pechora Basin (Belyakova & Stepanenko 1991). These data, taken in relation to the regional aeromagnetic anomaly maps, provide the basis for inferring the regional distribution of the main Timanide rock units. In particular, the magmatic rocks (mainly andesitic with gabbro-diorite intrusions) of the Pechora Zone are well defined by strong magnetic anomalies (Kostyuchenko 1994; Kostyuchenko et al. 2006) and can be followed from the Uralian mountain front northwestwards beneath the Pechora Basin and at least 300 km offshore into the Pechora Shelf; thick Palaeozoic and Mesozoic successions of the eastern Barents Shelf obscure their possible track further north. However, together with the fragmentary evidence from the Timanide basement of Novaya Zemlya (Korago et al. 1993, 2004), it is probable that Timanide complexes, including some reworked Mesoproterozoic 'Grenvillian' basement (Gee et al. 2000; Lopatin et al. 2001), extend northwards through the entire Barents Shelf to Franz Josef Land and the Eurasian shelf-edge.

Novaya Zemlya From the Polar Urals, east of Vorkuta, with their preservation of both Neoproterozoic (Engane-Pe) and Palaeozoic (Seum-Keu) ophiolites, the Uralide orogen swings northwards through Pai

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D.G. GEE E T A L .

Fig. 3. (a) Pre-Palaeozoic basement of the Timan-Pechora region; (b) diagrammatic profile (SW-NE) through the Timan-Pechora region in the Devonian. prior to Uralian orogeny (based on Gee et al. 2000).

Khoi and Vaigach into the Novaya Zemlya Archipelago. The 1500 km long Novaya Zemlya fold and thrust belt (Korago et al. 1992, 2004; Otto & Bailey 1995) is a direct northerly continuation of the frontal fold belt of the Polar Urals (Fig. 1). In southernmost parts of Novaya Zemlya and on Vaigach Island, greenschist-facies Neoproterozoic turbidites and dark phyllites provide the basement for shelf sandstones of Early Ordovician age, which pass up into a carbonate-dominated Palaeozoic succession (Bondarev 1989; Gee 2004). Turbidites appear in the Permian when the shelf collapsed and gave way, in the Early Triassic, to a molasse facies, prior to mid-late Triassic folding and thrusting. In contrast, in northern Novaya Zemlya, a turbidite succession of late Neoproterozoic age (Vendian and perhaps older) passes up, apparently without break, through the Cambrian, Ordovician, Silurian and Lower Devonian succession in deep marine, slope-rise facies. Recent observations (Gee 2005c) include a wide range of channel deposits, debrites and even olistostrome within the Palaeozoic succession. A recent seismic transect from the Barents to the Kara Shelf across central Novaya Zemlya (Sakulina et al. 2003) has indicated the presence of a Palaeozoic 'basin' beneath the Mesozoic sequence of the western shelf of the Kara Sea and it is possible that this complex contains representatives of the classical Uralian hinterland oceanic terranes of the NW Urals.

Taimyr and the North Kara Terrane

The NW-vergent, mid-Triassic foreland fold-and-thrust belt of Novaya Zemlya (Korago et al. 1992, 2004), swings eastwards in its northernmost parts and, on the basis of potential field data, is generally considered (e.g. Bogdanov et al. 1998) to connect southeastwards across the Kara Sea to the Taimyr Orogen (Fig. 4). However, the Taimyr thrust system (Bezzubtsev et al. 1983, 1986) is SE-vergent, with Palaeozoic and Triassic platform successions being transported towards the Siberian Craton, the latter being exposed to the south of the Khatanga trough Mesozoic successions. In the central part of the Taimyr orogen (Fig. 4), Lower Palaeozoic successions are basinal, in graptolite shale facies (Sobolevskaya et al. 1995, 1997, 1999). They rest with major Late Vendian unconformity on Neoproterozoic complexes, including both late Mesoproterozoic-Neoproterozoic continental fragments (Pease et al. 2001) and c. 700Ma ophiolites (Vernikovsky 1996). The Late Neoproterozoic (Vendian) accretion to the northern margin of Siberia has much in common with similar Baikalian orogenic activity elsewhere around the Siberian continent (Vernikovsky et al. 2004). Taimyr's central block of Neoproterozoic complexes, unconformably overlain by Palaeozoic basinal successions and flanked

OROGENS IN THE EURASIANHIGH ARCTIC

511

Fig. 4. Simplifiedgeological map of northern Taimyr and SevernayaZemlya geology (from Lorenz et al. 2006b, based on Bezzubtsev et al. 1986). (For regional context, see Fig. 1).

to the south by the Siberian platform, is separated from more northerly parts of the Taimyr orogen by a major north-dipping thrust zone. The latter emplaced metasedimentary rocks, widely intruded by syntectonic Carboniferous granites, onto the Siberian outer margin. The metasediments of Taimyr's northern belt are dominated by turbidites that have been inferred to be of Neoproterozoic, including Vendian, age. In northwestern Taimyr, the latter have been shown to pass up into Cambrian black shales (Sobolevskaya et al. 1997) and these have been correlated in facies and fauna with similar lithologies in the central block. The thrust zone separating the northern and central belts of the Taimyr Orogen (Fig. 4) is considered by most workers (e.g. Vernikovsky 1996; Bogdanov et al. 1998) to mark the suture between the Siberian margin and the North Kara Terrane, northern Taimyr being considered to form the southern rim of this fragment of accreted continental crust (Bogdanov et al. 1998). The main exposures of the North Kara Terrane are located further north on Severnaya Zemlya; it is in this region that a more complete Palaeozoic and Neoproterozoic history can be established for assessing the affinities of this terrane. The Severnaya Zemlya Archipelago (Gramberg & Ushakov 2000) is dominated by four main islands (Fig. 5): Bol'shevik, October Revolution, Komsomolets and Pioneer. The first of these has a bedrock of Neoproterozoic tttrbidites and shales that have yielded Riphean and Vendian acritarchs. October Revolution has a stratigraphy reaching from the Early Cambrian (probably Vendian) to the Permian, the succession being interrupted by a deformation episodes in latest Cambrian to earliest Ordovician time and in the Early Carboniferous. These episodes are referred

to as the Kan'on River deformation and the Severnaya Zemlya deformation, respectively. Carboniferous and Permian sandstones transgress the older strata with major unconformity. The other northwesterly and northerly islands are ice-dominated and expose parts of the Palaeozoic successions, locally. Early Carboniferous (perhaps latest Devonian) deformation influences the whole of October Revolution Island and involved east- to NE-vergent folding and thrusting (Lorenz et al. 2006b). Detachment is particularly prominent in the Ordovician successions, where gypsiferous marls and limestones fold disharmonically. The age of this Severnaya Zemlyan deformation is constrained by the age of the youngest overlying late Early Carboniferous strata (Dibner 1982), and also by Early Carboniferous (c. 345 Ma), post-tectonic granites (Lorenz et al. 2007) that cut the folds on Bol'shevik Island (Markovsky et al. 1988).

Baltica affinity of the North Kara Terrane Evidence favouring close affinity of the North Kara Terrane with Baltica ranges from some characteristic lithofacies and faunas, to the provenance of the sediments, the age of their source rocks, and the age of xenocrysts in Ordovician igneous rocks. The magnetic anomaly data (treated below) provide evidence of the regional extent of the North Kara Terrane as an appendage of Baltica, and its relationship to Siberia. The Severnaya Zemlya succession is summarized briefly below, with emphasis on those features that are particularly relevant to the question of the relationships of the North Kara Terrane to Baltica and Siberia.

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D.G. GEE ET AL.

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Devonian strata dominate much of the western parts of Severnaya Zemlya, locally overlain, with major unconformity, by Carboniferous to Permian sandstones. The Devonian succession is in Old Red Sandstone (ORS) facies (Nalivkin 1973), and Lochkovian red conglomerates and sandstones are separated by a minor stratigraphic gap from underlying P~dolf marls (Khapilin 1982). The ORS, some 2 km thick, has been described to thicken and coarsen westwards, a source in this direction being supported by a variety of measurements of cross-bedding and channelling (Kurshs 1982). Subordinate carbonates occur, in particular in the youngest part (Famennian) of the succession. Detrital muscovites in the sandstones have yielded c. 450 Ma ages by the single-crystal Ar/Ar method. Fish faunas have been described as being closely comparable with those in other ORS basins within the Caledonides of Svalbard and Scotland (Karatajute-Talimaa 1978; M~irss & Karatajute-Talimaa 2002). Gee & Bogolepova (2006) have proposed that this Severnaya Zemlya non-marine siliciclastic succession was deposited in a foreland basin, sourced from the Caledonide orogen to the west.

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Fig. 5. Geological map of the Sevemaya Zemlya Archipelago (modified from Egiazarov 1967)

The underlying Silurian succession is composed of thick (c. 2.5 km) limestones (Menner et aL 1979; Khapilin 1982; Matukhin & Menner 1999), which, with a minor stratigraphic gap at the base, also forms the Upper Ordovician (Ashgill) sequence. In its upper part it gives way to red marls in the P~'/dolf, below the ORS. M~innik et al. (2002) have drawn attention to the similarities of the Silurian vertebrate fauna to those of the Baltica platform, in areas in the vicinity west of the Sub-Arctic Urals. Interestingly, subsequent studies of the Silurian sequences on Taimyr (Tesakov et al. 1995, 1998) and Severnaya Zemlya (Menner et aL 1979; Matukhin & Menner 1999) suggested that they are most similar to those on the Siberian platform. Beneath the Ashgill limestones and basal sandstones (quartzites), the Ordovician on October Revolution Island is divided into three parts and rests with major unconformity (Proskurnin 1999; Lorenz et al. 2006a) on folded strata of Late Cambrian age. Lower Ordovician conglomerates and sandstones pass quickly up into carbonates that give way into red marls and sandstones, with a prominent bimodal volcanic component. The

OROGENS IN THE EURASIANHIGH ARCTIC volcanicity started already in the Tremadoc (Lorenz et al. 2007) at c. 490 Ma and appears to have been centred on a north-trending belt through the eastern part of the island, associated with granites and gabbros. The red beds are gypsiferous and contain salt pseudomorphs; these evaporites, which appear in the Lower Ordovician sequence, persist into the overlying Mid-Ordovician black shales and limestones which continue through the Caradoc. Of particular interest for Baltica-North Kara relationships has been the identification of c. 550 Ma xenocrysts within the Ordovician igneous rocks (Lorenz et al. 2006a), which suggests Timanide-related source rocks in the lower crust. The Ordovician faunas are rich and diverse, consisting of mainly benthic groups (brachiopods, trilobites, gastropods, bryozoans, tabulate corals and crinoids). A distribution analysis of macro- and microfossil assemblages (Bogolepova et al. 2006, and references therein) from Severnaya Zemlya shows that they contain elements typical of Baltica (60% of common taxa), Siberia (30% of common taxa) and Laurentia (30% of common taxa) biotic provinces. One example of these affinities can be shown with regard to bryozoans, which occur commonly in the Ordovician successions of October Revolution Island. Only trepostoms, including Halloporina and Nicholsonella, have been recorded there so far (Nekhorosheva 2002). Nicholsonella is common in the shallow-water benthic assemblages typical of the Middle Ordovician successions of Arctic Russia (e.g. Siberia, Taimyr, Polar Urals) and North America. The genus Halloporina, being present mainly in the Middle Ordovician successions of North America, has been described from the Middle Ordovician of Estonia, and is also known from Wales. Halloporina severozemelica from the Ozernaya Series of Severnaya Zemlya shows affinity with H. parva (Nekhorosheva 2002), which is known from the Middle Ordovician sequences of North America and Estonia. The similarity between the Middle Ordovician representatives of Halloporina from the North Kara Terrane, Laurentia and Baltica favours their palaeogeographical proximity. The major break (Kan'on River Unconformity), separating the Upper Cambrian from Lower Ordovician rocks on October Revolution Island (Lorenz et al. 2006a) is remarkable in that it involves only a very small time interval, separating beds with trilobites of the Agnostus to Peltura minor zones (Lazarenko 1982; Bogolepova et al. 2001; Rushton et al. 2002) from Tremadoc tufts, yielding c. 490 Ma ages (Lorenz et al. 2007). Conodonts from the uppermost limestone beds from the Kruzhilikha River section of the lowest Ordovician unit, the Kruzhilikha Series, which have been previously correlated with the Latorp and Volkhov horizons of Baltica, suggest a late Tremadoc-Arenig age for these strata (Abaimova, cited by Markovsky & Makar'ev 1982). The area over which the unconformity is exposed is restricted to the southern parts of October Revolution Island, where the underlying rocks, despite upright folding and a generally high-angle discordance to the overlying Ordovician, are all apparently of Late Cambrian age. This part of the Cambrian succession is largely in black shales facies with subordinate limestones and turbidite sandstone intercalations, the latter increasing upwards and dominating in the northern exposures of central October Revolution Island. Muscovite is a conspicuous detrital mineral in the graded sandstones and has yielded single-crystal A t / A t ages of c. 545 Ma (V. Vernikovsky, pers. comm.; see below). According to Lazarenko (1966, 1982), the Cambrian trilobite faunas from October Revolution Island and those in Novaya Zemlya (the northeastern margin of Baltica) are similar. Further analyses of faunas (Bogolepova et aL 2001; Rushton et al. 2002) and acritarchs (Raevskaya & Golubkova 2006) suggest that the North Kara assemblages have affinity with those of both Baltica and Siberia; for example, the evidence from the trilobites Kujandaspis ketiensis and Maladiodella aft. abdita indicates a connection between Severnaya Zemlya and Siberia, whereas, brachiopods are distinct from those of Siberia. However, because there are taxa

513

common to both domains, exchange between the continents was possible. This may have been facilitated either by proximity, with an ocean small enough (700-1000km) to allow dispersal of marine organisms, or by a substantial separation and the presence of 'bridges' via islands, providing a migratory route for such a distribution of fauna. The occurrence of the Silurian myodocope ostracode Entomozoe in Severnaya Zemlya (Siveter & Bogolepova 2006) and Baltica, as well as Laurentia, can likewise be explained either by proximity or by the 'island hopping' mechanism. Shallow marine, highly bioturbated mudstones, sandstones and sandy limestones make up the underlying Middle Cambrian succession. The lower part of the sequence includes limestone lenses with trilobites (Ellipsocephalus sp., Paradoxides sp., Solenopleura sp. and Xystridura sp.) and brachiopods (Nisusia sp.). These pass down into Lower Cambrian grey and black siltstones with concretions containing the trilobites Fallotaspis sp., Galahetes sp., Hebediscus sp., Nevadella sp., Pagetiellus cf. lenaicus and Palaeolenus sp. (Lazarenko 1982). The Neoproterozoic turbidite-dominated successions of Severnaya Zemlya, largely (perhaps entirely) confined to Bol' shevik Island, strike southwards into northern Taimyr and swing southsouthwestwards through Taimyr. Pease (2001) reported detrital zircon populations in these turbidites of c. 555 Ma and drew attention both to their similarity in age to the granites of the Timanides beneath the Pechora Basin (Gee et al. 2000) and to the importance of this evidence for assessing Baltica-North Kara relationships; she concluded that northern Taimyr was a 'sliver of Baltica proximal to the Pechora Basin in the latest Neoproterozoic-Cambrian'. Detrital zircon populations of similar age have been obtained by us recently from the turbidites ofBol' shevik Island.

Kara Magnetic Anomaly It can be concluded from this summary treatment of the North Kara Terrane successions that there is a substantial body of stratigraphic and isotopic data favouring its affinity to the Timanide domain of Baltica (Fig. 1). In addition, the faunal and even some of the palaeomagnetic (see below) evidence favours no great separation from Siberia. Potential field anomalies provide essential information on the probable connection of the North Kara Terrane westwards into the eastern parts of the Barents Shelf and Novaya Zemlya, where the Timanian rocks are overprinted by late Uralian orogeny (Fig. 1). A major positive magnetic anomaly (Fig. 6) dominates the northern and eastern parts of the Kara Shelf and is referred to here as the Kara Magnetic Anomaly. The North Kara Terrane is characterized by a remarkable, nearly circular, bulbous magnetic anomaly pattern (Fig. 6), the outer rim of which reaches from Bol'shevik Island southwards and westwards through northern Taimyr and then northwestwards as far as northern Novaya Zemlya. This outer rim coincides with the outcrops of Carboniferous granites (some demonstrably magnetite-bearing) on Bol'shevik Island and northern Taimyr. Another, inner belt of anomalies is related to the prominent zone of Ordovician igneous rocks in eastern October Revolution Island. The Kara Magnetic Anomaly is truncated in the north by the continental slope of the Eurasia Basin and in the south by the thrust zone marking the boundary of the North Kara Terrane in Taimyr; in the west, it reaches into northern Novaya Zemlya and fades beneath the thick Mesozoic successions of the eastern Barents Shelf, where it may also be influenced by the deformation related to the eastern parts of the Barentsian Caledonides. Very large positive anomalies in the central core of the anomaly, located over easternmost Severnaya Zemlya and to the west in the northern Kara Shell are comparable in intensity and form to those over the Neoproterozoic subduction-related, mafic Pechora Zone igneous complex, but this is only one of several possible correlations.

514

D.G. GEE ETAL.

Fig. 6. Kara Magnetic Anomalyin the eastern Barents and Kara seas, enlarged from the magnetic map of the Arctic in Glebovsky et al. (2002). Based on these magnetic anomaly data, it can be inferred that the North Kara Terrane extends into northern Novaya Zemlya. The Kara Magnetic Anomaly reaches from northernmost Novaya Zemlya along the entire western half of its northern island, coinciding approximately with the deeper structural levels of the thrust core of the Novaya Zemlya Anticlinorium, the major structure that dominates the entire length of the archipelago (Gee 2005c). It may provide a measure of the allochthoneity of the fold-and-thrust belt exposed on land. The magnetic anomaly (Fig. 6) provides no support for the interpretation that a major Palaeozoic suture (Fig. 1) separates the North Kara Terrane from Baltica. Stratigraphic and palaeontological evidence has been presented by Abushik et al. (1997), Sobolevskaya et al. (1997, 1999), Kaban'kov et al. (2003) and Korago et al. (2004) for the close affinity of the Severnaya Zemlya succession to that of the northern part (North Island) of Novaya Zemlya. These workers (e.g. Sobolevskaya and Korago) claimed fundamental differences in facies in the Lower Palaeozoic successions between the northern parts of Novaya Zemlya's northern island and the rest of the archipelago, and argued for the existence of a major fault (the Main Novozemel'sky Fault) separating a northern block (SeveroNovozemel'sky) from Timanian regions further south. However, the magnetic anomalies are not displaced by the latter, and it is proposed here that the differences in facies are essentially controlled by a change in depositional environment, from shallow shelf in the south, to deep, slope-rise environments in the north.

Palaeomagnetic evidence The only palaeomagnetic evidence obtained from the Palaeozoic succession of the North Kara Terrane was from samples collected from October Revolution Island (Metelkin et al. 2000, 2005). This evidence suggested that the North Kara Terrane occupied similar latitudes to Baltica and Siberia in the Devonian and Silurian, but was at a somewhat different latitude in the Early Ordovician.

However, the evidence for the Early Ordovician was based on samples collected from above and below the Kan'on River Unconformity; these ranged in age from mid-Cambrian to midOrdovician. The palaeomagnetic evidence has been used to support the interpretation that the North Kara Terrane existed as an independent microcontinent in the Palaeozoic (Cocks & Torsvik 2002, 2005). We favour a previous interpretation (Pickering & Smith 1995; Nikishin et al. 1996; Cocks & Fortey 1998; Torsvik 1998), which included the North Kara Terrane as a part of Baltica.

Discussion Two main subjects are discussed below: the relationships between Baltica and Siberia, and some similarities in the Neoproterozoic to Early Palaeozoic evolution of northeastern and southwestern Baltica. Baltica-Siberia

relationships

The evidence presented above for the close affinity of the North Kara Terrane and Baltica favours an interpretation with the former being a direct continuation of the latter (Fig. 7a), the Vendian and younger strata being sourced from and deposited across Baltica's Timanide margin. In one part of October Revolution Island, along the east coast, at the deepest structural levels, granites from two localities have been reported to yield Late Precambrian ages (Proskurnin 1995); information about the sampling and analytical data have not been presented and the ages are therefore regarded as preliminary. However, they do point to the possibility that the Severnaya Zemlya Neoproterozoic successions of Riphean and Vendian turbidites and shales-phyllites may be interrupted by a major unconformity in the Vendian. This interpretation is supported by the evidence (Proskurnin & Shul'ga 2000; Lorenz et al. 2006b) that the oldest part of the

OROGENS IN THE EURASIAN HIGH ARCTIC

Fig. 7. Recent interpretations of the North Kara Terrane (NKT) in the Early Ordovician. (a) The NKT is an integral part of Baltica, as favoured by the present authors; (b) the NKT is a microcontinent, separated from Baltica by a spreading ridge and from Siberia by a major transform fault zone (Metelkin et al. 2005); (c) the NKT is separated from Baltica by a subduction zone and both are widely separated from Siberia (Cocks & Torsvik 2005)

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succession on October Revolution Island, the Nekrasov Series, beneath fossiliferous Lower Cambrian units, is divisible into two parts, a lower turbiditic unit and an unconformably overlying conglomerate-sandstone unit. Lorenz et al. (2006b) reported that a similar unconformity might exist on Bol'shevik Island. An unconformity on Severnaya Zemlya in the late Vendian, or earliest Cambrian (c. 540 Ma), as proposed above, would be closely associated in time with the well-defined late Vendian major unconformity in the central belt of Taimyr. This unconformity, as in the Timanides of the Timan-Pechora area, transgresses across a Neoproterozoic accretionary complex of ophiolites, fragments of Mesoproterozoic continental crust (intruded by c. 900 Ma granites; Pease et al. 2001b), various volcanic rocks, and siliciclastic flysch and molasse formations. A direct connection between the Baikalian accretionary complex in Taimyr and the Timanides would be a natural inference were it not for the generally accepted view that Baltica and Siberia were separated by a major ocean (Uralian) at least from the Ordovician to the Carboniferous. The question thus arises as to whether a wide Uralian Ocean extended into the high Arctic, as proposed by most previous workers (e.g. Zonenshain & Natapov 1987; Seng6r et al. 1993). As noted above, the faunal evidence from the Ordovician and Silurian successions is not compelling; a large number of taxa are common to both continents. Sobolevskaya et al. (1995, 1997, 1999) have described similarities in lithofacies and fauna between the Middle-Upper Cambrian successions of October Revolution Island, northwestern Taimyr and the central Taimyr belt (Kaban'kov et al. 2003). Cocks & Modzalevskaya (1997) drew attention to similarities between the Late Ordovician fauna of Baltica and southern Taimyr. However, even these similarities and correlations would allow continental separation in the later Palaeozoic. Other alternatives have been proposed; for example, Golonka et al. (2003) suggested oceanic separation of Baltica and Siberia in the Early Palaeozoic and Caledonian closure, the evidence for which is obscure. What then is the evidence that the Uralian Ocean extended from the northernmost Urals Mountains into the high Arctic? The Uralides are characterized over a distance of c. 2000 km by some of the world's most spectacular ophiolites (Savelieva & Nesbitt 1996), island-arc and back-arc associations (Brown et al. 2002), and (in the footwall) high-P glaucophane-bearing eclogites (Dobretsov & Sobolev 1984). These allochthons can be traced into the Polar Urals and to the coast of the southernmost Kara Shelf, and no further (Fig. 1). A continuation of the Uralides northwards and eastwards is defined by the NW-vergent Novaya Zemlya fold belt and the SE-vergent Taimyr orogen. The absence of both the characteristic ocean-derived allochthons and evidence of deep underthrusting of continental margin crust north of the Polar Urals can be explained by the limited exposures of the orogens between the Kara and Barents seas (note also that seismic data indicate the possible presence of a Uralian hinterland complex beneath the western margin of the Kara Sea; Sakulina et al. 2003). However, the lack of these main features of the Uralide orogen on Taimyr is more difficult to explain. Their absence between the northern and central belts on Taimyr would require that the North Taimyr Thrust (possibly with substantial transcurrent displacement; Inger et al. 1999; Metelkin et al. 2000, 2005) cuts out all these units (Fig. 7b). However, the low greenschistfacies grade of metamorphism of the Palaeozoic successions in the Taimyr Central Belt, in the footwall of the North Taimyr Thrust, testifies against the previous emplacement of thick overlying oceanderived allochthons. Likewise, the geology of Severnaya Zemlya provides no support for the influence of Uralian orogeny in this region (Lorenz et al. 2006b). Similarity between the orogen in northern Taimyr (Fig. 4) and the Uralide orogen has mainly been advocated on the basis of the presence of Late Carboniferous deformation, metamorphism and syn- to post-tectonic (c. 3 0 0 - 2 6 0 Ma) granites, the youngest

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of these being mainly in Taimyr's Central Belt. The granite-intruded Neoproterozoic to Cambrian turbidites of the Northern Taimyr Belt strike northwards into Bol'shevik Island, but here both stratigraphic (Proskurnin 1999) and new isotope age (Lorenz et al. 2007) data indicate that the folding and thrusting were earliest Carboniferous or earlier (pre-345 Ma) in age, unrelated to the Uralian orogeny. Furthermore, on Severnaya Zemlya, there is no evidence of a north-vergent deformation zone that would be expected to have developed, complementary to the Taimyr south-vergent fold-and-thrust belt. A reasonable alternative hypothesis to explain the relationships discussed above is that the Uralian Ocean did not exist in the region of today's high Arctic. Many Russian colleagues (e.g. G. Kovaleva & E. Korago, pers. comm.) regard the Eurasia Basin as a suitable analogue for the Uralian Ocean. Thus, the termination of the Uralian ocean-derived rocks to the north of the Polar Urals is compared with the abrupt termination of the Gakkel Ridge related Cenozoic oceanic lithosphere of the Eurasia Basin (Drachev et al. 1998) at the Laptev Sea shelf-edge. Against this interpretation that the Uralian Ocean did not continue northwards beneath the western Kara Sea (or at least, narrowed and eventually terminated NE of Novaya Zemlya), is the evidence provided by some studies of both palaeomagnetism and fauna provinciality (Cocks & Torsvik 2005, and references therein); these suggest that both Baltica and the North Kara Terrane were separated substantially from Siberia in the Palaeozoic (Fig. 7c). Clearly, the different lines of evidence, influencing interpretations of Baltica (including the North Kara Terrane) and Siberia, remain controversial; more work is necessary to further explore the inconsistencies. It is particularly desirable, when interpreting terranes as independent (micro)continents, based on palaeomagnetic and/or biogeographical evidence, that the terrane boundaries can be defined by shelf-edge, slope and rise lithofacies, and sutures. The North Kara Terrane does not provide this evidence and its independence from Baltica is denied by the regional magnetic data (see Figs 1 and 6).

C o m p a r i s o n with the w e s t e r n m a r g i n o f Baltica

The synchroneity of the Timanian orogeny along the eastern margin of the East European Craton (EEC) and the Cadomian orogeny along the western margin has led some workers (e.g. Puchkov 1997; Scarrow et al. 2001) to propose that the two were physically related, one being a continuation of the other and both being part of the active margin of Gondwana in the Neoproterozoic. However, the Timanide orogen along the northeastern margin of the EEC is truncated by the Caledonide orogen of Scandinavia, within which there is a clear record of an extensional margin during the Neoproterozoic (Kumpulainen & Nystuen 1985; Siedlecka et al. 2004). The Scandinavian Caledonides meet the Caledonian terranes (Strachan & Dewey 2005) of the classical British Caledonides beneath the North Sea (Pharaoh et al. 2006). Along the western margin of Baltica, evidence from the Trans-European Suture Zone is unambiguous (Katzung et al. 1993). Cadomian-deformed Avalonian terranes of northern Germany and Poland were emplaced northeastwards onto the craton margin in the mid-Palaeozoic. These terranes apparently obtained their Neoproterozoic tectonothermal imprint along the Gondwana margin (Pharaoh et al. 2006; Winchester et al. 2006), prior to the closure of the Tornquist Sea. Further SE along this margin of the East European Craton, in the area of the Holy Cross Mountains and southwestwards towards Brno, Lower Palaeozoic successions with affinities to Baltica (Cocks 2002) are overlain by Old Red Sandstones containing clasts of Vendian granites (unpublished Pb/Pb zircon dating by F. Hellmann, 1996). Transgression of a Vendian complex

(perhaps with ophiolites, near Brno) by Cambrian successions with Baltica affinities indicates that accretion of some Vendian terranes to southwestern Baltica occurred prior to the Cambrian, perhaps contemporaneous with the evidence of volcanicity dated on the Lublin slope to c. 550 Ma (Compston et al. 1995). Further SE along Baltica's southern margin (Winchester et al. 2006), the occurrence of ophiolites and associated oceanderived units has been related to the emplacement of Cadomian accreted terranes, but the evidence is not unambiguous. In these parts of Baltica's southwestern margin, there remains the possibility that, as in the Timanide orogen, Vendian accretion occurred prior to rifting and development of an Early Palaeozoic passive margin. Continuation of these accretionary complexes eastwards, via the Scythian Platform to the Southern Urals, has been proposed by Stephenson et al. (2006). Accretion of Avalonian terranes with the closure of the Tomquist Sea (Cocks & Fortey 1998) occurred substantially later during the 'Caledonian' orogeny.

Conclusions Recent work on the high Arctic of Eurasia has indicated the following. (1) The Caledonides of northeastern Greenland, prior to the Cenozoic, extended northwards directly into the Svalbard Caledonides, forming the Laurentian flank of the orogen. (2) The Barentsian Caledonides dominate the basement of the western Barents Shelf, with suture zones, separating Baltica from Laurentia, located between Franz Josef Land and KvitCya. (3) The eastern Barents Shelf is dominated by the Timanide orogen, comprising the northern part of the Early Palaeozoic palaeo-continent Baltica. This orogen probably contains fragmented 'Grenvillian' complexes. (4) The North Kara Terrane was an essential component of the Timanide margin of Baltica, which continued northwards (present-day coordinates) an unknown distance across the Lomonosov Ridge and the Arctic Basin. The shape of Baltica in the Early Palaeozoic is not known; it was certainly strongly elliptical, not round as generally portrayed in palaeogeographical reconstructions. (5) The Uralian orogeny does not influence Severnaya Zemlya, and the north Taimyr thrust, separating Baltica (the North Kara Terrane) from the Siberian margin, was probably not a suture zone. (6) A wide (>_1000 km) Uralian Ocean, fragments of which are now preserved in the Urals from the far south to the Arctic, narrowed northwards, and may have terminated beneath what is now the western edge of the Kara Sea Shelf. This implies that the Uralian Ocean opened by rotation, in a way similar to that for the Cenozoic Eurasia Basin. (7) Relationships between Baltica (the North Kara Terrane) and Siberia, particularly as they are exposed in northern Taimyr, deserve closer field investigations and further analysis. (8) The Timanide evolution of northeastern Baltica (presentday coordinates) is similar to the Neoproterozoic evolution of parts of southwestern Baltica, in terms of both timing and the character of the Neoproterozoic accretionary processes. This paper is based on many years of international collaboration in the high Arctic, as part of EUROPROBE's TIMPEBAR project and with the support of INTAS, through three projects: HALE (No. 96-1941), Taimyr (No. 97-1139) and NEMLOR (No. 01-0762). We thank V. Vemikovsky for arranging 4~ analyseson detrital nmscovites, which were performed using a Micromass 5400 static mass spectrometer at the Institute of Geology and Mineralogyin Novosibirsk, Russia. We thank R. Cocks, R. Stephenson and T. Torsvik for helpful commentson the manuscript.

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Crustal structure and tectonic model of northeastern Baltica, based on deep seismic and potential field data S. K O S T Y U C H E N K O 1, R. S A P O Z H N I K O V 2, A. E G O R K I N 1, D. G. G E E 3, R. B E R Z I N 2 & L. S O L O D I L O V 1

iV. V. Fedinsky Center GEON, 32a, Marshala Tukhachevskogo, Moscow, 123154, Russia 2FGUP Spetsgeofizika, Povarovka, Solnechnogorsky region, Russia 3Uppsala University, Villaviigen 16, SE-752 36 Uppsala, Sweden

Abstract: The Early-Mid-Palaeozoic successions of the Pechora Basin along continent Baltica's northeastern margin unconformably overlie Neoproterozoic complexes of the Timanide orogen; further west, Palaeozoic platform strata cover the East European Craton basement. The latter is exposed in the northern parts of the Fennoscandian (Baltic) Shield and is also present beneath the Neoproterozoic (perhaps also Mesoproterozoic) and younger successions of the Mezen Basin. Geophysical data help define the crustal and upper mantle structures of this part of northwestern Russia, from the Kola Peninsula to the Uralide orogen. Within the Shield, seismic refraction and near-vertical reflection profiling allow the terranes of Archaean and Palaeoproterozoic age, defined at the surface (the Murmansk, KolaLiinakhamary, Central Kola, Keivy and Belomorian terranes, and the Pechenga-Imandra-Varzuga and Lapland-Kolvitsa Granulite Belts constituting the Lapland-Kola orogen, and the West Karelian, Central Karelian and Vodlozer terranes forming the Karelian province), to be traced into the deeper crust, promoting interpretation of various early Precambrian compressional and extensional regimes. Thrusting, transcurrent and extensional faulting have been identified in a crust that has been partly thickened during late Archaean and Palaeoproterozoic collision and locally extended both in back-arc basin settings and during syn-post-collisional collapse. The main Fennoscandian terranes can be traced southeastwards beneath the Mezen Basin based on gravity and magnetic data. Seismic profiling shows that the Phanerozoic successions of this basin overlie intracratonic rifts of Neoproterozoic and possibly Mesoproterozoic age. The rifts are composed of a number of half-graben, the details of which are well defined by vibroseis reflection profiling. During the development of Mezen Basin rifts, a Neoproterozoic passive margin was established in the vicinity of what is now the Timan Range. Late Neoproterozoic (Vendian) orogeny emplaced slope-rise turbidite-dominated successions southwestwards onto the platform. Towards the hinterland of the Timanides (also referred to as Baikalides), beneath the Phanerozoic Pechora Basin, a combination of potential field and seismic surveys, both shallow reflection and deep refraction profiling, have allowed analysis of the crust and uppermost mantle. Neoproterozoic magmatic arcs, Grenville-age(?) microcontinents and possible oceanic domains are inferred to exist beneath the early Palaeozoic unconformity. A vast region of Neoproterozoic accretion reaches from the western Urals, northwards beneath the Pechora Basin and the eastern Barents Shelf.

Northwestern Russia, from the Fennoscandian (Baltic) Shield and Mezen Basin, eastwards across the Timan Ridge and Pechora Basin to the Urals and northwards into the Barents Sea (Fig. 1), is a region of considerable economic importance, with mineral resources, including hydrocarbons, metals and diamonds. The bedrock is dominated in the west by the northeastern part of the East European Craton (EEC) with its Archaean and Palaeoproterozoic core cropping out in the Fennoscandian (Baltic) Shield. Defining the possible continuation of these old Precambrian complexes southeastwards beneath the Mezen Basin, and also eastwards towards the Urals and northwards beneath the Barents Shelf is of fundamental importance for understanding the younger (Neoproterozoic and Phanerozoic) evolution of Northern Europe. Accretion of Timanide (Baikalide) lithosphere to the craton margin in the Vendian influenced a wide region, and included both juvenile crust and microcontinents. Drilling (generally not to more than c. 5 km depth) provides local evidence of the character of this bedrock, but the main foundation for understanding the crustal evolution of all these areas that are covered by platform successions rests on the geophysical database, dominated by potential field surveys and seismic profiling. The tectonic evolution of this northern part of the EEC and its Neoproterozoic margin was given priority by E S F ' s E U R O P R O B E programme (Gee & Zeyen 1996). Two projects, SVEKAL A P K O and TIMPEBAR, focused on the Shield and the younger accreted terranes, respectively. In this volume, three papers concern the pre-Neoproterozoic evolution of the Fennoscandian Shield (Daly et al. 2006; Korja et al. 2006; Slabunov et al. 2006) and one (Stephenson et al. 2006) treats the Palaeozoic history. A recent memoir (Gee & Pease 2004) has provided an extensive treatment of the Neoproterozoic evolution of the Timanides and development of the Early Palaeozoic passive margin of this northeastern part of continent Baltica.

This paper presents much of the geophysical evidence for the composition and structure of the crust of northwestern Russia. It relates this evidence of deep structure to that of surface geology and deep drilling and presents interpretations of the tectonic evolution.

Geophysical framework: overview of the database The geophysical data include regional gravity and magmatic maps of northwestern Russia and seismic profiling, particularly wide-angle reflection and refraction deep soundings (so-called DSS profiles), and near-vertical reflection profiling (common midpoint; CMP), using vibrators for the source and recording to 2 0 - 2 5 two-way travel time (TWT). Integration of potential field and seismic data has been essential for the crustal modelling and tectonic interpretations.

Potential f e l d data Magnetic and gravity (Bouguer anomaly) maps of northwestern Russia provide fundamental information on the structure and composition of the crust. In the area of interest, state magnetic anomaly maps (Fig. 2) and state Bouguer gravity anomaly maps (Fig. 3) are available at scales of 1:200000, 1:1 0 0 0 0 0 0 and 1:2 500 000. Wybraniech et al. (1998) published a colour-shaded gravity anomaly map of Europe using the gravity reference system GRS1980. Korhonen et al. (1999) compiled uniform potential field data in digital form to make both magnetic and Bouguer anomaly (density 2.67 g cm -3) gravity maps of the Fennoscandian Shield. Traditional quantitative interpretations of the fields as well as 2D and 3D modelling of the sources of the anomalies have been made to improve the geophysical database for the area treated here. Buyanov et al. (1989, 1995), Glaznev

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 521-539. 0435-4052/06/$15.00

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Fig. 1. Index map of northwestern Russia showing the principal geological provinces of northeastern Baltica.

et al. (1989), Mints et al. (1996) and Glaznev (2003) have provided interpretations of the Fennoscandian Shield based on the potential field data. Berezovsky et al. (1993) and Kostyuchenko et al. (1999) presented several 2D gravity models of the crust along the wide-angle refraction and reflection (DSS) profiles for the Mezen Basin, and Kostyuchenko & Romanyuk (1997) also created a 3D gravity model for the northern segment of the Mezen Basin. Ismail-Zadeh et al. (1997) published a gravity model for the northern part of the Pechora Basin including calculations of the influence of the sedimentary cover on the gravity field, and the depth to the Moho boundary; they also defined internal discontinuities within crystalline crust.

Three-dimensional modelling of the magnetic sources in the area of interest has been carried out in the V. V. Fedinsky Center GEON, based on an algorithm developed by Bhattacharyya (1966). Kostyuchenko et al. (1999) presented the results of this modelling for the Mezen Basin. Magnetic bodies, each defined by a homogeneous magnetic signature and simplified shape, were used for this modelling. Only induced magnetization has been assumed and the values of magnetic susceptibility have been chosen to give the best fit to the measured magnetic field by application of a 'calculation, comparison and adjustment' method. High-gradient contours of magnetic anomalies were used to determine the plan-view shape of each magnetic body. Estimations of the depth to isolated magnetic sources in the 'magnetic basement' were obtained, by direct examination of the shapes of magnetic anomalies and application of 'Pyatnitsky's tangent technique' (Exploration Geophysics 1964). The magnetic map (Fig. 2) shows the complex structures of the Palaeoproterozoic and Archaean basement, from the well-exposed Fennoscandian Shield in the NW to the regions covered by younger Proterozic and Phanerozoic successions in the SE. A clearly defined, NW-trending boundary separates regions where the Early Proterozoic and Archaean complexes are near the surface from the thick turbidite-dominated successions of the Timanide orogen (Gee & Pease 2004) in the Timan Range and Izhma Zone. These non-magnetic successions are flanked to the east by an equally prominent boundary to a highly magnetic NW-trending belt beneath the Pechora Basin that is known from drillcores to be dominated by volcaniclastic successions and a Neoproterozoic magmatic complex. This 50-100 km wide linear zone gives way eastwards to a region with a variable 'cellular' magnetic signature. All these anomalies, related to the Neoproterozoic basement flanking the northeastern margin of the East European Craton, are abruptly truncated further east by the Uralide orogen, the strong north-trending Uralian magnetic highs being related to the Uralian ophiolites and island-arc volcanic rocks.

Fig. 2. Magnetic anomaly map of northwestern Russia (from Petrov et al. 2004a). Epoch T, 1965 (model VSEGEI). (For abbreviations see Figs 6 and 11.)

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Fig. 3. Gravity map (Bouguer reduction 2.67 g cm-3) of northwestern Russia (from Petrov et al. 2004b). Abbreviationsare as in Figures 6 and 1l.

Seismic data

For the area of study, there are many publications, mostly in Russian, in which seismic data are presented; however, many data remain unpublished and are available only in reports of state and industrial companies. Figure 4 provides the location of the main seismic profiles over the Russian part of the Fennoscandian Shield and Figure 5 does the same for the area of the Mezen Basin, Timan Range and Pechora Basin. State companies from the former Soviet Union, Spetsgeofizika and the Leningrad Mining Institute, carried out the first crustal-scale seismic studies of the Fennoscandian Shield in 1958. This and most of the other field experiments, until 1980, were based on wide-angle reflection and refraction seismic methods which, in combination, were known as deep seismic sounding (DSS). So-called 'continuous' DSS profiling has been applied, in which analogue seismic stations were moved along the profile to obtain a seismograph spacing of 7 5 - 1 0 0 m; shot point spacing varied from 2 0 k m to 75 km (Litvinenko 1963, 1984). In the early 1980s, the arrangement of the land array consisted of links of the profile about 5 km long spaced every 2-4.8 km, in which the recorder spacing was 200 m. Shot-points had intervals from 35 km to 80 km and were augmented by industrial explosions from local quarries. For most of the above-mentioned profiles, the most prominent arrivals are refractions from the Moho and wide-angle reflections from velocity discontinuities in the crust. From 1973 to 1994, the Geological Institute of the Kola Science Center of the Academy of Sciences (RAN) carried out many experiments based on the registration of industrial explosions from quarries (Sharov 1993) and succeeded in outlining the Moho depth on a regional scale. In the period from 1971 to 1990, the GEON Center carried out DSS surveys across the Fennoscandian Shield, Mezen Basin, Timan Range and Pechora Basin. The six-channel seismometers, with an analogue recording system, were spaced every 8 12 km, and three components of ground displacement (vertical,

radial and tangential) were recorded. Chemical shots were located in the line of the profiles at 3 0 - 6 0 km intervals, the charges ranging from 2 to 5 tonnes. Underground nuclear explosions (PNEs) were used in the Murmansk-Kizil profile (6 in Fig. 5). During 1998-2000, the GEON Center carried out DSS investigations in the area of the Fennoscandian Shield using digital recorders (Alfa-GEON and Delta-GEON) spaced every 2 - 3 km. For most of the shots, major compressive (P) and shear (S) seismic wave phases in a frequency range of 1.520.0 Hz were recognized. First arrivals for P and S waves were identifiable on trace-normalized records within a shot pointreceiver offset of 200-300 km. The first arrivals, the seismic wave phases Pg (refracted through the basement within the Mezen and Pechora basins), were well observed beyond about 20 km from the shot point (SP). Pn waves (refracted through the uppermost mantle) were sometimes seen beyond 175-200 km. PmP and SmS (wide-angle reflections from the Moho), and wide-angle reflections from the reflectors in the crystalline crust (PcP and ScS) are also prominent. A combination of frequency and r.m.s. velocity filtering (Dix 1965) in the t - x domain and velocity filtering developed in the GEON Center (Egorkin et al. 1985; Egorkin 1998, 1999) was used to improve the correlation of the reflected phases and the signal-to-noise ratio of the DSS data. The software was developed in the GEON Center, and the processing sequence included the following steps: (1) digitization of the analogue data; (2) band-pass frequency filtering; (3) normal move-out corrections; (4) constant r.m.s, velocity stacking for each offset-limited sub-gather. In addition to this processing sequence, a summing of traces with different r.m.s, velocities was applied, mainly to enhance reflections with time-distance curves that differ from the hyperbolae. Application of r.m.s, velocity filtering reveals both horizontal and low-angle dipping reflections. Interpretation of the seismic data has been carried out using 2D forward modelling as well as the methods developed by Zelt & Smith (1992).

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Fig. 4. Location of regional seismic profiles across the eastern part of the Fennoscandian Shield. Numbers indicate the profiles referred to in the text: 1, Murmansk-Khibiny; 2, Kalevala (Uchta-Kem; 3, KostomukshaSemipalatinsk (western part); 4, Pechenga-Lovno; 5, Keivy. Finland is shown as a light grey area. Ladoga and Onega are lakes.

Fig. 5. Location of seismic profiles across the Mezen Basin, Timan Range and Pechora Basin. Numbers indicate the profiles referred to in the text: 1, White Sea-Vorkuta; 2, Dvina Guba-Mezen river; 3, CMP experiment; 4, 22-PC; 5, 27-PC; 6, PNE profile. Principal tectonic units are as in Figure 11. Other explanation is as in Figure 4.

Converted wave data, derived from registration of remote earthquakes, supplement most DSS sections collected by the GEON Center. In addition, during 1976-2002, Nevskgeologia acquired several lines, using the converted wave method in the Kola Peninsula (Pechenga-Alarechian area) and southern parts of the Karelian region. Depths to Moho and major crustal discontinuities were the main features derived from these studies. Near-vertical reflection profiling, based on the 'common midpoint' seismic method (CMP), began in the Kola Peninsula in the beginning of the 1990s. In the northwestern part of the area under discussion, the Russian state geophysical company EGGI carried out the N o t t a - S a l m i y a r v i and L o t t a - V e r k h n e t u l o m s k i y lines. The record length was 15 s TWT. In 1992, an international group of geophysicists from Russia, USA, Norway and Scotland measured a 38 km long 25 s T W T line across the Kola super-deep borehole (Ganchin et al. 1998). During 1997-2000, Spetsgeofisika completed a long CMP profile, 1-EB, which transects the area of the Fennoscandian Shield, from the Kola borehole in the north to Petrozavodsk in the south, and a second profile from Kalevala (Uchta) to Kern, which stretches from the boundary between Russia and Finland in the west to the White Sea in the east (Berzin et al. 2001, 2002). In 2 0 0 0 - 2 0 0 2 , Spetsgeofizika conducted a vibroseis CMP experiment across the western and central parts of the Mezen Basin and along a line across the Timan Range. The total length of the available CMP profiles (both state and industrial) is about 2000 km. Figures 4 and 5 show only the statefinanced CMP lines, which are considered in this review. I n p u t Output-2 and Sercel (SN-388) seismic stations were used for the profiling. The maximum source-receiver distance was 10 kin. Groups of 12 geophones were spaced every 44 m. Four or five,

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA 10 tonne vibrators spread over a base of 50 m were used for the profiling. The record length was 25 s TWT and a 50-fold stack was obtained. Both standard processing on the basis of ULTRA-SPARC-2 and PROMAX systems as well as special processing by differential stacking (MDS) were used during the study. With these methods, many seismic images of the crust were obtained, which could be correlated with geological structures at the Earth's surface. In the CMP images, it has usually been possible to define the lower boundary of a reflective lower crust. Combination of CMP and DSS data along the Kalevala (Uchta)-Kem profile (Berzin et al. 2001, 2002) allowed a better resolution of the geological nature of seismic discontinuities. Across the Mezen Basin, Timan Range and Pechora Basin, the state geophysical companies from the former Ministry of Geology of USSR, mainly Zapadny Geophysical Trust and Pechorageofizica, carried out geophysical studies using conventional reflection and refraction seismic methods (Fig. 5). The main goals of the old conventional refraction measurements were to control the thickness of the sedimentary cover and the depth to surface of the underlying basement. During these studies, 4 8 - 6 0 channel stations, were used, usually spaced every 100-200 m. The seismograms were recorded on photographic paper by the wiggle-trace method. Shot-point spacing was from 10-20 km to 5 0 - 7 0 k m and shooting, as a rule, was direct and reversed. Refraction observations were shot with maximum offset up to 100-150km. A travel-time method, with plane-layered velocity models, was often applied to determine the depth variations of recorded refractors. The early reflection observations and subsequent shallow nearvertical profiling, which came into use in the discussed area after 1977, were carried out mostly in the Pechora Basin, with the aim of studying the structure of the sedimentary cover. The reflections were recorded for 2 - 3 s TWT (only in few cases for 4 - 5 s TWT). Initial 6-24- and later 48-fold stacks were acquired during these

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old experiments. Since the end of 1980s, several regional profiles (e.g. 15-PC, 19-PC) were acquired by Pechorageofizika (subsequently renamed Severgeofizika). In some short portions of the profiles, the recording time was extended to 10 and even 20 s TWT.

Fennoscandian Shield: crustal structure of key domains The Russian part of the Fennoscandian Shield (Fig. 6) can be subdivided into the Lapland-Kola orogen and Karelian provinces (Mints et al. 1996). The Lapland-Kola orogen comprises Archaean terranes (Murmansk, Kola-Liinakhamary, Central Kola, Keivy, Tersk and Belomorian) and Palaeoproterozoic belts (Lapland-Kolvitsa Granulite and Pechenga-Imandra-Varzuga). All the terranes are bordered by major fault boundaries. To the SW, the Karelian composite province is subdivided into West Karelian and Central Karelian domains of late Archaean age (2.7-2.8 Ma), and the oldest (3.5-2.85 Ma) Vodlozer domain (Lobach-Zhuchenko et al. 2000), in the southernmost part of the Shield near Lake Onega (Fig. 6). Several narrow Archaean greenstone belts, which generally strike approximately north-south, occur within and in the boundaries of the Karelian domains (Rybakov & Golubev 1999). In the eastern part of the Central Karelian domain, and just to the east of the Vodlozer domain, volcanogenic-terrigenous Palaeoproterozoic rocks occur (from NW to SE in the Karasyok, Sala-Sodankyulya, Pana-Kuolayarva, Shombozersky, Vetrenny Poyas and Shardozersky units) constituting the East Karelian part of Karelian province. The evolution of most of these units is still under discussion. Based on the results of this study, presented below, at least the Shombozersky,

Fig. 6. Index map of northeastern Fennoscandian Shield showing principal tectonic units and geologicalprovinces (compiled and modifiedfrom Mints et al. 1996; Lobach-Zhuchenkoet al. 2000). Bold straight lines show profilesreferred to in the text: 1, Murmansk-Khibiny; 2, Kalevala (Uchta)-Kem; 3, KostomukshaSemipalatinsk.

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Vetrenny Poyas and Shardozersky units are inferred to be palaeorifts. From reflection, refraction and converted wave seismic data, the crustal structure of the Fennoscandian Shield is seen to be strongly layered, consisting of several complexes of different composition and age. Interpretation of the deep seismic data leads to the conclusion that these tectonic units were generated and then interacted and were reworked during compressional and extensional tectonic regimes with repeated reactivation. The compressional history is proved by crustal shortening and stacking of thrust sheets. Figure 7b illustrates a profile across the Kola Peninsula, where the Murmansk and Kola-Liinakhamary upper crustal slices have been overthrust, from north to south, onto the Central Kola terrane. The northern part of the latter has been involved in the south-directed displacements. Based on interpretation of geological and geophysical data, Mints et al. (1996) argued for a model in which the Keivy terrane, in Palaeoproterozoic time, was thrust southwards beneath the northern part of Imandra-Varzuga Belt. Accordingly, the near-vertical fault between Keivy terrane and Central Kola terrane in Figure 7b can be interpreted as a major strike-slip fault. The latest Palaeoproterozoic stage of compression in the evolution of the Karelian province is illustrated (Fig. 8a) by a CMP cross-section from Kalevala to Kem (Mints et al. 2001; Berzin et al. 2002). Additional evidence of crustal shortening was obtained from reprocessing of old DSS data collected in 19581959 along the same line. Figure 8a shows the velocity model of the crust and location of wide-angle reflection seismic boundaries

in the upper and middle crust, which are gently inclined eastwards. Figure 8b presents the interpretation of both the DSS data and the CMP images, and suggests that the Belomorian terrane was overthrust from east to west onto the Shombozersky unit. The Shombozersky, Vetrenny Poyas and Shardozersky rifts are inferred to have evolved during Palaeoproterozoic extension and are represented by metamorphosed volcano-sedimentary successions in the uppermost crust (Sokolov et al. 1987; Rybakov & Golubev 1999). Based on interpretation of only the CMP data along the K a l e v a l a - K e m profile, Mints et al. (2001) proposed a model in which the Shombozersky units were a package of eastdipping tectonic slices between the Central Karelian granitegreenstone terrane and the Belomorian terrane. Subsequent combination of CMP and DSS data, as presented in Figure 8a, led to the recognition of the rift character of the Shombozersky structure. On the CMP profile, the seismic pattern from 5 to 8 s TWT (i.e. equivalent to a depth of 15-25 kin) is transparent; it coincides with velocity values of 6.5-6.6 km s -a in the middle crust. Together, these data suggest that a magma chamber exists beneath the volcanic and terrigenous rocks of the Shombozersky complex. A positive gravity anomaly occurs over this area. The signatures of upper crustal rifting in the Shardozersky unit, as seen in Figure 9a and b, can be traced into the deeper crust in the DSS cross-section. Based on these data the volcanic and sedimentary rocks of the Shardozersky rift fill a 5 km deep graben, beneath which there are increased seismic velocities in the middle and lower crust. The difference between the P-wave velocities in the rift and in the surrounding crust reaches

Fig. 7. Deep seismic cross-section (a) and interpretation (b) of the Lapland-Kola orogen along the Murmansk-Khibiny DSS line (1 in Figs 4 and 6). Velocity model in (a) after Zolotov et al. (unpubl. data). Numbers above the sections are selected recording stations. Faults are inferred from geological and seismic data. Colour in (b) emphasizes the major tectonic units. KhM, Khibiny massif; LVZ, low-velocity zone. Other abbreviations are as in Figure 6.

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA

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Fig. 8. Deep seismic cross-section (a) and interpretation (b) of the Karelian province along the Kalevala (Uchta-Kem line (2 in Figs 4 and 6). Seismic pattern is from Berzin et al. (2001, 2002).

1.0-1.5 km s - 1, and for S-wave values it reaches 0 . 2 - 0 . 3 km s- t These differences suggest that basic magmatic rocks dominate the crust beneath the rift. At depths between 30 and 41 km (Moho boundary), the P-wave velocity is about 7.4 km s -1 and S-velocity varies from 4.37 to 4.41 km s -1, suggesting that part of a rift-pillow, or crustal-mantle mixture, may be present beneath the rift.

To map the depth to the Moho boundary (Fig. 10) beneath the Fennoscandian Shield in Russia, data were used from DSS, CDP and converted wave profiles. Comparison was made of the available seismic data from different lines, based on arrival-time diagrams and the velocity structure of the crust. During this work, several parts of existing lines were recalculated and all available old data, for example from the P e c h e n g a - L o v n o profile (4 in

Fig. 9. Deep seismic cross-section (a) and interpretation (b) of the Karelian province along the Kostomukha-Semipalatinsk DSS line (3 in Figs 4 and 6). Velocity model in (a) after Egorkin et al. (unpubl. data). Abbreviations are as in Figure 7.

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Fig. 4) and Keivy profile (5 in Fig. 4), were taken into account. Because of the irregular distribution of the data, no graphical software was used to construct them. Figure 10 shows that within the Shield the depth to the Moho boundary varies from 36 to 52 km. There is no clear correlation between these depths and most of the geological domains. Thus within the Kola-Lapland orogen, only the Keivy terrane clearly coincides with a shallow Moho. Beneath the Lapland Granulite Belt and neighbouring northern portion of the Belomorian terrane the Moho deepens to more than 44 km. In the area between latitude 66~ and 67~ a c. 50 km deep depression of the Moho exists beneath the southern part of the Belomorian terrane and northern margin of the Karelian province. This depression connects westwards on the Finnish part of the Shield with a Moho depth of more than 45 km depth (e.g. beneath the Karasjok-Kitila Greenstone Belt along the Polar Profile and parts of the Fennolora profile (Luosto 1990). Moho depths of 3 8 - 4 0 km dominate within central and northern parts of the Karelian province, whereas 4 2 - 4 4 km depths are characteristic of southwestern parts of the province. These data suggest that only the Keivy terrane, with its Archaean accretionary history (Mitrofanov & Bayanova 2000), has retained its original crustal thickness, whereas the roots of other parts of the Shield have been changed during the subsequent evolution.

Fig. 10. Depth to Moho (thickness of the crust) within the northeastern part of the Fennoscandian Shield. Data for Finland are from Luosto (1990).

M e z e n Basin: crustal architecture and tectonic model The Mezen Basin is a c. 5 0 0 k m wide, NW-trending trough extending c. 1000 km along the northeastern margin of the East European Craton from the Kola Peninsula to the Sisola Dome (Fig. 11). The underlying basement is inferred to be comparable

Fig. 11. Index map showing the major geological and tectonic units in the MezenTiman-Pechora region. Major faults: WTF, West Timanian fault; CTF, Central Timanian fault; PFZ, Pre-Pechora fault zone. Graben: I-I, Onega-Dvina; II-II, LeshukonaPinega, III-III, Safonov.Numbered outcrops of Neoproterozoic successions within the Timan Range: 1, Kanin; 2, North Timan; 3, Chetlasky; 4, Vimsky; 5, Ochparma; 6, Dzhezhimparma. Tectonic units within the Pechora Basin: 7, Chernyshev Range; 8, Chernov Range. Square frame shows the area of Figure 12.

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA

Fig. 12. Map of depth to basement beneath the Mezen Basin within the area shown in Figure I 1. Half-grabenin the Leshukona-Pinega rift are numbered: 1, Pinega; 2, Kuloy; 3, Kimzha; 4, Ust-Mezen; 5, Karpogory; 6, Mid-Pinega; 7, Mid-Dvina, located in the Onega-Dvina rift; 8, Safonov rift.

with that of the Fennoscandian Shield, exposed to the NW. Because only a few drillholes penetrate the entire sedimentary fill to reach the basement, geophysical data, particularly magnetic anomalies (Gafarov 1963, 1985) are the primary source for interpreting the structure and composition of this basement.

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Dedeev et al. (1976), Kratz et al. (1979), Bogdanova (1993) and Bogdanova et al. (1996) also took into account borehole data, gravity anomalies and, where possible, seismic profiles to improve our understanding of the geological and tectonic setting of the basement. Most investigators accept that the major terranes and tectonic zones of the Fennoscandian Shield have their prolongation southeastwards beneath the Mezen Basin basement. Based on analysis of P-wave velocity values in the top of the basement and on interpretation of deep seismic data along DSS profiles, Kostyuchenko & Egorkin (1994) and Kostyuchenko (1995, 1998) defined several terranes (e.g. the North Mezen, Central Mezen and Plisetsk) to constitute the crystalline crust beneath the Mezen Basin. This was supported by the results of 3D gravity modelling of the crust (Kostyuchenko & Romanyuk 1997). Within the area of the Basin, the magnetic anomalies also provide evidence of internal features within the basement. Several negative anomalies reflect the location of graben in the basement, whereas local positive anomalies indicate the presence of mafic magmatic bodies that intruded into the basement along the edge of the graben (Kostyuchenko & Romanyuk 1997). Deposition of the sediments in the Mezen Basin is thought to have begun in the latest Mesoproterozoic (c. 1000 Ma). Three major stratigraphic units are readily distinguishable, which characterize the epi-cratonic sedimentation in the area. They are (from base upwards): (1) Late Mesoproterozoic(?) to Devonian siliciclastic successions deposited in continental and shallow marine environments; (2) Carboniferous to Early Permian carbonate-evaporite shallow marine and lagoonal strata; (3) Late Permian to

Fig. 13. Interpretion (a) of part of a seismic cross-section, by Sapozhnikov et al. (2003) along line 3 in Figure 5. The stages of graben formation in the Leshukona-Pinega rift are shown schematically in (b). The units interpreted as correlative from graben to graben are in the same symbols and assigned the same letter (A, B, C and D) in the diagrams. Numerals for the half-graben are the same as in Figure 12. (See text for further explanation.)

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Cenozoic, predominantly siliciclastic, shallow marine and continental strata. Three major NW-trending graben dominate the Mezen Basin, referred to as the Onega-Dvina, Leshukona-Pinega and Safonov rifts (Fig. 11). In 1985, these structures were identified by deep seismic wide-angle reflection and refraction experiments carried out by the GEON Center (Kostyuchenko 1995; Kostyuchenko et al. 1999). In 1999-2002, Spetsgeofizika, based on vibroseis CMP data, discovered that several of the rifts contained successions up to 10 km thick within their deeper parts (Sapozhnikov et al. 2003). Figure 12 shows the depth to the basement in the Mezen Basin. The isolines were drawn on the basis of a combination of borehole and seismic data (DSS, conventional and CMP) and interpretation of gravity and magnetic anomalies. A series of half-graben have been discovered, within which the depth to basement varies from 2 - 3 km to 10-12 km. The Leshukona-Pinega rift is composed of the Pinega, Kuloy, Kimzha, Ust-Mezen, Karpogory and Mid-Pinega half-graben (numbered 1 - 6 in Fig. 12). This major rift zone is up to 150 km wide and more than 300 km long. The Mid-Dvina graben of the Onega-Dvina rift is located 120 km to the west (7 in Fig. 12). Within the eastern portion of the Mezen Basin, the 4.5 km deep Safonov rift (8 in Fig. 12) has been proposed based on older data (Kostyuchenko 1995; Bogatsky et al. 1996; Bogdanov et al. 1996; Kostyuchenko et al. 1999). In general, SW-dipping faults control the eastern flanks of the graben in the area between 42~ and 45~ and north of latitude 64~ In the area close to 64~ and to the south, the faults are inclined both northeastwards and southwestwards. High-resolution CMP data image the detailed structure of the Leshukona-Pinega rift (Fig. 13). Figure 13 shows part of an interpreted CMP cross-section through the Kuloy and Kimzha graben (2 and 3 in Fig. 12) and also displays a reconstruction showing the different stages of graben formation from the Mezoproterozoic(?) to the Vendian. We infer that mostly clastic sedimentary rocks fill the graben and several distinct units can be traced in the sections. The units interpreted as correlative from graben to graben are assigned the same letter on the diagrams ( A - D in Fig. 13). Parallel reflectors dominate in the upper portion of each unit and indicate more widespread sedimentation, whereas onlap-like reflectors in the base of the units indicate the active stage of graben formation accompanied by faulting. There are, from bottom to top, four units (A, B, C and D) in the Kuloy graben and only three (B, C and D) in the Kimzha graben. We also recognize four units in the Pinega graben (1 in Fig. 12) and only the uppermost two (C and D) units in the Ust-Mezen graben (4 in Fig. 12). Located close to the CMP line, the Ust Nyafta G-1 and Middle Nyafta G-21 drillholes penetrated the entire Phanerozoic sedimentary cover of the Mezen Basin and reached the middle Neoproterozoic strata of the Ust-Mezen graben. These drillholes show that m i d - l a t e Neoproterozoic shallow-marine siliciclastic successions occupy the uppermost part of this graben; the age of the deeper sedimentary successions is not known. Late Vendian strata (Grazhdankin 2004), sourced from the Timanide orogen to the NW, unconformably overlie the older successions. The data presented above give evidence of extensional faulting which may include major subhorizontal detachments as well as fault wedges affecting the entire crust. The normal faults apparently migrated laterally eastwards through the Neoproterozoic, step by step, from the Pinega and Kuloy, through the Kimzha to the Ust-Mezen graben. Kostyuchenko & Egorkin (1994) and Kostyuchenko et al. (1999) recognized a high-velocity (7.0-7.1 km s -1) and highdensity (about 3.0 g cm -3) body in the middle crust of the Leshukona-Pinega rift. This body is located beneath a c. 100 reGal positive gravity anomaly in the northern part of the rift. Kostyuchenko & Romanyuk (1997) estimated that the highdensity body contributes about 50% of the value of this positive

Fig. 14. Seismic cross-section across the Mezen Basin along the Dvina Guba-Mezen river line, based on amalgamationof the DSS (line 2 in Fig. 5) and CMP (line 3 in Fig. 5) data. Other explanationis as in Figure 7. anomaly and inferred that it was of magmatic origin. Beneath the Leshukona-Pinega rift, thinning of the crust is accompanied by west-dipping faults, and both these features are clearly seen in the cross-section in Figure 14. There is also an east-dipping fault, which crosses from the top of the basement to the lower

Fig. 15. Depth to Moho (thickness of the crust) in the MezenTiman-Pechora region. Other explanationis as in Figures 10 and 11.

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA

crust. The lower P-wave velocity values of 6.0-6.1 km s -~ in the uppermost basement of the rift probably reflect the disintegrated condition of the faulted rocks. The high velocity (c. 6.7 km s -1) in the middle crust occurs between 10 km and 25 km depth below sea level (b.s.1.); it is very similar to those in the crosssection along the Murmansk-Kizil DSS line (Kostyuchenko et al. 1999), where it crosses the Leshukona-Pinega rift. The depth to Moho beneath the central portion of the Mezen Basin is about 36 km and increases to 4 2 - 4 4 km under adjacent regions to east and west (Fig. 15). In the northern portion of the Leshukona-Pinega rift, just in the area beneath the 100 mGal positive gravity anomaly, the depth to Moho decreases to 32 km. Figure 16 shows that velocities of over 8.4 km s -1 are located under the rifts in the uppermost mantle. These data suggest that Neoproterozoic (perhaps earlier) extension, leading to rift formation, was associated with horizontal intrusion and underplating at the crust-mantle boundary.

The Timan Range: main crustal structure To the NE of the Mezen Basin, the East European Craton is overthrust by the Timanian-deformed, turbidite-dominated successions of the Timan Range (Shatsky 1935; Getsen 1987). The latter is a SW-verging fold and thrust belt, unconformably overlain to the east by the Palaeozoic successions of the Pechora Basin (Figs 1 and 11). It stretches from Kanin Peninsula, in the NW on the Pechora Sea coast, southeastwards to the Polyudov Swell located in the Urals foreland between the Middle and Northern Urals. Tschernyshev (1901) correctly suggested that a 'Timanian mountain chain' extended northwestwards to the Varanger Peninsula and northern Norway. Siedlecka (1975), Olovyanishnikov

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et al. (1997) and Roberts & Siedlecka (1999) gave comprehensive evidence to support the existence of c. 1800 km long T i m a n Varanger Fold Belt. Its prolongation both further to the north beneath the Barents Sea and southwards to the southern Urals has been treated by Gee & Pease (2004). Several contrasting models for the origin of the Vendian Timanide orogen have been proposed. The first of them, the 'Geosyncline Model', suggested that the Timan Range successions constituted a western marginal miogeosyncline, which was flanked to the east in the Neoproterozoic by a broad eugeosynclinal basin (Gafarov 1963; Zhuravlev 1972; Getsen 1975; Dedeev & Zaporozhtseva 1985). The second, the 'Aulacogen Model', defined a major Timanide graben that was separated from the Polar Urals by a cratonic block (Shatsky 1946; Bogdanov 1961; Bogdanov & Khain 1968; Ivanov 1981). Tschzyu (1964) and Raznitsin (1968) suggested that a narrow intra-cratonic geosyncline developed on pre-Riphean basement. The last two models imply that East European Craton crust underlies the Pechora Basin. During the last decade, geodynamic models for the evolution of the Timan-Pechora region have developed further (Getsen 1991; Olovyanishnikov et al. 1996, 1997; Olovyanishnikov 1998; Gee et al. 2000), based on the combined study of both geological, particularly drillcore (Belyakova & Stepanenko 1991), and geophysical data (Kostyuchenko 1994). Within the Timan Range, predominantly siliciclastic successions are thought to range in age from late Middle Riphean (c. 1000 Ma) to Vendian (Olovyanishnikov et al. 1997; Olovyanishnikov 1998; Roberts & Siedlecka 1999). The folded and thrust sedimentary rocks, which are variously intruded by diabase-gabbro suites and, locally, by granites, are exposed in only a few locations ( 1 - 6 in Fig. 11) along the Timan Range. NE-dipping cleavages and low greenschist-facies metamorphism

Fig. 16. Compressional(P)-wave velocities in the uppermost mantle in the Mezen-Timan-Pechora region. Dark stripes indicate velocity anisotropy: 8.4 km s 1 in the direction along stripes, and 8.1- 8.3 km s- 1 across. Abbreviations for the provinces: FS, Fennoscandian Shield; MB, Mezen Basin; TR, Timan Range; PB, Pechora Basin.

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s. KOSTYUCHENKOETAL.

(Getsen 1987; Olovyanishnikov 1998) dominate the successions. Locally, on the Kanin Peninsula, the rock succession is more metamorphosed, up to high amphibolite facies (Lorenz et al. 2004) in a structural dome; elsewhere this grade has been recognized only in a couple of deep drillholes reaching basement below the Pechora Basin. The grade of metamorphism correlates well with the gravity data. Thus, strong positive gravity anomalies occur over the Kanin Peninsula, whereas much less positive anomalies cover the general area of the Timan Range. Roberts et al. (2004) have provided evidence that these predominantly siliciclastic turbidite-dominated successions were deposited in deep basinal (probably passive margin) environments before they were thrust southwestwards onto the sedimentary cover of the EEC. Two major thrusts, the West Timanian and Central Timanian fault zones, dominate the Timan Range (Fig. 11). The West Timanian zone emplaces the basinal facies successions westwards onto the epicratonic strata. Further east, the Central Timanian fault occurs within the Neoproterozoic rocks in the central part of the Timan Range, where it is seen to carry the high-grade metamorphic rocks in the hanging wall (Lorenz et al. 2004). Limestones and dolomites occur within these fault zones, which merge (Fig. 11) in the area south of the Dzhezhimparma Swell (Olovyanishnikov 1998; Timonin 1998). Based on the interpretation of near-vertical reflection data obtained by 'Pechorageophysika' along the 22-PC and 27-PC profiles, Olovyanishnikov et al. (1996) and Olovyanishnikov (1998) concluded that the Timan Range turbidites were overthrust southwestwards a distance at least 25 km. Timonin (1998) estimated this displacement to reach 4 0 - 5 0 km. Figure 17 shows

a cross-section through the Central Timanian fault within the central portion of the Timan Range. Reprocessing of wide-angle reflection and refraction seismic data, obtained along the White Sea-Vorkuta DSS line in 1985, using modern computer technology in the GEON Center (1 in Fig. 5) is illustrated in Figure 18. The cross-section starts in the Mezen Basin and extends eastwards via the Timan Range and neighbouring Pechora Basin to the foredeep of the Polar Urals. In this section beneath the Timan Range, P-wave velocities of about 6.0 km s-~ and S-wave velocities in the range from 3.5 to 3.65 km s- 1 occur in the uppermost crust, being related to Neoproterozoic rocks metamorphosed to greenschist and sub-greenschist facies. The seismic data show that these rocks reach depths of 7.58.0 km and extend eastwards a distance of 200-300 km beneath the Pechora Basin. Evidence, particularly on the Kanin Peninsula, that they increase in metamorphic grade downwards implies that interpretation of the higher velocities in the deeper levels further east is controversial and will need closer analysis. Crystalline basement beneath the Neoproterozoic successions of the Timan Range occurs at depths from 3 to 10 kin. P-wave seismic velocities from 6.15 to 6.35 km s-~ and S-wave velocities from 3.65 to 3.8 km s -1 characterize this basement. Based on petrological interpretations of seismic velocity distributions within continental crust world-wide (Christensen & Mooney 1995; Nikolas & Mooney 1995; Musacchio et al. 1997) and comparisons locally with data on the basement of the EEC, we infer that felsic igneous rocks and high-grade metamorphic complexes constitute the 'crystalline basement' beneath the range. Based on existing

Fig. 17. (a) Parts of seismic cross-section through the Central Timanianfault, line 22-PC (from Malyshev 2002). (b) Interpretation of line 27-PC, based on Olovyanishnikov(1998). Location of the lines is shown in Figure 5. R-V, Neoproterozoic; PZ, Palaeozoic. VI and VII, reflectors.

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA

533

Fig. 18. Seismic cross-section (a) of the northeastern margin of Baltica (White Sea-Vorkuta DSS line; 1 in Fig. 5) and interpretation (b). Other explanation is as in Figures 7 and 14. DSS data, the thickness of the Timan Range crust varies from 38 to 42 km.

The Pechora Basin: crustal architecture and tectonic model The Pechora Basin occupies the area between the Timan Range and the Northern and Polar Urals (Figs 1 and 11). The Phanerozoic sedimentary cover of the Pechora Basin (O'Leary et al. 2004) thickens eastwards from the range to 3.0-4.0 km within the western parts of the basin, 4.0-8.0 km in the central parts, and to at least 12 km in the easternmost area of the Uralian foreland basin (Fig. 19). After the culmination of Timanian Orogeny in the late Vendian, with uplift and erosion, a passive continental margin was established along this northeastern edge of Baltica in the Early Palaeozoic (Puchkov 1997; Bogolepova & Gee 2004). Rift-related deposition apparently started in the Late Cambrian and the passive margin was established by the Mid-Ordovician. Reactivation of the early rift regime in Late Silurian-Early Devonian time occurred in the central area of the Pechora-Kolva Zone. This was followed by predominantly epicratonic sedimentation and then renewed rifling and related magmatism during the Middle Frasnian and, subsequently, local inversion (Timonin 1998; Malyshev 2002). Many drillholes sample the entire sedimentary fill and penetrate the basement within western and southwestern portions of the basin; about 70 reach the pre-Palaeozoic basement in central parts (Belyakova & Stepanenko 1991). Based on these geological data, Belyakova & Stepanenko (1991), Olovyanishnikov et al. (1997) and Roberts & Siedlecka (1999) inferred that Timanian

Fig. 19. Map of thickness of the Palaeozoic sedimentary cover (i.e. depth to Neoproterozoic basement) beneath the Pechora Basin. Explanation is as in Figure 12.

534

s. KOSTYUCHENKOETAL.

(Baikalian) deformation of the Neoproterozoic pericratonic and basinal sedimentary successions resulted in formation of the basement of the Pechora Basin. Gee et al. (2000) presented previously unpublished K / A r age data and new single-zircon (Pb-evaporation) ages of c. 560 Ma from late tectonic intrusions in the basement beneath the Pechora Basin; they concluded that Timanian Orogeny ended in the late Vendian with late orogenic granite intrusion, waning compression, uplift and erosion. Pease et al. (2004) provided geochemical evidence that these granites were subduction related. The drillholes into the basement beneath the Pechora Basin (Belaykova & Stepanenko 1991) were particularly important for identifying the character of the major magnetic anomalies (Fig. 2) related to the Timanian complexes beneath the Palaeozoic unconformity. They demonstrated that the very strong magnetic anomalies of the Pechora Zone outlined by the 'Pre-Pechora' Faults (shown in Figs 11 and 18) coincided with a belt of volcanic and volcano-sedimentary rocks with major gabbro-diorite intrusions and granites. To the west of the Pechora Zone, in the Izhma Domain (Fig. 11), turbidites, similar to those in the Timan Range, were shown also to host some of the late tectonic granites referred to above. To the east of the Pechora Zone, the drillholes sampled acid volcanic rocks and granites of the Khoreyver Domain (Bol'shezemel'skaya Zone of Gee et al. 2000; Gee 2005). Xenocrysts in zircons in the granites suggested that Grenville-age crust was present beneath this area, supporting previous proposals for the presence of older basement fragments beneath the Pechora Basin. Based on a combination of drill-core data, conventional and wide-angle reflection and refraction seismic data and the interpretation of gravity and magnetic fields, Kostyuchenko (1994) proposed that the Timanian basement, lying to the east of the exposures in the Timan Range, can be divided into three major domains (the Izhma, Khoreyver and Pre-Urals in Fig. 11). As noted above, within the western, Izhma Domain, Neoproterozoic successions are similar to those in the Timan Range. The 5.96.0 km s-1 P-wave velocities are inferred to define the extent of these rocks beneath the Palaeozoic sedimentary cover (Fig. 20). A smooth, low magnetic anomaly field (Fig. 2) reflects the Timanian rocks in the Timan Range and Izhma Domain. The abrupt boundary to the Pechora Zone, so conspicuous in the magnetic signatures (Fig. 2), is marked in the seismic data (Fig. 118) by only a small increase in P-wave velocity to 6.1-6.2 km s - . Further east, in the central Khoreyver (Bolshezemel'sky) Domain in the upper crust, P-wave velocities vary between 6.1 and 6.3 km s -1. A wide range of magnetic anomalies, from positive to negative, occur above this domain. The magnetic data and upper crustal seismic data together suggest that the Khoreyver Domain consists of rocks that are intruded by magmatic suites with variable character and chemistry and, at least in part, are metamorphosed to amphibolite facies. The few drillholes penetrating this zone provide evidence of sandstones intruded by granites (Belyakova & Stepanenko 1991). Palaeoproterozoic and older complexes have been suggested for the Khoreyver Domain (Berlyand 1989; Olovyanishnikov et al. 1996; Olovyanishnikov 1998), but the only available isotope evidence suggests the presence of Grenville-age accreted terranes (Gee et al. 2000; Gee & Pease 2004). Further east in the so-called Pre-Urals Domain, beneath the Urals foreland basin, a P-wave velocity of 6.4-6.6 km s- 1 has been obtained for the rocks in the upper part of the pre-Palaeozoic basement, indicating that the Precambrian crust may contain rocks of oceanic affinity (Kostyuchenko 1994), as shown in Figure 18. The major igneous belt of the Pechora Zone (Pre-Pechora fault zone of Figs 11 and 18) is about 50 km wide in the line of the White Sea-Vorkuta DSS profile (Fig. 5), but wider and thicker to the NW and SE, as shown by the magnetic anomalies separating the Izhma from the Khoreyver domains. The magmatic rocks are inferred to be subduction related (Dovzhikova et al. 2004). Figure 20 shows that P-wave velocities higher than 6.4 km s -~

Fig. 20. Compressional(P)-wave velocityin the top of the basementbeneath the Pechora Basin. Abbreviationsof domains: A, Timanian; B, Izhma; C, Khoreyver;D, Pre-Urals.

are characteristic of this zone and indicate the presence of mafic igneous rocks. The positive gravity field and strong positive magnetic anomalies support this interpretation. The Pechora Zone is also characterized by a narrow linear depression in the basement topography in which the depth to the Lower Palaeozoic unconformity varies from 6.0 to 8.0 km, increasing towards the SE. The seismic data across the Pechora Basin along the White SeaVorkuta DSS line (Fig. 18a) provide a basis for interpreting the deeper crustal structure (Fig. 18b). Beneath the Neoproterozoic successions of the Izhma Domain, in the upper crystalline crust, P-wave velocities vary from 6.15 to 6.35 km s -1 and S-waves are 3.753.8 km s -1. In the underlyin~ middle crust, the P- and S-wave velocities are 6.37-6.53 km s- and 3.84-3.95 km s -1, respectively. The western part of this crystalline crust lies on a c. 10 km thick lower crust which extends eastwards from the Mezen Basin and has P- and S-wave velocities of 6.73 k m s -~ and 3 . 9 6 k m s -1, respectively. In general, the seismic data provide evidence that the 20 km thick crystalline basement beneath the Timan Range and Izhma Domain is composed of East European Craton complexes, underthrust beneath the Timanian allochthons (Fig. 18b). These rocks composed the basement of the Neoproterozoic passive margin. In the cross-section (Fig. 18a), there are several east-dipping

CRUSTAL STRUCTURE OF NORTHEASTERN BALTICA

seismic boundaries that penetrate both the Neoproterozoic complex and the crystalline basement. We suggest that these boundaries are thrust faults, which originated during Timanian compression and uplift in the foreland of the Mezen Basin. The evidence for thrusting along with the local evidence of high-grade metamorphism of the turbidites (e.g. on Kanin Peninsula) suggests that the crystalline crust beneath the Izhma Zone is imbricated and intercalated with Timanian metasediments. Figure 18 shows that the Neoproterozoic sedimentary rocks of the Khoreyver Domain, recorded in the drillholes referred to above, are too thin to be identified by the seismic profiling. The P-wave velocities from 6.05 to 6.35 km s -1 and S-wave values from 3.75 to 3.8 km s-1 reflect the crystalline character of the basement close beneath the Palaeozoic sedimentary cover, as in the Mezen Basin. Deeper beneath the Khoreyver Domain, in the middle crust, P-wave velocities are c. 6.5 km s -1 and S-wave values are c. 3.95 km s-l. In the lower crust, from a depth of c. 26 km down to the Moho, P-wave velocities vary from 6.7 to 7.1 k m s -1 and S-wave velocities range between 4.1 and 4.35 km s -l. The total thickness of the crust in the Khoreyver Domain and its velocity structure indicate the presence of continental crust, similar to platform crust in the global crustal models of Pavlenkova (1996) and Mooney et al. (1998). Thus, the Grenvilleage microcontinent hypothesis for the Khoreyver Domain, suggested by the zircon data, finds support in the geophysical data. As noted above, the Moho depth in the Pechora Basin area (Fig. 15) varies from 34 km to 46 km b.s.1., increasing to 48 km beneath the front of the Urals. There is a NW-trending linear depression in the Moho to 46 km along the Pechora Zone. Pn-wave phases, recognized during DSS studies along the Murmansk-Kizil line (2 in Fig. 5) and the Kineshma-Vorkuta line (5 in Fig. 5) provided evidence of high-velocity mantle (c. 8 . 4 k m s -1) occurring west of the Pechora fault zone, and lower-velocity mantle ( 7 . 8 - 8 . 0 k m s -1) underlying the crust further to the east; this evidence implies that the Pechora fault cuts through the whole crust (Fig. 16).

Summary and conclusions Many aspects of the tectonic evolution of northwestern Russia, from the Archaean-Palaeoproterozoic history of the Fennoscandian Shield to the Mesoproterozoic(?) and Neoproterozoic rifting of the Mezen Basin and the Vendian orogeny of the Timan Range and Pechora basement remain controversial. The existing geophysical data provide important constraints on the crustal structure and tectonics, both those related to Palaeoproterozoic and Neoproterozoic orogeny and the subsequent Palaeozoic basin evolution. Continuing wide-angle and deep near-vertical (vibroseis) profiling across the Mezen Basin to the Polar Urals will further contribute to the understanding of this part of continent Baltica's northeastern margin. A late Archaean and Palaeoproterozoic history of accretion of microcontinents, island arcs and oceanic terranes, interrupted by extensional events, was inferred by Berthelsen & Marker (1986), Gaal & Gorbatschev (1987) and Marker (1990); it is treated elsewhere in this volume by Slabunov et al. (2006) and Daly et al. (2006) and considered further below. Using only geophysical data, it is often not possible to determine whether the processes responsible for the structure of the crust were compressional or extensional. Nevertheless, the geophysical data do provide evidence of rock composition and the thin-skinned and thick-skinned nature of the crustal structure, with faults indicative of rifting and shortening; these data are essential for relating the deep structure to the surface geology and interpreting the tectonic evolution. In general, the Palaeoproterozoic evolution of the northeastern part of the Fennoscandian Shield, summarized by Mints et al. (1996), included the following stages: from 2 . 4 9 - 2 . 4 G a to 2.42-2.33 Ga, rifting of the Late Archaean protocraton, with

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opening of intracontinental basins between the Kola and Belomorian microplates and southward subduction, accompanied by continental margin or island-arc type volcanism; from 2.22.17 Ga to 2.02-1.87 (1.78) Ga, back-arc rifting and spreading, followed by subduction and shortening of the basins, development of active margins and uplift; from 1.78 Ga to 1.7 (1.67) Ga, collision of the Kola and Belomorian microplates, resulting in formation of the Pechenga-Imandra-Varzuga suture and Lapland back-thrusting. For the Kola Peninsula, Mints et al. (1996) inferred that the crust of the Central Kola microcontinent formed during Late Archaean collision, and collision of the Central Kola and Murmansk microcontinents took place at c. 2.8-2.7 Ga. Before the latter (in Archaean time), the Keivy microcontinent was accreted to the Central Kola unit (Mitrofanov & Bayanova 2000). The Titovka-Keivy suture zone that now separates the Murmansk microcontinent from the Central K o l a - K e i v y composite continent represents the remnants of Archaean island arcs and ocean floor. Based on the geophysical data, presented in this review, the Archaean Murmansk allochthon was overthrust southwards onto the Kola-Liinakhamary and Central Kola units. The Kola-Liinakhamary unit is also allochthonous and dips under the Pechenga Belt. Lenses of granulites, perhaps similar to those in Lapland Granulite Belt, are thought to occur in the middle crust of the Kola Peninsula, beneath the Central Kola slice, based on P-wave velocity data. A near-vertical fault separates the Keivy and Central Kola terranes (Fig. 7b). Based on interpretation of geological and geophysical data, the Keivy terrane is suggested to be thrust southwards in Palaeoproterozoic time, and underthrust beneath the northern portion of the Imandra-Varzuga Belt. In this context, the nearvertical fault mentioned above can be interpreted as a transcurrent structure. In the northeastern part of the study area, Mesoproterozoic uplift and erosion of the Archaean-Palaeoproterozoic core of the East European Craton was followed by Late Mesoproterozoic and Neoproterozoic rifting. These rifts were covered by late Neoproterozoic and younger successions, and their character has been revealed only recently by a combination of seismic profiling, analysis of potential field data and, locally, deep drilling. Synrift sediments were deposited in half-graben that developed in response to stretching and thinning of lithosphere apparently in a way similar to that in many well-studied continental rifts (e.g. Baldridge et al. 1995; Keller et al. 1995; Parsons 1995). Several of the faults in the upper and middle crust of the rifts appear to be listric and give way downwards to major subhorizontal detachment surfaces as well as fault wedges. Highly reflective lower crust beneath the rifts may be the result of ductile thinning of the lithosphere. The seismic velocities from these crustal layers, derived from wide-angle reflection and refraction profiling, are high in some places (7.0-7.2 km s-a), perhaps indicating the input of mafic magma into the middle and lower crust. High velocities (c. 6.7 km s -1 and above) in the middle crust are inferred to reflect mid-crustal magma chambers (Kostyuchenko & Romanyuk 1997; Kostyuchenko et al. 1999). Timanian compression, affecting the Mezen Basin in Late Vendian time, resulted in local inversion, with reverse movement on some of the extensional faults of the rifts. The growth of continent Baltica towards the east during the Neoproterozoic and Early Palaeozoic comprised three main stages (Gee & Pease 2004; Gee 2005): (1) a long period of development of a passive margin with deposition mainly of turbidites in deep marine outer shelf and slope-rise environments (Roberts et al. 2004); (2) Timanian (Baikalian) Orogeny and late Neoproterozoic accretion of outboard juvenile crust and microcontinents (Gee et al. 2000; Roberts & Olovyanishnikov 2004); (3) uplift, erosion and deposition of Palaeozoic passive margin successions (Puchkov 1997; Malyshev 2002; Bogolepova & Gee 2004), prior to local Late Devonian inversion and the far-field influence of the Late Palaeozoic Uralian Orogeny.

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Fig. 21. Early Neoproterozoic to Mid-Palaeozoic tectonic evolution of the northeastern margin of the East European Craton (EEC). (For explanation, see the text.)

The geophysical evidence presented here provides further constraints on the character of the Neoproterozoic and Palaeozoic evolution (Fig. 21). The Neoproterozoic metasedimentary rocks of the Timan and Izhma zones do not extend eastward much beyond the Pre-Pechora fault zone. They were thrust westwards over crystalline crust of about 20 km thickness (Fig. 18) and, together with these allochthonous units of the EEC margin, were emplaced onto the autochthonous edge of the Craton. To the east of the Izhma Domain, within the Pechora zone outlined by the PrePechora faults, a volcano-sedimentary complex of island-arc affinity (Belyakova & Stepanenko 1991; Dovzhikova et al. 2004) has been recognized, based on drillcore analysis. Interpretation of seismic, and particularly, potential field data (Kostyuchenko 1994) allowed better definition of this subduction-related complex, dominated by volcaniclastic rocks and volcanic arc magmatism. The seismic velocity data from the area east of the Pechora zone, with evidence of continental crust beneath the Khoreyver Domain, taken together with the Grenville-age xenocrysts in the Vendian plutons, suggest that this terrane was not originally a part of the Archaean-Palaeoproterozoic EEC. Whether it was accreted during early or late Neoproterozoic orogeny remains open to debate, but the long Neoproterozoic and perhaps late Mesoproterozoic history of extension in the Mezen Basin and Timan Range favours the latter. Outboard, NE of the Khoreyver microcontinent, the high seismic velocities in the basement beneath the Uralian foredeep suggest the possible preservation of ocean-derived terranes similar to those occurring locally in the Enganepe Anticline (Dushin 1997), in the front of the Urals. Thus, a model for the evolution of the Neoproterozoic-Early Palaeozoic margin of northeastern Baltica (Fig. 21) includes, first, a long period of passive margin development and perhaps an intra-oceanic magmatic arc (Pechora zone); thereafter, collision of the oceanic (Enganepe) outboard terrane(s) with the Khoreyver microcontinent, and of the latter with the Pechora arc in the early-mid-Vendian; followed by continued subduction and generation of granites in the late Vendian; and, finally, thrusting of the entire allochthon over the passive margin turbidites (Timan-Izhma) and the latter onto the EEC margin. Figures 18 and 21 show thickening of the base of the magmatic arc beneath the Pechora zone volcanic belt and subsequent

underplating related to Devonian rifting and widespread basalt magmatism. We would like to thank N. Bogdanov, V. Pease, V. Bogatsky, A. Pystin, N. Malyshev, V. Olovyanishnikov, M. Mints, A. Golubev, N. Filippov and N. Sharov for many discussions and comments concerning the geological and geophysical data and evolution of the area of study. We thank the staff of the GEON Center and Spetsgeofizika for providing the unpublished seismic data. The Ministry of Natural Resources of Russia financed the research and the INTAS project No. 1941 'High Arctic Lithosphere of Europe' (HALE) partially supported the investigations leading to this paper. We are also grateful to R. Stephenson and M.-F. Brunet for their reviews, which significantly improved the paper, and particular thanks go to O. K. Bogolepova for getting the paper into shape for publication.

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Electrical conductivity and seismic velocity structures of the lithosphere beneath the Fennoscandian Shield S.-E. H J E L T 1, T. K O R J A 1, E. K O Z L O V S K A Y A 1'2, I. L A H T I t'3, J. Y L I N I E M I 2 & BEAR* A N D SVEKALAPKO SEISMIC TOMOGRAPHY t WORKING GROUPS

1Division of Geophysics, Department of Physical Sciences, University of Oulu, POB 3000, FI-90014 Oulu, Finland 2Sodankylii Geophysical Observatory, Oulu Unit, University of Oulu, POB 3000, FI-90014 Oulu, Finland (e-mail: Elena.Kozlovskaya @ oulu.fi) 3Geological Survey of Finland, Northern Finland Office, Rovaniemi, Finland

Abstract: Key results from the BEAR (Baltic Electromagnetic Array Research) soundings and the SVEKALAPKO Seismic Tomography Experiment (SSTE), which explored geophysical properties of the upper mantle beneath the Fennoscandian Shield, are summarized here. These two projects formed the two key geophysical experiments in EUROPROBE's SVEKALAPKO project. The major result obtained from the teleseismic tomography indicates P- and S-wave velocity variations of up to 4%, as compared with the IASP'91 global model. The positive P-wave velocity anomaly seems to extend down to 300 km, without any indication of an asthenospheric low-velocity layer. This is corroborated by surface-wave analysis. The Archaean-Proterozoic suture zone has no continuation at upper mantle depths. Instead, a laterally and vertically heterogeneous structure of the subcontinental lithospheric mantle in the contact zone of Archaean and Proterozoic domains beneath the SVEKALAPKO area has been revealed by both teleseismic and local event studies. Comparison between stacked receiver functions from the TOR (Teleseismic Tomography of TORnquist Zone) and SSTE arrays indicates that the difference between arrival times of the converted P410 and P660 phases increases below the Shield as a result of a cooler upper mantle. The asthenosphere beneath Central Europe terminates at the southern edge of the shield and cannot be identified in the SVEKALAPKO data. The variations in Moho depths from 30 km to 50 km (locally, more), known from earlier deep seismic sounding (DSS) work, were corroborated by the receiver function analysis. The BEAR (Baltic Electromagnetic Array Research) experiment data indicate that an upper mantle conducting layer is required in some places beneath Fennoscandia. In the northern part of the Shield the upper mantle conductor is located at a depth of c. 170 km. Magnetotelluric data exhibit strong anisotropic behaviour, in particular in the central parts of the Fennoscandian Shield. The present state of modelling implies, however, that isotropic 3D crust and upper mantle explain nearly all observed anisotropic features and that very minor anisotropy of the upper mantle is required to explain the data. The crust is strongly variable electrically, ranging from resistive areas to highly conducting elongated structures. The conductances (= conductivity x layer thickness) obtained range from a few Siemens to several tens of thousands of Siemens. The conductive structures delineate boundaries between Archaean and Proterozoic crustal units and are inferred to image relics of subduction processes. In the central parts of the Shield, the lower crust is conducting, with dipping structures extending far to the NE below the Archaean-Proterozoic boundary. Thus, the present interpretation of electromagnetic BEAR and tomographic data from the Fennoscandian Shield demonstrates that the structure of the upper mantle beneath the Shield is much more heterogeneous than was supposed when the project started. Models of the evolution of lithosphere must be revised to accommodate lateral and vertical heterogeneity in the upper mantle.

The Precambrian Fennoscandian Shield (Fig. 1), with its thin sedimentary cover, is an ideal target for geophysical studies of the deeper parts of cratons. On the other hand, such research is challenging, as the long geological history of the Shield, with several orogenies, has resulted in a complicated internal structure (Gorbatschev & Bogdanova 1993; Nironen 1997; Korsman et al. 1999; Lahtinen et al. 2005). An integrated multi-method approach is essential to decipher the evolutionary history of the Fennoscandian Shield. This paper concentrates on the deep seismic and electromagnetic (EM) results of E U R O P R O B E ' s multidisciplinary S V E K A L A P K O project (Hjelt et al. 1996). Because processing and modelling of these data has not yet been finished, integration of the results will not be provided. Petrophysical, potential field and geological databases (Koistinen et al. 2001; Korhonen et al. 2002) exist for large parts of the Shield and they are essential in constructing surface structures while preparing models of the deeper parts (Kozlovskaya et al. 2004a). F E N N O L O R A and SVEKA81 were the first large-scale deep seismic sounding (DSS) profiles in the Fennoscandian Shield. These and later DSS profilings showed that the thickness of the crust in Fennoscandia varies significantly, from 27 to 65 km (Luosto 1991, 1997; Korsman et al. 1999; Malaska & Hyv6nen *The BEAR Working Group consists of scientists from 18 institutes in Finland, Germany, Russia, Sweden, UK and Ukraine (see Appendix). *The SVEKALAPKO Seismic Tomography Working Group (SSTWG) consists of scientists from 15 institutes in the Czech Republic, Finland, France, Germany, Poland, Russia, Sweden and Switzerland (see Appendix).

2000). The variations in the Precambrian parts are due to highvelocity lower crust of variable thickness (Korja et al. 1993; Korsman et al. 1999). Reflection seismology provides a high-resolution geophysical technique for obtaining details of tectonic processes that have affected the crust. The B A B E L seismic experiment (Fig. 1) demonstrated that high-quality images can be obtained for the entire crust in the Fennoscandian Shield ( B A B E L Working Group 1990, 1993; Korja & Heikkinen 1995, 2005; Snyder & Hobbs 1999). Two seismic reflection profiles were measured in the western parts of Kola Peninsula and Russian Karelia during S V E K A L A P K O . A planned profile in the central part of the Shield was later included in a massive four-profile experiment, FIRE (Finnish Reflection Experiment), presented elsewhere. Electromagnetic (EM) data provide models of the crust and upper mantle that are complementary to seismic images. The E M studies started with magnetotelluric sounding (MTS) along the DSS profile SVEKA81 (Adam et al. 1982), long-period E M soundings (Jones 1980, 1981; Rokityansky et al. 1981) and the use of magnetovariational (MV) arrays (Pajunpii~i et al. 1983) as described by Hjelt (2001). Several regions with upper or midcrustal conductors were identified (Pajunp~i/i 1987). Later, broadband MTS was completed across m a n y of the most significant crustal conductivity anomalies (e.g. Korja et al. 1986; Rasmussen et al. 1987). These findings have been summarized by Korja & Hjelt (1993, 1998) and Korja et al. (2002). The S V E K A L A P K O project (Hjelt et al. 1996) continued investigations of the evolution and deep structure of the Precambrian

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 541-559. 0435-4052/06/$15.00

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Fig. 1. Simplifiedgeologicalmap of the central FennoscandianShield, modified from Koistinenet al. (2001). Geological units traversed by GGT-SVEKA are as follows. Palaeoproterozoic Svecofennian: SaR, Satakunta Rapakivi Area; HaB, H~ne Belt; PiB, Pirkkala Belt; TaB, Tampere (Schist) Belt; CFGC, Central Finland Granitoid Complex; SaB, Savo Belt; PoB, PohjanmaaBelt; Karelian (Palaeoproterozoic rocks): KaB, Kainuu (Schist) Belt. Karelian (Archaeanrocks): IC, Iisalmi Complex;PC, Pudasj~vi Complex; EFC, Eastern Finland Complex. Other units: Skelleftea, SkellefehVolcanic Arc; WiR, Wiborg RapakiviArea; OA, Outokumpu Area (North Karelian Schist Belt); LGB, Lapland GranuliteBelt; IVB, Imandra-Varzuga Belt; PeB, Pechenga Belt. Inverted open triangles denote sites of 1D conductivitymodels shown in Figure 5. Open stars denote seismic shot points along the GGT-SVEKA seismic line (Korsman et al. 1999) as a reference for Figure 11. Continuous lines denote seismic reflectionlines: lines in sea areas are BABEL lines and a irregularline subparallel to GGT-SVEKAis for FIRE-1 and FIRE-2.

Fennoscandian (Baltic) Shield along a N E - S W transect from the Barents Sea to Central Finland. The shape, thickness and age of the lithosphere and the disposition of major lithospheric structures of the Shield were studied. The BEAR electromagnetic (magnetotelluric) array (Korja & BEAR Working Group 2000) and seismic tomography experiment (Bock & SVEKALAPKO Seismic Tomography Working Group 2001) were the largest efforts of the SVEKALAPKO research (Fig. 2). When the data had been processed and inverted, geophysical images of the major crustal structures, features of the lithosphere-asthenosphere system and comparison of the deep structures with surface geology became possible (Engels et al. 2002; Korja et al. 2002; Kovtun et al. 2002; Pajunp~i~i et al. 2002; Vanyan et al. 2002; Varentsov et al. 2002; Alinaghi et al. 2003; Sandoval et al. 2003, 2004; Bruneton et al. 2004a,b; Yliniemi et al. 2004; Lahti et al. 2005).

Electromagnetic studies Factors affecting electrical conductivity: crust and mantle

In tectonically active regions small percentages of partial melt and free saline fluids form conducting structures that are distinct for different tectonic environments. Continental rifts, mid-ocean ridges, oceanic subduction zones and convergent continental margins all have their own conducting fingerprints in the lithosphere (e.g. Hjelt & Korja 1993; Brown 1994). In stable regions ancient tectonic processes have often left electrically conducting traces of collisions of various crustal units revealing the location

of palaeosuture zones or even more specific tectonic assemblages such as foredeeps (e.g. Boerner et al. 1996). Enhanced electrical conductivity is in many cases caused by graphite- and/or sulphide-bearing rocks of sedimentary sequences that have undergone complex deformation and underthrusting in the deep crust (e.g. Korja et al. 1996). Graphite may also have precipitated from CO2-rich fluids into the shear zones (Haak & Hutton 1986; J6dicke 1992; Mareschal et al. 1992). Mineral composition, pore fluids and temperature also affect conductivity in the crust, but usually to a far lesser degree. In the continental mantle roughly at depths between 100 and 400 km, temperature and mineral composition control electrical conductivity, and high-PT laboratory experiments have made it possible to compile several laboratory-based conductivity-depth profiles, such as the dry adiabatic profile for olivine (Xu et al. 2000). The experimental conductivity-depth profiles serve as background models for field observations. An increased conductivity with respect to dry olivine conductivities in the upper mantle has been found in many studies (for a review, see Heinson 1999) and several sources to explain an enhanced conductivity have been proposed. Electrical conductivity is enhanced in the presence of partial melt, if the melt forms interconnected networks (Waft 1974; Roberts & Tyburczy 1999). Other possibilities include a grain boundary carbon phase at levels above the diamond stability field (Duba & Shankland 1982) or dissolved water (H + diffusion) in olivine (Karato 1990; Mackwell & Kohlstedt 1990). Unlike seismic or geothermal observations, EM observations are more sensitive to the presence of even small fractions of melt or volatiles (Heinson 1999). Below 400 kin, laboratory

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Fig. 2. The station map of the SVEKALAPKOBEAR and seismic tomography array experiments.BEAR data were acquired on June-July 1998 and seismic data on September 1998-April 1999. Geology is simplifiedfrom Koistinen et al. (2001). data indicate an increase in conductivity as a result of the phase change of olivine to wadsleyite and ringwoodite (Xu et al. 1998).

E l e c t r o m a g n e t i c techniques in lithospheric studies

The distribution of large-scale crustal structures can best be mapped by MV and MT measurements whereas investigation of the internal 'microstructure' of conductors requires methods with dense lateral coverage, such as airborne electromagnetic (AEM) and audiomagnetotelluric (AMT) soundings. An appropriate combination of large-scale and small-scale methods with adequate horizontal and vertical sampling and penetration provides a powerful tool and allows us to investigate in detail both the position of conductors with respect to geological and tectonic structures and the internal structure of the conductors themselves (Korja & Hjelt 1998). Methodological advances as well as the availability of multiparameter and spatially coinciding geophysical datasets have significantly improved the interpretation of geoelectrical models in more specific tectono-geological terms (Jones 1992, 1999; Korja et al. 1993; Korja et al. 1996).

The major advantages of MT studies are that (1) the data provide structural information complementary to seismic results, (2) the techniques for genuine 3D modelling exist, and (3) the texture and fillings of the subsurface affect the measured fields rather than volume, which is complementary to most geophysical methods. The three major challenges in the study of the deep conductivity structure by MT measurements in the Fennoscandian Shield in particular are (1) the source effect, (2) the presence of crustal conductors, and (3) static shift. The source consists of the currents flowing in the ionosphere. Their variations depend on the geomagnetic activity and include rapid temporal and spatial variations (Pulkkinen et al. 2003a,b). These variations affect, in particular in studies carried out at high latitudes (Mareschal 1986), the traditional inversion principles of deep EM soundings at long periods required for deep penetration. The effect is enhanced in regions of resistive crust. Close to the auroral zone careless interpretation may produce virtual conductors. Analysis of the BEAR data indicates that the inducing field may have a non-planar behaviour already at periods of a few hundreds of seconds. Yet the advanced and improved data-processing techniques developed in the BEAR research (Varentsov et al. 2003b)

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and elsewhere (Jones & Spratt 2002) have provided ways of reducing substantially the harmful source effects of long-period data. The analysis of the BEAR data (Varentsov et al. 2003b) suggests that local transfer functions (impedance and tipper) are generally consistent with plane wave source up to periods of c. 10 000 s. In the case of two-site transfer functions (e.g. horizontal magnetic data) that describe spatial relationships between the geomagnetic field components, the source effect contaminates the data from periods of a few hundreds of seconds, yielding a quasi-linear northern trend in transfer functions. In studies of the upper mantle crustal conductivity structure must be known, because long-period electromagnetic methods are sensitive to crustal conductors even though the structure cannot be resolved because of sparse or inadequate spatial sampling. Crustal conductors in shields, typically schist and graphite belts in crystalline bedrock, limit the depth of penetration of EM waves and 'leaks' into inverted models both at depth and at the sides of the sounding area. The Fennoscandian Shield, in particular, contains a number of long crustal conductors (Fig. 3). Fortunately, the crustal conductivity structure is relatively well known, as a result of a number of MT profiling and MV array measurements. A detailed knowledge of the location and geometry of the upper crustal conductors has been improved in the BEAR studies through the effective use of AEM datasets (Engels et al. 2002; Korja et al. 2002; Pajunp/i~i et al. 2002; Varentsov et al. 2002). The static shift of apparent resistivity is inherent in all MT data, and, if not corrected properly, the depths and conductivities

of conductors derived from MT data are biased upwards or downwards, depending on the sign of the static shift. Static shift is caused by near-surface conductivity structures, where electric charges accumulate at conductivity boundaries and consequently make electrical field piece-wise continuous. Thus, the possible static shift is introduced from electric field measurements whereas the magnetic field is free of static shift. Several methods exist to correct for static and some of them are discussed below, together with the analysis of the BEAR data.

C r u s t a l c o n d u c t i v i t y in F e n n o s c a n d i a

The set of Fennoscandian (Baltic) Shield EM data is unique both in volume and quality. The techniques used range from regionalscale MV arrays and MT soundings, across anomalous conductivity structures, to local controlled source data from direct current and VLF resistivity surveys, frequency soundings, and airborne electromagnetic mapping (e.g. Korja et al. 1996, 2002; Arkimaa et al. 2000). The great variety of methods allows study of structures at regional scales of hundreds of kilometres, as well as small local details of a few metres or even less (Korja et al. 1996). As a part of the BEAR research, the results of all previous crustal studies were compiled and a 3D model (S-map) of crustal conductivity was created (Korja et al. 2002, and reference to previous work therein). The work in the BEAR research was motivated mainly to provide a tool to incorporate crustal

Fig. 3. Integrated crustal conductance [fir(z)dz] of Fennoscandia for the uppermost 60 km as derived from the 3D model (S-map) of Fennoscandia (Engels et al. 2002; Korja et aL 2002). Geology simplified from Koistinen et aL (2001). KIR (Kiruna), KAR (Karelia) and B42 give the location of 1D models shown in Figure 4. Black continuous line shows the location of the magnetotelluric GGT-SVEKA profile (Fig. 5). Sea areas are masked out in the figure, although the 3D model covers both the Baltic Sea and Atlantic Ocean, because no data (models) are available from sea areas; the model for sea areas is an extrapolation of models for land areas. TESZ, Trans-European Suture Zone; TTZ, Teisseyre-Tornquist Zone.

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY,FENNOSCANDIANSHIELD conductivity into modelling and inversion of deeper structures (Engels e t al. 2002; Varentsov e t al. 2002), but the model can naturally be used to describe the large-scale conductivity structure of the Fennoscandian Shield, to compare it with other geophysical data and to discuss the geological significance of crustal conductivity. In Figure 3, we show the integrated crustal conductance derived from the S-map. It shows the total crustal conductance S (conductivity x thickness) from the surface to a depth of 60 km. It should be noted that, in the map, all crustal conductance is compressed into one map-sheet and all vertical variations are excluded. In Figure 4, we show one example of vertical variations of crustal conductivity across the central part of the Fennoscandian Shield along the GGT-SVEKA profile (Lahti et al. 2005). The electrical structure of the crust is highly variable in the Fennoscandian Shield (Fig. 3) with the resistivity variations ranging from 1 0 + 5 ~ m to below 10 -1 gl m. Several elongated belts of conductors ( S > 1000 Siemens), which delineate more resistive crustal units (S < 100 Siemens), characterize the electrical structure of the Shield. The most dominant conductor coincides roughly with the border between the Archaean and the Proterozoic crust in Finland, and in Russia under the Palaeozoic sediments of the East European Platform. Another branch cuts the Palaeoproterozoic Svecofennides in Southern Finland and continues through the Bothnia Belt in Finland towards the Skellefte~ region in Sweden. Prominent crustal conductors exist also around the Lapland Granulite Belt (S several thousand Siemens), in the Pechenga Belt, and in the Imandra-Varzuga Belt (Figs 1 and 3). The Caledonides also contain conductive material, at least in J~imtland, Sweden (Gee 1972; Gharibi 2000). In the following, we summarize the main features of the long, elongated conductors, although each of them also has specific features that provide information on the structure, properties and evolution of its host region (Korja & Hjelt 1993, 1998; Korja e t al. 2002). (1) The conductance of conductors is high (several thousand Siemens), which can be explained only by electronic conducting mechanisms. Hence the conductors are likely to consist mainly of graphite- and sulphide-beating rocks.

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(2) Graphite in conducting assemblages has an organic origin, suggesting that the conductors are composed mainly of sedimentary rocks. (3) The internal structure of the conductors is complex, as is evident from airborne electromagnetic data, containing extremely conducting graphite- and sulphide-beating metasedimentary layers, hosted by resistive rocks. (4) Dipping conductors often have an association with a band of seismic reflectors. (5) Most of the conductors are located in the upper and middle crust without penetration into the lower crust. However, there are some conductors that penetrate through the entire crust (e.g. Skellefte~ and Bothnian regions; Figs 1 and 3), suggesting a transportation of coffductive sedimentary material into the lower crust. (6) Virtually all conductors represent supracrustal Palaeoproterozoic and younger assemblages. It is noteworthy that the age of most of the conductors, if known, seems to concentrate around 2.1-1.9 Ga, with the notable exception of the conductors within the Caledonides, which mainly represent Cambrian metasediments. As a part of the BEAR work, old MT data from the SVEKA profile (Korja & Koivukoski 1994) have been reinverted together with new data from 112 MT soundings along a 750 km long GGT-SVEKA profile (Fig. 1). An improved conductivity model (Fig. 4) confirms that the upper crust is highly resistive and the lower crust is conductive (S ~ 2 0 0 - 5 0 0 Siemens) beneath the Palaeoproterozoic Central Finland Granitoid Complex (CFGC), whereas the entire crust is very resistive in the Archaean Karelian Domain to the east of the Kainuu Belt. Two dipping upper and mid-crustal conductors exist at both sides of the CFGC. One set of conductors is found beneath the Tampere, Pirkkala, Hame Belts and Satakunta Rapakivi area in southern Finland. The conductors are probably caused by two separate subduction and collision processes, as their dips are towards north and south. Minor, SE-dipping conductors beneath the Ladoga-Bothnian Bay Zone, and a major SW-dipping conductor beneath the Iisalmi Archaean unit and the Kainuu (Schist) Belt can be

Fig. 4. Smooth 2D-inversionmodel of the conductivityof the crust and uppermost mantle along the GGT-SVEKAprofile (lower panel) and integrated crustal conductance of the model from surface to 60 km depth (upper panel). The main geological units and their abbreviationsare given in Figure 1. KuB, Kuhmo Greenstone Belt; LBBZ, Ladoga-Bothnian Bay Zone. Moho boundary is from Korsman et al. (1999). Figure is modifiedfrom Lahti et al. (2002).

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explained by the presence of Palaeoproterozoic sedimentary rocks beneath the resistive Archaean rocks of the Iisalmi complex (Lahti et al. 2002). The lower crust is electrically rather heterogeneous in the Fennoscandian Shield. Beneath the belts of upper and middle crustal conductors, it is difficult to obtain any information on the lower crust because of the attenuation of EM fields in highly conducting overlying bodies. Reliable information on lower crustal properties can therefore be obtained only within the resistive regions. The lower crust in the Archaean Belomorides and Karelides is highly resistive, having a conductance below 10 S in many parts (Figs 3 and 4). The northwestern part of the Karelides is over 10 times more conductive (Fig. 3) than the southeastern part; yet the Archaean lower crust is in general much less conductive than the Palaeoproterozoic Svecofennian crust (Fig. 4), where lower crustal conductance is well over 100 S. The lower crust in Central Sweden is, however, highly resistive, having conductivities similar to the presumably Archaean lower crust ( < 100 S). Finally, the new results from the BEAR data from site B42 (Fig. 5) show that the Archaean middle to lower crust may have conductances of a few hundreds of Siemens, similar to the Proterozoic lower crust. In summary, there are areas of both resistive and conductive Archaean lower crust and areas of conductive and resistive Proterozoic lower crust. Consequently, there is no obvious correlation between the age (nominal age determined according to the surface lithology) and conductance of lower crustal rocks. Therefore, it is difficult to explain the enhanced lower crustal conductivity by some 'universal' causes (e.g. precipitation of carbon from CO2-bearing fluids from the mantle, or trapping of water in the lower crust). Explanations are likely to be related to 'local' tectonics; that is, to the style of subduction and following collision processes, which determine how much and where conductive sedimentary material are transported.

U p p e r m a n t l e c o n d u c t i v i t y : the B E A R p r o j e c t

The BEAR subproject was one of the key experiments of EUROPROBE's SVEKALAPKO project. The BEAR project focused on studying the electrical properties of the upper mantle

Fig. 5. Examples of 1D models of resistivity v. depth for northern Felmoscandia from BEAR data (B42) and older experiments (KIR, Jones 1982, 1983); KAR Kaikkonen et al. 1983; Korja & Koivukoski 1994; Korja et al. 2002). Left panel has a logarithmic and fight panel a linear depth scale. For comparison, a 1D model of Central Europe from Olsen (1998) and dry olivine resistivities (= 1/ conductivities) under a relevant continental geotherm _ 100 ~ from Xu et al. (2000) are shown. Moho depths (Korsman et al. 1999) are for site B42. Figure is modified from Lahti et al. (2005).

beneath Fennoscandia and thereby aimed at gaining a deeper insight into the structure, evolution and contemporary dynamics of the continental lithosphere beneath cratons and possibly deeper below lithosphere, and finally at correlating and interpreting the results jointly with other geophysical and geological data available from the Fennoscandian Shield. The BEAR experiment itself consisted of ultradeep electromagnetic sounding, with use of a Shield-wide MT and magnetometer array of simultaneous long-period recordings (Fig. 2). Time variations of the Earth's electromagnetic field were measured for 45 days at 46 MT and 20 magnetometer sites, having an average separation distance of c. 150 km. The time series data from the array recordings were processed by three methods (Varentsov et al. 2003a), which resulted in a number of different EM transfer functions (magnetotelluric impedance, tipper, horizontal magnetic, horizontal spatial gradients) for a wide period range of 10-100 000 s and for a number of remote reference approaches (Varentsov et al. 2003a). However, as mentioned above, because of the source effect (proximity of the source region of the magnetotelluric fields), the useful period range is limited to c. 10 000 s for magnetotelluric impedance data. Consequently, the following results are obtained using BEAR data from 10 s to 10 000 s and pre-BEAR A M T - M T data from 1000 Hz to 1000 s. The analysis and modelling of the BEAR data have resulted so far in the following three conclusions that will be discussed separately below: (1) conducting material is required somewhere below 100 km almost everywhere beneath Fennoscandia; (2) 1D inversion results from site B42 suggests a depth of c. 170 km for the upper mantle conductor in northern part of the Shield; (3) magnetotelluric data exhibit strong anisotropic behaviour, in particular in the central part of the Fennoscandian Shield. The 3D modelling has shown (Engels et al. 2002; Varentsov et al. 2002) that an excess of roughly 5000 S of conducting material (e.g. 50 km of 10 1)in) is required somewhere below 100 km. The 3D model used included 3D crust and 1D upper mantle. For 1D upper mantle, two model variants were used. The first was the Fennoscandian reference model (Korja et al. 2002), which is, in general, compatible with the dry olivine conductivity model (Xu et al. 2000) and the model for Central Europe (Olsen 1998). The model had no conductive layer in the upper mantle, but a monotonous increase as a result of rise in temperature. Comparison of the model responses with observations showed that the observed phases were systematically higher than the modelled phases. The second model contained a conductive layer in the upper mantle having a conductance of 5000 S. This model removed systematic bias between observations and model responses, yet this model, with a single conductive layer at a depth of 1 5 0 - 2 0 0 k m , could not produce a satisfactory fit between observations and model responses (one-third of the sites had a good fit, in another third of the sites observed phases were still higher than the model phases, and in the remaining sites they were smaller), indicating that a simple 1D model with a single conducting layer is not valid for the Shield. Dimensionality analysis of the BEAR data (Lahti et al. 2005) shows that in the northeastern part of the array, electrical structure is nearly 1D, whereas in other parts the data has rather strong 2D or 3D character. Therefore site B42 (Salla) was selected (Fig. 5) as an example, where simple 1D inversion will provide reliable information on upper mantle conductivity. For static shift correction, a commonly used method is to adjust the level of apparent resistivity curves at their long-period branch using some global reference curve obtained from magnetic data, which is not affected by static shift. An obvious reference curve would be that of Olsen (1998) for Central Europe, estimated from European geomagnetic observatory data. It turned out that phases from site B42 and Olsen's response do not coincide, which is a condition for the correction of static shift, at a common long-period interval

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD around 10 000 s. This indicates that the upper mantle structure at depths corresponding to 10 000 s is different in these two areas and longer periods would be needed. Longer period data, however, are difficult to obtain in the BEAR experiment, as a result of the source effect, as discussed above. Alternative approaches for the correction of static shift include the correction at short period intervals using magnetic data, e.g. from time domain EM methods, or the correction by spatial averaging. The latter was used for site B42 and the static shift was corrected by averaging six apparent resistivity curves from nearby sites (Lahti et al. 2005). Following this procedure, 1D inversion of magnetotelluric data from site B42 was accomplished using two approaches (Lahti et al. 2005). The resulting models (smooth Occam model and a fivelayer model with a minimum number of layers required by the data) and their comparison with some other models are shown in Figure 5. The three main features of the model are: (1) the presence of a middle to lower crustal conducting layer; (2) an abrupt increase of conductivity at a depth of c. 170 km; (3) rather low conductivity (c. 100 [l m) of the mantle lithosphere. The cause of the middle to lower crustal conductor at site B42 is difficult to explain because, according to geological mapping, the site is located in an area of Archaean crust. The result suggests that either the Archaean lower crust can be conductive, in contradiction to previous results (e.g. Jones 1992), or an unknown process has affected the lower crust since Archaean times and made the middle to lower crust conductive. The resistivity of the mantle lithosphere is c. 100 l) m beneath site B42, although it should be noted that the actual resistivity of a layer below a conducting layer (middle to lower crust in this case) is difficult to obtain. Yet the resistivity is compatible with the results from site KIR in northern Sweden (Jones 1982, 1983), but 10 times lower than in the central part of the Shield in Karelia (KAR, Korja & Koivukoski 1994). Similarly, both at sites B42 and KIR, an abrupt increase in conductivity is detected at depths of 170 and 150 km, respectively, whereas at KAR no such interface is found in the uppermost 200 km. This indicates, as pointed out above, that there exist considerable lateral variations in the electrical properties of the upper mantle in Fennoscandia. The olivine conductivity profile (Xu et al. 2000) is shown in Figure 5. Comparison of this with the models of B42 and KIR shows that the mantle lithospheric conductivities are higher in the northern part of the Shield than predicted by the dry olivine model, in particular in the region of enhanced conductivity below 170 km at B42 and 150 km at KIR. In contrast, at KAR in the central part of the Shield, the model resistivities are in agreement with the dry olivine model, at least at a depth of 200 km. Comparison of results from Fennoscandia with those from the Canadian Shield (Schultz et al. 1993; Hirth et al. 2000; Neal et al. 2000) indicates that the mantle lithosphere in Fennoscandia is roughly 10 times more conductive than the Archaean lithosphere beneath the Canadian Shield, whereas in the central part of the Fennoscandian Shield (KAR), the resistivity is similar to that in the Canadian Shield. In summary: (1) upper mantle conductivity is laterally heterogeneous in the Fennoscandian Shield; (2) there must be a layer of enhanced conductivity in the upper mantle beneath the entire Fennoscandian Shield, which has 10-100 times higher conductivities than predicted by the dry olivine model; (3) the depth to the top of the conducting layer (or a region of enhanced conductivity) is 1 5 0 - 1 7 0 k m in the northern part of the Shield, whereas in the central part of the Shield it must be deeper than 200 km; (4) the conductivity of the mantle lithosphere, above the conducting layer, is roughly 10 times higher than the dry olivine conductivity in the northern part of the Shield, whereas in the central part of the Shield the conductivity is comparable with the conductivities of the dry olivine model. Magnetotelluric data from the BEAR array are strongly anisotropic, in particular in the central part of the Fennoscandian Shield; that is, the data yield stable geoelectric strikes (50~

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and large phase split (30-45~ This led Bahr & Simpson (2002) to suggest that the upper mantle is electrically anisotropic in the Fennoscandian Shield. Similarly, using earlier data, Rasmussen (1988) and Korja & Hjelt (1998) have suggested that the deep crust and upper mantle might be electrically anisotropic in the Fennoscandian Shield. Three-dimensional modelling using the 3D crustal model compiled in the BEAR project (Korja et al. 2002) shows, however, that isotropic 3D crust and isotropic, layered mantle can explain nearly all the observed anisotropic features (Korja & BEAR Working Group 2003). The isotropic 3D model produces very stable strikes of c. 50 ~ NE, as observed in the field. Similarly, the 3D model produces phase splits nearly as large as observed (30-45~ The remaining part (i.e. the part that cannot be explained by the current isotropic 3D crustal and 1D mantle models) might be due to genuine anisotropy, or due to heterogeneities in the upper mantle. It is clear, however, that the observed strikes have no bearing on the azimuth of the anisotropy and that if the upper mantle is anisotropic (the unexplained part of phase split), then the anisotropy factor (proportional to phase split) is much smaller than estimated from the original phase split.

Seismological studies M o h o m a p ( D S S profiles)

Early DSS profile data indicated significant variations of the crustal thickness in the Precambrian parts of the Fennoscandian Shield. The thickness varies between 42 and 52 km but can reach greater thicknesses up to 65 km. The map of Moho topography (Fig. 6) by Luosto (1991, 1997) has become a seminal starting point for both seismic tomography and reflection studies. Malaska & Hyvrnen (2000) improved the crustal model by interpolating and smoothing the published 2D seismic models into a 3D model. Korsman et al. (1999) analysed the 160 km wide and 840 km long GGT-SVEKA transect using all existing geophysical and geological information. The transect covers the western part of the Archaean Karelian Province, crosses its boundary zone towards the Palaeoproterozoic Svecofennian arc complex, traverses the main tectonic units of the northern part of the Svecofennian complex, and ends in the Mesoproterozoic rapakivi granite area. The main conclusion was that thinner crust (with average crustal thickness of 45 km) is found in regions that have experienced one or more anorogenic extensional events, whereas the orogenic crust of the Svecofennian Domain has much greater thickness, on average 55 km, and orogenic collapse, normally producing a thinned crust, was apparently inhibited. The crust was thickened tectonically and by magmatic under- and intraplating. The entire Svecofennian crust equilibrated soon after magmatic underplating, between 1.885 and 1.800 Ga, and mafic magmatism increased the density of the crust, helping to preserve the thickened crust (Korsman et al. 1999). A systematic reanalysis of the old DSS data by Pavlenkova et al. (2001) more or less corraborated the previous findings by Luosto (1991, 1997) and Korja et al. (1993). They divided the crust into three layers with velocities of 6.0-6.4, 6.5-6.6 and 6.8-7.0 km s -1. The interface between the two upper layers is the most stable and its depth increases from 9 - 1 0 km in the Kola province to 16-18 km in southern Finland. The boundary separating middle and lower crustal layers is very stable throughout the region and is situated at depths of 27-30 km. The thickness of the lower crust varies from 10 to 12 km. In the region of thickened crust, an additional high-velocity lower crustal layer with velocities of 7.2-7.4 km s-1 was necessary to explain the observations. Reflection profiling

The BABEL seismic experiment, in 1989, revealed a great variety of structures in the crust of the Fennoscandian (Baltic) Shield.

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Fig. 6. The Moho depth map of Fennoscandia drawn from data collected by Luosto (1991, 1997), Korsmanet al. (1999) and Sandoval et al. (2003). Black dots show original data points (=sampling of velocity models) of Korsman et al. (1999).

Prominent reflections in the uppermost mantle were originally interpreted to define a mantle convergence zone where Proterozoic mantle underthrust Archaean lithosphere (BABEL Working Group 1990, 1993). A southward-dipping zone of less reflectivity was interpreted as the major strain zone accommodating horizontal shortening in the crust (Snyder e t al. 1996). Snyder (2002) has reinterpreted data from the northernmost BABEL profiles and concluded that the Archaean block forms a wedge of uppermost mantle rocks embedded in a Proterozoic block. The extent of the Archaean rocks is as great as 1 0 0 - 2 0 0 k m at Moho depths, suggesting that the Archaean lithosphere is laterally more extensive at depth than at the surface. In his model, Snyder (2002) suggested that the crustal convergence was partitioned between a wedge of weaker Archaean crust, thrusting higher in the crust to the south and channel flow within the lower crust. Altogether, he preferred a shear deformation origin to a magmatic enhancement of impedance contrasts for the 'bright Moho' reflector observed on the northernmost BABEL lines. Korja e t al. (2001, 2006) have reinterpreted the BABEL lines 1 and 6, along with new marine gravity data in the central part of the Gulf of Bothnia. The two parallel, north-south lines 1 and 6 have a transparent central area flanked by reflective structures dipping away from the centre. In the northern part of the profile, bright saucer-shaped reflectors have been interpreted as post-Jotnian diabase sills (BABEL Working Group 1993) that crop out on the sea bed (Korja e t al. 2001). Otherwise, the northern part has a complicated reflectivity pattern with a weakly

reflective upper crust and highly reflective, northward-dipping structures in the lower crust. Korja e t al. (2001) explained the reflectors in the lower crust as a double Moho structure formed by under- and intraplating of previously thinned lower crust in the northern part of lines 1 and 6. They concluded that the thickness and strength of the lithosphere were great enough to prevent the heat pulse from the mantle rupturing the crust; instead, minor extension and rifting took place. In the south, extensional shear zones are seen as a band of dipping reflectors levelling out horizontally at a depth of 40 km underneath weakly reflective areas interpreted as rapakivi granite batholiths. With these new interpretations of BABEL data it has become increasingly evident that a 3D approach to the structure of the Fennoscandian lithosphere is essential. Additional seismic reflection data crossing the geological structures of the central and thick part of the Fennoscandian Shield are necessary, and the reflection results must be supplemented with other geoscientific data, most notably with potential field data. Two major reflection profiles have been studied in the eastern parts of the Shield. Along the profile 1-EV of the International Global Geosciences Transects programme (GGT), the dataset incorporates geological, gravimetric and magnetic maps, compiled along a 100 km wide zone, and seismic (CDP and DSS), magnetotelluric and interpretational structural deep sections. The transect crossed the Karelian craton, the Palaeoproterozoic Central Russian region, and Belomorian fold belts and the Svecofenian Domain.

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD

The preliminary interpretation of profile 4B of SVEKALAPKO has been presented by Berzin et al. (2002). The profile runs almost east-west from the southern end of the White Sea to the RussianFinnish border at the northeastern end of the SVEKA profile. In the detailed CDP cross-sections obtained from wide-angle and near-vertical reflections several inclined boundaries are traced from the surface to 2 5 - 3 0 km depth. The boundaries correlate with the well-known fault zone between the Belomorian Mobile Belt and the Karelian Craton. The DSS data also show these inclined reflectors, as well as a near-horizontal boundary at a depth of 10-15 km, under a low-velocity zone. A lower crustal boundary at a depth of 30 km has no clear expression in the CDP reflectivity pattern. Strong PmP reflections from the Moho at 40 km coincide with the boundary between reflective lower crust and transparent upper mantle on the CDP section. Berzin et al. (2002) suggested that the near-horizontal crustal boundaries and the Moho are transition zones with high-velocity gradients and not sharp discontinuities. Further investigation and modelling of profile 4B are in progress. The continuation of 4B along the SVEKA line in Finland was originally a part of the SVEKALAPKO project plan, but was not measured until 2001, as line 1 (see Fig. 1) of the Finnish Reflection Experiment (FIRE). Along FIRE 1 the lower crust is weakly reflective, which has been suggested to indicate magmatic underplating in addition to tectonic thickening. The data along FIRE 1 are of good quality, and changes in the reflectivity patterns are correlatable to surface geology (Heikkinen et al. 2003).

S V E K A L A P K O seismic t o m o g r a p h y e x p e r i m e n t

The SVEKALAPKO seismic tomography experiment consisted of a network of 128 temporary stations (40 broadband and 88 shortperiod instruments). Data from 15 permanent seismic observatories were also used. The array covered the Shield from 59 to 68~ and 18 to 34~ (Fig. 2)~ The array was designed for maximum ray density of teleseismic sources at the depth range between 100 and 300 km. From August 1998 to May 1999 more than 1300 local, regional and teleseismic events were recorded. The first results of multidisciplinary seismic tomography, anisotropy and receiver function studies of the dataset have been presented (Bock & SVEKALAPKO Seismic Tomography Working Group 2001; Bruneton et al. 2002, 2004a,b; Sandoval 2002; Funke et al. 2003; Alinaghi et al. 2003; Sandoval et al. 2003, 2004; Yliniemi et al. 2004; Plomerovfi et al. 2006). One of the key targets of the experiment was the upper mantle. Variations in the crustal velocities, however, distort the teleseismic wave fronts, causing spatial travel-time variations. Unless appropriate corrections for the crustal effects are used, the latter are back-projected during the inversion and lead to artefacts in the derived structure of the upper mantle. Only the corrected teleseismic travel-time observations were inverted for mantle structure. Comparing the inversion results of the synthetic travel-time dataset, with and without crustal corrections, demonstrates the need to apply appropriate 3D crustal corrections in high-resolution regional tomography for upper mantle structure beneath the Fennoscandian Shield (e.g. Sandoval 2002; Bruneton et al. 2004b). Crustal models. Sandoval (2002) prepared a refined 3D crustal velocity model from existing DSS data; for example, Moho topography and lateral variations in average velocity as defined by Luosto (1991, 1997), Korja et al. (1993) and Korsman et al. (1999). Other geophysical information (e.g. gravity data) was not included in the model at this stage. The model was constructed by first determining the Moho interface and topography, followed by calculation of the 3D velocity structure. In the case of

549

SVEKALAPKO, the anomalous high-velocity lower crust demanded an additional step. The Svecofennian crust is thicker than the crust of the Karelian or Lapland-Kola realm. The thinnest crust, with a thickness between 38 and 42 km, surrounds the deep central Svecofennian Domain. The lack of sedimentary rocks on the surface is a major advantage when constructing a priori models. In the eastern parts of the Shield, where recent data were scarce, a uniform Moho depth of 42 km was assigned. Weighting and interpolating the data a Moho depth uncertainty of __%2 km at minimum was obtained for the highest-quality reflectors and a Moho uncertainty of +_ 10 km for the lowest-quality reflectors. The maximum crustal thickness of 64 km occurs beneath the surface contact region between the Archaean and the Svecofennian regions. A secondary maximum of crustal thickness exists beneath the western coast of Finland, with a value of 56 km. A narrow trough, with crustal thickness up to 52 km, stretches from the main maximum and becomes shallower (48 km on average) to the north. In central Fennoscandia, two interfaces in the 3D velocity model are the Moho and the upper limit of the lower crust. Sandoval et al. (2003) used a constant velocity of 8.3 km s -1 at the base of the model at 70 km and a constant value of 5.9 km s -1 at the surface, both values chosen as an average value derived from the DSS experiments. The top of the high-velocity lower crust under the central part of the Baltic Shield was defined as the depth at which the P-wave velocity reaches 7.0 km s -1 (Korja et al. 1993). Increased P-wave velocities are observed just above the Moho interface. Two velocity gradients were defined, the first between the surface and the top of the high-velocity lower crust and the second between the upper limit of the high-velocity lower crust and the Moho. The S model was derived from the P model by assuming the same Moho interface for both models and by applying Vp/V~ ratios of 1.71, 1.76 and 1.78 for the upper crust, lower crust and upper mantle, respectively (Luosto 1997; Korsman et al. 1999). The analysis of Sandoval (2002) indicated that stations situated in the centre of the SVEKALAPKO array show the largest positive delays, with 0.22 s for the P model and 0.24 s for the S model. Positive delays occur in areas with thicker crust, although the retarding effect of the crust is reduced by up to 50% by the presence of fast lower crust. Early arrivals are obtained in the surrounding areas, with minimum values of - 0 . 4 3 s for the P model and - 0 . 6 2 s for the S model. The average crustal thickness here is 42 km (7 km thicker than IASP'91), but the high-velocity lower crust increases the average crustal velocity to values close to 6.4 km s -1 (compared with 6.1 km s -1 in IASP'91). The P-wave delays relative to IASP'91 have a distribution centred at - 0 . 1 5 s with maximum and minimum delays of 0.37 and - 0 . 4 9 s, respectively. The S-wave delay distribution has a similar pattern and is centred at - 0 . 3 0 s with a maximum value of 0.38 s and a minimum delay of - 0 . 7 6 s. This is an expected result, as the S model is derived from the P model. In both distributions a secondary maximum can be observed between 0.00 and 0.20 s. This population of positive delays is caused by the points that lie in the area with thickest crust (Sandoval 2002). A further improvement of the crustal model has been prepared by Kozlovskaja et al. (2004a). They used the crustal Vp velocity model of Sandoval et al. (2003), petrophysical data on the density of bedrock in Finland and new velocity data for the crust obtained from SVEKALAPKO studies of local events, a priori, information for inversion of the observed Bouguer anomaly. A four density-layer model was obtained: the upper crust varying from 2610 to 2900 kg m -3, the middle and lower crust from 2800 to 3000 kg m -3, underlying high-velocity lower crust from 3050 to 3250 kg m -3, and the density beneath the Moho boundary varying from 3250 to 3245 kg m -3. The resulting model demonstrates that the Moho depression in central and southern Finland is not reflected in the observed

550

S.-E. H J E L T E T A L .

SVEKALAPKO P-velocity 250 km depth

SVEKALAPKO P-velocity 70 km depth I000

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Fig. 7. P-wave velocity structure beneath the Fennoscandian Shield from high-resolution teleseismic tomography. (a) Deviations from the0 IASP'91 velocity model (in percent); (b) a schematic illustration of the main structural tectonic elements of the crust and upper mantle beneath the study area (Sandoval et al. 2002).

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD

Bouguer anomaly. The depressions in central and southern Finland are fully compensated, or even overcompensated, by dense mafic rocks in the crust and hence no corresponding minimum of the Bouguer anomaly is observed. On the other hand, the Moho depression beneath the Gulf of Bothnia is compensated only in the southern part, resulting in a regional-scale minimum of the Bouguer anomaly in the northern part of the depression. The different degree of compensation may result from differences in the origin and age of the Moho boundary that is generally defined by the last major tectonothermal event in the area. Thus, formation of the thick crust and present-day Moho geometry in central and southern Finland was due to several consecutive tectonic processes during the Svecofennian orogeny between 1885 and 1800 Ga that were concluded by magmatic underplating (Korsman et al. 1999). The increased density in the upper and middle crust here resulted from both mafic magmatism and thrusting of highly metamorphosed crust toward the surface. On the other hand, the formation of the thick crust beneath the Gulf of Bothnia was most probably the result of accretion of two microcontinents (or terranes?), as proposed by Lahtinen et al. (2005), who suggested that the Svecofennian domain was formed as a result of five orogenic processes in the time period 1.92-1.88 Ga and that the whole Fennoscandian segment of the lithosphere was formed by the accretion of several microcontinents. Accretion of two crustal blocks with different densities explains the different degree of compensation of the Moho depression in the northern and southern parts of the Gulf of Bothnia. Upper mantle structure: receiver function studies. Bock & SVEKALAPKO Seismic Tomography Working Group (2001) and Alinaghi et al. (2003) stacked receiver functions to enhance converted P-to-S amplitudes. The arrival times of PS indicate a considerable thickening of crust across the Trans-European Suture Zone (TESZ) from 30 km in the German Basin to over 50 km in the Fennoscandian Shield. The change in crustal thickness across the TESZ was corroborated by previous seismic studies (EUGENO-S Working Group 1988; Gossler et al. 1999; Grad et al. 2002). The pronounced asthenosphere beneath the c. 100 km continental lithosphere of West-Central Europe abruptly terminates along the TESZ, altogether the TOR seismic experiment documents a surprisingly sharp boundary of the Baltic Shield along the TESZ in Denmark (e.g. Wilde-Pi6rko et al. 2002). The variations in Moho depths beneath the Baltic Shield ranging from 40 to 60 km and established by previous controlled-source seismic experiments are observed also in the receiver function data. The two major 410 km and 660 km upper mantle discontinuities are clearly observed both under the TOR profile and underneath the SVEKALAPKO network (Figs 8 and 9). Whereas the arrival times of converted P to S waves from 410 km and 660 km discontinuities undergo changes across the TESZ, at the southern edge of the Shield, the difference between the arrival times of P410s and P660s phases increases below the Shield. This is indicative of cooler upper mantle underneath the Precambrian Fennoscandian Shield than that of the Palaeozoic North German Basin. Across the SVEKALAPKO profile the thickness of the transition zone shows signs of increase in the eastern part whereas traces of some local anomalies can be found in the central parts (Bock & SVEKALAPKO Seismic Tomography Working Group 2001). However, generally the variations of depth to the 410 km and 670 km boundaries are small beneath the Shield. This result, together with thermal models of the lithosphere beneath the SVEKALAPKO area and thermobarometric data on mantle xenoliths in eastern Finland (Kukkonen et al. 2003), allows us to suppose that no significant temperature variations exist in the upper mantle beneath Finland. Upper mantle structure: teleseismic P-wave tomography. Using a non-linear teleseismic tomography algorithm, Sandoval (2002)

Moho? /

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551

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Fig. 8. Receiver function north- south traces from the TOR and SVEKALAPKO (19 station subset) seismic tomography arrays. Dark streaks are converted P-S waves emerging from boundaries with depth-increasing velocity. The change in distance scale and the break between the array profiles are indicated by the bold line (Bock & SVEKALAPKOSeismic Tomography Working Group 2001).

and Sandoval et al. (2003, 2004) found P-wave velocity variations of up to 4% throughout the SVEKALAPKO region. A positive velocity anomaly can be followed down to about 300 km depth beneath the centre of the array (Fig. 7) that correlates very well with the region of thickened crust. Sandoval et al. (2004) interpreted this as the signature of the deepest-reaching tectosphere beneath the Shield. The Archaean-Proterozoic suture zone does not show up as a perceptible structure in the mantle (Sandoval 2002; Sandoval et al. 2003, 2004). Because both thermal modelling of the lithosphere and receiver function studies revealed no significant temperature variations beneath the Shield, and the lithosphere has been in place since about 1.5 Ga, the high-velocity anomaly was interpreted as a continental 'keel' stabilized by compositional differences. Differences may have been created by extraction of melt during the formation of the thick lower crust in this region. However, interpretation of velocity variations revealed by teleseismic P-wave tomography in terms of composition is difficult, because the tomography reveals only relative values of P-wave velocities that cannot be directly compared with the values revealed by petrophysical studies of

552

S.-E. HJELT ETAL. Lsfitude (deg)

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The broadband part of the SVEKALAPKO seismic array showed clear signs of seismic anisotropy: time delays between the fast and slow shear split waves, and varying directions of the fast S polarization of the incoming teleseismic waves. In general, shear-wave splitting was detected at most stations and measured time delays of the slow shear waves were between 1 and 2 s, on average. Lateral variations of shear-wave splitting parameters indicate lateral variations of the anisotropic structure of the upper mantle beneath the SVEKALAPKO seismic array. The mantle lithosphere seems to consist of several large-scale domains with different orientation of anisotropy (Plomerovfi et al. 2002a,b, 2006). The splitting parameters are coherent within the individual lithospheric blocks, but vary from one block to another. According to Plomerovfi et al. (2002b), a subcrustal lithosphere of about 150 km is thick enough to accommodate this observed large-scale seismic anisotropy. Additional information about an anisotropic upper mantle structure was obtained from combined analysis of S-wave splitting parameters and direction-dependent P-wave residuals. In particular, strong anisotropy and uniform orientation of anisotropic material in the upper mantle was revealed beneath the Archaean domain. In contrast, the anisotropic pattern corresponding to the Proterozoic domain is more heterogeneous and weakly anisotropic (Plomerovfi et al. 2006; Kozlovskaya et al. 2006). Both large Archaean and Proterozoic tectonic units of the eastern part of the Shield seem to be composed of several smaller lithospheric domains with different orientation of large-scale mantle fabric that may result from different geological history.

59

O10

Fig. 9. Time domain (a) and migrated (b) sections of move-out corrected receiver functions along the TOR and SVEKALAPKOarrays, stacked in 50 km wide windows with moving intervals of 10 krn providing an 80% overlap between adjacent windows. The dark positive amplitudes represent an increase of the S-wave velocities with depth (Alinaghi et al. 2003).

mantle rocks. In addition, Sandoval et al. (2003, 2004) inferred that the anomalous upper mantle velocities are well defined horizontally, but vertically smeared both upwards and downwards. Sandoval et al. (2003) also assessed the influence of the crustal correction on the resolution of the upper mantle structure. Two inversion tests were used, first a synthetic dataset was inverted (1) for crustal and mantle structure combined and (2) for mantle structure only, after correction for crustal effects. They demonstrated a strong 'leakage' of crustal effects down to 200 km. The effects can still be noticed at 450 km depth. However, introducing crustal corrections allows significant reduction in the effect of the crust on upper mantle velocities. Therefore, improved crustal models will be essential in improving upper mantle models beneath the Fennoscandian Shield.

Upper mantle structure: seismic anisotropy. Anisotropy of seismic velocities can result for various reasons; for example, from oriented fractures in the upper crust, from alternating layers with different isotropic velocities, or because of the alignment of crystals of rock-forming minerals in a stress field. A highly heterogeneous crust can contribute only up to about 10% of the observed large-scale anisotropy. According to the analysis of Plomerovfi et al. (2001, 2002a,b, 2006) the main source of the large-scale anisotropy beneath Scandinavia has to be in the upper mantle, caused especially by its large-scale fabric owing to preferred orientation of olivine.

Upper mantle structure: surface-wave investigations. Surface-wave studies have an advantage over teleseismic body-wave tomography, because they allow estimation of the absolute values of S-wave velocity in the upper mantle that are directly comparable with values estimated by studies of upper mantle xenoliths from Finland (Kukkonen et al. 2003) and from other Precambrian areas (Weiss et al. 1999; Griffin et al. 2003). Bruneton et al. (2002, 2004a,b) used data from the 2D grid of 46 broadband stations of the SVEKALAPKO array. Fundamentalmode Rayleigh wave arrival times with periods between 10.5 and 190 s were used to investigate the S-wave velocity in the upper mantle beneath the SVEKALAPKO array. Joint inversion for the S-wave velocity model under the array and the shape of incoming wave fronts reduced the artefacts caused by structure outside the study region (Fig. 10). The results of inversion for the upper mantle seem to be very well constrained to a depth of 150 km and weakly dependent on crustal thickness (Bruneton et al. 2004a,b). A regional average 1D shear-wave velocity model for the SVEKALAPKO area to a depth of 300 km (Bruneton et al. 2004b) has S-wave velocities that are c. 4% faster than in standard Earth models. The model lacks a substantial low-velocity layer that could define the base of the lithosphere. This indicates a cold upper mantle beneath the SVEKALAPKO array and agrees with the results of Sandoval (2002), Alinaghi et al. (2003) and Sandoval et al. (2004). The 3D S-wave velocity model (Bruneton et al. 2004a) shows both lateral and vertical S-wave velocity variations ( -t- 3%) that can be explained by variations of composition of upper mantle peridotites; for example, different modal proportions of rock-forming minerals (mainly olivine and orthopyroxene) and differences in M g / ( M g + Fe) ratio (Weiss et al. 1999; Griffin et al. 2003). The model obtained by Bruneton et al. (2002, 2004a,b) is in agreement with the result of surface-wave studies (Fig. 11) by Funke et al. (2003). The 3D model has a mean crustal thickness of 52 km. It reveals positive and negative S-wave velocity variations, but no perceptible low-velocity zone in the upper 300 km. The absolute values of S-wave velocities in the upper mantle beneath the SVEKALAPKO area vary from 4.6 to 4.8 km s - t .

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIAN SHIELD

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Similar values were obtained by Griffin et al. (2003) by analysis of xenolith samples from Proterozoic and Archaean upper mantle around the world. However, the models of Bruneton et al. (2002, 2004a,b) and Funke et al. (2003) do not show S-wave velocities that are systematically higher beneath Archaean domains than beneath the Proterozoic, as one could expect from global xenolith analysis (Griffin et al. 2003). Instead, both S-wave velocity models revealed a laterally and vertically heterogeneous structure for the upper mantle. Because the upper mantle temperatures beneath the Shield are uniform (Alinaghi et al. 2003; Kukkonen et al. 2003), the S-wave velocity heterogeneities identified by Bruneton et al. (2002, 2004a,b) and Funke et al. (2003) can be explained by either compositional variations or anisotropy of the seismic velocity. Both models revealed a layer of very high S-wave velocity (down to c. 120 kin) that is widespread beneath the boundary of

553

Archaean and Proterozoic domains (Fig. 11). The velocity in this layer (up to 4.8 km s -1) agrees with the value estimated from highly depleted lherzolite and harzburgite xenoliths from eastern Finland that contain about 70% olivine with a high Mg/ (Mg + Fe) ratio of about 0.9 (Kukkonen et al. 2003; Bruneton et al. 2004a). The velocity beneath this layer is 4.65-4.7 km sand lower than that estimated from xenoliths from Finland, which may be due to a less depleted composition. However, these values are also slightly lower than the velocity in Archaean mantle xenoliths reported by Griffin et al. (2006). This can be explained by seismic anisotropy revealed beneath the Archaean domain by Plomerovfi et al. (2002b, 2006) and Kozlovskaya et al. (2004b). Similar stratification of the upper mantle was revealed also beneath the Slave craton, where an ultradepleted upper mantle layer is underlain by a more typical depleted Archaean mantle with a higher orthopyroxene/olivine ratio. Griffin et al. (1999) proposed that the ultradepleted layer of the Slave mantle was generated in a collisional setting. In the Proterozoic area in Finland, the surface-wave studies revealed several domains with slightly varying stratified velocity structure. Generally, the velocity beneath the Proterozoic domain to a depth of 100-120 km varies from 4.6 to 4.7 km s -1, which is lower than the values estimated from xenoliths from eastern Finland for the same depth range. These values agree with those estimated from Proterozoic xenoliths worldwide (Griffin et al. 2003), which may indicate a more fertile composition. Beneath this layer, the velocity values agree well with those estimated from xenoliths from Finland (Bruneton et al. 2004b), indicating highly depleted mantle there. However, such a distribution of mantle material (e.g. a high-density fertile layer over a low-density depleted layer) would not be gravitationally stable. Therefore, the more feasible explanation is that the low velocity in the upper layer is due to contamination by the crustal material. Upper mantle structure: local event studies. Yliniemi et al. (2004) presented results of forward raytrace modelling of reflected and refracted P waves of the strongest local events registered by the SVEKALAPKO array. They reported two types of mantle reflections: subhorizontal and gently dipping reflectors below the Moho at a depth of 70-90 km, and phases originating from a depth of 100-130km. Based on the irregular character of reflectors of the first group, on their different spatial orientation and on a correlation with Moho offsets, they interpreted the boundaries of the first group as relicts of ancient subduction and collision processes. This explanation is in accord with that of Alinaghi et al. (2003), who did not identify any upper mantle discontinuities except for global ones at 410km and 660km. This can be explained by both masking effect of multiples of the Moho conversions and the irregular nature of boundaries at a depth of 70-80 km that do not produce coherent P-to-SV conversions. The position of the reflectors from the first group beneath the SVEKA profile coincides with the location of a highly depleted upper mantle layer (Fig. 11). The reflectors of the second group coincide spatially with an area of slight change of both P-wave velocity revealed by teleseismic tomography (Sandoval et al. 2004) and S-wave velocity revealed by surface-wave studies (Funke et al. 2003) (Fig. 11). Therefore it can be attributed to a lithological contact between a highly depleted upper mantle layer and more typical Archaean mantle. This boundary correlates also with the estimated depth to the lower boundary of the mechanically strong lithosphere (i.e. the depth at which the ductile strength is reduced to 50 MPa; see Fig. 11) (Kaikkonen et al. 2002). The rheological weakening and deviatoric stresses may result in reorientation of anisotropic minerals (mainly olivine and orthopyroxene) and a horizontal foliation, which would explain the high reflectivity at this depth (Weiss et al. 1999). The heterogeneous velocity structure of the subcontinental lithospheric mantle (SCLM) beneath the contact of Archaean

554

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-0.16 0.0 0,16 Variation of Vs (km/s)

and Proterozoic domains in the SVEKALAPKO area, together with traces of former subduction and collisional processes, suggests a complex history of formation and stabilization of the lithosphere in the region. This is in agreement with the recent analysis of xenoliths representing Archaean and Proterozoic SCLM by Griffin e t al. (2003). They proposed that formation of SCLM beneath Archaean and Proterozoic domains was the result of two very different tectonic regimes. Most of the Archaean SCLM was formed by high-degree melting at sub-lithospheric depth producing thick, highly depleted volumes of buoyant upper mantle material that formed the roots of continents. This tectonic regime seems have operated until c. 2.5 Ga, after which it ended as a result of secular cooling of the Earth. Subsequently, a regime similar to that of modern plate tectonics was established, which included moderate depletion at spreading centres, subduction and cyclic delamination and replacement. The regime was operating during the late Archaean and Proterozoic, resulting in progressive modification and a heterogeneous structure of former Archaean SCLM. C o m p a r i s o n o f s e i s m i c a n d e l e c t r i c a l models. We can compare the results of the SVEKALAPKO seismic experiment and the BEAR experiments only in the area covered by both arrays (Fig. 2) in the central Fennoscandian Shield. Such a comparison makes sense only in the case when both the seismic and electromagnetic datasets are sensitive to the same structures or properties of the SCLM. Interpretation of the SVEKALAPKO seismic data has demonstrated, however, that the velocity heterogeneities beneath the cold and stable central Fennoscandian Shield are explained mainly by compositional variations in upper mantle peridotites formed as a result of ancient tectonic processes. Such compositional variations have no effect on electrical conductivity and cannot be detected by MT data. However, the partially molten asthenosphere would decrease both seismic velocities and electrical conductivity and would have an effect upon both seismic and MT data, as is observed in young and active regions. Therefore, comparison of the results of the SVEKALAPKO seismic experiment and the BEAR experiments can be used to answer the question about the possible presence of

Fig. 11. The subcrustal structure along the SVEKA profile based on 3D models of S-wave velocity in the upper 300 km of the mantle (reproduced by courtesy of Funke & Friederich 2003). Fine black lines show intracrustal boundaries and bold black line shows the Moho boundary (Korsman et al. 1999; Kozlovskaya & Yliniemi ! 999; Kozlovskaya et al. 2004a). Black triangles indicate location of the shot points of the SVEKA profile. Bold yellow lines indicate position of the upper mantle reflectors derived by Yliniemi et al. (2004). Position of the boundary of mechanically strong lithosphere (after Kaikkonen et aL 2002) is indicated by dashed blue line. The boundary of anisotropic Archaean mantle (red dot-dashline) is adopted from Plomerovfi et al. (2002b) and Kozlovskaya et al. (2004b).

partially molten asthenosphere beneath the central Fennoscandian Shield and to explain the origin of the upper mantle conductivity in this area. Seismic velocity models obtained by various seismic techniques demonstrate that there is no low-velocity layer that can be attributed to partially molten asthenosphere down to the depth of c. 300 km. This generally agrees with the interpretation of the BEAR data, showing that in the central part of the Shield the conductivity is comparable with the conductivities of the dry olivine model. However, the MT data also indicate an abrupt increase in conductivity somewhere below 200 km in the area, which is covered by both datasets. The exact depth and geometry of this feature is uncertain, and cannot be retrieved from MT data alone. Therefore, the conducting region could be caused by partial melting, or dissolved water in olivine or the 410 km phase transition. Partial melting is less plausible because seismic methods do not detect a layer that could be associated with such phenomena. Therefore this conducting feature can be caused by dissolved water, if it proves to be shallower than 410 km. In the northern part of the Fennoscandian Shield (site B42 of BEAR and KIR of Jones 1982, 1983), the enhanced conductivity is observed at much shallower depth ( 1 5 0 k m at KIR and 170 km at B42). This enhanced conductivity cannot be explained by graphite, because the graphite-diamond transition takes place at shallower depths in the region (for further discussion, see Lahti et al. 2005). Therefore, it may be caused either by partial melting (asthenosphere) or by dissolved water. However, we cannot distinguish between these two alternatives, because the area was not covered by the SVEKALAPKO array and the velocity structure of the upper mantle is poorly known at present.

Conclusions

The analysis and interpretation of the latest large-scale seismic and EM arrays on the Fennoscandian Shield are far from completed. Work is in progress on S-wave tomography of both teleseismic and local events, for anisotropy studies and

ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIAN SHIELD

improvements of the crustal velocity models. The preparation of a 3D conductivity model of the Fennoscandian lithosphere has been painstakingly complicated, but the final tests are under way. The centre of the Fennoscandian Shield is characterized by thickened crust. This is accompanied by seismic velocity anomalies that extend to at least 250 or 300 km depth. The Pand S-wave velocities seem to be up to 4% higher than in the global Earth models for the upper mantle down to 200 km. The difference between the 410 km and 660 km arrival times increases beneath the Shield, which corroborates the interpretations of a thick, cold and early stabilized lithosphere. The depth to the 410 km and 670 km boundaries is very stable, which implies that no significant temperature variations exist in the upper mantle beneath the Shield. No evidence for a substantial mantle low-velocity layer (LVL) has been obtained so far. P- and S-wave velocity inhomogeneities in the mantle lithosphere are most probably explained by compositional variations and/or by seismic anisotropy. According to electromagnetic investigations: (1) upper mantle conductivity is laterally heterogeneous in the Fennoscandian Shield; (2) there must be a layer of enhanced conductivity in the upper mantle beneath the entire Fennoscandia Shield, which has 10-100 times higher conductivities than predicted by dry olivine; (3) the depth to the top of the conducting layer (or a region of enhanced conductivity) is 150-170 km in the northern part of the Shield, whereas in the central part of the Shield it must be deeper than 200 km; (4) the conductivity of mantle lithosphere (above the conducting layer) is roughly 10 times higher than the dry olivine conductivity in the northern part of the Shield, whereas in the central part the conductivity is compatible with the conductivities of the dry olivine model. The interpretation of BEAR and SVEKALAPKO data has demonstrated that the structure of the upper mantle beneath the shield is heterogeneous; this supports the major conclusion obtained already (e.g. from the interpretation of BABEL reflection experiments and previous magnetotelluric soundings), namely, that the structure of the Fennoscandian lithosphere, in general, is highly variable and complicated. Therefore, models of the lithosphere evolution must be revised to accommodate lateral and vertical heterogeneity. The architecture of the Fennoscandian deep lithosphere is not yet known, because of inadequate spatial sampling. The SVEKALAPKO seismic tomography array was relatively small compared with the size of the Shield, although the lateral sampling interval was small. As a result, the tomography array provided detailed images of the upper mantle, but from only a rather limited region. The BEAR array, on the other hand, covered the entire Shield, but the distance between sites was large, making it difficult to define the exact location of the borders of lithospheric units. Total coverage and denser spatial sampling is therefore required for the detailed understanding of the structure of deep lithosphere-upper mantle in this craton. Among the most interesting problems remaining is to define the structure and geometry of the transition between the cratonic and oceanic lithosphere or a transition from oceanic lithosphere to cratonic tectosphere. In addition, seismic reflection data are needed especially across the crustal structures of the thick central part of the Fennoscandian Shield, and reflection results have to be complemented with other geophysical data, most notably with potential field and electromagnetic data, to obtain improved 3D understanding of the crust and upper mantle. The top-quality seismic images provided by the FIRE project as well as joint inversion of potential field and seismic data will certainly contribute to a better understanding of the birth and structure of the Fennoscandian lithosphere.

555

Appendix Participating organizations of the SVEKALAPKO Seismic Tomography Working Group CZECH REPUBLIC Geophysical Institute of CAS, Prague GERMANY GFZ Potsdam University of Stuttgart FINLAND University of Oulu University of Helsinki FRANCE University of Grenoble University of Strasbourg NETHERLANDS Utrecht University POLAND Warsaw University Institute of Geophysics of PAS RUSSIA Kola Scientific Center RAS Apatity Institute of the Physics of the Earth Moscow St. Petersburg University Spetzgeofisika MNR Moscow SWEDEN University of Uppsala SWITZERLAND Institute of Geophysics, ETH Zurich The SVEKALAPKO Seismic Tomography Working Group consists of following individuals: U. Achauer, A. Alinaghi, J. Ansorge, G. Bock, M. Bruneton, W. Friederich, M. Grad, A. Guterch, P. Heikkinen, S.-E. Hjelt, T. Hyv6nen, E. Isanina, J.-P. Ikonen, E. Kissling, K. Komminaho, A. Korja, E. Kozlovskaya, M. V. Nevsky, N. Pavlenkova, H. Pedersen, J. Plomerovfi, T. Raita, O. Riznichenko, R. G. Roberts, S. Sandoval, I. A. Sanina, N. V. Sharov, J. Tiikkainen, S. G. Volosov, E. Wieland, K. Wyegalla, J. Yliniemi and Y. Yurov.

Participating organizations and individuals of the BEAR Working Group FINLAND Finnish Meteorological Institute, Geophysical Research, Division, Helsinki, and NurmijSxvi Geophysical Observatory, Helsinki and Nurmijarvi, Finland Team leader: A. Viljanen Team members: K. Pajunp~i~i,H. Nevanlinna University of Oulu, Institute of Geosciences, and Geological Survey of Finland, Oulu and Espoo, Finland Team leader: T. Korja Team members: S.-E. Hjelt, P. Kaikkonen, I. Lahti, I. Silvola, J. Tiikkainen, E. Kozlovskaya

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S.-E. HJELT ETAL.

GERMANY Technical University of Braunschweig, Institute for Geophysics, Braunschweig, Germany Team leader: K. Roden University of Goettingen, Geophysical Institute, Goettingen, Germany

This paper is dedicated to our SVEKALAPKO colleague Dr. GiJnter Bock, who tragically lost his life in an airplane crash in November 2002. The following members of the Working Groups have made valuable contributions to the manuscript: A. Korja, S. Sandoval, A. Alinaghi, M. Bruneton, M. Engels, W. Friederich, S. Funke, V. HaRk, V. Kobzova, A. Kovtun, N. Palshin, H. Pedersen, L. B. Pedersen, J. Plomerovfi, M. Smimov, E. Sokolova, I. Varentsov, and A. Zhamaletdinov. We wish to express our thanks to two anonymous reviewers for fruitful comments.

Team leader: K. Bahr Team member: E. Steveling

References

GeoForschungsZentrum-Potsdam, Potsdam, Germany

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The Svecofennian orogen: a collage of microcontinents and island arcs A N N A K A I S A KORJA l, RAIMO LAHTINEN 2 & MIKKO NIRONEN 2

1Institute of Seismology, P.O. Box 68, FI-O0014 University of Helsinki, Helsinki, Finland (e-mail: Annakaisa.Korja@ helsinki.fi) 2Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland

Based on an integrated study of geologicaland geophysicaldata, a tectonic model for the Palaeoproterozoic evolutionof the Svecofennianorogen within the FennoscandianShield at the northwestern corner of the East European Craton is proposed. The Svecofennian orogen is suggested to have formed during five, partly overlapping, orogenies: Lapland-Savo, Lapland-Kola, Fennian, Nordic and Svecobaltic. The Svecofennianorogen evolved in four major stages, involving microcontinentaccretion (1.92-1.88 Ga), large-scale extensionof the accreted crust (1.87-1.84 Ga), continent-continent collision (1.87-1.79 Ga) and finally gravitationalcollapse (1.79 and 1.77 Ga). The stages partly overlapped in time and space, as different processes operated simultaneouslyin different parts of the plates. In the Lapland-Savo and Fennian orogenies, microcontinents (suspect terranes) and island arcs were accreted to the Karelianmicrocontinent,which itself was accreting to Laurentiain the Lapland-Kola orogeny. The formation of the Svecofennian orogen was finalizedin two continentalcollisionsproducingthe Nordic orogen in the west (Fennoscandia-Amazonia)and Svecobaltic orogen in the SSW (Fennoscandia-Sarmatia).The collisionswere immediatelyfollowed by gravitationalcollapse. Abstract:

Orogeny is, by definition, a process of creation of mountain belts by tectonic activity (Bates & Jackson 1995). Orogenic belts are characterized by folding, faulting, regional metamorphism and igneous activity. In terms of plate-tectonic theory, orogenic belts mark sites of continent-continent or continent-island arc collision zones at convergent, destructive plate margins. At first, convergence is accommodated by subduction and later by tectonic thickening of one or both of the plates (Fig. la). When converging plates are moving at oblique angles, major strike-slip faults parallel to the subduction zone will develop. These faults may develop into transform plate boundaries where pieces of the margin are transported along strike. If such smaller fragments bordered by fault zones are later recognized in an orogen, they are interpreted to be suspect, exotic or translated terranes. Good examples are found along the west coast of North America (Jones et al. 1983). Oblique convergence also initiates transtensional regimes where microplates, including remnants of the continent and/or island arcs, are formed (e.g. Woodlark Basin in Indonesia; Hall 2002). The fate of the microplate is to collide either with the parent continent or with another continent on the other side of the nascent ocean. Good examples of the latter are found in the Tethysides (Stampfli & Borel 2002; von Raumer et al. 2003). Suspect terranes can also be formed by escape tectonics where smaller fragments of colliding plates are pushed aside from the main collision front into areas of thinner crust and lithosphere, and thus the orogen spreads laterally via major strike-slip faults. Orogenies are referred to as either collisional or accretionary (Windley 1993) depending on the dominant type of colliding plates. Collisional orogenies occur when large continental plates collide and in these the crust is mostly reworked. The formation of an accretionary orogen is more diverse, as it may involve the accretion of arc terranes formed along long-lived convergent margins, of exotic terranes split from neighbouring continents, and of oceanic seamounts. Lateral growth of continental plates mainly takes place in accretionary orogens. In both types of orogenies, the thickening of the crust and lithosphere may spread to the adjacent areas by escape tectonics. In general, all major orogenies begin in the accretionary modes at convergent margins and some of them evolve into collisional ones. Plate-tectonic theory accounts poorly for the effect of gravity and gravitational instabilities produced by the thickening of the crust. Thermal instabilities are also induced during orogenies. These anomalies are the driving force of gravitational collapse (Rey et al. 2001), whereby thicker orogenic crust is thinned, leading to the thickening of the adjacent crust (Fig. l b - d ) .

Another stabilizing phenomenon that may take place is lithospheric delamination, in which the cooler and denser parts of the thickened lithosphere detach and sink and are replaced by warmer and lighter asthenospheric material (e.g. Platt & England 1994). This process leads to increase in heat flow, magmatic underplating of the crust, and regional extension (Fig. l e - h ) . The orogenic cycle includes pre-collisional, syncollisional and post-collisional tectonic and magmatic stages in the plate-tectonic framework. Emphasizing gravity, the orogenic cycle is characterized by thickening of the crust, thermal maturation, partial melting, and syn- to post-convergence gravitational collapse (Vanderhaeghe & Teyssier 2001). The pre-collisional tectonics includes subduction of oceanic material and obduction of ophiolites. The syncollisional tectonics involves the colliding of accreted terranes or continents and associated thickening of the crust as well as lithospheric mantle, and post-collisional tectonics involves continued indentation of the colliding terranes or continents, and finally gravitational collapse. Although new continental crust is produced during the pre- and syncollisional phases of accretionary orogeny, the post-collisional phase is particularly important because, during this phase, many of the lithological associations are exhumed to higher structural levels. The major Phanerozoic orogens of the world are linear belts (Appalachians-Caledonides, Western Cordilleras, Alps, Himalayas), which have formed over long periods (of the order of 100 Ma). In detail, however, the orogens were formed in sequential short-lived (10-20 Ma) tectono-metamorphic events (such as the Acadian, Grampian and Laramide events), often referred to as separate orogenies. As in the Phanerozoic, the major Proterozoic orogenic belts were also formed over long periods and during semi-continuous processes (Windley 1993). One example of the proposed long-lasting orogenies is the Svecofennian (100 Ma; Ga~il & Gorbatschev 1987; Fig. 2). Although the Svecofennian has been classified as an accretionary orogeny (Windley 1993), the synorogenic, late-orogenic and postorogenic terminology above suggests that one continuous orogenic process formed the Svecofennian orogen. The first plate-tectonic model for the Svecofennian orogen was created by Hietanen (1975), who compared it with the Western Cordilleras of the USA. Since then, most of the plate-tectonic models have concentrated on the evolution of the ArchaeanProterozoic boundary, an important suture zone (Fig. 2; Bowes & Ga~il 1981; Koistinen 1981). Three models have been used to explain the development of the suture: (1) continental arc-continent collision (e.g. Gafil 1990; Lahtinen 1994);

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 561-578.0435-4052/06/$15.00 9 The Geological Society of London 2006.

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(g)

Gorbatschev & Bogdanova (1993) in outlining continuations of the Fennoscandian Palaeoproterozoic lithologies beneath areas covered by Phanerozoic rocks. They also defned Fennoscandia's southern boundary towards Sarmatia. This paper focuses on the processes associated with the formation of the Svecofennian orogen. It is a complementary paper to that by Lahtinen et al. (2005), which emphasized the plate kinematics forming the Fennoscandian Shield. It is suggested here that the so-called Svecofennian orogeny involved a sequence of accretionary and collisional events or orogenies that partly overlap in time and space, but resulted in different structural grains. It is also suggested that Svecofennian orogeny began with accretionary tectonics and was followed by continental collision tectonics. An attempt is made to define the colliding exotic or suspect terranes (microcontinents), arcs and continents. The evolution of the Svecofennian orogen is divided into five sequential events, here called orogenies. An attempt is also made to recognize extensional stages alternating with collisional ones during the Svecofennian evolution.

Geological background

Fig. 1. Orogenicprocesses forming or reworking the continental crust. (a) Thickening by stacking of continental slivers (after Coward 1994); (b-d) orogenic collapse (after Rey et al. 2001); (e-g) lithosphericdelamination (after Dewey 1988); (h) magmaticunderplating.After thickening, via stacking of continental slivers or suspect terranes (a), the crust is in gravitational potential disequilibriumwith its surroundings (b). The thickened crust may be thinned by upper crustal extension and sliding of material to the sides (e), or the lower crust may flow sideways(d). If the thickenedlithosphere is denser than its surroundings (e), then it may delaminate (f), and it may be replaced by asthenosphericmaterial from the sides (g). The asthenospheric material may initiate partial melting of the upper mantle. Rising melts may cause mafic underplating of the crust.

(2) back-arc-retro-arc basin development related to NE-directed subduction, occurring further to the SW of the suture (e.g. Hietanen 1975; Ga~il, 1986); (3) strike-slip movement, where all parts of the Svecofennian orogen are considered exotic (e.g. Park 1985). Wilson (1982) suggested an Andean-type plate-tectonic model for the Swedish part of the Svecofennian orogen. The increasing number of isotopic and geochemical datasets in the 1990s allowed more detailed and more complex models for the Svecofennian evolution. Lahtinen (1994) defined three arc complexes and three collisional events at 1.91-1.90 Ga, 1.89-1.88 Ga and 1.86-1.84 Ga and Nironen (1997) presented the first kinematic model for the Svecofennian orogen. Tectonic models for the contemporaneous Lapland-Kola orogen, situated in the northeastern part of the Fennoscandian Shield, fall into two groups: (1) models with the suture zone within the Lapland Granulite Belt (LGB) and subduction towards the NE (Fig. 2; Barbey et al. 1984; Krill 1985; Daly et al. 2001; Daly et al. 2006); (2) models with the suture zone within the Imandra Varzuga-Pechenga Belt (IVB and PeB in Fig. 2) and subduction towards the SW (Berthelsen & Marker 1986a; Marker 1990). Crustal-scale geophysical data in the 1980s added the vertical dimension and inspired correlation with modem analogues (BABEL Working Group 1990). In an integrated geologicalgeophysical study, Korja et al. (1993) attempted to locate the sutures and terrane boundaries within the Svecofennian orogen and proposed mantle underplating to account for the thick crust in central Finland. Later, Korja (1995) suggested that orogenic collapse may have played a role in the crustal evolution of southern Finland. Drill-core and geophysical data guided

The East European Craton (EEC) is composed of the FennoscandJan, Sarmatian and Volgo-Uralian crustal segments (Gorbatschev & Bogdanova 1993), of which the last two are mainly covered by Phanerozoic platform sediments. The Fennoscandian segment is exposed in its northern and central parts (Fennoscandian Shield), and covered by platform sediments in the south and by the Caledonides in the west (Fig. 2). This study concentrates on the Finnish and Swedish parts of the Fennoscandian Shield, with less emphasis on the Kola area, which has been described by Daly et al. (2006). The evolution of the other parts of the EEC has been described by Bogdanova et al. (2006) and Claesson et al. (2006). Ga~il & Gorbatschev (1987) have divided the Fennoscandian Shield into the Karelian, Belomorian and Kola Provinces, Svecofennian Domain, Transscandinavian Granite-Porphyry belt, Southwest Scandinavian Domain and Caledonides (Fig. 2a). Traditionally, the Archaean bedrock in the Fennoscandian Shield includes two cratonic nuclei, the Karelia and Kola Provinces, dispersed as fragments and subsequently reassembled during the Palaeoproterozoic (Ga~il & Gorbatschev 1987). Based on the existence of the Baltic-Bothnian megashear along the Swedish-Finnish national boundary (Berthelsen & Marker 1986b), lithological differences across the boundary, and especially the existence of the ophiolite-bearing Kittil~i allochthon (KA; Fig. 2), Lahtinen et al. (2005) have proposed that another unit, the Norrbotten craton, in the western part of the Karelian Province, may be a separate block (Fig 2b). In this paper, the Karelian Province is divided into the Karelian and Norrbotten cratons and intervening Proterozoic terranes. The Karelian craton (see Fig. 5b) encompasses the eastern part of the Karelian Province and the Belomorian Province (Fig. 2a). The Karelian craton consists of Archaean granitoid-gneiss complexes and supracrustal rocks (e.g. greenstones) ranging in age between 3.2 and 2.5 Ga (Sorjonen-Ward & Luukkonen 2005). In the Palaeoproterozoic, the Archaean lithologies were intruded by layered mafic intrusions (2.5-2.1 Ga), A-type granitoids (2.52.4 Ga) and mafic dykes (2.4-1.97 Ga). Autochthonous supracrustal rocks ranging from quartzites to pelites and mafic volcanic rocks were deposited on the Archaean basement from 2.45 Ga onwards. Allochthonous younger units comprising greywackes, the c. 1.95 Ga ophiolites at Outokumpu (O) and Jormua (J), as well as the Kittil~i allochthon (KA in Fig. 2) composed partly of oceanic crust (Koistinen 1981; Kontinen 1987; Peltonen et al. 1996; Hanski & Huhma 2005), were thrust onto the craton and its cover at 1.9 Ga. The enigmatic Central Lapland Granitoid Complex (CLGC; 1.85-1.77Ga) covers large areas of the

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Fig. 2. Simplified geological map of the Fennoscandian Shield, based on Koistinen et al. (2001). The shear zones are mainly interpreted from magnetic and gravity maps (Korhonen et aL 2002). (a) major geological units of the Fennoscandian Shield, after Gafil & Gorbatschev (1987). N, Northern Svecofennian Subprovince; C, Central Svecofennian Subprovince; S, Southern Svecofennian Subprovince. (b) Archaean cratonic terranes of the Shield. Archaean units: Norrbotten craton, Kola craton and Karelian craton, including Belomorian. Palaeoproterozoic units in Kola peninsula: IA, Inari area; PeB, Pechenga Belt; IVB, Imandra Varzuga Belt; UGT, Umba Granulite Terrane; TT, Tersk Terrane. Palaeoproterozoic units in Finland: LGB, Lapland Granulite Belt; KA, Kittili allochthon; CLGC, Central Lapland Granitoid Complex; SB, Savo Belt; CFGC, Central Finland Granitoid Complex; TB, Tampere Belt; HB, Hime Belt; UB, Uusimaa Belt. Palaeoproterozoic units in Sweden: SD, Skellefte district; BB, Bothnian Basin; BA, Bergslagen area; S6B, S6rmland Basin; OJB, Oskarshamn-J6nk6ping Belt; TIB, Transscandinavian Igneous Belt. J, Jormua; K, Knaften; O, Outokumpu; R, Revsund. BBZ, Baltic-Bothnia Megashear; HSZ, Hassela Shear Zone; LBZ, Ladoga-Bothnia Bay Zone.

564

A. KORJAETAL.

Karelian craton. Palaeoproterozoic (1.9-1.8 Ga) plutonic rocks intruded large areas of Palaeoproterozoic cover sequences in Northern Sweden and Northern Finland (Haapala et al. 1987; Ohlander & Ski61d 1994; Perttunen et al. 1996; Bergman et al. 2001; Nironen 2005). The Belomorian Belt and its boundary zones were strongly reactivated in the Palaeoproterozoic (Gafil & Gorbatschev 1987; Bibikova et al. 2001). The area between the Karelian and Kola cratons comprises a mixture of Archaean, reworked Archaean and Palaeoproterozoic terranes. The following Palaeoproterozoic terranes are identified in the Lapland-Kola area (Fig. 2): Inari area (IA), Lapland Granulite Belt (LGB), Umba Granulite Terrane (UGT), and Tersk Terrane (TT) (Korsman et al. 1997; Daly et al. 2001; Koistinen et al. 2001). The Archaean Kola and Karelian cratons have also been affected by Palaeoproterozoic magmatic activity and deformation. These are expressed as calc-alkaline arc magmatism in the Inari area (1.94-1.93 Ga; Barling et al. 1997) and Tersk Terrane (c. 1.96 Ga; Daly et al. 2001) and as strong crustal reworking and metamorphism in the Lapland Granulite Belt, the Umba Granulite Terrane, and the Belomorian Belt (Bibikova et al. 2001). The largest Palaeoproterozoic rift-related belt in the Kola craton, the Imandra Varzuga-Pechenga Belt (IVB and PeB in Fig. 2; e.g. Berthelsen & Marker 1986a), which displays a long evolutionary history from 2.5 to 1.8 Ga (Melezhik & Sturt 1994), is located near or at the Kola-Karelia contact zone. Based on lithological associations, the Svecofennian Domain (Fig. 2a) has been further divided into Northern, Central and Southern Subprovinces (Ga~l & Gorbatschev 1987). At the southern rim of the Northern Svecofennian Subprovince, the oldest lithological units are a few remnants of older volcanic rocks and granites (> 1.95 Ga; Knaften (K)) south of the Skellefte district (SD; Wasstr6m 1993, 1996; Eliasson & Str~ing 1998), and 1.92 Ga tonalites interlayered with volcanites and turbidites in the Savo Belt (SB; Lahtinen 1994; Korsman et al. 1997). The 'primitive' arc complex of Lahtinen (1994), or the Savo oceanic island arc, was later intruded by synkinematic granitoids between 1.89 and 1.88 Ga, and by post-kinematic pyroxene-bearing granitoids starting at 1.885 Ga (H61tt~ et al. 1988; Nironen & Front 1992; Kousa et al. 1994). The Skellefte district (SD) in the northwestemmost part of the Northern Svecofennian Subprovince is composed of two groups of calc-alkaline metavolcanic and metasedimentary rocks intruded by a variety of granites. The older volcanic rocks (1.89-1.88 Ga) were deposited in a marine environment and the younger volcanic rocks (1.88 Ga) in a continental extensional environment. The volcanic sequences were intruded by granitoids at 1.89 Ga, 1.881.86 Ga and 1.80-1.78 Ga, and deformed in three stages, at 1.87, 1.8 and 1.79 Ga (Weihed et al. 1992, 2002; Allen et al. 1996b; Bergman Weihed, unpub, data). In the eastem part of the Central Svecofennian Subprovince, a more continental arc environment is found in the Central Finland Granitoid Complex (CFGC), comprising mainly calc-alkaline I-type granitoids (1.89-1.88Ga) with minor amounts of mafic plutonic rocks as well as remnants of deformed sedimentary and volcanic rocks. Later, the CFGC was intruded by a younger group of hybabyssal rocks as well as post-kinematic granitoids at 1.88-1.87 Ga (Huhma 1986; Elliott et al. 1998; Nironen et al. 2000; Nironen 2003). Based on Sm-Nd (~N~(1.9~ -- 1.6 to +0.6) and geochemical data, an older protolith (c. 2.1-2.0Ga) for the 1.89-1.87 Ga granitoids in the CFGC has been proposed (Lahtinen & Huhma 1997; R~im6 et al. 2001), indicating an older crustal nucleus. At the southern rim of the CFGC, calc-alkaline granitoid rocks and arc-type volcanic rocks (1.90-1.88Ga) are found in the Tampere Belt (TB), which has been interpreted as a mature arc, or to have been formed close to a continental margin. Migmatites with tonalite leucosome (1.89-1.88 Ga) south of the TB have been interpreted as remnants of an accretionary prism (K/ihk6nen 1987; Lahtinen 1994; Korsman et al. 1999).

The Bothnian Basin (BB), in the western part of the Central Svecofennian Subprovince, is composed of psammitic metagreywackes interbedded with black shales and minor mafic volcanic rocks as well as 1.89-1.87 Ga calc-alkaline granitoids. The peak of metamorphism and deformation was associated with the formation of migmatites and granites at 1.82-1.80 Ga (Claesson & Lundqvist 1995; Lundqvist et al. 1998). Inherited Archaean zircons and Sm-Nd (eNd(1.5) --8.5 to --5.7) data from Mesoproterozoic rapakivi granites (Andersson 1997; Andersson et al. 2002) indicate the existence of an older Archaean to Palaeoproterozoic crustal source beneath the BB. The Southern Svecofennian Subprovince includes the 1.901.89 Ga Bergslagen area (BA) and Uusimaa Belt (UB), partly formed in an intra-arc basin of a mature continental arc (e.g. K~ihk6nen et al. 1994; Allen et al. 1996a). Crustal-type Pb-isotopic composition in sulphides and Sm-Nd data (/3Nd(1.9) c. 0) indicate older (>2.0 Ga) crust in the southernmost part of the UB (Lahtinen & Huhma 1997; R~im5 et al. 2001). Similar results have been obtained from the BA (Valbracht et al. 1994). Less evolved island-arc volcanic rocks are found in the H~ime Belt (HB in Fig. 2; K~ihk6nen 2005). The subprovince is transected by a swarm of roughly east-west- to SW-NE-oriented shear zones. Typical lithologies are volcanic rocks with variable tectonic affinities, pelite-dominated sedimentary rocks, quartzites and carbonates. Plutonism shows age groups of 1.89-1.85Ga, 1.84-1.82 Ga and 1.81-1.79 Ga. The S-type late orogenic granites (1.84-1.82Ga) and migmatites with granite leucosome form a belt that extends from southeastern Finland (UB) to central Sweden (BB; e.g. Lundqvist et al. 1998; Korsman et al. 1999). To the south of the Bergslagen area lies the S6rmland Basin (S6B; Fig. 2) composed of several groups of juxtaposed supracrustal rocks. The sedimentary and volcanic rocks were formed in environments ranging from terrestrial to shallow water or marine (Beunk & Page 2001). Further south, volcanic and plutonic rocks, formed in a continental volcanic arc environment, are found in the 1.83 Ga Oskarshamn-J6nk6ping Belt (OJB; Fig. 2; Mansfeld 1996; Mansfeld & Beunk 2004). A possible continuation of the OJB is found in western Lithuania (Mansfeld 2001; Skridlaite & Motuza 2001). Southern and western Sweden is dominated by a c. 1400 km long, north-south-trending batholithic belt, the Transscandinavian Igneous Belt (TIB in Fig. 2; Patchett et al. 1987). Three age groups of volcanic and plutonic rocks (Larson & Berglund 1992; Ah~ill & Larson 2000) have been found: TIB1 (1.811.77 Ga), TIB2 (1.7 Ga), and TIB3 (1.68-1.65 Ga), of which TIB 1 are by far the most voluminous and constitute the southernmost part of the belt. Andersson (1991) and Gorbatschev & Bogdanova (1993) also included the Revsund granitoid intrusions in central Sweden in the TIB. Based on deep borehole samples, Sundblad et al. (1998) have suggested that the TIB continues through the Baltic Sea into the Baltic States. The Fennoscandian Shield becomes younger towards the west, where Gothian evolution took place between 1.75 and 1.55 Ga. Rocks in the westernmost part were reworked during the Sveconorwegian-Grenvillian orogeny at 1.15-0.9 Ga (e.g. Gorbatschev & Bogdanova 1993; Ah~ill & Larson 2000). Distribution of terranes within the Fennoscandian Shield There is growing evidence that Palaeoproterozoic crustal growth older than 1.92 Ga occurred in Fennoscandia. Geochemical and isotopic data (Valbracht et al. 1994; Anderson 1997; Lahtinen & Huhma 1997; R~im6 et al. 2001) as well as the occurrences of 1.95-2.1 Ga ages in detrital zircons (Huhma et al. 1991; Claesson et al. 1993; Lahtinen et al. 2002) indicate that continental nuclei, now seen as crustal domains or suspect terranes, had already started to form at 2.1-2.0 Ga.

SVECOFENNIAN OROGEN

Geophysical markers

It is suggested here that changes in the orientation of the Moho depth isolines (Fig. 3) broadly indicate terrane boundaries within the Fennoscandian Shield. The trends are north-south in the Karelian Province and Southwest Scandinavian Domain (SSD), and east-west in most parts of the Svecofennian Domain, but change to N W - S E south of the Bergslagen area (BA). The large crustal thickness variations within the Svecofennian Domain indicate a heterogeneous block structure (Korja et al. 1993). The thinnest parts have been correlated with the Mesoproterozoic Baltic Sea aborted rift (Korja et al. 2001). The inference that crustal thickness variations are related to suspect terrane geometries is supported by the existence of dipping wide-angle mantle reflections (Fig. 3). These moderately dipping mantle reflections, which are interpreted as frozen subduction zones, serve as criteria for suspect terrane identification within the Svecofennian orogen. Figure 4 shows a 1200 km long crosssection of the Svecofennian orogen along BABEL profiles (Korja & Heikkinen 2005). The seismic data suggest that the crust is composed of crustal terranes that are currently c. 100 km in width. The age of the crustal units decreases from Archaean in the north (Karelia) to late Palaeoproterozoic in the south (TIB). The Mesoproterozoic, extensional rapakivi event was superimposed on the collisional structure (Korja et al. 2001; Korja & Heikkinen 2005). The cross-section displays the end result of plate-tectonic processes that in detail are governed by the stiffer rheology of crustal indentors. Deep conductivity anomalies (Hjelt et al. 2006) are interpreted to represent closed basins between older crustal blocks or indentors. At present, they define major terrane boundaries. In a search for other types of terrane boundaries, lineaments from potential field data have been interpreted. The trends on both Bouguer anomaly and aeromagnetic maps (Korhonen et al. 2002a, b) show regional variation in intensity, anomaly patterns and lineament strikes. The changes are most apparent between the Karelian Province and the Svecofennian Domain, as well as between the Svecofennian Domain and TIB. An abrupt change also takes place at the southern margin of the Southern Svecofennian Subprovince.

Terrane outline

Based on lithological, geochemical, isotopic and geophysical data, it is suggested that there are several Palaeoproterozoic crustal fragments or suspect terranes within the Svecofennian orogen (Tables 1 and 2). These terranes are outlined in Figure 5. The older terranes ( > 1.92 Ga) that took part in the Svecofennian orogeny fall into three categories: Archaean cratonic terranes (Karelia, Kola and Norrbotten), Palaeoproterozoic (>2.0 Ga) microcontinents (Keitele, Bergslagen and Bothnia) and Palaeoproterozoic island arcs (Kittil~i (c. 2.0 Ga), Savo, Knaften, Inari and Tersk (c. 1.95 Ga)). Keitele and Bothnia are hidden and have no identified surface expressions. The approximate extent of the microcontinents at depth is shown in Figure 5a, and the lithological and isotope data defining the crustal terranes are given in Table 2. Later during the Palaeoproterozoic, additional new terranes were formed (e.g. Tampere Belt (TB), H~ime Belt (HB), Uusimaa Belt (UB)) and the abovementioned terranes were modified. The surface extent of the exposed terranes is outlined in Figure 5b. The Ume~ area is distinguished from the Bothnian microcontinent and Skellefte districts (SD) based on seismic reflection data, where the uppermost 15-25 km of crust is interpreted to be detached from the lower and middle crust (Fig. 4; Lahtinen et al. 2005). At the surface, an isotopic boundary delineating older rocks (> 1.9 Ga) in the northern part of the Bothnian Basin has been suggested by Rutland et al. (2001).

565

Lahtinen et al. (2005) correlated the rocks to the south of the Bergslagen area (BA) with those in the Baltic States that have similar NW-SE-striking geophysical patterns and lithological similarities, and called them the Svecobaltic area.

Model Lahtinen et al. (2005) suggested a plate-tectonic model accommodating the geophysical, geochemical and geological observations. In the model, it was assumed that the current extent of terranes was not significantly affected by collisions, implying that no account was taken of transportation along strike-slip faults, escape tectonics, or major internal stacking of the microcontinents. In the following, the model is elaborated with emphasis on processes forming the continental crust during each orogeny at a given time. The model distinguishes five stages: (1) continental rifting; (2) microcontinental accretion; (3) extension of the accreted crust; (4) continent-continent collision; (5) extensional collapse. C o n t i n e n t a l rifting

At the beginning of the Palaeoproterozoic, the Archaean Karelian continental crust was subject to large-scale extension leading to emplacement of mafic layered intrusions, granitoid intrusions and bimodal volcanism at around 2.5 Ga. This was followed by several generations of mafic dyke swarms between 2.2 and 1.97 Ga. Formation of oceanic crust in marginal basins is indicated by the 1.97-1.95 Ga Jormua and Outokumpu ophiolites. Juvenile oceanic island arc lithologies in the Savo Arc, at the ArchaeanProterozoic boundary zone, indicate opening of an ocean. To explain the abundant oceanic island arc lithologies in the Kittil~i allochthon (KA), a major rift is also suggested to have developed between the Karelia and Norrbotten cratons. An ocean between Karelia and Kola is proposed by island arc type lithologies in the Tersk Terrane and Inari Area. M i c r o c o n t i n e n t a l accretion

A microcontinental accretion stage is suggested to explain the complex assemblage history between 1.92 and 1.88 Ga in Fennoscandia (Fig. 6). After the rifting of the Archean continental plate at 2.1 Ga, new juvenile arcs were initiated. Arcs of different evolutionary stages started to assemble and to form microcontinents. By 1.89 Ga, several small plates carrying continental fragments had accreted to the Karelian continent from several directions. Oblique collision caused some of the shear zones at or close to the terrane boundaries. The accretionary processes and the dimensions of the accreting blocks are schematically illustrated by vertical cross-sections (see Fig. 8). The amalgamation of Fennoscandia started from the NE, where the Kola and Karelian cratons as well as intervening Palaeoproterozoic terranes (IA, TT, LGB, UGT) merged together during the Lapland-Kola orogeny (Fig. 6a). Simultaneously, at the western margin of the Karelian craton, the Lapland-Savo orogeny commenced as the result of the approaching Norrbotten craton and the Keitele microcontinent. Juvenile crust that had formed in island arcs (KA, SB, K) close to the continental margins was accreted before the final collision of the microcontinental nuclei. The main phase of the Lapland-Savo orogeny took place when the Keitele microcontinent collided with the Karelian craton (Fig. 6b). The westward growth of the continental collage continued with the docking of the Bothnia microcontinent (Figs 6c and 8a). The collision caused a change in the plate motions and led to a subduction switchover, with the onset of northward subduction at the southern edge of the Keitele-Karelia collage (TB; Figs 6c and 8b). The northward subduction ended when the southern ocean was consumed; the Bergslagen microcontinent was accreted to Keitele, starting the Fennian orogeny.

566

A. KORJA E T AL.

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SVECOFENNIAN

OROGEN

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A. KORJA ETAL. Table 1. Crustal units of the Fennoscandian Shield older than 1.92 Ga Dominant rock types

Age (Ga)

Tectonic environment

Granodiorite, tonalite, tholeiite, gabbro, quartzite Metavolcanic and metasedimentaryrocks Mafica and felsic volcanic rocks, ophiolite Tonalite, dacite, basalt Basic-intermediate volcanic rocks, metasediments, granitoids

1.95-1.93 1.96 2.02-2.01 1.93-1.92 1.95-1.94

Evolvedarc/active continental margin Evolved arc Oceanicisland arc Oceanicisland arc Islandarc

References*

Arc terrane

Inari Tersk Kittil~i Savo Knaften

1 4 8, 16 9, 10 19, 23

Hidden terranes Crustal type

Bothnia Keitele Bergslagen

Archean lower crust/mature crust? Mature crust Mature crust

2.7/2.1 ? 2. 1-2.0 2.1-2.0

28 27 34

*List of references is given as legend to Table 2.

The Fennian orogeny resulted in considerable shortening of the accretionary prism and volcanic belts (TB, HB, UB) between the Bergslagen and Keitele microcontinents, thrusting at the western margin of the Karelian craton, and basin inversion in Lapland. In the west, the shortening was delayed and less severe between the Bothnia and Bergslagen microcontinents. E x t e n s i o n o f the a c c r e t e d crust

At the end of the microcontinent accretion stage, a larger but unstable continental plate (Fennoscandia) had developed. The newly formed continental interior was thermally and gravitationally unstable, and as soon as active compression shifted towards the margins, extension took over in the interior (Figs 7a and 8c). Minor gravitational collapse around the Keitele microcontinent occurred at 1.881.87 Ga, expressed as post-kinematic granitoid intrusions and mafic dykes (Nironen 2005), but coeval north-south compression at the southern margin inhibited large-scale extension. Overall, gravitational collapse was taking place in the centre of the continent while subduction was commencing at its margins. A more widespread collapse was possible only at 1.86-1.85 Ga, when the subduction systems had retreated further to the south and west leaving the central part in an extensional back-arc setting (Figs 7a and 8c). This extension caused the formation of large-scale extensional basins in the hinterland, within the newly formed continent. The basins were filled with psammites and pelites (Lahtinen et al. 2002) and intruded by mafic dykes and granites (Sfftlhts 1976; Suominen 1991). C o n t i n e n t - c o n t i n e n t collision stage: formation of a supercontinent ?

Lithospheric convergence brought Fennoscandia in contact with Kola craton (Laurentia) in the NE, Amazonia in the west, Sarmatia in the SE, and an unknown continent in the SW (Fig. 7b). The net result was that at the peripheral margin of Laurentia, the Fennoscandian nucleus was surrounded by several continents. The continent-continent collision stage began with the final docking of Fennoscandia to Laurentia (Kola craton), which resulted in the uplift of the Lapland Granulite Belt and the reactivation of the Belomorian Belt between 1.84 and 1.81 Ga. This was followed by the docking of Sarmatia. Oblique collision in the east started the Svecobaltic orogeny (1.84-1.80 Ga), whereas along the southwestern margin, an Andean-type subduction regime still prevailed. Transpressional tectonics, resulting from oblique collision, can be observed as large-scale thrusting of the foreland fold-and-thrust belts and margin-parallel (east-west) shear zones in central Sweden and southern Finland (Fig. 2; St~lhts 1976; Ehlers et al. 1993). It is suggested that, during the Svecobaltic

orogeny, the crust was reworked as the extended Fennian sequences were stacked back to their original positions. On the lithological map (Fig. 2), these areas are distinguished by 'late orogenic' S-type granites and migmatites in the Bothnian Basin, Bergslagen area and Uusimaa and H~ime Belts. Thickening of the preheated crust led to local migmatization and formation of minimum-melt S-type granites from sediments. In BABEL seismic sections (Fig. 4), thrusting is shown as stacking of folded sequences and crustal wedges. The third continent-continent collision was between Amazonia and Fennoscandia at 1.82-1.80 Ga. This produced the Nordic orogeny, which affected the central and northern parts of the western edge of Fennoscandia. The core of the now exposed part of the Nordic orogen is the northern part of the TIB (Fig. 2). The TIB1 (1.81-1.77Ga) intrusions (Larson & Berglund 1992) represent magmatism in a convergent margin environment (e.g. Wilson 1982; Gorbatschev & Bogdanova 1993; Andersson 1997; ~gdaS.11 & Larson 2000). It is suggested that the TIB1 rocks are related to the waning phase of subduction, continent-continent collision and the following orogenic collapse (Fig. 7). The N W - S E to east-west collision also produced the second major deformation phase described from northern Sweden and northern Finland. The deformation caused north-trending, open to tight folds and north-south compressional deformation zones. Yotmger metamorphism and associated S-type granites at 1.82-1.80 Ga are correlated with this stage (Claesson & Lundqvist 1995; Allen et al. 1996b; Billstr6m & Weihed 1996; Perttunen et al. 1996; Lehtonen et al. 1998; Bergman et al. 2001). To explain the different characteristics of upper and lower crust in BABEL profiles 3 and 4 (Fig. 4) and 2 (Korja & Heikkinen 2005), we propose a model of basin inversion and thick-skinned, large-scale thrusting of a migmatite complex to the ESE at 1.83-1.81 Ga. This thrust slice, named the Umeh allochthon (Fig. 5), coincides with the pre-l.9 Ga sequence outlined by Rutland et al. (2001) in the Bothnian Basin, and it probably extends to western Finland. The large-scale thrusting could explain the tectonic discordance between the Skellefte district (SD) and the migmatite area south of it (BB in Fig. 2; Rutland et al. 2001). The bottom of the subhorizontal conductive layer (Rasmussen et al. 1987; Hjelt et al. (2006, Fig. 3) could also be interpreted as a basal thrust plane. At the southern margin of the collage, the Svecobaltic orogeny continued with the docking of a yet unknown continent (Fig. 7c); Abramovitz et al. (1997) called this continent an 'intermediate terrane'. Partly gneissic granitoid rocks that trend W N W - E S E to west-east in the southeastern corner of Sweden, recently dated to 1.81-1.77 Ga (Gorbatschev 2001), may be related to this continent. This continent filled a gap between Amazonia and Sarmatia and caused the closure of the S6rmland basin, the accretion of the Oskarshamn-J6nk6ping Belt (Fig. 8d), and the

SVECOFENNIAN OROGEN

569

Table 2. Exposed Palaeoproterozoic terranes in the F e n n o s c a n d i a n Shield Terrane

Inari area, IA Lapland granulite belt, LGB

Dominant rock types

Tectonic environment

Age of magmatism (Ga)

Granodiorite, tonalite, tholeiite, gabbro, quartzite Felsic granulites and enderbites

1.95-1.93 1.93

Evolved arc/active continental margin Back-arc/active continental margin

Age of deformation (Ga) < 1.93 1.91-1.90 1.88-1.87

Umba granulite terrane, UGT Tersk terrane, TT Kittil~ allochthon, KA

Felsic granulites and enderbites Metavolcanic and metasedimentary rocks Mafic and felsic volcanic rocks

<1.91 2.02-2.01

Evolved arc Oceanic island arc

1.89-1.87

Continental island arc/back-arc

1.96

1.92

References

1 2, 3, 4, 5, 6, 7 4

36 4 8, 16

Northern Svecofennian

Skellefte district, SD

Knaften, K

Savo belt, SB Ume5 allochthon, UA

Andesites Granites and migmatites Granitoids (TIB) Basic-intermediate volcanic rocks, metasediments, granitoids Granitoids (TIB) Tonalite, dacite, basalt Tonalite, granodiorite, px-granitoids Migmatites (turbidites) Granites and migmatites

1.80-1.79

1.80-1.79 1.79-1.76 1.95-1.94

Island arc

1.80-1.77 1.93-1.92

Oceanic island arc

1.89-1.88 > 1.90 1.84-1.82

Accretion prism?

1.89-1.87

Accretion prism

1.85-1.82

14,35 33 46 19, 23 14, 15 9,10 18 13,38 42

Central Svecofennian

Bothnian basin, BB

Central Finland granitoid complex, GFGC

Turbidites, basalts Granites, migmatites Granitoids (TIB) Granodiorites and granites; andesites and dacites

1.84- 1.82

1.84-1.82

1.80-1.77 1.90-1.88

Active continental margin

12, 13,15, 29 14, 15 14 17

Southern Svecofennian

Tampere belt, TB

Granites; andesites and dacites; turbidites

1.90-1.88

Active continental margin

H~ime belt, HB

1.89-1.88

Uusimaa belt, UB

Andesites, basalts, pelites, granodiorite, tonalite Andesites, dacites, rhyolites; quartzites

Continental island arc and rift basin Rifted active continental margin

Bergslagen area, BA

Granodiorite, tonalite, Granites, migmatites Andesites, dacites, rhyolites; quartzites

1.89-1.85 1.84-1.82 1.90-1.87

Granites, migmatites

1.84-1.82

1.90-1.88

< 1.89

10, 11,17, 21 39 22, 28, 37 27, 43

1.84-1.82

21 44 20, 41

1.84-1.82

45

Rifted active continental margin

Svecobaltic

S6dermanland belt, SOB

OskarshamnJ6nk6ping belt, OJB TIB

Gothian

Quartzite, mafic and felsic volcanic rocks, turbidites Granitoids and gabbros, migmatites Granitoids Mafic-felsic plutonic and volcanic rocks, arkose Granitoids (TIB) Granites, gabbros and felsic volcanic rocks Granites, gabbros and felsic volcanic rocks Granites, gabbros and felsic volcanic rocks Granites, gabbros and felsic volcanic rocks

> 1.88 1.85-1.81 < 1.81 1.83

Terrestial shallow water Back-arc environment Active continental margin

1.80-1.77

1.80-1.76

29, 30

Active continental margin

1.71-1.69 1.67-1.65 1.73-1.55

< 1.83

30, 31 30 25, 30, 40 25 32 32 32 24

*The table is based on the following references: 1, Barling et al. 1997; 2, Meril~iinen 1976; 3, Bernard-Griffiths et al. 1984; 4, Daly et al. 2001; 5, Barbey et al. 1982; 6, Sorjonen-Ward et al. 1994; 7, Tuisku & Huhma 1998; 8, Hanski & Huhma 2005; 9, Lahtinen 1994; 10, Kousa et al. 1994; 11, KSlak6nen & Nironen 1994; 12, Lundqvist et al. 1998; 13, Rutland et al. 2001; 14, Allen et al. 1996b; 15, Claesson & Lundqvist 1995; 16, Lehtonen et al. 1998; 17, Lahtinen et al. 2002; 18, Korsman et al. 1997; 19, Wasstr~Sm 1993; 20, Alien et al. 1996a; 21, KfihkSnen et al. 1994; 22, KShk6nen 2005; 23, Wasstr6m 1996; 24, ,~h~ill & Larson 2000; 25, Mansfeld 1996; 26, Dobbe et al. 1995; 27, Lahtinen & Huhma 1997; 28, Andersson et al. 2002; 29, Claesson et al. 1993; 30, Beunk & Page 2001' 31, Andersson & Wikstr6m 2001" 32, Larson & Berglund 1992; 33, Bergman Weihed (unpubl. data); 34, Valbracht et al. 1994; 35, Billstr6m & Weihed 1996; 36, Glebovitsky et al. 2001' 37, Hakkarainen 1994; 38, Mellqvist et al. 1999; 39, Vaasjoki & Huhma 1999; 40, Mansfeld & Beunk 2004; 41, Beunk & Valbracht 1991; 42, Weihed et al. 1992; 43, V~iis~inen& M~intt/iri 2002; 44, Huhma 1986; 45, St~h6s 1991; 46, 0hlander & SkiiSld 1994.

f o r m a t i o n of m a r g i n - p a r a l l e l shear zones, such as the H S Z . TIB 1 granitoid intrusions are related to the w a n i n g p h a s e o f this collision and f o l l o w i n g o r o g e n i c collapse; thus the TIB is t h o u g h t to result f r o m two orogenies: S v e c o b a l t i c in the south and N o r d i c in the north. After the final d o c k i n g o f the ' u n k n o w n c o n t i n e n t ' , Fennosc a n d i a was p o s i t i o n e d in the m i d d l e o f a supercontinent.

Gravitational

collapse

stage

After the c o n t i n e n t - c o n t i n e n t collisions, F e n n o s c a n d i a u n d e r w e n t a m a j o r stabilization period, w i t h gravitational collapse, t h e r m a l resetting and late t e c t o n o - m a g m a t i c episodes (Figs 7c and 8e). This p e r i o d is c h a r a c t e r i z e d by v o l u m i n o u s granitoid m a g m a t i s m

570

A. KORJA ETAL.

~raton

m

orweg

(a)

(b)

Go,high ~

~176

~

and pegmatite intrusions around the Svecofennian nucleus (Larsson & Berglund 1992; Romer & Smeds 1994, 1997; Eklund et al. 1998; Alviola et al. 2001; Bibikova et al. 2001). The gravitational collapse following the docking of Fennoscandia to Laurentia reset U - P b ages and caused the intrusion of pegmatites in the Belomorian Belt (Bibikova et al. 2001; Corfu & Evins 2002). In the northern parts of the Nordic orogen, the 1.80-1.77 Ga granite-syenite-gabbro association, together with granitepegmatite association or migmatites, marks orogenic collapse. Perttunen et al. (1996) have suggested that the late metamorphism and migmatization might have occurred during decompression. Further south, some of TIB1 granitoids (including Revsund granites, R), intruded in an extensional setting in an intracratonic regime at 1.80-1.78 Ga (e.g. Anderson 1997; ~X,hSll & Larson 2000), are here considered to represent the collapse stage after collision of the 'unknown continent' and Amazonia, referred to above. Rapid uplift of the Svecofennian lithologies has been dated between 1.81-1.79 Ga in metamorphic studies (Korsman et al. 1999; V~iis/inen et al. 2000). 'Post-collisional' mantlederived granitoids intruded during this time period (Eklund et al. 1998). The youngest zircons (1.8 Ga) from mantle and lower crustal xenoliths in eastern Finland indicate mantle activity up to this time (H61tt~ et al. 2000; Peltonen & M/intt/iri 2001). All these features are interpreted to mark gravitational collapse that took place between 1.80 and 1.75 Ga. It was possible only after all the longterm compression at the margins of Fennoscandia had ceased. The general collapse was a combination of several smaller episodes and perhaps lithospheric delamination of Amazonia after the Nordic orogeny.

Discussion

Plate boundary forces create variable tectonic environments that overlap in time and space. In size and complexity, Precambrian Fennoscandia is comparable with Phanerozoic Europe and to the present Indonesian archipelago, where many processes operate simultaneously and where the change from one tectonic setting to another is rather abrupt (Lee & Lawyer 1995; Stampfli & Borel 2002). The model presented here tries to visualize the complexity of overlapping plate-tectonic environments that could have prevailed at the same time in the geographically restricted area of Fennoscandia in the Palaeoproterozoic. In this model, the Svecofennian orogeny initiated when the passive Karelian margin developed into a convergent margin. Island arcs began to form at the periphery and suspect terranes, here called microcontinents, were attached to them, first causing the Lapland-Savo accretionary orogeny. After this collision, the plate movements readjusted and suspect terranes arrived from the south, causing the Fennian accretionary orogeny.

s.... ~ 7 ~

:

5ookm :

Fig. 5. Distributionof microcontinental nuclei, island arcs and terrane boundaries in the FennoscandianShield. Abbreviations are as in Figure 2. (a) Older than 1.92 Ga, hidden and exposed suspect terranes found in the SvecofennianOrogen. (b) Major Palaeoproterozoic terranes of the Fennoscandian Shield.

At convergent margins, oceans separating major continents are consumed and new crust is formed in island arcs and accretionary wedges. The accretionary orogenies, during which this new material is attached to the initial continents, can be regarded as preludes to continental collision, in which the previously accreted material is merely reworked. In this respect, the Lapland-Savo and Fennian orogenies are preludes to the Svecobaltic and Nordic orogenies. The Lapland-Savo and Fennian orogenies can be compared with, for example, the formation of the Tethysides, which were later involved in continent-continent collision between India and Eurasia, producing the Alpine-Himalayan orogeny (Stampfli & Borel 2002; von Raumer et al. 2003; Golonka 2004), or with the formation of the Neoproterozoic Avalonia-Cadomian orogeny (Nance et al. 2002; Gutierrez-Alonso et al. 2003). The evolution of the Svecobaltic orogeny is rather complex. Based on geophysical data a transform fault is placed between the eastern and western parts of the area. The transform implies, however, that either another piece of Sarmatia or another unknown continent arrived from the south along the transform fault. It also implies that both the approaching continents collided obliquely with Fennoscandia, inducing margin-parallel strike-slip faults. This could explain the wealth of strike-slip faults in southern Finland and central Sweden. The initiation of this transform is, however, not clear. Skridlaite & Motuza (2001) have suggested that this plate boundary is an accretionary one. Bogdanova et al. (2006) and Claesson et al. (2006) have further discussed the accretionary processes at the Sarmatian and Volgo-Uralian margins. In the model preferred here, the Karelian, Kola and Belomorian Provinces, as well as the Svecofennian Domain and Transscandinavian Granite-Porphyry Belt (TIB) as defined by Gafil & Gorbatschev (1987), were all active during the Svecofennian orogeny. We have split the Svecofennian orogeny into five partly overlapping subordinate orogenies: Lapland-Savo, Fennian, Lapland-Kola, Svecobaltic and Nordic orogenies. The Northern and Central Svecofennian Subprovinces of Gafil & Gorbatschev (1987) were formed mainly in the Lapland-Savo orogeny. The northern Svecofennian comprises more juvenile island arc material of the Knaften, Savo and Kittil~i arcs whereas the Central Svecofennian is cored by older continental crustal blocks: the Keitele and Bothnia microcontinents. The Southern Svecofennian Subprovince was initiated in the Fennian orogeny, during which the Bergslagen microcontinent was accreting to the newly formed Lapland-Savo orogenic belt. The Southern Svecofennian terrane was, however, heavily reworked by Svecobaltic orogeny, induced by the collision of Sarmatia with Fennoscandia. TIB was mainly formed in the Nordic orogeny. Each of the collisional events was immediately followed by gravitational collapse. One of the peculiarities of the Fennoscandian Shield is the thick lithosphere (>200 km), with thick crust and mantle (Fig. 3; Korja

SVECOFENNIAN OROGEN

/,~//P~.

571

Kola i .................... 1.....................

] I

Mostly Archean crust Archean and Palaeoproterozoic crust 2.1-2.0 Ga crust 2.0-1.95 Ga crust

5OO k m

Active plate boundaries (a) 1.93 Ga Active terrane boundaries i.........;;;;...........J

Terrane boundary Direction of relative plate motion

i==t> }i Direction of compression '..................

Direction of extension (b) 1.91 Ga

(c) 1.8! Fig. 6, A schematic plate-tectonic model for the Fennoscandian Shield modified from Lahtinen et al. (2005), showing the microcontinent (mc) accretion stage at 1.93-1.88 Ga. Schematic cross-sections of the plate-tectonic processes along lines a - a ' and b - b ' are shown in Figure 8. Abbreviations are as in Figure 2. (a) Subduction (IA, TT) and back-arc rifling (LGB, UGT) in the Lapland-Kola area, westwards subduction under the Keitele mc (Savo arc, Savo Belt) and Norrbotten mc (Kittilii arc, Kittil~i allochthon), and subduction to the NE under the Norrbotten mc (Knaften arc) and east under the Keitele mc. (b) Peak of the Lapland-Kola and Lapland-Savo orogenies. Initial stage of collision of the Bothnian mc with the Norrbotten and Keitele mc. Initiation of Northern and Central Svecofennian Subprovinces. (e) Beginning of the Fennian orogeny. As a result of the locking of northward subduction under the Keitele mc (Tanpere Belt), the southern ocean is consumed by subduction to the south under the combined H~ime island arc (Harne Belt) and Bergslagen mc (Bergslagen Area, Uusimaa Belt). After amalgamation of the Bothnian mc, another subduction zone is formed at its southern margin. Initiation of the Southern Svecofennian Subprovince.

572

A. KORJA E T A L .

['-

Mostly Archean crust Mostly Palaeoproterozoic crust Terrane b o u n d a r y

Active plate boundaries . . . . . 'j Active t e r r a n e b o u n d a r i e s

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

[--~-~i

Direction of c o m p r e s s i o n

i ~':~> ! Direction of e x t e n s i o n

~a) 1 .tSb L~a -

(b) 1.80 Ga

(e) 1.78 Ga Fig. 7. A schematic plate-tectonic model for the Fennoscandian Shield modified from Lahtinen et al. (2005), showing the continent-continent collision stage at 1.87-1.79 Ga. The Fennoscandian continental plate, formed in the accretionary stage (Fig. 6), has been divided into Archaean and Palaeoproterozoic parts. Cross-sections of lines c - c ' d - d t and e - e ' are shown in Figure 8. (a) Subduction to the SE and to the NE initiates at the southern margin and large-scale extension takes place in the hinterland. (b) The amalgamation of Laurentia, Fennoscandia, Amazonia, Sarmatia and an unknown continent in the SW ends at 1.81-1.79 Ga, and a Palaeoproterozoic supercontinent is formed. (c) Orogenic collapse and lithospheric delarnination stabilizes the Fennoscandian Shield between 1.79 and 1.77 Ga.

SVECOFENNIAN OROGEN

a-a'

573

1.89Ga Knaften arc

Bothnian mc

b-b'

1.89 Ga Bergslagen mc UB

TB

HB

Keitete SB ~,~

c-c"

1.86-1.85 Ga

Bergslagen

BA

mc

UB HB

TB

mc ~

~.~

.......

Kareliancraton ........

Keitele SB ~

mc ~

~-~,

Karelian craton

~ . . . . . . .

0

d-d" 1.81-1.80 Ga Unknown mc

e-e"

OJB

BA Bergslagen mc

S~SB

1.78Ga

Amazonia

TIB

0

t

m

"

|

Ume~ allochthon

Bothnia mc

Ume~ allochthon

Knatten ac

Keitele mc SB

Karelian craton

O

O

50 Vertical ~ km

HSZ

'

~ I

I

Mainly Archean crust

Sedimentary rocks

pre-2.0 Ga crust

Oceanic crust

2.0-1.95 Ga crust

Remnants of oceanic crust

Island arc

Direction of plate movement

Back-arc basin

Oceanic lithosphere

Ume& allochthon

Continental lithosphere

Magmatic underplate

Scar in lithosphere

Mafic magmatism

Melting of lithospheric mantle

Felsic to intermediate magmatism

Fig. 8. Vertical cross-sections of the tectonic processes taking place at lines a - a r and b - b r in Figure 6 and c - c r, d - d r and e - e I in Figure 7. Vertical exaggeration 1:2. The lithospheric thickness is not to scale and it is underestimated for drafting purposes. Abbreviations are as in Figure 2. a - a r, Microcontinent (mc) collision stage. During the Lapland-Savo orogeny, the Norrbotten craton, Knaften arc and Bothnian microcontinent were accreted to the Karelian craton, while a subduction system was developing at the southern margin of the newly accreted terrane, b - b r, Microcontinent collision stage II. The docking of the Keitele mc with the Karelia craton led to locking of northward subduction under the Keitele mc (TB). This led to a subduction switchover and to the onset of southward subduction beneath the approaching Bergslagen mc (UB, HB). c-c', Extension of the accreted crust. Collapse of the accreted orogen in an extensional back-arc setting. Large-scale extensional basins form in the hinterland, especially within the newly formed Fennian orogen. The basins are filled with immature sediments and bimodal magmatism, d - d r, Continentcontinent collision. Following subduction, a continent collision took place at the southwestem margin of the accreted Fennoscandia. In this final collision the Strmland basin (StB) was closed and the Oskarshamn-Jtnktping Belt (OJB) was accreted to the unknown continent, e - e r, Orogenic collapse and lithospheric delamination. After continent-continent collision, the lithosphere was overthickened and unstable, and gravitational collapse took over. Both crust and lithosphere were thinned. The lower part of the Amazonian plate may have delaminated after the Nordic orogeny.

574

A. KORJA ETAL.

et al. 1993; Peltonen & MS_ntt/iri 2001; Hjelt et al. 2006). The

Svecofennian orogen seems to have been accreted from Palaeoproterozoic terranes and intervening island arcs and basins. The plate sizes and crustal thickness were probably comparable with those found in the Indonesian and Mediterranean areas today. Because the size of the colliding microcontinents was small, the lithosphere could not have been extremely thick, as the thickness of the continental lithosphere is dependent on the width of the plate. The collisions allowed both the growth of the plates and the increase in the lithospheric thickness in Fennoscandia. The reason for the preservation of the thick lithosphere is controversial. Because the crustal roots are not topographically compensated, they must be isostatically compensated within the crust. Around 10 km of exhumation has been proposed to have taken place by 1.8 Ga in southern Finland (Eklund et al. 1998; V/iis~inen et aL 2000), indicating rapid compensation after the Svecofennian orogeny. The preservation of the thick crust was explained by Korja et al. (1993) and Korsman et al. (1999) by thick ( 1 0 - 3 0 k m ) high-velocity (Vp > 7 . 0 k m s - 1 ) , highdensity mafic lower crust that more or less compensates the crustal thickness variations and accounts for the flat topography. Early arrival times of earthquake waves and recent results of the SVEKALAPKO tomographic array (Bock & SVEKALAPKO Seismic Tomography Working Group 2001; Hjelt et aL 2006) indicate high velocities of the mantle lithosphere. High velocities indicate high densities and low temperatures that could explain the quasi-isostatic equilibrium of the lithospheric keel. It is inferred that the microcontinental collision stage resulted in thickened lithosphere in the core area. Collapse took place at 1.87 Ga, when the post-kinematic granitoids (Nironen et al. 2000) and undeformed mafic dykes (Marmo 1963; Aro 1987) intruded in central Finland. The collapse spread towards the orogenic fronts where its effect was more profound. In southern Finland, the orogenic collapse area suffered from simultaneous back-arc extension. The increased heat flow may have resulted in mafic underplating and/or upwelling of the lower crust (Dewey 1988; Vanderhaeghe & Teyssier 2001). This may have initiated partial melting and migmatite formation. Traces of a hot extensional basin are found as 'intra-orogenic' mafic dykes (Stfilh6s 1976; Ehlers et al. 1993), tonalite dykes and granitoid rocks (1864 Ma; Suominen 1991) as well as sporadic 1.86 Ga quartzites (Lahtinen et al. 2002). During the collisional Svecobaltic orogeny, the extensional basin was inverted, folded and thrust onto the continent. The collapse stage of the Nordic orogeny is seen as voluminous magmatism in central Sweden and Lapland. Two types of magmatism are found: lithospheric mantle-derived gabbros to monzodiorites, and crust-derived monzogranites. The mantle-derived magmatism could have been the heat source for the crustal magmas. The NNE-SSW-trending belt is associated with a linear belt of positive magnetic and Bouguer anomalies. A similar belt of magnetic anomalies that trends N N W - S S E (reworked during the Sveconorwegian orogeny) is found in southern Sweden. This belt is associated with mantle-derived TIB1 intrusions (1.8-1.75Ga) (Andersson 1997). Reflection data (Fig. 4) suggest the presence of mantle-derived intrusions beneath the TIB. These intrusions may be associated with the collapse of the Svecobaltic orogeny. Few occurrences of c. 1.8 Ga mantle-derived post-kinematic intrusions are found at the periphery of the Svecobaltic orogen in Southern Finland and Russia. These features and mantle and lower crustal xenoliths, recording a magmatic event at around 1.8 Ga (H61tt~i et al. 2000; Peltonen & M~inttSxi 2001) along the Archaean-Proterozoic contact zone, suggest mantle reactivation in the collapse processes. The thick lithospheric keel, combined with the small volume and lithospheric nature of the mantle magmatism, does not favour asthenospheric upwelling in the central parts of the Fennoscandian Shield. The volume of the mantle magmatism increases to the

west (TIB), indicating that mantle delamination may have taken place in this area. If so, it probably belonged to the subducted Amazonian plate. In the areas where gravitational collapse was enhanced by lithospheric delamination, the crustal thickness attained its 'normal' thickness of 40 km, whereas thick crustal roots (50-60 km) were preserved in the previously stabilized areas.

Conclusions Based on an integrated study of geological and geophysical data, it is concluded that the Svecofennian orogen was formed during five partly overlapping orogenies: Lapland-Savo, Lapland-Kola, Fennian, Nordic and Svecobaltic. The Svecofennian orogen evolved in four major stages, involving microcontinent accretion (1.92-1.88 Ga), large-scale extension of the accreted crust (1.87-1.84Ga), continent-continent collision (1.87-1.79Ga) and, finally, gravitational collapse (1.79 and 1.77Ga). The stages partly overlapped in time and space, as different processes operated simultaneously in different parts of the plates. It is concluded that, even in Palaeoproterozoic Fennoscandia, the orogenic belts have been formed in short and distinct orogenies, each involving alternating collisional and extensional stages. The complexity arose from interacting processes, small plates and different tectonic environments existing on opposing margins of a plate at a given time. Orogenies evolved from accretionary to collisional stages in the Palaeoproterozoic, as they do today. F. F. Beunk and R. W. R. Rutland are thanked for critical reviews that greatly improved an early draft of the manuscript.

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SVECOFENNIAN OROGEN

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VAASJOKI, M. & HUHMA,H. 1999. Lead and neodymium isotopic results from metabasalts of the Havefi Formation, southern Finland: evidence for Palaeoproterozoic enriched mantle. Bulletin of the Geological Society of Finland, 71, 143-153. VAISANEN, M. & MANTTARI, I. 2002. 1.90-1.88 Ga arc and back-arc basin in the Orij~irvi area, SW Finland. Bulletin of the Geological Society of Finland, 74, 185-214. V~,ISANEN, M., MANTTfi~RI, I., KRIEGSMAN, L. M. & HOLTTA, P. 2000. Tectonic setting of post-collisional magmatism in the Palaeoproterozoic Svecofennian Orogen, SW Finland. Lithos, 54, 63-81. VALBRACHT, P. J., OEN, I. S. & BEUNK, F. F. 1994. S m - N d isotope systematics of 1.9-1.8-Ga granites from western Bergslagen, Sweden: inferences on a 2.1-2.0-Ga crustal precursor. Chemical Geology, 112, 21-37. VANDERHAEGHE, O. & TEYSSIER, C. 2001. Crustal-scale rheological transitions during late-orogenic collapse. Tectonophysics, 335, 211-228. VON RAUMER, J. F., STAMPFLI, G. M. 8,= BussY, F. 2003. Gondwanaderived microcontinents--the constituents of the Variscan and Alpine collisional orogens. Tectonophysics, 365, 7-22.

WASSTROM, A. 1993. The Knaften granitoids of Viisterbotten County, northern Sweden. In: LUNDQVIST, T. (ed.) Radiometric dating results. Geological Survey of Sweden, Reports, C823, 60-64. WASSTROM, A. 1996. U - P b zircon dating of a quartz-feldspar porphyritic dyke in the Knaflen area, V~isterbotten County, northern Sweden. In: LUNDQVIST, T. (ed.) Radiometric dating results 2. Geological Survey of Sweden, Reports, C828, 34-40. WEIHED, P., BERGMAN,J. • BERGSTROM,V. 1992. Metallogeny and tectonic evolution of the Early Proterozoic Skellefte district, northern Sweden. Precambrian Research, 58, 143-167. WEIHED, P., BILLSTROM, K., PERSSON, P.-O. 8r BERGMAN WEIHED, J. 2002. Relationship between 1.90-1.85 Ga accretionary processes and 1.82-1.80Ga oblique subduction at the Karelian craton margin, Fennoscandian Shield. Geologiska FOreningens i Stockholm FOrhandlongar, 124, 163-180. WILSON, M. R. 1982. Magma types and the tectonic evolution of the Swedish Proterozoic. Geologische Rundschau, 71, 120-129. WINDLEY, B. F. 1993. Proterozoic anorogenic magmatism and its orogenic connections. Journal of the Geological Society, London, 150, 39-50.

The Lapland-Kola orogen: Palaeoproterozoic collision and accretion of the northern Fennoscandian lithosphere J. STEPHEN D A L Y 1, V I C T O R V. B A L A G A N S K Y 2, M A R T I N J. T I M M E R M A N 3 & M A R T I N J. W H I T E H O U S E 4

1School of Geological Sciences, University College Dublin, Belfield, Dublin 4, Ireland (e-mail: stephen.daly @ ucd.ie) 2Geological Institute, Kola Science Centre, Russian Academy of Sciences, 14 Fersman St., Apatity, 184209, Russia 3Institut fiir Geowissenschaften, Universitiit Potsdam, Postfach 60 15 53, 14415 Potsdam, Germany 4Swedish Museum of Natural History, Box 50007, 10405 Stockholm, Sweden

Abstract: A tectonic model is proposed for the Palaeoproterozoic Lapland-Kola orogen (LKO) in the northern Fennoscandian Shield. Although long regarded as an Archaean craton, integrated geological, geochemical and geophysical observations show that the Lapland-Kola orogen is a Palaeoproterozoic collisional belt containing both Archaean terranes and an important component of juvenile Palaeoproterozoic crust. Rifting, from 2.5 to 2.1 Ga, began under the influence of a mantle plume (> 1000 km diameter), related to the break-up of the Kenorland supercontinent. Two linear suture zones within the orogenic core mark the sites of continental separation, ocean formation and closure. One of these is identified as a belt of 1.98-1.91 Ga juvenile crust of both arc magmatic and sedimentary origin, marked by the Lapland Granulite, Umba and Tersk terranes. Palaeomagnetic data and ancient sedimentary detritus within these terranes suggest limited oceanic separation. Collision of juvenile terranes with the surrounding Archaean took place mainly between 1.93 and 1.91 Ga, resulting in a Himalayan-scale mountain belt, manifest by a thick-skinned region of high-P granulite-facies metamorphism, including the classical Lapland Granulite Belt and a broad zone of compressional deformation extending southwards into the Belomorian Mobile Belt. Protracted cooling and exhumation, possibly related to the buttressing effect of surrounding lithosphere, culminated in the intrusion of 1.80-1.77 Ga post-tectonic granites.

The Fennoscandian (Baltic) Shield is one of the best known of the Earth's Precambrian regions. It is well exposed and has been the focus of a number of national and international geoscience programmes including a wide-ranging network of deep seismic investigations (Fig. 1). The Fennoscandian Shield exhibits a broad general trend of decreasing age of geological activity towards the SW. The northeastern part of the Shield is dominated by Archaean rocks, whereas the major part is made up of the Palaeoproterozoic 1.8-2.0 Ga Svecofennian Province (Korja et al. 2006) with the 1.65-1.8 Ga Transscandinavian Igneous Belt and the 0.9-1.2 Ga Sveconorwegian Province farther to the SW. This paper focuses on the Lapland-Kola orogen (Hjelt et al. 1996) located between the Murmansk and Karelian cratons in the northeastern part of the Shield, originally known as the Lapland-Kola mobile belt (Bridgwater et al. 1992). Although long regarded as an Archaean craton, recent investigations have shown that the Lapland-Kola orogen (LKO, Fig. 1) centred on the Kola Peninsula, is a Palaeoproterozoic collisional belt (e.g. Daly et al. 2001). Volumetrically, the LKO is made up mainly of Archaean material, but contains an important component of juvenile Palaeoproterozoic crust identified through isotope geochemistry combined with geochronology. The LKO provides a well-exposed section across a deeply eroded Palaeoproterozoic collisional belt, which exposes all of the critical geotectonic elements, thus providing insight into the operation of plate-tectonic processes from Neoarchaean times to the late Palaeoproterozoic. The LKO was the locus of major plume activity (Amelin & Semenov 1996; Lobach-Zhuchenko et al. 1998; Sharkov et al. 2000) in the late Neoarchaean to early Palaeoproterozoic, which led to the rifting and break-up of the Kenorland supercontinent (Williams et al. 1991; Pesonen et al. 2003; Mints & Konilov 2004; Balagansky et al. 2006) and then to the formation of oceanic crust, subduction and generation of Palaeoproterozoic juvenile continental crust (Daly et al. 2001, Balagansky et al. 2006). Later, collision formed the Palaeoproterozoic Lapland-Kola orogen, a Himalayan-scale, high-pressure, collisional belt traceable across the Atlantic to Greenland and Labrador (Bridgwater et al. 1992).

This paper is based on the authors' own work, together with a synthesis of published material from both Russian and western sources and involving a wide range of disciplines including petrology, structural geology, geochemistry, geochronology and geophysics, especially seismic reflection data. Bringing this material together was greatly facilitated by the Europrobe SVEKALAPKO project, of which the senior author was project co-leader. Following a brief outline of the geological framework, this paper sets out a working geodynamic model for the evolution of the LKO.

Geological background The LKO is structurally bounded to the north and south by the Murmansk and Karelian cratons, respectively. It is separated from the Karelian Craton to the south by the SW-dipping Northern Karelia suture and from the Murmansk Craton to the north by a NE-dipping Neoarchaean suture, which was reactivated during Palaeoproterozoic collision and is best known in the KolmozeroVoron'ya Belt (Fig. 2; Mints et al. 1996, and references therein). The Murmansk Craton comprises Neoarchaean tonalitetrondhjemite-granodioritic orthogneisses (TTG gneisses) and minor supracrustal rocks (Fig. 2; Mitrofanov et al. 1995b), intruded by Palaeoproterozoic basic dykes (Mitrofanov & Smol'kin 1995). It forms the northern foreland of the LKO and is largely free of Palaeoproterozoic compressional deformation. The Karelian Craton is a classic Neoarchaean granitegreenstone province, containing a few remnants of Mesoarchaean crust. It is cut by a series of Palaeoproterozoic rifts and related layered mafic intrusions (Slabunov et al. 2006). Balagansky et al. (1998a) divided the LKO into dispersed and accreted terranes. The dispersed terranes (Murmansk Craton, Kola Province, Belomorian Mobile Belt, Inari Terrane and Strel'na Terrane, Fig. 2) comprise fragments of a rifted Neoarchaean craton, reassembled in the Palaeoproterozoic. The accreted terranes include the Lapland Granulite Terrane (usually known as the Lapland Granulite Belt, LGB), Umba Granulite Terrane and Tersk Terrane, all composed of Palaeoproterozoic juvenile

From: Gzz, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 579-598. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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crust generated in an island-arc setting (e.g. Daly et al. 2001). These three terranes, together with the Tanaelv and Kolvitsa tectonic m~langes, and the Inari and Strel'na terranes make up the NW-trending orogenic core of the LKO between the Belomorian Mobile Belt and the Kola Province (Fig. 2). Collisional deformation is strongly developed in the orogenic core and also extends southwards through the Belomorian Mobile Belt into the Karelian Craton (Fig. 1). The Kola Province is a composite entity comprising the Kola-Norwegian Terrane, Keivy Terrane and the little known Sosnovka Terrane. The Kola-Norwegian Terrane is a typical Neoarchaean granulite-gneiss region made up mainly by

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Fig. 1. Schematicgeological(a) and tectonic (b) maps of the Fennoscandian Shield (data compiled from Gorbatschev & Bogdanova 1993; Balagansky2002; Glaznev 2003) showing the location of the principal seismic refraction lines and the BABEL seismic reflectionlines. Tectonic boundaries are shown as bold black lines. Dashed outlines of unexposedrapakivi granites (GR) are based on geophysicaldata (Glaznev 2003).

tonalite-trondhjemite-granodiorite (TTG) and diorite gneisses and peraluminous metasediments (Mitrofanov et al. 1995b). The Kola-Norwegian Terrane underwent only slight structural and metamorphic/thermal reworking in the Palaeoproterozoic (Dobrzhinetskaya 1989; de Jong et al. 1999; Balagansky 2002). The Keivy Terrane contains a rare example of Neoarchaean alkali granite magmatism (Batiyeva 1976; Mitrofanov et al. 2000; Zozulya et al. 2005) and spectacularly coarse-grained kyanite, staurolite and garnet schists (Bel'kov 1963), which are unknown elsewhere in the Fennoscandian Shield. The Sosnovka Terrane is composed of TTG gneisses of uncertain age (Balagansky et al. 1998a).

L A P L A N D - K O L A OROGEN

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The Belomorian Mobile Belt is a composite terrane comprising regional-scale nappes, which formed during Neoarchaean subduction and collision (Slabunov et al. 2006). Lithologically, these are made up of TTG gneisses, garnet and kyanite gneisses and mafic rocks. The Belomorian Mobile Belt experienced extensive Palaeoproterozoic structural and metamorphic reworking within discrete shear zones (Balagansky 2002) as well as more widespread thermal effects (Bibikova et al. 2001). The Lapland and Umba Granulite terranes are Palaeoproterozoic high-pressure granulite-facies metamorphic belts. The former is composed of a norite-enderbite ('charnockite') series and garnet, sillimanite and cordierite paragneisses, traditionally classified in the literature on the LGB as 'khondalites'. The latter consists of enderbite, charnockite and granite as well as khondalites (Balagansky 2002, and references therein). Deformation under high-grade conditions increases progressively downwards towards the footwall of the LGB (Ga~il et al. 1989). The Lapland and Unba Granulite terranes overlie the mainly mafic Tanaelv and Kolvitsa m61anges respectively, the upper parts of which experienced metamorphism cofacial with the overlying granulites. In addition, the Umba Granulite Terrane is closely related to the amphibolite-facies Tersk Terrane, which comprises TTG gneisses and subordinate supracrustals. The two terranes share a common regional variation in metamorphic grade, with the underlying Tersk Terrane showing an inverse metamorphic gradient (Belyaev et al. 1977; Daly et al. 2001). The Lapland Granulite Terrane and the Tanaelv M61ange are grouped together as the 'Lapland Granulite Belt' sensu lato, and the others are linked as the 'Kolvitsa-Umba-Tersk Belt'. The Lapland Granulite and Umba Granulite terranes and the Tersk Terrane are composed entirely of Palaeoproterozoic juvenile material (Huhma & Meril/iinen 1991; Daly et al. 2001, Fig. 3). Daly et al. (2001) emphasized the importance of belts of juvenile crust in identifying collisional sutures, and used this approach together with structural and deep seismic data to delineate the Lapland-Kola Suture (LK, Fig. 2) as one of the bounding structures of the orogenic core. The Inari Terrane was previously interpreted as mainly Archaean in age, but both Archaean (2.50-2.75Ga) and Palaeoproterozoic (c. 1.9 Ga) juvenile granitoid suites have been distinguished (Meril~iinen 1976; Vetrin et al. 1987; Barling et al. 1996, 1997; Skuf'in et al. 2003). The Strel'na Terrane, like the Inari Terrane, contains both Archaean and Palaeoproterozoic granitoids and various supracrnstal suites (Fig. 4). It is distinguished from the Tersk Terrane by its geological and geophysical structure, particularly its magnetic and

Tectonic packages of c.2.7, 2.4-2.5 and 1.9-2.0 Ga rocks TTG orthogneisses and supracrustals, 2.5-3.0 Ga

581

Fig. 2. Tectonic map of the Kola Peninsula and adjacent regions of northern Fennoscandia (modified after Balagansky 2002). Profiles 10 and 1EB are seismic reflection lines.

gravity (Balagansky et al. 1998a) and electric conductivity (Lyubavin et al. 1999) character. The Polmak-Pasvik-Pechenga-Imandra-Varzuga Greenstone Belt ('PV Belt'; Fig. 2), is a major Palaeoproterozoic rift structure, largely preserved on its northern side but substantially reworked on its southern side where it forms the northern boundary of the orogenic core. Berthelsen & Marker (1986) interpreted the PV Belt as the Palaeoproterozoic Kola collisional suture (now Pechenga-Imandra-Varzuga suture, PIV; Fig. 2) and a locus of ocean closure. The mid-ocean ridge basalt (MORB)-like geochemistry of the PV Belt tholeiitic basalts combined with island-arc geochemistry of andesites in the South Pechenga and Tominga groups (Fig. 2) stimulated a further development of this concept (Melezhik & Sturt 1994; Mints et al. 1996). The Pechenga Zone (Fig. 2) is widely known for its economic C u - N i deposits and as the site of the world's deepest borehole (Kozlovsky 1984).

G e o p h y s i c a l constraints on the m a j o r structures

The geometry and scale of the two major trans-crustal structures, the LK and PIV sutures (Fig. 2), that delimit the orogenic core and define the deep structure of the entire orogen have been imaged on various seismic refraction and reflection lines (Figs 1-3; Kozlov etal. 1995; Berzin etal. 1998; Pilipenko etal. 1999; Sharov 1997). The PIV suture is well constrained by seismic studies in the Pechenga Zone (Sharov 1997). Thrusts and reverse faults in the central and southern portion of the Pechenga Zone dip to the SSW and are traced by seismic reflection images southwards down to depths of 12-15 kin. These structures display a listric geometry, becoming subhorizontal at the deepest levels. A listric reverse fault coinciding with the SSW-dipping Polmak-Pasvik portion of the PV Belt was imaged by the POLAR profile (Fig. 1) down to c. 15 km (Gafil et al. 1989). A set of faults marking the PIV suture on the surface can be traced from the Pechenga Zone to the SE where they coincide with the southern boundary of the Imandra-Varzuga Zone (Fig. 2; Zagorodny et al. 1982; Mints et al. 1996; Mitrofanov 1996). Near the north-south-trending Main Ridge Anorthosite (just west of Monchegorsk, Fig. 2), the faults swing from northsouth (dipping east) to N W - S E (dipping NE), the latter illustrated by seismic reflection images along profile 1EB in the Monchegorsk area (Mitrofanov & Sharov 1998; Fig. 2). The NNE dip is typical of the reverse faults that delimit the Tominga Group

J.S. DALY ETAL.

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Fig. 4. end V. time plot for representative rock suites in northern Fennoscandia based on Balagansky (2002), using data from Bernard-Gfiffiths et al. (1984), Huhma & MeriRiinen (1991), Timmerman & Daly (1995), Amelin & Semenov (1996), Barling et al. (1997), Balaganskyet al. (1998b), Daly et al. (2001) and this paper (Table 1). (Fig. 2) in the western Imandra-Varzuga Zone. Similar dips are seen in the easternmost part of the zone. In the central part of the Imandra-Varzuga Zone, reverse faults dip towards the SSW along its northern and southern boundary. This geometry is supported by 3D density modelling, which indicates that relatively dense PV Belt metabasalts continue far to the south beneath lighter Archaean TTG gneisses (Mints et al. 1996). The geometry of the LK suture and associated structures bounding the LGB is well displayed on the POLAR profile (Fig. 1). NE-dipping listtic thrusts and highly reflective bands underlie the belt (Gafil et al. 1989) and can be traced to Moho depths (Walther & Fliih 1993; Pilipenko et al. 1999). The average dip of the lower boundary of the entire belt varies from 11 ~ (Gafil et al. 1989) to 14 ~ (Pilipenko et al. 1999; Fig. 3a and b). At the surface, the northern boundary of the belt steeply dips the SW on the POLAR profile (GaA1 et al. 1989). However, combined gravity and magnetic modelling suggests that it becomes vertical at deeper levels and, further downwards, dips to the NE (Marker et al. 1990). Profile 10 (Nota-Lotta-Rajakoski-Salmijarvi) along the RussianFinnish border (Figs 2 and 3c) shows that the LK suture dips gently to the NW near the surface and has an overall subhorizontal geometry (Kozlov et al. 1995; Sharov 1997). The northern boundary also dips gently to the NW, defining an acute wedge-like morphology of the belt, a geometry also indicated by gravity modelling (Marker et al. 1990; Buyanov et al. 1995). Farther east, along profile 1EB (Berzin et al. 1998; Fig. 2), the reflection data cannot be interpreted unambiguously close to the LK suture. However, they are consistent with a gentle dip to the NNE in conformity with field observations (Mitrofanov & Pozhilenko 1991). A general image of the crustal architecture of the northern Fennoscandian Shield constrained by field observations and seismic data is given in Figure 3d.

Tectonic model

The following sections set out a working tectonic model for the evolution of the Lapland Kola Orogen between c. 2.5 and c. 1.7 Ga, dealing with tiffing, oceanic separation, subduction and crustal growth, collision and post-orogenic relaxation. Rifting and break-up (2.5-2.1 Ga)

Rifting and break-up of the Kenorland supercontinent (Pesonen et al. 2003) in the early Palaeoproterozoic is manifested by

583

large-scale extensional mafic and ultramafic magmatism across the entire Fennoscandian Shield and elsewhere (e.g. Vogel et al. 1998). In northern Fennoscandia, widespread rifting of Archaean crust, as indicated by the emplacement of layered mafic and anorthositic intrusions and the initiation of rift basins (Zagorodny et al. 1982; Ga~il & Gorbatschev 1987), occurred between c. 2.35 and 2.5 Ga (Amelin et al. 1995; Mitrofanov et al. 1995a; Bayanova 2004). Rifting probably occurred in three stages, at c. 2.50-2.35 Ga, c. 2.35-2.2 Ga and c. 2.2-2.1 Ga (Melezhik & Sturt 1994), although the two latter intervals are less well documented. Supracrustal expressions of rifting are found within the Strel'na Group of the PV Belt (Melezhik & Sturt 1994) and the Kandalaksha Sequence of the Kolvitsa M61ange (Balagansky et al. 1998b, 2001b). Early stage (c. 2 . 5 - 2 . 3 5 Ga). The earliest rifting in the PV Belt

at c. 2.5-2.35 Ga (Melezhik & Sturt 1994) is represented by subaerial tholeiitic basalts and minor basaltic andesites and rhyolites (from bottom to top) of the Strel'na Group (Purnach, Kuksha and Seidorechka Formations) accompanied by shallow-water sedimentation within a linear clastic sedimentary basin (Zagorodny et al. 1982). The Purnach and Kuksha formations were deposited prior to the emplacement of the 2.5 Ga Pana-Fedorova mafic layered intrusions (Zagorodny et al. 1982). The youngest volcanism in the Strel'na Group (Seidorechka Formation) is constrained only by the crystallization age of 2.44 Ga for the Imandra Mafic Lopolith (Amelin et al. 1995; Bayanova & Balashov 1995), which cuts it. (Unless specified otherwise, all quoted ages are U - P b zircon ages determined by thermal ionization mass spectrometry (TIMS).) Precise U - P b zircon and baddeleyite ages of c. 2.44 Ga from a Seidorechka Formation granophyre (originally termed 'Imandrite'), identical to the age of the lopolith, have been interpreted as dating the volcanism (Amelin et al. 1995; Bayanova & Balashov 1995). However, Galimzyanova et al. (1998) have shown that the granophyres formed as a result of melting of the rhyodacitic host rocks to the lopolith. Thus the c. 2.44 Ga age is only a minimum estimate for the age of volcanism. The age of the rift-related Kandalaksha sequence in the Kolvitsa M61ange is better known (i.e. 2467 4- 3 Ma, Balagansky et al. 1998b, 2001b). Rifting initiated under dextral transtension (Balagansky et al. 1998a, 2001b) as a result of the effects of a large-scale plume (Amelin & Semenov 1996; Lobach-Zhuchenko et al. 1998; Sharkov et al. 2000). The location and orientation of the rift zones follows ancient structures within the Archaean basement (Balagansky et al. 2001a), possibly a response of structurally anisotropic lithosphere to plume activity. Structural anisotropy favoured dyke emplacement subparallel to the regional strike of both ancient structures and newly formed rifts rather than a radiating dyke swarm (Tommasi & Vauchez 2000). A large-scale plume is indicated by similar dextral transtensional kinematics over the entire Archaean outcrop within the Fennoscandian Shield (Balagansky et al. 1998a) and by the spatial distribution of mafic layered intrusions. Based on the latter, Amelin & Semenov (1996) suggested a plume diameter of at least 1000 km; that is, from the southern Karelian Craton to the Murmansk Craton (Fig. 1). Rifting resulted in the emplacement of large (kilometre-scale) layered gabbro-norites into Archaean host rocks throughout the LKO (e.g. Mt Generalskaya intrusion, Bayanova 2004) and within the Karelian Craton (e.g. Burakovka intrusion, Amelin & Semenov 1996), whereas smaller (100 m scale) dykes intruded the Murmansk Craton (Mitrofanov & Smol'kin 1995, and references therein). Nd and Sr isotopic data on mafic layered intrusions, gabbro-anorthosites, anorthosites and minor coronitic gabbros (known as 'drusites' in the Russian literature), and mafic volcanic rocks (eNd(t) from +0.5 to --3; Amelin & Semenov 1996; Balagansky et al. 1998b; Lobach-Zhuchenko et al. 1998, Mitrofanov & Smol'kin 2004) indicate a mildly enriched mantle

584

J.s. DALYETAL.

source consistent with a mantle plume model with small amounts of crustal contamination (Amelin & Semenov 1996). Mantlederived mafic magmatism was accompanied by anorogenic potassic granites (Lobach-Zhuchenko et al. 1998) dated at 2.402.45 Ga (Bogdanova & Bibikova 1993; Kaulina & Bogdanova 2OOO). The Strel'na Group along the entire PV Belt and all the 2.5 Ga intrusions (Mt Generalskaya, Monchegorsk and Pana-Fedorova) define a NW-trending axial zone of earliest extension (Fig. 2). Later extensional magmatism occurred over a wider area and is interpreted as an episode of magmatic underplating as the plume spread northeastwards and southwestwards. This is manifested by the intrusion of gabbronoritic dykes in the Murmansk and Karelian Cratons at c. 2.45 Ga (Mitrofanov & Smol'kin 1995; Mertanen et al. 1999; Hanski et al. 2001), corona gabbros in the Belomorian Mobile Belt dated at 2.43-2.46 Ga (Bogdanova & Bibikova 1993; Lobach-Zhuchenko et al. 1998; Kudryashov et al. 1999; Sharkov et al. 1999, 2004, Slabunov et al. 2001) and the younger, c. 2.44 Ga mafic layered intrusions (Amelin et al. 1995; Bayanova 2004) throughout the Fennoscandian Shield. Anorthosites dated at 2.45 Ga intruded the 2.5 Ga Mt. Generalskaya and Pana-Fedorova intrusions in the same interval (Bayanova 2004). Magmatic underplating has been also suggested by petrological and isotopic data from xenoliths in Devonian intrusions interpreted to represent Palaeoproterozoic lower crust under the Belomorian Mobile Belt (Downes et al. 2002; Vetrin 2006). Underplating was probably widespread and resulted in the development of voluminous lower crustal magma chambers, accounting for the numerous drusites throughout the Belomorian Mobile Belt (Amelin et al. 1995; Lobach-Zhuchenko et al. 1998; Sharkov et al. 1999, and references therein). The latest drusite magmatism has been dated at 2.36 Ga (Kudryashov et al. 1999), coeval with the formation of potassic granitic veins at 2.38 Ga (Kaulina & Bogdanova 2000). The development of abundant and widespread layered mafic intrusions and yet no development of an ocean between c. 2.5 and 2.35 Ga suggests relatively low extension rates and quiescent tectonics (i.e. little motion of the Kenorland supercontinent). This is supported by the common observation of repeated injection of mafic magmas over protracted time intervals, up to 100 Ma (Kudryashov et al. 1999; Balagansky et al. 2001b; Bayanova 2004; Mitrofanov & Smol'kin 2004). The restriction of the oldest magmatism to the PV Belt may indicate the position of the plume centre as suggested by Zagorodny & Radchenko (1988). The initial rifting probably occurred under quiescent conditions above the central region of the rising plume, which may account for the emplacement of plagioclase-rich cumulates and the development of gabbro-anorthosites, such as the Kolvitsa (Balagansky et al. 2001b) and Main Ridge (Mitrofanov & Smol'kin 2004) bodies, within the core zone of the LKO. In this respect the early stage of rifting is distinctive. The early stage was terminated by rift inversion accompanied by uplift, erosion and weathering. Inversion may be explained by crustal thickening by underplating (see above) and the subsequent isostatic response. Renewed rifting resulted in the deposition of the Ahmalahti and Polisarka Formations of the Pechenga and Varzuga Groups, respectively (Melezhik & Sturt 1994). Underplating resulted in crustal thickening and subsequent uplift and erosion. As a result, the Ahmalahti conglomerates unconformably overlie the early rift-related layered intrusions, such as the 2.5 Ga Mt. Generalskaya intrusion (Zagorodny et al. 1982, Melezhik & Sturt 1994), and the Seidorechka Formation was also inverted. Main stage (c. 2 . 3 5 - 2 . 1 Ga). The main tiffing stage is well recorded

in sedimentary-volcanic sequences of the PV Belt (Zagorodny et al. 1982; Melezhik & Sturt 1994; Mitrofanov & Smol'kin 1995; Mints et al. 1996) and less well documented by plutonic and dyke rocks. Shallow-water protoevaporitic 'red beds' and subaerial volcanic rocks were deposited in restricted basins, giving

rise to the Ahmalahti and Kuetsjarvi Formations of the Pechenga Group and the Polisarka and II'mozerka Formations of the Varzuga Group. Terrigenous and carbonate sediments display light rare earth element (LREE) enrichment, suggesting a continental crustal source. Alkaline volcanic rocks have a clear sodic trend, high TiO2 and Fe/Mg > 1, also consistent with a continental setting (Wilson 1989). Native copper mineralization is another characteristic feature of these volcanic rocks. Age determinations are available only for the Pechenga Group and are open to interpretation, being based on the whole-rock Rb-Sr isochron method (Mitrofanov & Smol'kin 1995). The Ahmalahti basaltic andesites and dacites yielded an age of 2324 _+ 28 Ma, identical to a Sm-Nd isochron age for basalts from the Per~ipohja Schist Belt in northern Finland (Huhma et al. 1990). It should be noted that basal conglomerates in both these cases overlie the 2.45-2.5 Ga layered mafic intrusions and contain pebbles derived from these plutonic rocks. The Kuetsj/irvi subalkaline volcanic rocks gave an age of 2214 • 54 Ma. Numerous, mainly NW-trending, dykes were emplaced in both Finland and Russia (Koistinen et al. 2001) during the main stage of rifting and have been correlated geochemically with coeval volcanic rocks (Vuollo 1994; Mitrofanov & Smol'kin 1995). In the Belomorian Mobile Belt the dykes comprise high-Fe tholeiitic garnet-coronitic gabbros comparable with continental flood basalts, and are dated at 2115 +_ 25 Ma (Stepanova et al. 2003). In the Finnish Belomorian and in the Karelian Craton, c. 2.1 Ga doleritic dykes are correlated with the Jatulian volcanic rocks (Vuollo 1994; Hanski et al. 2001). Extensive dyke magmatism throughout the Fennoscandian Shield clearly implies an extensional setting. The orientation of dykes, oblique to the major NW-trending Palaeoproterozoic structures, suggests transtensional conditions inherited from the earlier stage of rifting. Sheet-like potassic granitic bodies with an anorogenic chemistry (Terekhov & Levitsky 1995) intruded the 2.4-2.5 Ga mafic granulites in the Kolvitsa M61ange at 2289 +_ 20 Ma (Kaulina & Bogdanova 2000). K-rich granitic veins also occur among TTG orthogneisses in the Belomorian Mobile Belt, and these are almost coeval with the granites from the Kolvitsa M61ange (2266 ___ 12 Ma, Kaulina & Bogdanova 2000). Metamorphism related to rifting. Dextral transtension controlled

the 2.4-2.5 Ga mafic magmatism in the Belomorian Mobile Belt and in the Kolvitsa belt, resulting in extensional shear zones (Balagansky et al. 2001b) and rift-related metamorphism (Ivanov & Rusin 1997). Metamorphic P - T conditions, between c. 2.46-2.42 Ga (Balagansky et al. 2001b) associated with extensional deformation at deep crustal levels, are well documented at several locations within the Belomorian; for example, c. 725~ and c. 9.8 kbar in the Pongoma area (Alexejev et al. 2001), 700-710~ and l l - 1 2 k b a r in the Tolstik area (Bogdanova 1996), and 666-734~ and 8.3-9.8kbar in the Kolvitsa M~lange (Alexejev 1997). Extensional deformation and metamorphism in the Kolvitsa belt terminated by 2387 ___4 Ma (Kislitsyn 2001). However, there remain considerable problems in distinguishing between the early rift-related metamorphic events and those associated with collision during the Lapland Kola Orogeny at 1.90-1.92 Ga (see below). O c e a n i c s e p a r a t i o n (c. 2 . 1 - 1 . 9 7

Ga)

The first tectonic models to apply the Wilson cycle concept to the Palaeoproterozoic of northern Fennoscandia were proposed by Barbey et al. (1984) for the LGB and by Berthelsen & Marker (1986) for the PV Belt (Fig. 2). Although more data are now available, much uncertainty exists and some of the conclusions in the following sections are controversial. In particular, both the number of locations where tiffing led to full-scale oceanic spreading and the width of the resulting oceans are uncertain. Discrete

LAPLAND-KOLA OROGEN ophiolites have yet to be identified but there is compelling evidence for oceanic separation from basalt geochemistry, structure and the presence of arc-derived juvenile crust. The current model envisages two oceans, both possibly of modest width, marked by the PIV and LK sutures. The Pechenga-Imandra-Varzuga and Lapland-Kola oceans were geographically separated and, as suggested by geochronology (see below), the former opened at c. 1.99-1.97 Ga and the latter at c. 2.05 Ga. Both oceans closed by c. 1.9 Ga. Pechenga- Varzuga ocean. Berthelsen & Marker (1986) developed

a model in which the entire PV Belt was interpreted as the site of a collisional suture (PIV suture in Fig. 2) marking the locus of oceanic closure. This model was further developed by Mints et al. (1996) and Sharkov & Smol'kin (1997), who emphasized the MORB-like tholeiites in the upper part of the PV Belt as evidence for oceanic spreading. Mints et al. (1996) envisaged rifting (c. 2.5-2.4 Ga) leading to the formation of intracontinental oceans at c. 2.4 Ga and c. 2.1-2.0 Ga and a back-arc ocean at c. 2.0-1.9 Ga. Sharkov & Smol'kin (1997) suggested that oceanic spreading started at 2.2 Ga in a back-arc basin. Melezhik & Sturt (1994) suggested an intracontinental origin for the lowermost and middle formations of the PV Belt and the existence of a short-lived, Red Sea type ocean at 1.991.97 Ga. This time interval saw the eruption of picritic lavas and the emplacement of Cu-Ni-bearing gabbro-wehrlite intrusions. Crucial for Melezhik & Sturt's model was the MORB-like chemistry of the uppermost basic lavas (Pilguj~irvi Formation) of the North Pechenga Subzone coeval with the gabbro-wehrlite intrusions and dated at 1.96-2.00 Ga (TIMS Pb-Pb zircon and Rb-Sr, Pb-Pb, Sm-Nd and Re-Os whole-rock isochrons, Melezhik & Sturt 1994, and references therein; Mitrofanov & Smol'kin 1995, and references therein) and the subductionrelated geochemical signatures of the South Pechenga volcanic rocks, tentatively dated at c. 1.86 Ga (Rb-Sr whole-rock isochron, Mitrofanov & Smol'kin 1995) as well as the Tominga Group in the Imandra-Varzuga Zone. In addition, Melezhik & Sturt (1994) suggested that the Kolosjoki Formation in the North Pechenga Subzone also may have had a MORB-like mantle source, and its age of c. 2.1 Ga (Rb-Sr whole-rock isochron, Mitrofanov & Smol'kin 1995) might thus date the oldest oceanic crust. Minor acid volcanic rocks low in the Pechenga Group stratigraphy contain numerous partially melted xenoliths of Archaean granites and granite-gneisses (Melezhik & Sturt 1994), showing that this stage of rifting occurred within a continental environment, with crustal melting probably triggered by mafic underplating. A dramatic change in ~13C ratios recorded in carbonate rocks of the Kuetsj~irvi and Kolosjoki formations (Karhu 1993) dated at c. 2.2 Ga, and c. 2.1 Ga, respectively (Rb-Sr whole-rock isochron, Mitrofanov & Smol'kin 1995), may be related to changes of depositional environment (from closed to open basins) consistent with seafloor spreading (Melezhik & Sturt 1994). The narrow Red Sea scale of the resulting Pechenga-Varzuga ocean is consistent with the paucity of deep-water sediments (Melezhik & Sturt 1994; Sharkov & Smol'kin 1997) and the limited volume of juvenile crust produced by subduction (see below). Lapland-Kola ocean. Some of the mafic and intermediate rocks of

the Tanaelv M61ange dated at c. 1.9-2.0 Ga (TIMS U-Pb zircon, Rb-Sr, Pb-Pb and Sm-Nd whole-rock, Bernard-Griffiths et al. 1984) are characterized by trace element signatures clearly different from those of subduction-related and intraplate magmas (Barbey et al. 1986) and may represent remnants of oceanic crust. So far, the oldest U-Pb zircon ages of 2041 _+ 10 Ma and 2056 _ 28 Ma have been obtained from mafic and intermediate rocks, respectively (Kaulina et al. 2004), and may date oceanic

585

separation. In the Kolvitsa M61ange, the same age (2056 _+ 3 Ma, Kaulina & Bogdanova 2000) was obtained for a pegmatite vein interpreted to have intruded during oceanic separation. This vein cross-cuts mafic granulites that were deformed and metamorphosed at c. 2.4-2.5 Ga under extensional conditions. Dunite, harzburgite and pyroxenite bodies (up to 12 km long and 700 m across) are scattered along the southern margin of the Russian part of the Tanaelv M~lange, forming the Notozero Ultrabasic Belt (Vinogradov 1971). These are the best candidates for the remnants of an ophiolite complex, but require detailed study. A rather significant volume of metasedimentary and meta-igneous juvenile, subduction-related rocks may indicate a rather wide oceanic basin. On the other hand, the presence of detritus from Archaean sources (Sm-Nd model ages as old as 2.5 Ga, Daly et al. 2001; detrital zircon ages varying from 2.0 Ga to 3.6 Ga, Bridgwater et al. 2001; Tuisku & Huhma 2005; and geochemical constraints, Barbey et al. 1984, 1986) suggests a narrow ocean. Palaeomagnetism. High-quality palaeomagnetic data are available only for 2.45 Ga dykes from the Karelian Craton (Pesonen et al.

2003), which limit estimates of continental drift relative to the Murmansk Craton between 2.5 and 2.0 Ga. Tentative results from the 2.5-2.45 Ga layered mafic intrusions (Arestova et al. 1999; Khramov et al. 2006) indicate no major relative displacements of these landmasses since 2.45 Ga. These data suggest that neither the Pechenga-Varzuga nor the Lapland-Kola ocean were of very great size. S u b d u c t i o n a n d crustal g r o w t h (c. 2. O - 1.86 Ga)

Subduction-related magmatic and sedimentary rocks have been recognized in the core zone of the orogen, in the Lapland, Umba, Tersk and Strelna terranes and in the PV Belt. Lapland and Umba granulite terranes. A sedimentary origin for the

khondalites is supported by sedimentary structures such as grading (Meril/iinen 1976), geochemistry (Barbey et al. 1986; Kozlov et al. 1990; Bibikova et al. 1993), carbon (Korja et al. 1996), oxygen (Bibikova et al. 1993) and helium (Avedisyan et al. 1998) isotopic data, as well as the presence of detrital zircons with ages varying from 1.95 Ga (ion microprobe secondary ion mass spectrometry (SIMS) U-Pb, Tuisku & Huhma 2005) to 3.6 Ga (laser-ablation microprobe inductively coupled-plasma mass-spectrometry, Pb-Pb, Bridgwater et al. 2001). Compositionally, they represent a turbiditic flysch suite (Barbey et al. 1984, 1986). Sm-Nd isotopic and detrital zircon ages show that they contain both Palaeoproterozoic and Archaean detritus, with the former being markedly dominant (Barbey et al. 1984, 1986; Bibikova et al. 1993; Bridgwater et al. 2001; Daly et al. 2001; Tuisku & Huhma 2005). The norite-enderbites originated as island-arc calc-alkaline and high-Mg andesitic magmas (Barbey et al. 1986; Kozlov et al. 1990). They play a minor role in western part of the belt (c. 22% by volume in Finland, Korja et al. 1996) but become dominant in the east, in Russia (Fig. 2). Enderbites from the Umba Granulite Terrane display an island-arc geochemical signature (Glebovitsky et al. 2001), whereas Sm-Nd isotope data suggest derivation from a depleted mantle source (eNd at 1910 Ma from +0.5 to +1.2, Daly et al. 2001). Sedimentary and magmatic protoliths within the Lapland and Umba terranes formed between 1.91 and 2.1 Ga (Bibikova et al. 1993; Kaulina & Bogdanova 2000; Daly et al. 2001; Glebovitsky et al. 2001; Kislitsyn 2001; Kaulina et al. 2004; Glebovitsky 2005; Fig. 4). The youngest U-Pb ages of 2.17Ga (Kaulina & Bogdanova 2000) and 2.14 Ga (Sorjonen-Ward et al. 1994) from multi-grain detrital zircon fractions, together with the youngest Sm-Nd model age of 2.12 Ga (Daly et al. 2001) define the

586

J.S. DALY ETAL.

lower age limit of sedimentation. Tuisku & Huhma (2005) suggested that some of the sediments were deposited at 1.95 Ga. A few khondalite samples have positive eNd(t) values varying from +0.1 to +1.1 (Fig. 4), which along with the S m - N d model age of 2.12 Ga (Daly et al. 2001) confirm that the sedimentary protoliths are dominated by detritus derived from rocks identical in chemistry and close in age to the norite-enderbite series (Sorjonen-Ward et al. 1994). A younger limit on the age of sedimentation of 1.91 - 1.94 Ga is provided by felsic and intermediate intrusions that cut the khondalites (Bibikova et al. 1993; Glebovitsky et al. 2001; Kislitsyn 2001; Tuisku & Huhma 2005). A U - P b zircon age for a garnet-bearing quartz diorite in the Lapland Granulite Terrane (Meril~iinen 1976) suggests that some sedimentation took place by c. 1.98 Ga. Northeastward subduction (Daly et al. 2001) is indicated by the occurrence of subduction-related plutonic rocks NE of the Lapland Granulite Terrane (Barling et al. 1996, 1997). The northeastward dip of lithological layering and foliations within the Lapland Granulite Terrane and the Tanaelv M61ange and the NE-dipping seismic layering, which can be followed from the surface to the lower crust (Walther & Fltih 1993; Pilipenko et al. 1999), are consistent with this conclusion. The enderbites and diorites of the Umba granitoid complex dated at 1944 _+ 19 Ma (Kislitsyn 2001) display subductionrelated geochemical signatures (Glebovitsky et al. 2001; Balagansky 2002). These rocks intruded the Umba khondalitic metasediments, following a metamorphic event possibly caused by advective heating from subduction-related melts. Inari Terrane. As well as Archaean rocks, the Inari Terrane con-

tains Palaeoproterozoic TTG gneisses as well as Palaeoproterozoic supracrustal rocks identified by S m - N d studies in Finland (Barling et al. 1996, fig. 14.1) and Russia (the Vyrnim and Tal'ya Formations, Table 1). According to Barling et al. (1996, 1997), 1.94 Ga Palaeoproterozoic calc-alkaline TTG rocks in Finnish Lapland (dioritic to granodioritic intrusions with mafic enclaves) are characterized by large ion lithophile element (LILE) enrichment, negative Nb anomaly and slight high field strength element (HFSE) enrichment, similar to modern calc-alkaline island arcs. The 1.94 Ga intrusions have initial SVSr/86Sr in the range 0.7021-0.7029 and eNd(1940) +0.71 to +3.82 with S m - N d model ages of 2.07-2.47 Ga. These

TTG suites formed in a moderately evolved arc setting. TTG granitoids dated at 1.94 Ga and quartz diorites of the Kaskel'yavr Complex in Russian Lapland also display subduction-related geochemical signatures (Vetrin et al. 1987; Skuf'in et al. 2003). The Kaskel'yavr Complex occurs immediately south of the PIV suture, suggesting SW-directed subduction in the PV Belt. All these granitoids and quartz diorites have been thrust northeastwards onto the South Pechenga Subzone. Tersk and Strel'na terranes. Supracrustal rocks geochemically similar to those in the Lapland and Umba terranes occur in the Tersk Terrane (Daly et al. 2001), with island-arc magmatism at 1961 _+9 Ma (SIMS U - P b zircon; Fig. 4). To assess the eastward continuation of the Tersk Terrane and the proportion of Palaeoproterozoic crust in the Strel'na Terrane, a traverse was undertaken along the Strel'na River (Fig. 5). S m - N d data are presented for 12 samples (Table 1, Fig. 5). Ion microprobe U - P b zircon ages (Table 2, Fig. 5) are presented for four samples from three locations along the Strel'na River. Near the northern end of the traverse, a felsic gneiss, interpreted as volcanic, has a S m - N d model age of 2922 Ma (Table 1, Fig. 5) consistent with an Archaean age. A dioritic orthogneiss from P'yany Creek (7/00-9) yields a concordia age (Ludwig 1998) of 2771 _+ 13 Ma (Fig. 6a), determined on magmatic, oscillatoryzoned zircons and indistinguishable from its S m - N d model age of 2724 Ma (Table 1). Farther south within the Sergozero Unit, close to the boundary with the Tersk Terrane, felsic volcanic rocks have model ages of c. 2.7 Ga (Table 1, Fig. 5; Timmerman & Daly 1995). These rocks are undated, but their late Archaean model ages suggest that Palaeoproterozoic volcanism involving a mixed ArchaeanPalaeoproterozoic source. Two orthogneisses with Palaeoproterozoic S m - N d model ages yield similar U - P b zircon ages of c. 1975 Ma, interpreted as igneous crystallization ages, and are interpreted as juvenile crustal additions related to arc magmatism. A granitic orthogneiss from the Tersk Terrane (sample 7/00-30, near the mouth of the Strel'na River, Fig. 5) contains oscillatory-zoned zircons interpreted as magmatic. These gave concordant to slightly discordant U - P b zircon data (Table 2) and a 2~176 age of 1975 _+ 11 Ma (Fig. 6b). A dioritic gneiss from the Sergozero Unit of the Strel'na Terrane (sample 7/00-23, Fig. 5), collected 8 km south

Table 1. S m - N d data for samples from the Strel'na Terrane, the Inari Terrane and the Pechenga- Varzuga Belt Sm

Nd

Psammite Pelite Felsic orthogneiss (tuff) Granitoidorthogneiss Pelite Granitoidclast Psammitematrix Felsic orthogneiss Felsic orthogneiss Felsic orthogneiss Granitoidorthogneiss Granitoidorthogneiss

3.50 4.94 3.26 2.05 5.48 1.70 5.14 6.06 5.30 2.60 5.16 4.18

15.90 26.71 16.65 11.22 30.99 9.46 29.63 28.82 24.45 13.29 29.15 25.99

0.1328 0.1117 0.1185 0.1104 0.1069 0.1086 0.1049 0.1272 0.1310 0.1182 0.1071 0.0972

Pelite, Tal'ya Fm Metagreywacke,Vyrnim Fm

6.78 2.05

38.05 11.26

4.67

22.34

Sample

Lithology

1478m/144Nd

143Nd/la4Nd

2o-

tDM (Ma)

eNa (1960)

0.511640 0.511578 0.511226 0.511200 0.511327 0.511108 0.511036 0.511853 0.511888 0.511367 0.511604 0.511470

22 12 12 16 10 14 10 10 8 16 22 16

2652 2183 2922 2724 2446 2814 2819 2089 2121 2688 2049 2050

-3.45 0.68 -7.92 -6.39 -3.01 -7.75 -8.22 2.16 1.88 -5.16 2.37 2.23

0.1078 0.1099

0.511372 0.511470

18 14

2401 2304

-2.36 - 0.97

0.1264

0.511880

14

2022

2.89

Strel'na Terrane

7/00-1 7/00-2 7/00-5 7/00-9 7/00-16 7/00-19 7/00-20 7/00-23 7/00-12 7/00-14 7/00-25 7/00-30 Inari Terrane

B-903 B-872-II

PV Belt, Tominga Group

S- 195

Metarhyodacite,Panarechka Fm

tDM, depleted mantle model age after DePaolo (1981). Samples were analysed at University College Dublin following methods described by Menuge (1988), as modified by Menuge & Daly (1990). All 143Nd/144Nd ratios have been corrected to a value of 0.511850-t-5 for the La Jolla standard.

LAPLAND-

KOLA

~ 7/00-1, psm, 2652 7/00-2, pel, 2183 2144, grw, 2222*

~

Gabbro-amphibolite,

~

Alkali

~

F + +++++++++t).f

F++++ + +++ ,~ 1+_1+ i +~+~ tJ~

- t

S t r e l ' 4-n a4 - J f ~,,'~Te r r a n e

meta-ultrabasite granite

granodiorite

Granodiorite, tonalite,

Psephite, psammite. sericitic quartzite (Pestsovaya Tundra Fm) Micaceous. garnet-mica paragneiss and schist (Vysokaya Zemlya Fm) Acid and intermediate metavolcanics (Bezymyannaya Fro)

7100-9, gr-gn, 2771 +_ 13, 2724 7/00-19, pbl, 2814 7100-20, psm, 2819

Basic metavolcanics, metakomatiite (Pyalochnaya Fm) Cong{omerate, psammJte, basic and acid metavotcanics, quart2ite (Pyatochnaya Fm) Micaceous, garnet-mica gneiss and schist, conglomerate, quartzite (Peschanoozerskaya Fm) Micaceous, garnet-mica gneiss • sillimanite (Chapoma Unit) Gneiss. amphibelite, caic-si!icate (Sergozero Unit)

7/00-16, pel, 2446 7/00-23, fls-gn, 1974 • 8, 2089 7/00-22, peg, 1896 + t0

"

Leucogranite,

~

trondhjemite Biotite, hornblende and ] pyroxene-biotite gneiss, tonalite, granodiorite

'x"="~i~-

..... ' :

587

Leucogranite (c.1.8 Ga, Pushkarev 1990) Peridotite, pyroxenite, abbro-norite (1.97 Ga, Kuz'min et al. 2005)

~

7/00-5, fls-tuff, 2922

OROGEN

!

i ii!ili!i ii i iiiii !i ii,!iii

Imandra-Varzuga rift-belt, Palaeoproterozoic (undivided)

Tersk erran

,,~ 7/00-14, fls-gn, 2688 2292, dac, 2703*

7100-25, gr-gn, 2021

Biotite, hornblende and pyroxenebiotite gneiss, granitegneiss, tonalite, granodiorite

FauR

~

t'~ Sample Thrust v Iocafity

Generalized position of Laptand-Kola suture 7/00-30 = sample number; rock types: dac = metadacite, fls = felsic, fls-gn = felsic gneiss, gr-gn = 9ranitoid gneiss~ grw = metagreywacke, pbt = clast, peg = pegmatite, pet = pelite, psm = psammite, ttg = tonatite-trondhjemite-granodiorite; ~

7/00-30, gr-gn, 1975 -+ 11, 2050

1975+_11 = U-Pb zircon age (Ma); t 0 km White Sea

of the dated Archaean orthogneiss, yielded a U - P b zircon concordia age of 1974 _+ 8 Ma (Fig. 6c). Both ages were obtained from magmatic zircons displaying oscillatory zoning. The foliation in this rock is cut by a granitic pegmatite (sample 7/00-22), which yielded a U - P b zircon concordia age of 1896 _+ 10 Ma (Fig. 6d, see below). These U - P b zircon and S m - N d ages (Tables 1 and 2) indicate a Palaeoproterozoic age (c. 1.97 Ga) for the eastern portion of the Sergozero Unit, which along with isotopic ages from the western portion (Daly et al. 2001) defines the entire unit as Palaeoproterozoic. Palaeoproterozoic metasediments from the Strel'na Terrane (Pestsovaya Tundra Formation) adjacent to the eastern Imandra-Varzuga Zone have yielded S m - N d model ages of 2.18-2.22 Ga and 2.65 Ga (Table 1, Fig. 5). These data suggest two independent sources of detritus, late Palaeoproterozoic and Neoarchaean. A similar provenance is suggested by S m - N d model ages of 2.45 Ga (pelite), 2.81 Ga (granitoid clast) and 2.82 Ga (psammite matrix) for the Vysokaya Zemlya Formation (Table 1, Fig. 5). Both these formations, like the Sergozero Unit, have proved to be Palaeoproterozoic in age. Peridotite, pyroxenite and gabbronorite intrusions are widespread in the Sergozero Unit within the Strel'na terrane (Fig. 6) and are essentially coeval with the felsic orthogneisses, as suggested by a U - P b zircon age of 1 9 6 6 _ 6 Ma (Kuz'min et al.

2005). Northeastward thrusting is characteristic of the boundary between the Tersk and Strel'na terranes (Balagansky et al. 1998a). This boundary can be traced westwards using geoelectric data (Lyubavin et al. 1999) where it merges with the tectonic

2050 = Sm-Nd model age (Ma);

*= from Timmerman & Daly 1995.

Fig. 5. Geological map of the Strel'na River area (simplified and modified after Mitrofanov 1996) showing the sampling traverse, ion microprobe U-Pb zircon and TIMS Sm-Nd model ages.

collage separating the Umba Terrane and Kolvitsa M61ange (Fig. 2). Because the m61ange dips in general northeastwards, all these tectonic units may constitute a composite klippe (see Fig. 3). Extrapolation of the structural dip to the subsurface to suggest southward or southeastward subduction (Daly et al. 2001) is unproven and further seismic studies are required to establish the deep structure. The presence of subduction-related igneous rocks to the NE in the Sergozero Unit of the Strel'na Terrane (Fig. 5) permits NE-directed subduction; that is, the polarity is identical that in the Lapland Granulite Belt. P V Belt. Subduction-related volcanic assemblages of the South

Pechenga and Tominga Groups comprise rhyolites, dacites, andesites, picrites, boninites and subordinate normal MORB (N-MORB)-like basalts (Melezhik & Sturt 1994). R b - S r wholerock isochrons indicate an age of 1.86 Ga for the South Pechenga Group (Mitrofanov & Smol'kin 1995). However, their large uncertainties (up to 60 Ma) demand further geochronology. Bimodal picritic-andesitic volcanism is characteristic of these rocks and many andesites are similar to modern boninites from island-arc settings. They are associated with black shales with an andesitic volcaniclastic matrix together with grey and black cherts (Melezhik & Sturt 1994). The Panarechka Formation, dated at 1.91-1.94 Ga (Gavrilenko et al. 2005) and composed of 95% andesitic volcaniclastic rocks and 5% 'granodioritic' sandstones, is thought to be the youngest unit of the Tominga Group. One of the volcanic rocks (S-195, Table 1) has yielded a S m - N d model age of 2023 Ma (eNd ( 1 9 6 0 ) = § indicating a depleted mantle source and a lack of older crustal components.

588

J . S . D A L Y ETAL.

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Geochemical features of the South Pechenga and Tominga groups indicate an island-arc setting (Melezhik & Sturt 1994; Mints et al. 1996). Intense NE thrusting along the southern margin of the PV Belt (Melezhik & Sturt 1994; Mints et al. 1996), seismic data (Sharov 1997) and 3D density modelling (Mints et al. 1996) all suggest subduction towards the SW. Collision

Thrusts, providing unambiguous indicators of crustal shortening and thickening, have been well established in the LGB (including the underlying Tanaelv M61ange), Kolvitsa M~lange, Umba Granulite and Tersk terranes and in the PV Belt (Figs 1 and 2). In the LGB, crucial evidence is the occurrence of hightemperature, high-pressure granulites on top of amphibolite-facies rocks, separated by an intense mylonite zone (H6rmann et al. 1980; Gafil et al. 1989). NE-directed thrusts and reverse faults in the central and especially the southern part of the Pechenga Zone have long been well documented (V~iyrynen 1959). The Imandra-Varzuga Zone displays a complicated pattern (Mitrofanov 1996), with thrusts dipping both northwards and southwards at the southern boundary. Three-dimensional modelling strongly suggests north-directed thrusting. The northward dip of the thrusts is due to later doming along this boundary (Mints et al. 1996). The Kolvitsa M61ange (Fig. 2) is a spectacular series of NE-dipping thrust sheets (Balagansky 2002). The base of this tectonic pile is made up of remnants of the 2.47 Ga Kandalaksha Sequence, the basal conglomerates of which sit on a basement of c. 2.7 Ga granitic gneisses. In one thrust sheet, made up of the 2.43-2.46 Ga Kolvitsa Layered Gabbro-Anorthosite Massif, located a few kilometres from the conglomerates, the intrusion recrystallized immediately post-magmatically as a high-P granulite-facies assemblage at 990~ and 12.4kbar (Fonarev 2004). It was later sheared and metamorphosed under extensional conditions by 2.39 Ga. Therefore the emplacement of the gabbroanorthosites occurred at depths of c. 42 km at the same time as the deposition of the conglomerates at the surface. One of the upper thrust sheets comprises 2.7-2.8 Ga granitic gneisses derived from the basement. The uppermost sheet represents a collage of tectonic lenses varying in size from tens of metres to kilometres,

55 ;~O'tPbl2~5 U

5,Z

59

Fig. 6. U-Pb concordia diagrams.

Data are plotted as 2o- error ellipses.

made up of components of the c. 2.45 Ga Kolvitsa volcanoplutonic complex and the 1.9-2.0 Ga Umba khondalites and granitoids. The roof thrust of this tectonic collage is the lower boundary of the Umba Terrane. The occurrence of all these contrasting rocks originating at very different depths in different environments can be explained only by tectonic juxtaposition involving considerable crustal shortening. Seismic reflection images (Berzin et al. 1998) indicate that near-horizontal TTG gneisses underlie the Kandalaksha Sequence within the Kolvitsa M61ange. Gravity data (Buyanov et al. 1995) suggest that this terrane and the entire Kolvitsa M61ange and the Umba Terrane form horizontal sheets resting on less dense TTG rocks. Assuming a maximum dip of 14~ similar to that of the LGB (see above), a rough estimate of the horizontal component of displacement can be calculated. For example, raising the Kolvitsa Massif c. 42 km up a 14~ ramp by orthogonal thrusting requires a horizontal displacement of c. 170 km. The oblique dextral character of collision (Balagansky et al. 1998a) indicates an even larger horizontal displacement. Collisional structures a n d kinematics. Mineral lineations within the LGB exhibit two distinct orientations (see Gafil et al. 1989 fig. 4). One is NE plunging and has a constant orientation regardless of the regional strike swing (Fig. 2). It is thought to have developed during the thrusting of the LGB southwestwards when the belt acquired its arcuate shape (Ga~il et al. 1989). The other lineation is parallel to the strike of the belt and thus varies regionally. This strike-parallel lineation must have preceded the thrusting and probably formed during the transpressional stage of collision (Balagansky 2002). The lineation in strongly sheared rocks in the South Pechenga Subzone of the Pechenga Zone displays a symmetrical flower-like pattern. Interpreting the lineation as a tracer of displacement, shortening occurred northnortheastwards, accompanied and/or followed by lateral displacements to either side of the zone of maximum strain (Balagansky 2002). This deformation took place under low-temperature and moderate-pressure amphibolitefacies conditions (Mitrofanov & Smol'kin 1995). Kilometre-scale sheath folds complementary to c. 1.9 Ga nappes occur in the northern Belomorian Mobile Belt, and their orientation unambiguously indicates SW-directed thrusting

590

J.S. DALYETAL.

(Glebovitsky 2005). These nappes have resulted from the large-scale thrusting of the LGB to the SW. The southern boundary of the central part of the ImandraVarzuga Zone is characterized by steep lineations that formed under low amphibolite-facies conditions (Zagorodny et al. 1982) and is consistent with reverse faulting to the NNE. The Kolvitsa Mrlange displays a subhorizontal, NW-trending, penetrative lineation that developed during dextral transpression (Balagansky 2002). The lineation is represented by orthopyroxene and sillimanite crystallized at 825-845~ and 9.7 kbar in mylonites surrounding tectonic lenses of pelitic rocks in the mrlange, which unambiguously defines the lineation as collision related (Kozlova et al. 1991). As in the Tanaelv Mrlange, it is parallel to the general NW strike of the Kolvitsa Mrlange, and is attributed to strike-slip movements (Balagansky 2002). Two principal thrusting episodes have been distinguished in the Strel'na Terrane. Complementary to the NE-directed thrusting, NE- and SW-directed thrusts have been established in the Black River area (Fig. 5). Leucosomes with garnet and hornblende often mark the thrust surfaces, indicating thrusting under hightemperature and high-pressure amphibolite-facies conditions. Younger thrusting occurred only eastwards, also with leucosomes along the thrust surfaces, but only rarely with garnet and hornblende. This east-directed thrusting has been discovered in a TTG dyke cross-cutting a NE-directed thrust, thus unambiguously defining the sequence of events. Thus, compression during the second thrusting was oblique to the NW-trending Lapland-Kola suture, and this is linked with the transpressional stage of collision in northern Fennoscandia (Balagansky et al. 1998a). As in the Kolvitsa Mrlange, dextral transpression has been demonstrated. Shearing is widespread in the Keivy Terrane, the most characteristic structural feature of which is a ubiquitous penetrative lineation that plunges to the NNE near the Murmansk Craton and elsewhere mainly to the SSW and NNE (Bel'kov 1963; Batiyeva 1976; Petrov & Glazunkov 1987). In pelitic schists the lineation formed under amphibolite-facies conditions (550-560~ 4.0-5.3 kbar; Petrov & Glazunkov 1987). The NNE-plunging lineation also occurs in the Serpovidny Ridge Belt, which is composed of rocks correlated with the PV Belt (Mitrofanov et al. 1995b, and references therein). This indicates a Palaeoproterozoic age for the deformation, the development of which probably reflects thrusting of the Murmansk Craton onto the Keivy Terrane to the SSW (Mints et al. 1996). The internal structure of the Serpovidny Ridge Belt, in which all rocks dip northwards, is clearly asymmetrical and discordant to the surrounding pelitic schists of the Keivy Group (Balagansky 2002). This is consistent with the idea that the belt is an east-westtrending tectonic sheet (Milanovsky 1984) rather than a syncline as usually believed. Furthermore, the entire terrane may represent a tectonic collage as suggested by Mints et al. (1996) and Bridgwater et al. (2001). In the northern Belomorian Mobile Belt, high-P-T amphibolite shear zones display a gentle, locally horizontal, lineation, subparallel to the NW-trending regional strike (Balagansky 2002). The shear zones are commonly associated with spectacular sheath folds in which the lineation, defined both by field observations and petrofabric analysis, is parallel to the sheath axis. Kilometre-scale sheath folds occur in the Seryak Lake area (Balagansky 2002), the geometry of which, together with kinematic indicators observed at the northeastern (Kislitsyn 2001) and southwestern boundaries of the Belomorian Mobile Belt, suggests dextral movements. Relationships between steep and gentle lineation. Collision has

resulted in steep and gentle lineations in upper and lower crustal levels, respectively, the steep lineation being oriented generally perpendicular to the regional strike and the gentle one parallel to the strike (Balagansky 2002). This lineation pattern conforms to the transpression model of Jones & Tanner (1995), which

suggests that (sub)vertical movements of compressed rock masses dominate in the upper, more rigid, crust whereas (sub)horizontal movements, parallel to the strike, prevail in the lower, more ductile, crust. Similar rheological controls have been demonstrated for gently and steeply dipping Laxfordian shear zones in the lower and upper crust, respectively, in Scotland (Coward 1990) and is to be expected given that the crust is rheologically heterogeneous on a large scale. Multistage history of collision. At least three stages of collision can

be distinguished (Balagansky 2002). The first stage is recognized in the Strel'na Terrane and is related to the main NE-SW compression, which may be interpreted as the result of collision between the oldest island arcs (dated at 1.98 Ga) or between these and a continent. Exposure-scale structures, formed during this event, were almost completely reworked during the second stage, when movements took place parallel to the regional NW strike, and high-grade (locally, orthopyroxene-sillimanite) lineated mylonites and blastomylonites formed. This stage is dominantly transpressional and is thought to be linked to continentcontinent collision. All available data suggest dextral movements, first pointed out by Balagansky et al. (1998a). The third stage is so far recognized only in the Finnish part of the LGB, and is characterized by SW-directed thrusting during which the earlier fabrics acquired their arc-like geometry. Age of collision. Crucial for determining the timing of collision are

isotopic ages from magmatic rocks, whose intrusive relationships bracket the collisional deformation. In addition, as the granulitefacies metamorphism is attributable to crustal thickening and collision, direct dating of petrographically distinct metamorphic zircons provides an independent constraint. In the Tanaelv M~lange, collisional shearing under granulitefacies conditions took place between the intrusion of an anorthosite at 1945 • 10 Ma and the subsequent intrusion of a cross-cutting basic dyke dated at 1928 _ 10 Ma (Kaulina et al. 2004). An age of 1943 • 3 Ma for metamorphic zircons from the Abvar Anorthosite Massif within the mrlange (Mitrofanov et al. 1995a) may be dating one of the earliest stages (between island arcs), provided that the granulite-facies metamorphism is correctly attributed to collision (for an alternative model for the metamorphism, see Mints & Konilov 2004). In the LGB in Russian Lapland, a 1925 • 12 Ma quartzhypersthene diorite belonging to the norite-enderbite suite (Meril~iinen 1976) is cut by a synmetamorphic enderbite, dated at 1925 • 1 Ma (Bibikova et al. 1993), suggesting that the metamorphic peak was reached immediately after the main (continent-continent) collision. Ages of 1925-1930Ma, attributed to the metamorphic peak (Bibikova et al. 1993; Kaulina et al. 2004), also should be related to the main collision. U-Pb ages of detrital zircons from the khondalite suite indicate that deposition continued until at least 1950 Ma, whereas intrusion of norite-enderbite bodies took place at 1910-1930 Ma (SIMS, U-Pb, zircon, Tuisku & Huhma 2005). These rocks are affected by the final stage of SW-directed thrusting, which therefore took place after 1910 Ma. In the Kolvitsa-Umba-Tersk Belt, the older age limit is defined by ages of 1944 _ 19 Ma (Kislitsyn 2001) for the Umba enderbites and 1961 _+ 9 Ma (SIMS, U-Pb, zircon, Daly et al. 2001) for the Sergozero meta-andesites, both displaying island-arc geochemistry. The 1912 • 8 Ma collision-related Umba charnockites (Glebovitsky et al. 2001) contain a xenolith of orthopyroxenesillimanite gneiss (Vinogradova & Vladimirov 1990), indicating that the metamorphic peak and the main collisional event occurred between 1944 + 19 Ma and 1912 + 8Ma. In the Kolvitsa M~lange, the younger limit is constrained by an age of 1912 + 2 Ma from a leucosome cross-cutting an orthopyroxenesillimanite felsic granulite with a gentle SE-plunging lineation (Kislitsyn 2001); that is, this post-dates the second, transpressional

LAPLAND-KOLA OROGEN stage. In the Tersk Terrane, a cross-cutting vein that also postdates the transpressional deformation is dated at 1906-t-9 Ma (SIMS, U-Pb, zircon, Daly et al. 2001). A late leucosome from the Voche-Lambina area in the northern Belomorian Mobile Belt, affected by an east-plunging, transpressional lineation is dated at 1898_+2 Ma (Kislitsyn 2001), and thus provides the youngest limit on the age of collision. To the south, in the Chupa-Loukhi area, the latest collisional metamorphism and migmatization took place between 1840 and 1875 Ma (SIMS, U-Pb, zircon, Bibikova et al. 2004). Post-tectonic granites in the core of the LKO yield U-Pb ages of metamorphic zircon, titanite, rutile and magmatic zircon and Ar-Ar hornblende ages in the 1.87-1.89 Ga interval (see review by Glebovitsky 2005). Thus, in the orogen core, three stages of collision occurred from c. 1945 Ma to c. 1860 Ma. It is possible that the earliest collisions between island arcs or island-arc-continent collisions could have taken place a few tens of million years earlier. The main collision in the Kola region took place from 1930 Ma to 1905 Ma. Younger ages of 1840-1875 Ma from the underlying Belomorian Mobile Belt suggest that, this being the footwall of the orogenic core, gravitational upfift and cooling was delayed. Additionally, some compressional deformation could have occurred as a far-field effect of the younger collision in the Svecofennian orogen (Korja et al. 2006) to the south (Fig. 1).

granulite- to amphibolite-facies rocks of the Tanaelv M61ange (Fig. 7; Gafil et al. 1989, Mints et al. 1996, 2000; Perchuk & Krotov 1998; Perchuk et al. 1999). Mafic granulites, within the Lapland Granulite Terrane and within the upper part of the Tanaelv M61ange, show the same co-facial multi-stage metamorphic history, which indicates that both units shared at least the latest metamorphic and deformational events, having formed a single composite tectonic structure, the LGB (Mints et al. 1996, 2000). Within the Lapland Granulite Terrane and the upper part of the Tanaelv M~lange, thermobarometric data from petrographically determined paragenetic sequences (M1, M2, etc.) define a series of steep arrays on the P - T plane (Fig. 7). Each array spans a wide range of pressure, but only a limited range of temperature. However, these arrays do not represent P - T - t paths. Individual rocks define retrograde P - T - t paths with slopes approximating to classic isobaric cooling regimes (Fig. 7). Importantly, when analysed spatially, the P - T arrays correspond closely to metamorphic field gradients. Significantly, maximum pressures are recorded near the structural base of the Lapland Granulite Terrane, close to the Tanaelv M61ange, and pressures consistently decrease upwards (Mints et al. 2000). This strong depth dependence of the observed pressure values (Fig. 7) provides clear evidence for the controlling influence of lithostatic loading within this terrane. Moreover, the thickness of the Lapland Granulite Terrane calculated from these data (20-23 km) corresponds well to the thickness estimated from the seismic data (18-20 km; Mints et al. 2000). Thus, although

Collision-related m e t a m o r p h i s m . Metamorphism of the LGB is a classic example of inverted metamorphism, with high-P granulites of the Lapland Granulite Terrane overlying lower-grade

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592

J.s. DALY ETAL.

individual rocks record a range of retrograde P - T values, in general, 'high-pressure' and 'moderate-pressure' metamorphic assemblages reflect simultaneous events in lower and upper parts of the crust rather than events of different ages. Nevertheless, the idea that some moderate-pressure granulites had preceded the high-P granulites also has been developed (e.g. Belyaev & Kozlov 1997). The earliest (M1) stage, which is recognized only in mafic granulites from the upper part of the Tanaelv M~lange (Mints et al. 1996, 2000) developed at 860-960 ~ and 10.3-14.0 kbar. P - T values for M2 in the khondalites range from 860 ~ at 12.4 kbar (depth c. 45 kin) to 800 ~ at 5.8 kbar (depth c. 21 km). M3 parageneses are the most widespread and record temperature variations (from bottom to top, structurally) from 770 ~ to 640~ corresponding to pressures from 10.7 kbar to 4.8 kbar, (i.e. a depth interval from 39 to 17 kin). M 4 P - T variations range from 650 ~ to 550 ~ and 8.4 kbar to 4.5 kbar, corresponding to depths of 30-16 km. Khondalites within the Umba Terrane define a clockwise P - T t path (Alexejev 1997). As in the LGB, maximum pressures (7.6-8.1 kbar at c. 850 ~ are recorded near the boundary with the underlying Kolvitsa M61ange. In contrast to the metamorphic pattern in the Lapland Granulite Terrane, the metamorphism of the Tanaelv M~lange is characterized by an inverted metamorphic gradient (Perchuk & Krotov 1998; Perchuk et al. 1999). Granulite-facies rocks in the upper part of the m61ange are underlain by amphibolite-facies rocks. Pelitic rocks record a prograde metamorphic path best seen in the Korva Tundra Sequence. These rocks display a single-stage amphibolite-facies metamorphism and a clockwise P - T - t path (Fig. 7), whose maximum P - T values converge with the lower P - T values for the M4 parageneses from the overlying upper part of the Tanaelv M61ange and Lapland Granulite Terrane. The precise age constraints indicate a rather long duration for the metamorphism associated with the Lapland-Kola collision in the orogenic core, between c. 1.95 and 1.87 Ga (Bibikova et al. 1993; Daly et al. 2001; Kislitsyn 2001; Alexejev et al. 2003; Kaulina et al. 2004; Tuisku & Huhma 2005). However, the present state of knowledge allows two principal metamorphic events to be identified. The first involved heating the tectonic pile making up the LGB to granulite-facies conditions. This occurred after the assembly of the thrust sheets and is dated between c. 1.95 Ga (youngest detrital zircon age) and the metamorphic peak at 1924 _ 2 Ma (Bibikova et al. 1993), recorded by the khondalites (M2 stage). Ignoring the age uncertainties, the first metamorphism thus took place over a maximum time interval of 26 Ma. Packages of khondalite and norite-enderbite are cut by pre-metamorphic thrusts (see Korja et al. 1996, fig. 1), implying that the norite-enderbites cannot have been the principal heat source for the early metamorphism. The second metamorphism (M3-M4) of the khondalite and norite-enderbite series occurred between 1.92 and 1.87 Ga (i.e. over a maximum interval of 50 Ma). Knowledge of this event is much more detailed. It records the cooling history of the hot granulites (from c. 850 ~ and the simultaneous prograde heating to 600 ~ (Fig. 7) of the underlying amphibolite-facies part of the Tanaelv M61ange (Korva Tundra Sequence), the hot overthrust slab of granulites acting as the heat source. Major crustal shortening took place between the two (M1-M2 and M3-M4) events resulting in the tectonic juxtaposition of deep-seated granulites, metamorphosed at 40-45 km depth, with surface deposits (the Korva Tundra conglomerates). Within the Belomorian Mobile Belt, the Lapland-Kola collision resulted in a regional set of compressional shear zones, which operated under amphibolite-facies conditions and were accompanied by migmatization (Mitrofanov & Pozhilenko 1991; Zinger et al. 1999; Balagansky 2002; Bibikova et al. 2004). Collisional leucosome developed near the orogenic core at

1 8 9 8 _ 2 M a (Kislitsyn 2001), whereas leucosomes formed throughout the Belomorian Belt were dated at 1840-1875 Ma (SIMS, U-Pb, zircon, Bibikova et al. 2004). P o s t - c o l l i s i o n a l stage

Post-collisional cooling and gravitational collapse is thought to result from erosion above the orogenic core and subsequent uplift and exhumation of high-grade orogenic roots. The time of cooling can be determined by conventional mineral isotopic dating (e.g. Cliff 1985; Mezger et al. 1992; Bibikova et al. 2001; Jenkin et al. 2001), although complications are evident within the LKO because of the widespread effects of excess radiogenic argon (e.g. de Jong et al. 1999). Orogenic core. An age of 1887 + 8 Ma (Alexejev et al. 2001) for an unmetamorphosed post-deformational granitic dyke crosscutting khondalites in the Umba Terrane provides a younger limit on the termination of regional metamorphism and deformation. At the same time the Umba granitoids cooled to c. 500 ~ as shown by a R b - S r hornblende-plagioclase isochron age of 1882 • 15 Ma and an A r - A r hornblende age of 1889_+8 Ma (Cliff et al. 1997). A r - A r hornblende ages of 1902_+3 Ma and 1875 • 3 Ma were determined on supracrustal rocks from the western portion of the Tersk Terrane (Daly et al. 2001). Similar cooling ages have been obtained for the Tanaelv MElange underlying the LGB. Metamorphic zircon U - P b ages of 1906 + 13 Ma and 1885 + 11 Ma from an ultrabasic dyke in the Russian part of the m61ange date the end stages of metamorphism, consistent with metamorphic zircon ages of 1866-1895 Ma from a 1928 _ 10 Ma basic dyke (Kaulina et al. 2004). The same time interval is characteristic of amphibolite-facies zircons, titanites and rutiles (1.85-1.92Ga; Nerovich 1999; Kaulina et al. 2004). In the Finnish part of the LGB, a S m - N d garnet age of 1870_+7 Ma (Daly et al. 2001) possibly indicates uplift of the granulites into the upper crust by this time (Tuisku & Huhma 2005). This age seems to characterize the third, SW-directed thrusting stage, following the second, transpressional stage dated at 1.90-1.91 Ga (see above). Late extensional events, thought to be indicative of gravitational collapse, are known in only a few places so far. The earliest extensional structures are conjugate granulite-facies shear zones, which developed after the peak of metamorphism and deformation in the Umba Terrane. WNW-ESE-oriented extension took place under P - T conditions of 7.9___0.25 kbar and 845+13 ~ as recorded by pelitic assemblages (Alexejev 1997), indicating that the transition to extension (i.e. the onset of gravitational collapse) started under granulite-facies conditions. In the Black River area in the Strel'na Terrane (Fig. 5), the final extension is represented by two deformational events. E N E WSW-orientated horizontal extension resulted in a set of normal faults filled with thin granite pegmatite veins. The second stage involved north-south-oriented extension and is represented by greenschist- and low-temperature amphibolite-facies shear zones affecting the veins. The orientation of the dated pegmatite (Figs 5 and 6d) suggests that it was emplaced during this event (i.e. at 1896 Ma). Footwall and hanging wall of the orogenic core. A prolonged thermal history has been established for the Belomorian Mobile Belt, in the lower part of the LKO allochthons (i.e. the footwall of the orogenic core). Thermal reworking that resulted in the growth of metamorphic zircons and in the complete resetting of all mineral isotopic systems at 1.8-1.9 Ga (apart from U - P b in zircon) is characteristic of the entire belt (see Bibikova et al. 2001). The northeastern marginal zone of the belt underlying the orogenic core has somewhat older titanite ages of c. 1.94-1.87 Ga. These

LAPLAND-KOLA OROGEN ages are consistent with the main collision and cooling in the orogenic core. Differences in the titanite and rutile U - P b ages suggests a cooling rate between 2 and 4 ~ Ma -1. In contrast, the southwestern marginal zone, including the contact with the Karelian Craton, is characterized by newly grown titanite and rutile, the ages of which vary from c. 1.78 Ga to 1.75 Ga and with considerably smaller age differences between the titanite and rutile than elsewhere in the belt. The same younging of cooling ages from NE to SW across the Belomorian Mobile Belt is also indicated by A r - A r hornblende ages (de Jong et al. 1999). In the central part of the Belomorian Mobile Belt, in the Pongom Navolok area, titanite yielded a U - P b age of 1 8 1 1 + 4 M a (Levchenkov et al. 2000). Extension there is indicated by the emplacement of pegmatite veins at c. 1.83 Ga (Alexejev, Balagansky et al. unpubl, data). In contrast to the footwall, the hanging wall of the orogenic core (i.e. the Kola Province; Fig. 2) has suffered insignificant thermal effects during the Lapland-Kola Orogeny. This is indicated by the local preservation of Neoarchaean A r - A r ages (de Jong et al. 1999). The younging of cooling ages from the orogenic core to the southwestern margin of the footwall (Belomorian Mobile Belt) may be explained by earlier exhumation of the hotter and more tectonically thickened core of the orogen. As a result, during the uplift of the core and its footwall, the latter seems to have been tilted slightly to the SW. If so, the latest movements along the northeastern boundary of the orogenic core (in particular, along the southern boundary of the PV Belt) should have resulted in normal faulting. Post-orogenic magmatism and implications for cooling and exhumation of the orogenic core. Post-orogenic granites known as 'Nattanen

granites' mark the final stage of the Lapland-Kola Orogeny. They intruded episodically into the Finnish part of the orogen at 1772 Ma, 1790 Ma and 1798 Ma (Front et al. 1989). The LitsaAra Guba complex in the northwestern Kola Region intruded at 1.76 +_ 0.01 Ga (Vetrin et al. 2002), whereas the Strel'na Granite Massif in the SE (Fig. 5) is dated at c. 1.8 Ga (Pushkarev 1990). One of the Nattanen granites, the Juvoaivi Massif, intruded the Lapland Granulite Terrane at depths of 2 - 4 km (Dubrovsky 1969). Assuming a temperature of 60 ~ at a depth of 4 km (Glaznev 2003), an average cooling rate of 5.3 ~ Ma-1 can be calculated using maximum P - T values for the c. 1.95Ga anorthosites of c. 12 kbar and 924 ~ and a 1.925 Ga age for the metamorphic peak (see above). Correspondingly, the average exhumation rate from a depth of 39km is 0.23 mm year -1. A faster cooling rate of 7.4 ~ Ma -1 can be calculated using rutile U - P b ages of c. 1860 Ma, assuming that these correspond to a closure temperature of c. 450~ In both cases, these calculations show that the orogenic core cooled faster than its footwall, where the cooling rate varies from 2 to 4 ~ Ma -1 (see above). Protracted cooling and exhumation may be explained by the interior position of the LKO within a newly assembled supercontinent, Nena (Gower et al. 1990), which includes the Svecofennian orogen to the south. The final Palaeoproterozoic magmatism in the orogen is represented by lamproites dated in the South Pechenga Subzone at 1.71 + 0.01 Ga (Skuf'in et al. 1999) and in the Por'ya Guba area (the Kolvitsa M61ange) at 1.72 Ga ( S m - N d whole-rock isochrom, Nikitina et al. 1999). Conclusions The preponderance of Archaean crust in the northern part of the Fennoscandian Shield led in the past to the misconception

593

that the region comprises a Neoarchaean craton. However, the data reviewed and presented in this paper show that the evolution of the Lapland-Kola orogen spans the break-up of the Neoarchaean Kenorland supercontinent and its reassembly as a collage of Neoarchaean crustal fragments and intervening belts of juvenile Palaeoproterozoic crust. Protracted cooling and exhumation may be explained by the buttressing effect of the interior position of the LKO within a newly assembled supercontinent, Nena, which includes the Svecofennian orogen to the south. In northern Fennoscandia, widespread rifting of Neoarchaean crust started at 2.5 Ga and continued until 2.1 Ga, associated with the break-up of the Kenorland supercontinent. Within the Lapland-Kola orogen, the major sites of rifting include the Tanaelv-Kolvitsa, Pechenga-Imandra-Varzuga and Northern Karelia rifts. Rifting is manifested by large-scale mafic and ultramafic magmatism, the development of rift basins, as well as high grade-metamorphism and extensional deformation. Rifting involved dextral oblique extension (transtension) under the influence of an ascending mantle plume. Rifting led to sea-floor spreading and the development of the Pechenga-Varzuga and Lapland Kola oceans, which later became the sites of collisional sutures. Subduction of oceanic crust within the Pechenga-Varzuga ocean led to the growth of juvenile crust, making up the South Pechenga Subzone and the Tominga Group, whereas larger-scale arc magmatism within the Lapland-Kola ocean led to the development of the Lapland Granulite, Umba and Tersk terranes, and parts of the Inari and Strel'na terranes. Juvenile Palaeoproterozoic crust mainly developed between 1.98 and 1.91 Ga and is manifest as tonalite-trondhjemite-granodiorite plutonic, sometimes charnockitic, orthogneisses, felsic orthogneisses of probable volcanic origin, and voluminous, often peraluminous, metasediments deposited at or before 1.95 Ga. Nd isotopic and U - P b zircon data show that although the majority of these materials are juvenile, both sediments and magmatic rocks were also derived from a component of Archaean material, suggesting that the oceanic system involved was of restricted extent, a conclusion also supported by palaeomagnetic data. Oceanic closure resulted in collisional orogeny initially between arc systems but eventually involving the adjacent continents, especially the protoliths of the Belomorian Mobile Belt and the Karelian Craton. Collision took place diachronously between 1.95 and 1.87Ga and mainly between 1.93 and 1.91 Ga. Collision resulted in a Himalayan-scale mountain belt manifest by the widespread development of high-pressure granulite-facies metamorphism, and the overthrusting of large slabs of deep-seated crust such as the Lapland Granulite Terrane over structurally shallower units such as the lower part of the Tanaelv M61ange. An inverted metamorphic field gradient in the footwall of the orogenic core (Tanaelv M61ange and Tersk Terrane) is co-facial with a normal gradient in its hanging wall (Lapland and Umba granulite terranes). Thermobarometric data show that the resulting metamorphic pressures were essentially lithostatic, with a close correspondence between recorded pressure and structural depth. Some of the latest compressional deformation may have been a far-field effect of the Svecofennian Orogeny taking place to the south, on the other side of the Karelian Craton. Gravitational collapse of the orogen resulted in high-grade shear zones associated with W N W - E S E extension, which changed to north-south extension as the orogen cooled. The P - T conditions and time of intrusion of post-orogenic 'Nattanen' granites, emplaced between 1.80 and 1.77 Ga, together with a wide variety of mineral cooling isotopic ages, place some constraints on the late thermal history and show that the orogenic core cooled (5-7 ~ year - i ) and was uplifted more rapidly than its footwall ( 2 - 4 ~ year-i).

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This paper is dedicated to the memory of Lena Balaganskaya. The authors acknowledge financial support from RFBR (project 00-05-65468), INTAS and RFBR (grant INTAS-RFBR 95-1330), ISF, RFBR and the Russian Government (grants NM 1000 and NM 1300). Field work in the Tersk and Strel'na terranes in 2000 was supported by De Beers Centenary and in 2004 by project ONZ 6.13 'Nanoparticles in the nature: conditions of formation, ecological and technological effects' (subprojects 1082-2004 and 1083-2004). Discussions with M. Mints greatly improved our understanding of many aspects of the Lapland-Kola orogen. D. Gee, T. Brewer and an anonymousreviewer are thanked for constructive comments and suggestions.This paper is a contribution to IGCP Project 509 'Palaeoproterozoic Supercontinents and Global Evolution' and programme ONZ-6 'Geodynamics and mechanics of lithosphere deformation', and is NORDSIM publication 132. The NORDSIM facility is financed and operated under an agreement between the research councils of Denmark, Norway and Sweden, the Geological Survey of Finland and the Swedish Museum of Natural History. We also thank M. Murphy (UCD) for skilled technical assistance with Sm-Nd analyses.

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TOMMASI, A. & VAUCHEZ, A. 2000. Continental rifting parallel to ancient collisional belts: an effect of the mechanical anisotropy of the lithospheric mantle. Earth and Planetary Science Letters, 185, 199-210. TUISKU, P. & HUHMA, H. 2005. Generation of the norite-enderbite series of the Lapland Granulite Belt: implications from SIMS U-Pb-dating of zircons. Geophysical Research Abstracts, 7, 08022. V,~YRYNEN, H. 1959. The crystalline basement of Finland. Foreign Literature Publishers, Moscow [in Russian, translated from Finnish: Suomen kallioperd sen synty ja geologinen kehitys. OTAVA Kustannusosakeyhtir, Helsinki]. VETRIN, V. R. 2006. Deep xenoliths. In: MOROZOV, A. F. & PAVLENKOVA, N. I. (eds) Structure and dynamics of the lithosphere of Eastern Europe: Results of EUROPROBE programme studies, in press [in Russian]. VETRIN, V. R., PUSHKAREV, Yu.D., RYUNGENEN, G. I., SHLAIFSTEIN, B. A. & SHURKLNA,L. K. 1987. Geological position and age of granitoids in the southern frame of the Pechenga Belt. In: BALAGANSKY, V. V., ZAGORODNY,V. G., PETROV, V. P. & RADCHENKO,A. T. (eds) Structure and metamorphic evolution of major structural zones of the Baltic Shield. Kola Branch of the USSR Academy of Sciences, Apatity, 83-93 [in Russian]. VETRIN, V. R., BAYANOVA, T. B., KAMENSKY, I. L. & IKORSKY, S. V. 2002. U - P b age and isotope geochemistry of helium in rocks and minerals of the Litsa-Ara Guba diorite-granite complex (Kola Peninsula). Transactions of Academy of Sciences, Moscow, 387, 85-89 [in Russian]. VINOGRADOV, L. A. 1971. The Alpine-Type Ultrabasitic Formation in the southwestern Kola Peninsula (Notozero Ultrabasic Belt). In: BEL'KOV, I. V., SHURKIN, K. A., IVANOVA, T. N., ZAGORODNY, V. G., TYUREMNOV,V. A. & BOGACHEV,A. I. (eds) Problems ofmagmatism of the Baltic Shield. Nauka, Leningrad, 147-153 [in Russian]. VINOGRaDOVA, G. V. & VLAOIMIROV,A. G. 1990. Xenoliths in intrusive charnockites of the Umba Complex (Kola Peninsula). In: VETRIN, V. R. (ed.) Deep inclusions and the petrogenesis of intrusive charnockitoids. Kola Science Centre, USSR Academy of Sciences, Apatity, 28-47 [in Russian]. VOGEL, D. C., VUOLLO, J. I., ALAPIETI, T. T. & JAMES, R. S. 1998. Tectonic, stratigraphic, and geochemical comparisons between ca. 2500-2440 Ma mafic igneous events in the Canadian and Fennoscandian Shields. Precambrian Research, 92, 89-116. VUOLLO, J. 1994. Palaeoproterozoic basic igneous events in eastern Fennoscandian Shield between 2.45 and 1.97 Ga, studied by means of mafic dyke swarms and ophiolites in Finland. Acta Universitatis Ouluensis, Oulu, A250, 1-47. WALTHER, C. & FLOH, E. R. 1993. The POLAR Profile revisited: combined P- and S-wave interpretation. Precambrian Research, 64, 154-168. WILLIAMS, H., HOFFMAN,P., LEWRY, J. F., MONGER, J. W. H. & RIVERS, T. 1991. Anatomy of North America: thematic geologic portrayals of the continent. Tectonophysics, 187, 117-134. WILSON, M. 1989. Igneous Petrogenesis. Unwin Hyman, London. WHITEHOUSE, M. J., CLAESSON, S., SUNDE, T. & VESTIN, J. 1997. Ion microprobe U - P b zircon geochronology and correlation of Archaean gneisses from the Lewisian Complex of Gruinard Bay, northwestern Scotland. Geochimica et Cosmochimica Acta, 61, 4429 -4438. WHI'rEHOUSE, M. J., KAMBER, B. S. & MOORBATH, S. 1999. Age significance of U - T h - P b zircon data from Early Archaean rocks of west Greenland: a reassessment based on combined ion-microprobe and imaging studies. Chemical Geology (Isotope Geosciences Section), 160, 201-224. WIEDENBECK, M., ALLt~, P., CORFU, F., ETAL. 1995. Three natural zircon standards for U - T h - P b , Lu-Hf, trace element and REE analysis. Geostandards Newsletter, 19, 1-23. ZAGORODNY, V. G. ~ RADCHENKO,A. T. 1988. Tectonics of the Karelides in the northeastern Baltic Shield. Nauka, Leningrad [in Russian]. ZAGORODNY, V. G., PREDOVSKY, A. A., BASALAEV, A. A., ETAL. 1982. The Imandra- Varzuga zone of the Karelides (geology, geochemistry, evolutionary history). Nauka, Leningrad [in Russian]. ZECK, H. P. & WmTEHOUSE, M. J. 1999. Hercynian, Pan-African, Proterozoic and Archean ion-microprobe zircon ages for a

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EUROBRIDGE: new insight into the geodynamic evolution of the East European Craton S V E T L A N A B O G D A N O V A 1, R. G O R B A T S C H E V 1, M. G R A D 2, T. J A N I K 3, A. G U T E R C H 3, E. K O Z L O V S K A Y A 4, G. M O T U Z A 5, G. S K R I D L A I T E 6, V. S T A R O S T E N K O 7, L. T A R A N 8 & E U R O B R I D G E A N D P O L O N A I S E W O R K I N G GROUPS*

1Department of Geology, Lund University, Srlvegatan 12, SE-223 62 Lurid, Sweden (e-ma il: Svetlana. Bogdanova @geol. lu. se) 2Institute of Geophysics, University of Warsaw, Pasteura 7, 02-093, Warsaw, Poland 3Institute of Geophysics, Polish Academy of Sciences, Ks. Janusza 64, 01-452 Warsaw, Poland 4Department of Geophysics, University of Oulu, FI-90014 Oulu, Finland 5Department of Geology and Mineralogy, Vilnius University, Ciurlionio 21/27, LT-2009 Vilnius, Lithuania 6Institute of Geology and Geography, T. Sevdenkos 13, LT-2600 Vilnius, Lithuania 7Institute of Geophysics, NAS Ukraine, Palladin Ave., 32, 03680 Kiev, Ukraine 8Institute of Geochemistry and Geophysics, Kuprievich 7, 220141 Minsk, Belarus

Abstract: The Palaeoproterozoic crust and upper mantle in the region between the Ukrainian and Baltic shields of the East European

Craton were built up finally during collision of the previously independent Fennoscandian and Sarmatian crustal segments at c. 1.8-1.7 Ga. EUROBRIDGE seismic profiling and geophysical modelling across the southwestern part of the Craton suggest that the Central Belarus Suture Zone is the junction between the two colliding segments. This junction is marked by strong deformation of the crust and the presence of a metamorphic core complex. At 1.80-1.74 Ga, major late to post-collisional extension and magmatism affected the part of Sarmatia adjoining the Central Belarus Zone and generated a high-velocity layer at the base of the crust. Other sutures separating terranes of different ages are found within Sarmatia and in the Polish-Lithuanian part of Fennoscandia. While Fennoscandia and Sarmatia were still a long distance apart, orogeny was dominantly accretionary. The accreted Palaeoproterozoic terranes in the Baltic-Belarus region of Fennoscandia are all younger than 2.0 Ga (2.0-1.9, 1.90-1.85 and 1.84-1.82 Ga), whereas those in Sarmatia have ages of c. 2.2-2.1 and 2.0-1.95 Ga. Lithospheric deformation and magmatism at c. 1.50-1.45 Ga, and Devonian rifting, are also defined by the EUROBRIDGE seismic and gravity models.

The East European Craton (EEC) is the coherent Precambrian (mainly Archaean and Palaeoproterozoic) part of Europe that occupies the northeastern half of the continent. Geodynamic research in this region, however, is hampered by the presence of an extensive younger sedimentary cover. Geological study of the EEC therefore commenced in the two shields of exposed Precambrian crust, the Baltic (also Fennoscandian) Shield in the NW and the Ukrainian Shield in the SW. From the mid-1930s onwards, study of these shields was complemented by geophysical surveys of the Russian Platform, and subsequently in the 1940s and thereafter by deep drilling into the basement. By the 1970s, large parts of the Baltic Shield had been studied in sufficient detail to allow the first attempts at plate-tectonics interpretation (Hietanen 1975). Despite the success of this approach, it was evident that the shields alone were not large enough to fit the scales of plate-tectonic processes. Attention therefore shifted increasingly to the covered parts of the EEC, particularly the region between the Baltic and Ukrainian shields. In that region, geophysical reconnaissance had indicated the presence of large arcuate, mainly NE-trending rock units and structures in *EUROBRIDGE & POLONAISE Working Groups: V. N. Astapenko, A. A. Belinsky, R. G. Garetsky, G. I. Karatayev, V. V. Terletsky, G. Zlotski (Belarus); S. L. Jensen, M. E. Knudsen, H. Thybo, R. Sand (Denmark); K. Komminaho, U. Luosto, T. Tiira, J. Yliniemi (Finland); R. Giese, J. Makris (Germany); A. Ce~ys,J. Jacyna, L. Korabliova,V. Nasedldn, G. Motuza, A. Rimsa, R. Serkus (Lithuania); W. Czuba, E. Gaczyfiski,M. Grad, A. Guterch, T. Janik, P. Sroda, M. Wilde-Pirrko (Poland); E. Bibikova (Russia), S. Bogdanova, R. Gorbatschev, S. Claesson, S.-A. Elming, C.-E. Lund, J. Mansfeld, L. Page, K. Sundblad (Sweden); J. J. Doody, H. Downes (UK); V. B. Buryanov, T. P. Egorova, T. V. II' chenko,O. M. Kharitonov,D. V. Lysynchuk,O. V. Legostayeva,I. B. Makarenko, V. D. Omel'chenko, M. I. Orlyuk, I. K. Pashkevich,V. M. Skobelev,L. M. Stepanyuk (Ukraine); G. R. Keller, K. C. Miller (USA).

the crystalline basement beneath the Phanerozoic and Mesoto Neoproterozoic sedimentary cover. The latter is c o m m o n l y 1 - 2 k m thick, but locally reaches c. 1 0 k m or more; for example, in the Pripyat Trough and along the margin of the EEC towards the Trans-European Suture Zone (TESZ). Major integrated geophysical and geological projects in the southwestern part of the EEC, discussed since the late 1980s, were given new impetus by discoveries concerning the structure and evolution of the entire EEC that were first presented at the EUROPROBE symposium in Jablonna, in 1991 (Bogdanova 1993). The new work demonstrated that the system of Neo- to Mesoproterozoic rifts, which subdivides the craton into three parts, was superimposed on previously unknown late Palaeoproterozoic sutures located where three independent crustal segments (Fennoscandia, Sarmatia and Volgo-Uralia) were inferred to have collided to form the EEC (Bogdanova 1993; Gorbatschev & Bogdanova 1993b; Bogdanova et al. 1996, 2005; Khain & Leonov 1996). Thus, the EUROBRIDGE project was designed to test fundamental hypotheses regarding the formation of the EEC with a focus on the junction zone between Fennoscandia and Sarmatia. Since the E U R O B R I D G E project mostly focused on a region covered by sedimentary deposits and the Baltic Sea, geophysical studies along a transect extending from the Baltic Shield in southeastern Scandinavia to the vicinity of the Black Sea came to serve as its backbone. The work comprised seismic refraction profiling associated with wide-angle reflection and other, mainly potentialfield, geophysical studies. These were integrated with extensive geochronological, geochemical and geological investigations, the last focusing particularly on the conditions of metamorphism and their variation in time. In this review, emphasis is placed mainly on the integration of geology and geophysics. Although reports on the various EUROBRIDGE seismic profiles, other geophysical studies, and most of

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 599-625. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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the isotope-geological data have been published independently (Bogdanova et al. 2001a), this synthesis also identifies crustal structures and upper mantle irregularities that can be related to post-collisional processes and the subsequent thinning of the crust as a result of post-collisional extension and magmatism. The sites and roles of suture zones and other boundary faults are assessed and discussed, and isotope geochronology, geochemistry and the conditions of metamorphism are employed to define major accretionary and collisional events.

to tectonic stacking of the Palaeoproterozoic crust (Gorbatschev & Bogdanova 1993a). Throughout this region, lower crustal granulites have been juxtaposed with upper and mid-crustal amphibolite-facies rocks. The Palaeoproterozoic ages and juvenile nature of this crust have been assessed by geochronological reconnaissance studies across the various tectonic units. Several isotopic methods were employed, including U-Pb on zircons and monazites, Ar/Ar on amphiboles, and Sm-Nd model ages (Puura & Huhma 1993; Bogdanova et al. 1994, 2001b; Bibikova et al. 1995, 2001; Claesson & Ryka 1999; Valverde-Vaquero et al. 2000; Claesson et al. 2001; Mansfeld 2001; Dtrr et al. 2002; Puura et al. 2004; Soesoo et al. 2004; Krzeminska et al. 2005). The results demonstrate that virtually no Archaean crust was involved in the Palaeoproterozoic processes in the Baltic-Belarus region,

Geological background The assembly of several rock belts in the region between the Baltic and Ukrainian shields (Figs 1 and 2) has been interpreted to be due

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Fig. 1. Major tectonic subdivisions of the crust in the western part of the East European Craton: CBSZ, Central Belarus Suture Zone; KP, Korosten Pluton; LLDZ, LoftahammarLinktping Deformation Zone; MLSZ, Mid-Lithuanian Suture Zone; O-J, Oskarshamn-Jtnktping Belt; PDDA, Pripyat-Dniepr-Donets Aulacogen; PKZ, Polotsk-Kurzeme fault zone. The dashed light yellow line delimits the Volyn-Orsha Aulacogen. Red lines show the positions of the EUROBRIDGE (EB'94, EB'95, EB'96 and EB'97), Coast and POLONAISE (P4 and P5) seismic profiles. The inset shows the three-segment structure of the East European Craton (Bogdanova 1993; Khain & Leonov 1996).

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except in the form of Archaean detrital zircons in Palaeoproterozoic metasediments. Thus, the Palaeoproterozoic crust of the Baltic Shield appears to continue southwards to the vicinity of the Meso- to Neoproterozoic V o l y n - C e n t r a l Russian Aulacogen, where Fennoscandia meets the Palaeoproterozoic margin of Sarmatia (Fig. 1).

Major terrane boundaries Majo~ f~ulVshear zones

Fig. 2. Major lithotectonic units of the crust in the EUROBRIDGE study area. The location of the refraction and wide-angle reflection DSS profiles are also indicated. (a) Magnetic anomaly patterns in the region (modified after a map provided by S. Wybraniec, Polish Geological Institute). (b) Tectonic domains and belts. B-I, Borisov-Ivanovo Belt; BPG, BelarusPodlasie Granulite Belt; BZ, Berdichev Zone; CBSZ, Central Belarus Suture Zone; CnZ, Ciechantw Belt; DD, -Dobrzyfi Domain; EL, East Lithuanian Belt; FSS, Fennoscandia-Sarmatia suture; KP, Korosten Pluton; MD, Mazowsze Domain; MLSZ, Mid-Lithuanian Suture Zone; Mz, Mazury plutonic rocks; Ok, Okolovo terrane; OMB, Osnitsk-Mikashevichi Igneous Belt; PD, Podolian Domain; PDD, PripyatDniepr-Donets Aulacogen; Tt, Teteriv Belt; VD, Volyn Domain; VG, Vitebsk Granulite Domain; WLG, West Lithuanian Granulite Domain. The dashed light yellow line delimits the Volyn-Orsha Aulacogen. Black lines show the position of the EUROBRIDGE (EB'94, EB'95, EB'96, EB'97) and POLONAISE (P4 and P5) seismic profiles. The inset shows the three-segment subdivision of the East European Craton (Bogdanova 1993; Khain & Leonov 1996), and the EUROBRIDGE study area.

The recently established age patterns suggest multistage deformation and metamorphism during several accretionary and stacking events, the latter affecting and reactivating already existing Palaeoproterozoic crust. It consists of several Palaeoproterozoic terranes (Figs 1 and 2), belonging both to Fennoscandia (the Okolovo, L i t h u a n i a n - B e l a r u s and P o l i s h - L i t h u a n i a n

602

S. BOGDANOVAETAL.

terranes) and to Sarmatia (the Osnisk-Mikashevichi Igneous Belt and Teterev-Belaya Tserkov belt). The Fennoscandia- and Sarmatia-related terranes participating in the wide Central Belarus Suture Zone are separated by the Minsk Fault, which is inferred to be the major lithospheric discontinuity in the EEC (Bogdanova et al. 1996; Taran & Bogdanova 2001). The precise ages of the latest stages of formation of the beltshaped tectonic pattern in the region were assessed by Ar/Ar work on newly grown amphiboles in mylonites along shear zones that separate the various rock belts. The results suggest a surprisingly late stage of metamorphism at 1.71-1.67 Ga throughout the study region (Bogdanova et al. 2001b). Similar ages have recently also been obtained by the S m - N d method for garnets from the granulites in southern Estonia (Puura et al. 2004). Numerous faults transect and complicate the collisional structures (Figs 1 and 2). The dominant NE to NNE trends of the metamorphic belts and their bounding faults were mostly formed by collisional tectonics before 1.80 Ga, but were reactivated later, between 1.8 and 1.7 Ga. Another set of NW- to WNW-striking faults may be related to similarly oriented, c. 1.8 Ga shear zones in southern Finland and central Sweden (Beunk & Page 2001; Bergman et al. 2004). A third important group are west-east-oriented faults and deformation zones, some of which are accompanied by Mesoproterozoic mafic and gran~toid intrusions (Beunk & Page 2001; Bogdanova et al. 2001b; Ce~ys et al. 2002; Skridlaite et al. 2003b; Ce6ys 2004). These zones clearly cut the various rock belts and offset the pre-existing tectonic patterns. Bogdanova (2001) proposed to refer this deformation and the attendant magmatism to the 'Danopolonian' orogeny, which affected the southwestern margin of the EEC at c. 1.50-1.40 Ga, roughly coinciding in time with the 'Hallandian' event of Hubbard (Hubbard 1975; S6derlund et al. 2002) in southwestern Sweden. Lithotectonic units

In this section, the lithotectonic units (domains and belts) and bounding deformation zones along the EUROBRIDGE transect are considered in order from NW to SE. The metamorphic records ( P - T - t paths) of the various terranes along this transect are important keys to the histories of their formation and interaction (Fig. 3). The latter aspect is particularly relevant in the tectonically complex Central Belarus Suture Zone between Fennoscandia and Sarmatia.

F e n n o s c a n d i a n terranes o f the B a l t i c - B e l a r u s region

Potential gravity and magnetic fields show that the Precambrian bedrock of southeastern Sweden continues far to the east across the Baltic Sea (Wybraniec et al. 1998; Wybraniec 1999; Bogdanova et al. 2005). Related lithologies have also been found in drill-holes into the basement of the island of Gotland (Sundblad et al. 1998; Sundblad & Claesson 2000). Still farther east and SE, the crust is subdivided into several belts and domains characterized by varying grades of metamorphism. These units form the Polish-Lithuanian, Lithuanian-Belarus and Okolovo terranes, which differ in age and tectonic position (Figs 1 and 2). The Polish-Lithuanian terrane comprises the West Lithuanian Granulite Domain, which continues southwards into the Mazowsze Domain, the Ciechan6w Belt and probably also the Dobrzyfi Domain in the crystalline basement of central Poland (WLG, MD, CnB and DD in Fig. 2). With regard to their evolution and crustal ages, between 1.85 and 1.80 Ga, these units are in many ways similar to each other; they also resemble the rock complexes in southeastern Sweden.

To the east of the Polish-Lithuanian terrane, the somewhat older (c. 1.90-1.87 Ga) Lithuanian-Belarus terrane (Fig. 2), is composed of amphibolite-facies East Lithuanian Belt (EL) and the granulitic Belarus-Podlasie Belt (BPG). They developed nearly simultaneously with the crust in the classical Svecofennian Domain of the Baltic Shield. Between the Polish-Lithuanian and Lithuanian-Belarus terranes is the c. 50 km wide Mid-Lithuanian Suture Zone (MLSZ in Figs 1 and 2), across which the gravity and magnetic patterns differ greatly, the tectonic grain trending more or less west-east in the Polish-Lithuanian terrane and NNE-SSW in the LithuanianBelarus terrane (Skridlaite & Motuza 2001). Substantial differences are also found with regard to the thickness of the crust and its P-wave velocity and density images. Some rocks within the MLSZ are strongly deformed equivalents of those in the adjoining terranes, and there is a rather sharp metamorphic break between moderatepressure western granulites and high-pressure amphibolite-facies rocks in the east. Thus, the MLSZ represents a major deformation zone along which the West Lithuanian Domain and the entire Polish-Lithuanian terrane appear to have been thrust eastwards over the Lithuanian-Belarus terrane (Skridlaite et al. 2003a). The Polish-Lithuanian terrane: the West Lithuanian Granulite Domain (WLG). According to evidence from numerous drill-cores

(Skridlaite & Motuza 2001; Motuza 2005), the WLG is made up of granulite-facies para- and orthogneisses, metamorphosed in the lower crust at depths of 30-40 km between c. 1.85 and 1.80 Ga. The gravity and magnetic fields as well as the geometries of rock distribution suggest the dominance of W N W - E S E structures, but locally these trends have been rotated to align with large NE-SW- and west-east-striking lineaments (Skridlaite & Motuza 2001; Motuza 2005). In general, the structural patterns of the crust in the WLG resemble those in southeastern Sweden, but its northernmost part, and farther north in Latvia, is occupied by the major AMCG-type, c. 1.6 Ga, Riga pluton (R~im6 et al. 1996). (Here and elsewhere, AMCG is the abbreviation for anorthositemangerite-charnockite-(rapakivi)-granite magmatic suites (after Emslie et al. 1994).) To judge from chemistry, the metasedimentary granulites have been mostly formed from marine pelites, whereas the protoliths of the orthogneisses were mostly intermediate and felsic calc-alkaline, island-arc type magmatic rocks. The isotopic ages of the detrital zircons and the S m - N d isotopic characteristics of the metasedimentary rocks (Claesson et al. 2001) suggest provenance from Palaeoproterozoic sources with ages between c. 2.4 and 2.0 Ga. The deposition of the sediments can have taken place at any time between 2.0 and 1.84 Ga, the latter age being that of the oldest charnockitic intrusions (Motuza 2005). Whereas charnockitic and somewhat younger granitoid rocks are fairly common in the WLG, they totally dominate its continuation in the Mazowsze Domain (MD) of Poland, where a granodiorite and some metavolcanic rocks have been dated at c. 1.8 Ga (Valverde-Vaquero et al. 2000; Krzeminska et al. 2005). The S m - N d model ages of the MD rocks (Claesson & Ryka 1999) allow correlation with similar rocks in the Transscandinavian Igneous Belt (TIB) of Sweden (H6gdahl et al. 2004). The major, roughly north-south-trending, positive magnetic anomaly, which accompanies this belt in Scandinavia, appears to turn towards the SE beneath the Baltic Sea (Wybraniec 1999; Bogdanova et al. 2005). This supports its continuation into Poland. In the western part of the WLG, charnockitic and enderbitic magmatism, and related peak metamorphism of metapelites and felsic granulites, took place at c. 1.85 to 1.80 Ga (U-Pb zircon ion probe (NORDSIM) ages. Temperatures and pressures of 850-900~ and 0.8-1.0GPa, respectively, suggest burial depths of 3 5 - 4 0 k m (Fig. 3a, path 1). In the same area, a second stage of high-grade metamorphism at c. 1.79 Ga (U-Pb

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Fig. 3. P - T metamorphic records of the tectonothermal evolution of the Palaeoproterozoic terranes in the EUROBRIDGE region (after Taran & Bogdanova 2001, 2003; Skridlaite et al. 2003a and new P - T work). The pressure-temperature-time ( P - T - t ) paths of the terrane evolution represent the following. (a) The WLG metasedimentary and meta-igneous granulites, which demonstrate post-accretionary uplift with several steps of isobaric reheating and cooling, indicating magmatic inter- and underplating. (b) Various records from the MLSZ. The metasediments in the east have recorded near-isothermal burial, near-isobaric heating related to granite intrusions and uplift. In contrast, meta-igneous rocks in the west display two types of paths with different peak temperatures at the same depths. This may be a result of their tectonic juxtaposition. Prominent steps of near-isobaric cooling and reheating were related to magmatic emplacements. (c) Collision and subsequent uplift of the EL metasediments adjacent to the BPG, and uplift overprinted by a step of isobaric cooling. The latter was related to granite intrusions. (d) Subduction-triggered accretion, collision and uplift of the interior and marginal part of the BPG as meta-igneous and metasedimentary rocks show. (Note a minor step of near-isobaric cooling in the marginal parts of the BPG, which was related to Mesoproterozoic granitic magmatism.) (e) Accretion followed by collision, near-isothermal decompression as recorded by the Okolovo metavolcanic rocks. (f) Subduction-triggered accretion followed by magmatic emplacements, uplift and the near-isobaric cooling as traced by metasediments of the VG Domain.

monazite age) was related to partial melting of the metasedimentary and metavolcanic granulites during subsequent uplift. The temperatures were 7 3 0 - 8 5 0 ~ at pressures of c. 0.8 GPa (Fig. 3a, path 1). A later reheating-cooling step at c. 1.641.61 Ga is prominent in all the P - T records from the western WLG. During that stage, temperatures varied between 550 and 700 ~ at pressures around 0.6 GPa (Skridlaite et aL 2004). In the northern part of the WLG, peak metamorphism at 760 ~ and 0.7 GPa (Fig. 3a, path 2) occurred at c. 1.62 Ga, this age estimate being derived from metamorphic overgrowths on igneous zircon cores in charnockites (Claesson et al. 2001). Major tectonothermal activity at roughly the same time is indicated also by the c. 1.6 Ga ages of the Riga pluton. In the central part of the WLG, P - T conditions of c. 600~ (Fig. 3a, path 2) at 0.5 GPa were reached during the metamorphism of volcanic and sedimentary rocks (Fig. 3a, path 3), but no isotope age data are as yet available for that event.

The tectonic setting of the W L G at 1.85-1.80 Ga resembles that of the Svecofennian crust close to the L o f t a h a m m a r - L i n k 6 p i n g Deformation Zone (LLDZ in Fig. 1), where back-arc rifting occurred coevally with the formation of the juvenile O s k a r s h a m n - J 6 n k 6 p i n g Belt (Beunk & Page 2001; Mansfeld et al. 2005). However, the W L G differs from the rocks in the Baltic Shield in that all the W L G sediments appear to have been subjected to granulite-facies metamorphism and charnockitic magmatism. The Mid-Lithuanian Suture Zone. In addition to reworked rocks from the adjoining W L G and EL terranes, the M L S Z contains numerous c. 1.84 Ga porphyritic, predominantly andesitic and dacitic volcanic rocks; gabbroic, dioritic and tonalitic intrusions are prominent in this zone. Such rocks also compose the southeastern part of the W L G and probably continue into the Mazowsze Domain. Island-arc type geochemical characteristics have been

604

S. B O G D A N O V A E T A L .

reported (Rimsa et al. 2001; Motuza 2005), and these rocks could therefore represent a c. 1.84 Ga rim of crustal growth in the Polish-Lithuanian terrane, similar to the 1.83-1.82 Ga Oskarshamn-J6nk6ping Belt in southeastern Sweden (Mansfeld 1996; Mansfeld et al. 2005). The former presence of an oceanic basin between the Polish-lithuanian and Lithuanian-Belarus terranes is thus very likely. A probable northern continuation of the MLSZ in Latvia is the In~ukalns Zone, where mantle-related supracrustal rocks have been intruded by similarly juvenile granitoids (Mansfeld 2001). The complex collisional structure of the MLSZ is characterized by the presence of a variety of P - T trends in its different parts (Fig. 3b). In its eastern part (Fig. 3b, path 1), garnet-bearing felsic gneisses have undergone amphibolite-facies metamorphism at c. 500-580 ~ and 0.6 GPa (Skridlaite & Motuza 2001), and subsequent heating to 750 ~ is evidenced by migmatite-related anatectic granites yielding ages as young as c. 1.50 Ga (Skridlaite & Whitehouse, pers. comm.). These ages are similar to the emplacement ages of the late- to post-collisional AMCG rocks in the Mazury plutonic complex in northern Poland (D6rr et al. 2002; Skridlaite et al. 2003b). In the western part of the MLSZ, inclusions of felsic granulites in enderbitic and charnockitic intrusions record peak metamorphism at c. 850-900 ~ and 1.0 GPa (Fig. 3b, path 2). Together with their host rocks, these inclusions underwent a second granulite-facies event at c. 800-850 ~ and 0.8 GPa, followed by decompression and near-isobaric cooling from c. 800 to 600 ~ at about 0.6 GPa. Adjacent coherent units of garnet-, biotite- and hornblende-bearing metavolcanic gneisses were initially probably also buried to similar depths (Fig. 3b, path 3); however, their metamorphic hornblendes, yielding ages of c. 1.71-1.66Ga (Bogdanova et al. 2001b), were formed at only 550-600~ and 0.4-0.5 GPa, which suggests sites in the middle crust. In consequence, these gneisses appear to have been uplifted 10-15 km after their peak of metamorphism. By analogy with conditions in the eastern part of the MLSZ, the isobaric stage of their metamorphic evolution was probably due to the intrusion of the 1.5 Ga AMCG igneous suite. A subsequent event of shearing~ uplift and cooling at 1.48-1.42 Ga is well documented by the 0 A t / 39 Ar ages of hornblendes throughout the entire southern part of the MLSZ (Bogdanova et al. 2001b).

The L i t h u a n i a n - B e l a r u s terrane. The East Lithuanian Belt (EL) and Belarus-Podlasie Belt (BLG) (Fig. 2) making up the LithuanianBelarus terrane probably represent different depth levels of the same age unit. The EL comprises mafic metavolcanic rocks as well as sedimentary rocks, particularly metagreywackes and marbles (Motuza 2005). Although the metagreywackes contain some Archaean detrital zircons, Palaeoproterozoic zircons are dominant. S m - N d isotopic modelling indicates Palaeoproterozoic ages of deposition (Mansfeld 2001). The supracrustal rocks of the EL are marked by banded, NNE-striking, magnetic anomalies that are clearly discordant in relation to the more or less west-east-trending anomaly patterns in the WLG (Skridlaite & Motuza 2001). From the NW to the SE across the EL there are steep metamorphic gradients from high-P rocks, presumably buried beneath SE-vergent thrusts, to lower-pressure amphibolite-facies units and, ultimately, to granulites similar to those in the neighbouring BPG (see below). In the easternmost EL, greywackes and pelites were buried at depths as great as 30 km and metamorphosed at c. 680 ~ and 0.8 GPa at about 1.9 Ga (Fig. 3c, path 1). From about 1.8 Ga onwards, they underwent decompression and retrogression similarly to the rocks of the neighbouring BPG (see Fig. 3d, paths 3 and 4). The P - T evolution of the easternmost EL therefore appears to have occurred roughly simultaneously with that of the BPG. None of these rocks seem to have been affected by later thermal events.

In the southern part of the central EL, interbedded metasedimentary and metavolcanic rocks reached their metamorphic peak at c. 700 ~ and 0.6GPa, thereafter undergoing near-isobaric cooling from 580 to 480~ at 0.2GPa (Fig. 3c, path 2). Because these rocks are in contact with the 1.50 Ga Kabeliai granite (Sundblad et al. 1994), their final metamorphism was probably under the influence of this intrusion. The Belarus-Podlasie Granulite Belt (BPG) occupies most of the territory of what was previously described as the 'BelarusBaltic Granulite Belt', a concept originally proposed by Aksamentova et al. (1982). Subsequent studies of the granulites in Latvia and Estonia (H61tt~i & Klein 1991; Mansfeld 2001; Puura et al. 2004; Soesoo et al. 2004) have shown, however, that the Latvian and Estonian granulites differ substantially in age and provenance from the granulites in western Belarus and the Podlasie region of Poland. Previously, it had been suggested that the east-westtrending Polotsk-Kurzeme belt of faulting (PKZ in Fig. 1) separates the two groups of granulites from each other (Garetsky et al. 2004). In the light of current insight, however, an en echelon configuration of two different rock belts appears more probable. The BPG is mostly between 100 and 200 km wide and extends for more than 600 km in a S W - N E direction from southeastern Poland across the western part of Belarus. It is made up of several large lensoid bodies of granulites, separated from each other by fault zones and mylonites (Aksamentova & Naydenkov 1990). The BPG rocks are mostly Palaeoproterozoic granulitic orthogneisses of mafic, enderbitic and charnockitic compositions. They belong to two igneous suites. The older one has an age of c. 1.89 Ga (Claesson et al. 2001) and is calc-alkaline in composition, whereas the younger (c. 1.80 Ga), is chemically more variable, alkali-calcic and bimodal (Bogdanova et al. 1994; Bibikova et al. 1995; Taran & Bogdanova 2003). Metasedimentary gneisses and migmatites are relatively subordinate, but still occur in many drill-cores. Their S m - N d isotopic characteristics, as well as those of the intrusive rocks, indicate a minor contribution of older materials (Claesson et al. 2001). The P - T history of the BPG, adjacent to the northwestern part of the Central Belarus Suture Zone (CBSZ), suggests a sequence of tectonothermal events resembling that found in the Okolovo rocks of the latter (see below). Early prograde metamorphism, recognized in the metasedimentary rocks, occurred in amphibolite facies at conditions varying from 530-550 ~ and 0.3-0.4 GPa to 650-670 ~ and 0.6-0.7 GPa. It was associated with deformation and tonalite-trondhjemite-granodiorite (TTG)-type magmatism, apparently related to subduction to depths of c. 30 km between 1.89 and 1.87 Ga (Fig. 3d, paths 1 and 2). Granulite-facies metamorphism at 1.8 Ga (750 ~ and c. 0.8 GPa) was superimposed onto these amphibolite-facies rocks, being caused by the input of numerous mafic and monzonitic to charnockitic intrusions, following late to post-collisional extension of the crust (Bogdanova et al. 1994; Taran & Bogdanova 2003). Rapid tectonic uplift to upper crustal levels and cooling took place between 1.78 and 1.74 Ga. Between 1.71 and 1.66 Ga, there was continued extension associated with retrograde metamorphism, strong deformation, and transtensional rearrangement of the tectonic grain of the BPG into lensoid and anastomosing patterns (Bogdanova et al. 1994). The last major tectonothermal event in the BPG involved reactivation along previously formed zones of faulting and granitic intrusions at 1.6-1.5 Ga (Bogdanova et al. 2001b). This evolution is indicated by the stepwise configuration of some of the P - T paths (e.g. path 4 in Fig. 3d). Comparison of the P - T paths from the different parts of the BPG suggests that the rock units in its interior (Fig. 3d, paths 1 and 2) were uplifted and cooled more rapidly than those situated along the margins. Two instances of the latter are the Rudma rocks along the boundary towards the Okolovo terrane (Fig. 3e, path 2) and the rock units in the westernmost part of the BPG, along its boundary with the EL (Fig. 3d, paths 3 and 4).

EUROBRIDGE Altogether, the available data suggest that the BPG and EL were both formed at c. 1.9 Ga, but in widely different tectonic settings. Whereas the BPG represents a mature island arc, the formation of the EL most probably occurred in a back-arc environment. The Central Belarus Suture Zone (CBSZ)

The CBSZ extends nearly 600 km along the northwestern margin of Sarmatia. Although it coincides geographically with the traditional 'Central Belarus Belt (CB)', EUROBRIDGE studies have shown that it is heterogeneous in structure and comprises several tectonically different rock complexes separated, in particular, by the major, suture-like, Minsk Fault (see below). To the NW of this fault is the Okolovo terrane; to the SE are located the Vitebsk Domain and the southeasternmost part of the former CB, referred to in the following as the Borisov-Ivanovo Belt (VG and B-I, respectively, in Figs 1 and 2).

605

c. 610 ~ and 0.6-0.7 GPa. Subsequently, but still before the intrusion of the granites, some uplift occurred, which caused retrogression (Fig. 3f). This records the burial of metapelites to depths close to 25 km and their subsequent uplift. A later event of contact metamorphism, which was associated with the 1.96 Ga granites, involved heating to c. 670 ~ Substantial cooling and decompression followed the initial stages of this magmatism, possibly still coinciding with the upward migration of the granitic melts. For the final stage of the metamorphic evolution, the recorded P - T - t path indicates cooling from c. 530 to 420 ~ at pressures dropping from 0.4 to 0.2 GPa. The metamorphic and igneous histories of the rocks in the VG agree in many respects with those recorded from the adjacent OMB. Within the latter, some supracrustal rocks were metamorphosed in the high-T range of high amphibolite facies at temperatures up to 700 ~ and pressures of 0.3-0.4 GPa (Khvorova et al. 1982). This event may have been caused by elevated heat flow at the active margin of Sarmatia and recurrent igneous activity in the OMB.

The Okolovo terrane. The c. 2.0 Ga Okolovo terrane (Bibikova et al. 1995; Claesson et al. 2001) forms a c. 10km thick,

WNW-dipping, tectonically delimited complex (Aksamentova et al. 1994). Along its borders towards the overlying BPG in the NW, the rocks have been metamorphosed in the granulite facies. At the base of the Okolovo terrane in the area immediately adjacent to the Minsk Fault, geophysical evidence suggests that there may exist a separate lens of granulites, isolated structurally from the rest of the Okolovo terrane. The Okolovo terrane is built up of metamorphosed komatiitic and tholeiitic, basalts, andesites, dacites and rhyolites of oceanic-arc affinities. Intercalated metasedimentary rocks include characteristic black shales and ferruginous as well as siliceous volcanogenic deposits. Although these igneous rocks are present throughout the Okolovo terrane, they are particularly abundant in its northernmost part. The metavolcanic rocks of the Okolovo terrane (Fig. 3e, path 1) were exposed to prograde metamorphism with peak temperatures ranging from 640 to 725 ~ at depths of 30-35 km. These conditions were attained at c. 1.9 Ga, concomitantly with the emplacement of TTG-type juvenile melts (Claesson et al. 2001), most probably related to subduction-triggered metamorphism (Taran & Bogdanova 2001). The subsequent rapid, tectonic uplift of some crustal blocks to depths of only 12-15 km was followed by extension and high-T metamorphism associated with the inferred intrusion of mafic and monzonitic-charnockitic melts at c. 1.8 Ga. This is reflected by the deviation of some P - T paths into the field of higher temperatures and by their nearly isobaric trends during metamorphism in amphibolite facies. Retrogression lasted from c. 1.77 to 1.67 Ga (Fig. 3e, paths 2 and 3). The close relationships of the Okolovo and BPG (i.e. Lithuanian-Belarus) terranes are of great importance, militating against a previous view of the Okolovo terrane as part of the Sarmatian plate (Bogdanova et al. 2001b). All these terranes shared similar geodynamic evolutions after 1.9 Ga. This period, in contrast, is not characteristic of the Palaeoproterozoic history of Sarmatia (see below). The Vitebsk Domain (VG) and the Borisov-Ivanovo (B-I) belt. The VG and the B - I belts, lying to the SE of the Minsk Fault (Fig. 2), are very different in age and evolution from the Okolovo terrane. These two units are characterized by tectonic settings and development histories very similar to those of the Osnitsk-Mikashevichi Igneous Belt (OMB), farther within Sarmatia (see below); they are interpreted as probable equivalents of the OMB, which developed at deeper crustal levels. In both units, juvenile metasediments and andesitic-dacitic metavolcanic rocks were formed at c. 1.98 Ga (Bibikova et al. 1995). Regarding metamorphism, pelitic xenoliths enclosed in 1.96 Ga high-temperature granites in the B - I and the VG document an early episode of regional metamorphism at P - T conditions of

West Sarmatian terranes

The Sarmatian crustal segment, which is exposed in the Ukrainian Shield and the Voronezh Massif, is built up of several Archaean proto-cratonic terranes and intervening belts of Palaeoproterozoic rocks. The continental crust of the Archaean units was formed between c. 3.7 and 2.7 Ga, whereas that in the Palaeoproterozoic belts was accreted to the Archaean cores mainly between 2.2 and 2.1 Ga, and again between 2.0 and 1.9 Ga. The EUROBRIDGE'96 and '97 profiles traverse three major lithotectonic units: the Palaeoproterozoic OMB, the similarly Palaeoproterozoic, but somewhat older Teterev Belt in the Volyn Domain (also known as the North-Western Domain in the Ukraine), and the northern part of the mostly Archaean Podolian Domain (Figs 1 and 2). The Palaeoproterozoic belts of Sarmatia differ in age from those of Fennoscandia by featuring 2.22.1 Ga crust, which is seemingly absent in northeastern Europe (see Claesson et al. 2006). The Osnitsk-Mikashevichi Igneous Belt (OMB). This 150-200 kaaa wide belt occupies the northwestern margin of the Sarmatian crustal segment. From the Trans-European Suture Zone along the southwestern limit of the EEC, it extends northeastwards to Moscow (Bogdanova et al. 2004a); that is, for a distance of more than 1000 km. It is buried almost entirely beneath Phanerozoic sedimentary rocks (Aksamentova 2002), cropping out only in an area in the northwestern part of the Ukrainian Shield and in a few horst-type elevations within the Pripyat Trough (the western continuation of the Devonian Dniepr-Donets Aulacogen). Nevertheless, the OMB can be traced fairly confidently in the magnetic fields, where it is marked by numerous rounded, mostly positive anomalies associated with large batholiths of granodiorites and granites, and intrusions of diorites and gabbros. The most reliable ages of these emplacements range from 1.98 to 1.95 Ga. Between and inside the plutonic massifs are 'septa' and inclusions of mafic and felsic metavolcanic and hypabyssal rocks, for which an age of 2.02 Ga has been obtained (Skobelev 1987; Shcherbak & Ponomarenko 2000). In addition, there also exists a later, c. 1.80-1.75 Ga, generation of sub-alkaline plutonic and volcanic rocks, associated with some sedimentary complexes deposited in minor basins. These are nearly coeval with the AMCG Korosten Pluton in the Volyn Domain and other similar intrusions in the Ukrainian Shield. In general, the lithologies and structures of the OMB suggest formation in an Andino-type active continental-margin environment, created by c. 2.0-1.95 Ga subduction of oceanic crust beneath the edge of Sarmatia. A period of apparent quiescence then followed, for which age and other information are almost

606

S. BOGDANOVAETAL.

totally lacking. This period lasted until the beginning of collision between Sarmatia and Fennoscandia at c. 1.84-1.82 Ga. Kinematic analysis of faults within and in the neighbourhood of the OMB (Gintov 2004) indicates their formation or activation simultaneously with the emplacement of the Korosten pluton between c. 1.80 and 1.74 Ga (Amelin et al. 1994). During this period, the present structural patterns of the crust in northwestern Sarmatia were generally established (Bogdanova et al. 2004b). The Volyn Domain (VD). The VD consists mainly of Palaeoproterozoic rocks (Figs 1 and 2). Dominant granitoids were emplaced at 2.06-2.02 Ga, concomitantly with high-T amphibolite facies metamorphism (Khvorova et al. 1982) and migmatization of older (2.2 Ga), strongly deformed, sedimentary and volcanic rocks of the Teterev Belt and its Belaya Tserkov continuation into the adjacent Ros'-Tikich Domain (Stepanyuk et al. 1998; Claesson et al. 2000; Claesson et al. 2006). Palaeogeographical reconstructions of the stages of formation of these supracrustal units (Lazko et al. 1975; Ryabenko 1993; Shcherbakov 2005) indicate that at c. 2.2 Ga the entire northwestern part of the Ukrainian Shield was characterized by intense igneous activity and sedimentation in coastal and marine settings. The sedimentary successions are terrigenous and include tuffitic and graphitic units and various types of turbidites. Basaltic, andesitic, dacitic and rhyolitic volcanism also occurred. Altogether, this suggests mature island-arc conditions and possibly derivation of detritus from Archaean sources. However, strong deformation, involving SE-vergent thrusting and associated transcurrent faulting during the formarion of the OMB at 1.98-1.95 Ga, prevents more detailed palaeotectonic interpretation. Numerous intrusions of mafic to monzonitic sub-alkaline rocks were emplaced into zones of extension at 2.02-1.98 Ga; that is, almost simultaneously with the early stages of igneous activity in the OMB. A most remarkable feature of the Volyn Domain is the giant AMCG-type Korosten pluton (Figs 1 and 2) that was formed between 1.80 and 1.74 Ga by the successive emplacement of many pulses of basic and acidic melts (Zinchenko et al. 1990; Bukharev 1992; Amelin et al. 1994; Verkhogliad 1995). Traditionally, the Korosten pluton has been considered as anorogenic, formed by mantle underplating. However, it has recently been related to the late and post-collisional tectonic regimes prevailing farther NW, in the CBSZ (Bogdanova et al. 2004b). It is also notable that, according to recent Sm-Nd isotopic data, none of the Korosten igneous rocks, with eNd~T) values of - 0 . 8 to -1.8, were derived from a depleted mantle (Dovbush et al. 2000). This supports the idea that the Korosten magmas originated by remelting of the OMB lower crust (Bogdanova et al. 2006). The nearly 100 km wide Berdichev Boundary Zone (BZ in Fig. 2) separates the Teterev Belt from the Archaean Podolian Domain farther south. Originally, this zone may have belonged to a Palaeoproterozoic collisional boundary belt between the Podolian Domain and crustal units to the north. It is mostly made up of high-T, upper crustal, garnet- and cordierite-bearing S-type granites, which, towards the north, pass into the 2.062.02 Ga granitoids of the Volyn Domain (Shcherbakov 2005). In the southern BZ, Neoarchaean charnockites intrude older, Archaean granulites (Stepanyuk et al. 1998). P - T conditions within the Berdichev Zone, in its deepest crustal sections in the south, reached c. 850 ~ at c. 8 GPa (Kurepin 2003). The Podolian Domain (PD). The PD is one of the principal Archaean units of the Sarmatian crustal segment (Claesson et al. 2006). Archaean as well as Palaeoproterozoic granulites prevail. Major zones of faulting subdivide it into the Vinnitsa region in the north and the Gayvoron region in the south. The boundary in the east, towards the Palaeoproterozoic Kirovograd Domain, is defined by the Golovanevsk Suture Zone, which features

numerous, nearly 3.0 Ga ultramafic and mafic rocks (Gornostayev et al. 2004). EUROBRIDGE' 97 seismic profiling particularly concerned the northern part of the PD, where granulites of sedimentary and mafic volcanogenic derivation are the dominant rocks. These were complicated structurally by doming associated with the formation of c. 2.1-1.9 Ga Palaeoproterozoic charnockites, granites and migmatites. Archaean charnockitic intrusions are widespread in the SE. The supracrustal granulitic rocks of the Podolian Domain belong to two associations. The oldest granulites are mafic, ultramafic and intermediate, and have been intruded by enderbites (Shcherbak et al. 2005). Sm-Nd data and ion-probe zircon studies suggest ages of the crust around 3.7 Ga and an event of granulite-facies metamorphism at c. 2.8 Ga (Claesson et al. 2006). Kalyaev et al. (1984) have proposed that the oldest mafic rocks could represent early Archaean oceanic crust. The apparently younger metasedimentary rocks include ferruginous quartzites as well as highly aluminous, partly graphitic schists, and carbonates, all metamorphosed in granulite facies, presumably during the Palaeoproterozoic. This suggests that the early Proterozoic tectonothermal evolution was fairly uniform throughout the entire western part of the Ukrainian Shield. The now-exposed rocks in western Sarmatia may represent a 10-15 km thick slice of Archaean- to Palaeoproterozoic crust. The granulites in the southern part of the PD were exhumed from depths of more than 35 km (Kurepin 2003; Shcherbakov 2O05).

Major stages of the Proterozoic crustal evolution A c c r e t i o n a r y stages

The pre-metamorphic composition of the lithotectonic units and their P - T evolution, as described above, suggest that, during the time before the final assembly of the EEC, the terranes along the northwestern margin of Sarmatia evolved differently from those related to Fennoscandia. The Palaeoproterozoic metasediments in the southeastern, that is Sarmatia-related part of the main suture zone, the CBSZ, were metamorphosed during subduction at c. 1.96 Ga; neither this part of the suture zone, nor northwestern Sarmatia in general, has yielded consistent evidence of tectonothermal activity during the period between c. 1.95 and 1.80Ga (Taran & Bogdanova 2001; Shcherbak et al. 2003). Thus, the metasedimentary rocks in the Vitebsk Domain (VG) were deformed and metamorphosed at c. 1.96 Ga in the middle crust beneath the Osnitsk-Mikashevichi Belt (OMB) along the Sarmatian continental margin. The prograde part of their P - T path (Fig. 3f) can be attributed to southeastwards (in present-day coordinates) subduction, which took place simultaneously with the OMB magmatism at some time after 1.98 Ga, the earliest age defined by the youngest detrital zircons in the VG (Claesson et al. 2001). In the supracrustal rocks of the Teteriv Belt farther inside Sarmatia, however, an early stage of high-Tmetamorphism, intense migmatization and anatectic remelting occurred already at c. 2.1-2.0 Ga. This major tectonothermal event in western Sarmatia locally also influenced the evolution of the Archaean Podolian Domain, causing repeated granulite metamorphism and associated high-T magmatism. In contrast, the Fennoscandia-related terranes, farther NW, mostly developed between 1.9 and 1.8 Ga, simultaneously with the formation of the Svecofennian Domain proper and the oldest, 1.83-1.82 Ga rocks in southeastern Sweden. An exception is provided by the Okolovo terrane within the CBSZ, which contains c. 2.0 Ga oceanic-arc supracrustal rocks, which have not been identified elsewhere in the Svecofennian region. As seen from their P - T - t path (Fig. 3e), the Okolovo supracrustal rocks were subducted northwestwards (present-day coordinates) beneath the

EUROBRIDGE

Lithuanian-Belarus terrane and, after 1.9 Ga, evolved together with the Belarus-Podlasie Belt (BPG) and East Lithuanian Belt (EL) of this tectonic unit. The prograde parts of the metamorphic paths of the Okolovo terrane and the BPG, and probably also some of the supracrustal rocks in the EL, imply a period of burial and deformation between 1.90 and 1.87 Ga associated with TTG magmatism. These data contradict a previous tectonic interpretation of the EL supracrustal rocks as an accretionary prism (Motuza 2005). If the latter was correct, subduction should have occurred toward the present SE; that is, in a direction opposite to that suggested by the evidence presented above. In the West Lithuanian Granulite Domain (WLG) and the part of the Mid-Lithuanian Suture Zone (MLSZ) closest to the PolishLithuanian terrane, the earliest orogenic events are difficult to discern. These lithotectonic complexes are characterized particularly by a stage of accretionary tectonics in the Baltic-Belarus region between 1.85 and 1.82 Ga. During that subduction and the formation of volcanic island arcs and back-arcs, TTG magmatism and burial of supracrustal rocks at great depths occurred. Their evolution is described by the preserved prograde parts of the P - T - t paths of some rocks of the MLSZ rocks (Skridlaite et al. 2003a). Comparison of the tectonothermal events during the period between c. 2.2 and 1.84 Ga (Fig. 4) thus suggests the following evolution. (1) Terranes related to Sarmatia and Fennoscandia, along the EUROBRIDGE transect, evolved separately until these two crustal segments docked with each other between 1.84 and 1.80 Ga. (2) At c. 1.90-1.87 Ga, the Lithuanian-Belarus terrane and the 2.0 Ga Okolovo oceanic arc were assembled to form a single unit, possibly a microcontinent. This process took place simultaneously with the formation of the classical Svecofennian Domain in the Baltic Shield. (3) The development of the Polish-Lithuanian terrane was nearly coeval and possibly even closely connected spatially with the accretion of new crust in southeastern Sweden. In and close to the O s k a r s h a m n - J r n k r p i n g Belt, subduction-related magmatism, the formation of an island arc and a back-arc (Mansfeld et al. 2005), and nearly coeval metamorphism and partial melting of meta-supracrustal gneisses (Beunk & Page 2001) occurred at c. 1.83-1.78 Ga.

607

is 'thick-skinned' and thus characteristically collisional, but was complicated by subsequent extension (see below). The giant Korosten Pluton in the Volyn Domain (VD) and the presence of a thick high-velocity layer in the lower crust beneath the CBSZ (see below) both suggest that high-T granulite-facies metamorphism, anatectic melting and near-isobaric retrograde cooling of most of the studied rocks were related to late and post-collisional tectonism between 1.80 and 1.75 Ga. The newly grown c. 1.80 Ga metamorphic zircons and monazites in the granulites of the WLG and MLSZ indicate that a coeval tectonothermal event also affected the Fennoscandia-related terranes farther NW. This was associated with 1.82-1.81 Ga charnockitic intrusions (Skridlaite & Motuza 2001). However, it is still conjectural whether all the terranes in the Baltic-Belarus region had been assembled by 1.82-1.80 Ga. Possibly an oceanic 'gap' may still have existed at that time between the Polish-Lithuanian and the Lithuanian-Belarus terranes (Skridlaite et al. 2003a). Major NW-trending transpressive shearing at 1.83-1.78 Ga also occurred in the crust of central and southeastern Sweden (Beunk & Page 2001; H6gdahl et al. 2004), coinciding with an early event of magmatism in the Transscandinavian Igneous Belt (TIB). In that region, subduction-related magmatism and attendant metamorphism along the present southwestern and western margins of the Svecofennian orogen may have occurred (Andersson et al. 2004). The subsequent, 1.71-1.67 Ga tectonothermal event (Fig. 4) may have been due either to the terminal collision and amalgamation of the Fennoscandian, Volgo-Uralian and Sarmatian crustal segments to form the East European Craton, or to within-craton deformation and reactivation of the pre-existing fault systems caused by continuing convergence (Bogdanova et al. 2001b). At roughly the same time (c. 1.73-1.67 Ga), a large north-southstriking belt of juvenile crust was formed as a result of eastward subduction in southwestern Sweden and central Norway (Andersson et al. 2004; Gorbatschev 2004). Notably, this event did not significantly affect the Svecofennian Domain in Finland, where major deformation and metamorphism ceased at 1.79-1.77 Ga (Lahtinen et al. 2005).

Mesoproterozoic intra-continental deformation and magmatism

Collision and post-collisional stages o f extension

In all the 'Fennoscandian' terranes in the EUROBRIDGE area and in southeastern Sweden there are numerous steep, e a s t west-trending zones of shearing, with associated c. 1.6-1.45 Ga AMCG and A-type granitoid intrusions (Bogdanova et al. 2001b; Skridlaite et al. 2003b; Ce~ys 2004). These may have been a far-field effect of continuing or renewed accretionary

Metamorphism in the Baltic-Belarus region largely took place at c. 1.8 Ga and was broadly coeval with widespread mafic and monzonitic magmatism, represented particularly well in the BPG, the EL and the Okolovo terranes. From geophysical data, the structure of the upper lithosphere in all these tectonic units

2

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FENNOSCANDTA: Baltic Shield (SE Sweden) Polish-Lithuanian terrane (WLG+Cn+MD+DD+MLSZ) Lithuanian-Belarus terrane (EL + BPG) Okolovo terrane WESTERN SARMATIA

Time, Ma

Fig. 4. Timing of Proterozoic crust-forming processes in the southwestern part of the East European Craton. The principal orogenic events in each of the terranes are shown in black. The cross-hatched lines (1, 2, 3 and 4?) indicate 'stitching', simultaneous events during the assemblies of the terranes and their shared evolution as discussed in the text. Line 2 shows the approximate time of the major collision of Fennoscandia and Sarmatia. The AMCG- and A-type granitoid igneous events (graded white-grey fill) are interpreted as indicators of late or post-collisional tectonic regimes. (See the text for the geochronological references.)

608

S. BOGDANOVA ETAL.

orogeny in the westernmost Baltic Shield (.~h~i11 et al. 2000), but they may also have been related to collision with other plates, for example, Amazonia (Bogdanova 2001). Somewhat unexpectedly, the P - T - t histories of metamorphism (Fig. 3) demonstrate the strong effects of this Mesoproterozoic tectonothermal activity. Some of the 'steps' and indications of near-isobaric cooling, which complicate the P - T paths, suggest a relationship between this younger metamorphism and the emplacement of AMCG-type magmas, such as the c. 1.6 Ga Riga pluton and the 1.54-1.50 Ga Mazury complex (see Figs 3 a - c and 4). The outer rims of zircons in the metamorphic lithologies have been dated by ion probe to around 1.50-1.45 Ga (Skridlaite et al. 2004). The AMCG magmatism at this time also triggered growth of amphibole in mylonites along the major shear zones and resetting of the Ar isotope systems of older amphiboles in various rocks (Bogdanova et al. 2001b).

Seismic and density images of the crust and upper mantle in the southwestern EEC

The seismic models of the southwestern part of the EEC presented here are based on EUROPROBE, EUROBRIDGE and POLONAISE deep seismic sounding (DSS) profiles that were acquired in the 1990s (Giese 1998, Guterch et al. 1998, 1999; EUROBRIDGE Seismic Working Group 1999; Czuba et al. 2001, 2002; Yliniemi et al. 2001; Grad et al. 2003b; Thybo et al. 2003). Data from FENNOLORA and the Coast Profile (Lund et al. 2001) have also been employed (Figs 1, 2 and 5). The study area has 12~

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Seismic profiles: acquisition of data and the observed wave field. The EUROBRIDGE and POLONAISE'97 profiles, altogether c. 2300 km in length, were carried out using modern digital seismic recorders spaced 1.2-4.0 km apart along the profiles. Shot points with charges of 300-1000 kg of TNT were located at intervals of 30-40 km. In the Coast Profile project, a shipborne airgun array was used to generate the seismic waves. The wave field recorded in the southwestern part of the EEC is, in general, of very high quality (Figs 6 and 7). Because of the commonly thin sedimentary cover (<2000 m), the refracted waves diving into the crust (Pg) produced clear first arrivals with apparent velocities ofc. 6 - 7 km s -1. Strong reflected waves from the Moho boundary (PmP), starting from the offsets at 80-120 km, and substantial differentiation of the arrival times, exceeding 2 s for Pg, ProP and Pn phases, were observed. This reflects differentiation of the crust- and upper mantle structure. Models of the crust and uppermost mantle. The seismic data for all

E U R O B R I D G E and P O L O N A I S E seismic profiling

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Fig. 5. Location of the refraction and wide-angle reflection DSS profiles in the area of southwestern margin of the East European Craton. Open stars indicate the shot points of the EUROBRIDGE (EB'94, EB'95, EB'96 and EB'97) and POLONAISE'97 (P5 and northern part of P4) profiles. The numbered stars refer to the locations of shot points for which examples of the record sections are shown in Figures 6, 7 and 9. The bold dashed line shows the southwestern edge of the craton; the fine dashed line shows the Fennoscandia and Sarmatia suture. TESZ, Trans-European Suture Zone.

the POLONAISE'97 and EUROBRIDGE profiles were modelled using 2D tomographic and ray-tracing techniques (Cerven~ & P~enfil'k 1983). P-wave velocity models of the crust and uppermost mantle along the profiles in the southwestern part of the EEC are shown in Figures 8 and 9. The S-wave velocity model and the Vp/Vs ratio distribution for the EUROBRIDGE'97 profile are shown in Figure 9. For the tectonic units related to Fennoscandia and Sarmatia, respectively, the particulars of the structures based on the P- and S-wave velocities are summarized in Tables 1 and 2. In general, the P- and S-wave velocity models both show similarity to the results previously obtained from Scandinavia (Grad & Luosto 1987, 1994; BABEL 1993; Guggisberg et al. 1991). The crust and upper mantle of the southwestern part of the EEC can be characterized as follows. (1) The thickest Phanerozoic cover deposits in the various parts of the Pripyat Trough and along the Trans-European Suture Zone correspond to seismically defined layers with P-wave velocities between 2 and 4 km s-]. (2) The crystalline crust in the surveyed area can generally be divided into three parts, with P-wave velocities of 6.1-6.4, 6.56.8 and 6.9-7.2 km s -1 for the upper, middle and lower crust, respectively. Relatively low velocities of c. 5.7 km s -1 in the uppermost crystalline basement were found locally in the MD and WLG, and in the BPG (Figs 1 and 2). The upper crystalline crust is commonly inhomogeneous, with low-velocity zones and high-velocity bodies alternating along some parts of the profiles. Normally, the rather weak, low-velocity zones reach c. 5 km in thickness and have velocity contrasts of 0.1-0.2 km s-a; mostly, they occur at depths between 4 and 15 km. Low-velocity layers in the upper crust have been found in the TIB, WLG, and in the VD and PD of Sarmatia, whereas the region of the Mazury intrusions (Mz), the CBSZ, and the VD with the Korosten Pluton (KP) feature high-velocity bodies in the upper crust. The middle crust has P-wave velocities of 6.5-6.8 km s -a, which increase to between 6.9 and 7.2 km s -a in the lower crust. The lowermost crust is marked by high P-wave velocities, reaching a maximum of 7.5 km s-a in the part of the Volyn Domain underlying the KP in Sarmatia. Characteristically, that region lacks the high reflectivity in the lower crust that is otherwise common in the VD. The crystalline crust in the EUROBRIDGE region mostly has low velocity gradients and small velocity contrasts at the seismic boundaries. Only in some places, such as, for instance, in the CBSZ, the OMB, and parts of the VD and the PD, has high reflectivity been observed. (3) The average values of the Vp/Vs ratios in the crystalline crust are 1.69, 1.70 and 1.76 in its upper, middle and lower parts, respectively. It follows that the S-wave velocities in the

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upper and middle crust are relatively high in comparison with the P-wave velocities, whereas in the lower crust V~ is relatively low. This may explain the strong SmS reflections from the Moho, seen, for instance, in the central part of the EUROBRIDGE'97 profile (Fig. 9). (4) High-velocity plutonic bodies in the upper crust along the EUROBRIDGE transect coincide with geologically well-known intrusions such as the Mazury and Korosten plutons of rapakivigranitic and gabbro-anorthositic rocks (Figs 1 and 2). In the Mazury igneous complex, a high-velocity body with P-wave velocities between 6.4 and 6.7 km s-x coincides with a gabbroanorthosite massif (Fig. 8). The Vp/V~ ratio in this body is estimated to be 1.75. In the VD, the Korosten Pluton is imaged as coinciding with a high-velocity anomaly of 6.35-6.7 km s -1, extending to depths of at least 11 km. At still greater depths, this anomaly appears to link up with another high-velocity anomaly in the lower crust where the Vp/Vs ratios are as high as 1.77-1.79. (5) Most of the crust in the EUROBRIDGE study area has a thickness between 40 and 50 km. Moho depths of c. 55 km have been found in the PD, whereas the shallowest Moho (c. 30 km) is that in the VD, beneath the Korosten Pluton. Mantle P-wave velocities immediately beneath the Moho are generally 8.28.35 km s - i ; lower velocities (8.0-8.15 km s -1) have been found only in the marginal zones of the EEC such as, for instance, in the Dobrzyfi Domain in Poland (Figs 1 and 2). The average Vp/V~ ratio for the uppermost mantle, determined from the Pn and Sn waves, is 1.75, with Fennoscandia having a lower average of c. 1.72 and Sarmatia a somewhat higher value of c. 1.80 (Fig. 9). (6) The uppermost mantle features numerous subhorizontal reflectors beneath both the Baltic Shield and the East

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European Platform (Grad 1992; Sroda & POLONAISE Profile P3 Working Group 1999; Czuba et al. 2001; Lund et al. 2001). These reflectors often are c. 10-15 km below the Moho. A major, SSW-dipping reflector has been recognized in the uppermost mantle beneath the EUROBRIDGE'97 profile (Figs 8 and 9). It extends from the Moho to depths of c. 75 km (Thybo et al. 2003). This reflector coincides with a subhorizontal reflector on the EUROBRIDGE'96 profile, close to its crossing point with the EUROBRIDGE'97 profile, in Sarmatia. We note, however, that the quality of data on the EUROBRIDGE'97 profile is higher than that on the EUROBRIDGE'96 profile, and the error of the Moho reflection positions at the crossing point is c. 10 km. The formation and age of this reflector are discussed below.

Gravity-seismic modelling of rock compositions along the EUROBRIDGE profiles The lateral and vertical variations of P-wave velocity in the crust and upper mantle along the EUROBRIDGE transect agree, in general, with the global compilation of data on seismic velocities and composition of continental crust presented by Christensen & Mooney (1995). Seismic velocities and densities along all the EUROBRIDGE profiles, which increase gradually with depth, can be explained in terms of rock compositions that change from felsic to mafic, and also with increasing grades of metamorphism. The P-wave velocities compiled in Table 1 indicate that the silica content in the rocks of the EUROBRIDGE area decreases from over 70% SiO2 in the upper crust to about 47%

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at the Moho boundary (Kern et al. 1993; Christensen & Mooney 1995). However, the interpretation of P-wave velocities in terms of rock compositions is non-unique, as a given velocity value can usually characterize a number of different lithologies (Christensen & Mooney 1995). Additional constraints on rock composition can be obtained from Vp/Vs ratios (or Poisson's ratios); these are generally sensitive to the chemical composition of the rocks (Kern et aL 1993; Sobolev & Babeyko 1994). However, correlation of Poisson's ratios with SiO2 contents is valid only for rocks with 5 5 - 7 5 % SIO2, that is for felsic and intermediate compositions (Christensen 1996). For mafic rocks (SIO2 < 55%), the Poisson's ratios vary greatly and there is no correlation with SiO2 content. In consequence, estimation of the compositions of the lower crust from the Vp/Vs ratio is not reliable. Because of the strong dependence of rock density on composition (Birch 1961), the composition of the crust can be inferred from the former. Thus gravity-field modelling and inversion can provide additional information for the interpretation of seismic results in terms of rock compositions. Another important advantage of gravity data is that they are more sensitive to vertical and subvertical crustal boundaries than DSS data. In conventional techniques of gravity data interpretation, the density models are calculated from seismic velocity models using density-velocity relationships estimated under laboratory conditions (Woollard 1959; Birch 1961; Ludvig et al. 1970; Christensen & Mooney 1995). For the EUROBRIDGE profiles, however, the integrated velocity-density models were calculated using a technique of joint interpretation of seismic and gravity data that is based on a

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Fig. 7. P-wave record sections (for location see Fig. 5). The large differentiation of the arrival times (exceeding2 s) of the Pg, PmP and Pn phases, which reflects differentiationof the structures in the crust and uppermost mantle, should be noted.

stochastic, quasi-linear relationship between density and seismic velocity, which is not assumed a priori, but rather obtained from the inversion of gravity data (Kozlovskaya et aL 2002). This approach makes it possible to obtain density models of the crust and also analyse the relationships between density and seismic velocities in the major crustal units traversed by the EUROBRIDGE profiles. The advantage of analysing the density-velocity relationship in detail is illustrated by Figure 10, which shows the relationships between density, Vp, Vs and V p / V s estimated by Monte-Carlo simulation for several types of rocks from the Fennoscandian and Ukrainian Shields. These represent major crustal lithologies and contain various proportions of the main rock-forming minerals. The rocks considered and their modal mineral compositions have been described in a previous paper (Kozlovskaya et al. 2004). As can be seen from these data, both the igneous and the metamorphic lower crustal rocks have similar values of Vp/l,7~, whereas their densities are significantly different. This indicates that the density-velocity relationships (and, in particular, the density-Vp relationship) depend on the varying mineral compositions of the rocks (Fig. 10c). In general, therefore, the density-velocity relationships of actual rocks with various chemical and mineral compositions will differ from each other and deviate also from those of anhydrous igneous rocks with averaged chemical compositions (Sobolev & Babeyko 1994) that are used as the basis for standard curves of reference and shown by star signs in Figure 10c. Thus, for example, rocks with high contents of plagioclase and low contents of amphibole have higher Vp values for any given density value

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200

250

Belarus-Podlasie Granulite Belt

I

Central Belarus SZ

I

300

350

400

450

Osnitsk-Mikashevichi Igneous Belt

!

500

I

550

600

650

700

Distance [ km ]

SE

Mazury BBR

I

TIB

I

KP

Mazowsze

Complex

Svecofennian Domain

i

Domain

! KAM o

10-4 E

- 10

20-4

- 20

30 -4

7"30

40-4

- 40

"i

50:]i60 0 SW

50

100

150

200

Osnitsk-Mikashevichi Igneous Belt I

250

300 350 Distance [km ]

Volyn Domain I

400

450

500

550

NE

I

Korosten Pfuton

Dobrzyfi , Domain

i

0 N

50

100

150

200

250 300 Distance [ km ]

350

50 100 150 Distance [ km ]

SE

Podolian Domain TESZ

Pripyat Trough ~,

"5~

' ' ~"' ! ~ ' ~ I ~ ' ~ ~ 1'~ J..... 60

0 NW

6oo

m~ 1 8

400

450

Ciechan6w Zone

Mazury

EL

Complex

500

Fig. 8. Crustal and uppermost mantle models for the E U R O B R I D G E transect (EB'94, EB'95, E B ' 9 6 and EB'97 profiles), the Coast profile and the POLONAISE'97 profiles P4 (northem part) and P5. P-wave velocities are given in km s -1. 1, sedimentary cover; 2, upper crust; 3, high-velocity zone in the upper crust; 4, low-velocity zone in the upper crust; 5, middle crust; 6, lower crust; 7, high-velocity lower crust; 8, uppermost mantle; 9, elements of seismic boundaries obtained from reflected and refracted waves; 10, zones of high reflectivity in the uppermost mantle; 11, Moho boundary; 12, mantle reflector or zone of high-velocity gradient; 13, zone of anomalously high velocity, probably associated with the Korosten Pluton; 14, zones of rapid lateral change of the seismic structure, probably indicating contact zones of crustal blocks. V.E., vertical exaggeration.

than the rocks of the standard reference curves. The opposite, of course, applies to rocks rich in amphibole, but poor in plagioclase. In another instance, mafic lower crustal rocks metamorphosed in granulite facies have densities around or below 3.0 g c m - 3 and their density-velocity relationships follow closely the standard curves of reference. With the appearance of substantial amounts

of garnet in high-pressure mafic granulites and eclogites, however, the densities increase to above 3.0 g cm -3 and the velocities to over 7.0 km s - I . As can be seen from Figure 9c, the density-velocity relationships for these rocks differ significantly from those of other mafic lower crustal rocks, whereas the Vp/Vs ratios are all similar.

612

S. BOGDANOVA ET AL. i,,,~,,,:,,~,i ~......

i r-~UR0-BRIDGE'97, sP06, Z-comp]

It

t~llll ', HRll~!

(,t)

i I~ i~ I~

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f

i

lOO Offset [km ]

0

.,

200

Volyn Domain

~ ---'1 II

.....

II

300

Podolian

Domain

Korosten Pluton I i

Vs [km/s] 4,50 4.40 4.30 4.20 4.10 4.00 3.90 3.80 3.70 3.60 3.50 3.40 3.30 3.20 3.10 3.00

/

25

50

IV.E.

3:1l

S - w a v e velocity model

75

0

25

._25 VpNs E 2.00 195 190 1.85 1.80 1.75 1.70 ~ 165

II

-

' t-" ~'50 a

75

t 0

50

....

v0Nsrato strbutonp 75

50

100

150

200 250 300 Distance [ km ]

Fig, 9. Two-dimensional seismic models along the EUROBRIDGE'97 profile developedby forward ray-tracing. Top: example of seismic record section for SP06 with P and S waves (for location see Fig. 5). Middle: S-wave velocity model. The parts of the discontinuities that have been constrained by reflected and/or refracted S waves are marked by bold lines. Arrows show the positions of the shot points. Bottom: Vp/Vs ratio distribution. The parts of the discontinuities that have been constrained by reflected and/or refracted P or S waves are marked by bold lines. V.E., vertical exaggeration.

Figure 1 l a and b shows the density-Vp relationships for the E U R O B R I D G E EB'95 and E B ' 9 6 profiles. Obviously, substantial scatter is present, but deviations from the reference curve are small. This indicates that the upper crust consists mainly of low-grade felsic rocks, whereas the middle crust is made up of compositionally intermediate rocks, their metamorphic grade increasing with depth from amphibolite to granulite facies. The scatter in the density-Vp plots is explained best by compositional differences of the rocks, which affect the relationships between density and P-wave velocity in the various units crossed by the EB'95 and EB'96 profiles, as follows. (1) The upper crust of the W L G domain is composed of aluminous metasedimentary granulites, some pyroxene-bearing

350

400

450

500

gneisses and charnockites, whereas the MLSZ contains partly mylonitized granulites, charnockites, metavolcanic rocks and plagioclase-microcline granites. The EL domain, in turn, is made up of rather dense biotite-plagioclase gneisses and amphibolites (Skridlaite & Motuza 2001). This explains the scattering on the Vp-density plot for EB'95 (Fig. 11). (2) In the region of the CBSZ, the P-wave velocities and densities in the upper crust of the Okolovo terrane ( 6 . 1 6.2 km s -1 and 2.7 g cm -3, respectively) are higher than those beneath the southeastern half of this zone (5.8-6.0 km s - t and 2 . 6 - 2 . 6 7 g cm-3). This reflects a distinct difference between amphibolite- and granulite-facies oceanic-arc rocks in the N W and a terrane of large plutons of quartz syenites and sub-alkaline

EUROBRIDGE

613

dU 9

od

.

>4

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o',

~.,~

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I

I ~

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~

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~

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9

d~ ~ ~V

~ ~

I

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I~

r~

~

,"a~

9~ r0 0

[..

S. BOGDANOVA ET AL.

614

Table 2. Modal mineralogy for selected rock types used for gravity-seismic modelling (values in per cent) Qtz

Igneous rocks Granite Rapakivi granite Anorthosite Gabbro-norite Gabbro Metamorphic rocks Amphibolite Mafic granulite Mafic garnet granulite Eclogite

Kfs

Pl(An20)

Pl(An60)

Cpx

Opx

O1

30-40 17-27 0 0-2 0

30-40 46-53 0-2 0-10 2-3

20-40 14-24 0 0 0

0-4 0-4 0 0

0 0 0 0

30-40 0 0 0

Amph

0 0 90-95 43-65 50-60

0 0-1 0-2 10-24 7-10

0 0 0-2 0-20 6-14

0 0-1.5 0-2 0-23 10-30

0 2-6 0-2 0-5 3-6

5-10 40-50 17-41 0

0 4-20 5-10 5-41

0 8-12 4-10 6-16

0 0 0 3-15

55-60 20-30 19-30 13-57

Mineral abbreviations (after Kretz 1983): Qtz, quartz; Kfs, K-feldspar; P1, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; O1, olivine; Amph, amphibolite; Grt, garnet; An, anorthite. An in P1 is shown in parentheses.

(a)

E

vr

(D E3

50

100 150 200 250 300 350 400 450 500 550 600 650 700

Distance (km)

t'~&AD

(b)

Density (g/cm 3) 3.4 3,2

E

3.0

v

r2.8

Q. (9

s

2.6

0

100

50

150 200 250 300 350 400 450

500

Distance (km)

(c) IGNEOUS ROCKS &5 &2587.757.5" 7.25-

/

~

+ Granite + Rapakivi granite Anorthosite + Gabbro-norite ..... Gabbro

METAMORPHIC ROCKS Amphibolite + Mafic granulite + Mafic garnet granulite + Eclogite

~. 6.7s, 6.56.25~

1'7i

5.5-

~Z"

{

16~4 I

5.75~ 2,6

2,8 3 3.2 3.4 Density (g/cma)

-i6 2.8

3.3%,4

Density (g/cm 3)

26

4 2,8 3 3.2 3.4 Density (g/crn 3)

Fig. 10. Simplified density models along EUROBRIDGE'95-96 profiles (a) and EUROBRIDGE'97 profile (b) (after Kozlovskaya et aI. 2001, 2002, 2004). The positions of the major tectonic units are indicated as in Figure 2. (c) Relationships between density and seismic velocities obtained by Monte-Carlo simulation for selected types of igneous and metamorphic rocks from the Ukrainian Shield and other Precambrian areas. The modal mineralogy of rocks is given in Table 2. Left: density- Vp relationship; centre: density-Vs relationship (for Vs, the axis is scaled by a factor of 1.73); right: density-Vp/Vs relationship. The reference density-velocity relationships are shown by open stars.

EUROBRIDGE

(a)

615

(b)

EB' 95

7.47 -

EB' 96

~ ~ ;

7.21

i

8i e 6.6] 6.

~41P'

i

~6.8i 6.6i

5~"

6.41

~" 6.~

6.2i~, . ~ 2 f 6.~ .~

Ii t

2.6 2.7 2.8 2.9 3.0 3.1 3.2

I

i !

i

621,

I

6.0La,, ~, 5.8r ,-----4 2.6 2.7 2.8 Z9 310 3.1 3.2

Density (g/cm3)

granites with subordinate metasedimentary and volcanogenic gneisses and migmatites in the SE. The principal difference between the EB'95 and EB'96 profiles is in the composition of the lower crust. Along profile EB'95, the lower crust has P-wave velocities of 6.8-6.9 km s-1 and densities of 2.9-3.0 g cm 3, which are values typical of mafic granulites. Profile EB'96, however, in addition, shows a major high-velocity layer at the base of the crust, which has P-wave velocities of 7.1-7.4km s -1 and densities of 3.0-3.1 g cm -3, and extends beneath the southern parts of the BPG, the CBSZ, and the OMB. Beneath the WLG and EL areas, such a high-velocity layer is missing. The high-velocity layer crossed by the EB'96 profile is also visible in the density-velocity diagram of Figure 11. Its S-wave velocity and density values are typical of garnet-bearing mafic rocks formed under high-pressure conditions. As seen in the figure, some deviation from the reference curve towards lower velocity values may be due to elevated plagioclase contents. As already considered, joint interpretation of seismic and gravity data based on P-wave velocities and densities meets the problem that rocks with high contents of calcium-rich plagioclase have higher P-wave velocities and thus high Vp/density ratios, which do not plot on the standard reference curves. For this reason, interpretations based on the relationship between density and P-wave velocity have not been applied to the EB'97 profile that intersects the Korosten gabbro-anorthosite-rapakivi granite pluton (KP) with its commonly extremely plagioclase-rich rocks. Because isotropic S-wave velocities appear to be generally better correlated with density than P-wave velocities, and are less affected by high contents of plagioclase, the interpretation of the gravity data along EB' 97 was performed using the relationship connecting density to both the P- and the S-wave velocities (Thybo et al. 2003; Kozlovskaya et al. 2004). The combined velocity and density model of EB'97 demonstrates pronounced lateral variations of these properties, which can be related spatially to the extents of the geological units crossed by this profile, that is the OMB, the VD and PD, and the KP (Table 1). In addition, the relationships between density and the seismic velocities Vp and Vs in the geological units crossed by the EB'97 profile have been obtained. As can be seen from Figure 12, these differ from each other and deviate also from the corresponding reference curve relationships. Generally, the density-Vp and density-Vs relationships for the OMB are close to those of the reference curves for the upper and middle crust, but are shifted upwards from the reference curve for the lower crust (Fig. 12), indicating that the entire crust of the OMB probably consists of igneous rocks (i.e. granites, granodiorites and

Density (g/cm3)

Fig. 11. Comparison of the relationships between density and Vp for the EUROBRIDGE'95 and '96 profiles.

gabbros in different proportions). Beneath the OMB, the additional lower-crustal high-velocity layer with Vp, Vs and density values of 6.9-7.2 km s- 1 , 3.95-4.0 km s- 1 and 2.953.1 g cm- 3 , respectively, is seen in both the EB 96 and EB' 97 profiles. This layer is probably composed of mafic garnet granulites and eclogites. The most complicated structure of the crust along the EB'97 profile is found in the Palaeoproterozoic VD, which is composed of rocks of very different character and origin. In this case, the density-Vp plots for the VD and KP scatter around the reference curve (Figs 10 and 12). Furthermore, the points appear to shift upwards from that curve for densities less than 3.0 g cm -3, which characterize most of the igneous rocks of the KP (i.e. the rapakivi granites, gabbro-norite-anorthosites and gabbronorites). The values of Vp, Vs and density determined for the upper crust of the KP at depths to 5 km suggest that this body consists of both rapakivi granites and anorthosites, but the resolution of the data for the EB'97 profile is not good enough to model the structure in detail. The points closest to the reference curve in the density-Vp plot (Fig. 12) have been obtained for a high-velocity body in the northern part of the KP that has high values of both the P-wave and S-wave velocities (6.4-6.46 km s -1 and 3.663.72 km s -1, respectively), and densities of 2.75-2.80 g cm -3 at depths of 5 - 1 2 km. This combination of Vp, Vs and density values can be attributed to metamorphic rocks (Lebedev & Korchin 1982; Lebedev et al. 1983), which probably belong to the host rocks of the KP. The lowermost crust along the EB'97 profile (Figs 9 and 10) has P-wave velocities of 7.0-7.6 km s -1 and densities over 3.0 g c m - 3 ; that is, values similar to those of the high-velocity layer at the base of the crust beneath the OMB and the CBSZ. This suggests that it may be partly composed of mafic garnet granulites. However, the density-Vp relationship in this crustal unit (Fig. 12) indicates high contents of plagioclase, as a result of the presence of igneous gabbro-anorthositic rocks in the lower crust. For the upper and middle crust, the density-velocity ratios in the PD plot close to the reference curves. For the lower crust, however, they shift slightly upwards (i.e. towards higher velocities). The latter indicates granulite-facies conditions of metamorphism despite the relatively great crustal thicknesses of more than 50 km. These would necessitate pressure-temperature conditions in the lower crust that correspond to eclogite facies. However, the density values of the lower crust in the area are below 3.0 g cm -3, suggesting only minor contents of garnet but relatively high contents of plagioclase, and implying that the lower crust had not been eclogitized extensively. The reason may have been its anhydrous condition (Austrheim et al. 1997).

S. B O G D A N O V A E T AL.

616

Osnitsk-Mikashevichi Igneous Belt

(a) l

9.0

I

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8.5-

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2.9 3.1 3.3 Density [ g/cm 3 ]

'o ~

~

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2.5

2.7

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Density [ g/cm 3 ]

Podolian Domain I

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3.1 3.3 Density [ g/cm 3 ] I

I

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3.5 ...... I

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3.1 3.3 Density [ g/crn 3 ]

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Density [ g/cm 3]

I

> 3.5

3.5

l

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I f I

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2.5

Integrated geological-geophysical interpretation of the structure of the upper lithosphere along the EUROBRIDGE profiles G e n e r a l structural characteristics o f the crust

The interpretative models of the Earth's crust along the EUROBRIDGE EB'95, EB'96 and EB'97 seismic profiles presented in Figures 13 and 14 have been compiled from the overall geophysical and geological information, referred to in the previous sections. Collision between Fennoscandia and Sarmatia was decisive in determining the seismic characteristics of the lower crust and upper mantle in the study region, and the distribution of the magnetic and gravity anomalies (e.g. Garetsky et al. 2002). Pre-collisional terrane tectonics, in contrast, is reflected best by

I

2.7

I

2.9

l

I

3.1 3.3 Density [ g/cm 3 ]

........

I

3.5

Fig. 12. Relationshipsbetween density, Vp and Vsfor the major geologicalunits crossed by the EUROBRIDGE'97 profile (after Kozlovskayaet al. 2004). The left panel shows the density- Vprelationship, the right the density-Vs relationship(for Vs, the axis is scaled by a factor of 1.73). The reference density-velocity relationshipsare shown by open stars. (a) OsnitskMikashevichiIgneous Belt; (b) Podolian Domain; (c) Volyn Domain and Korosten Pluton.

structures in the upper and middle crust. Tectonically, the rock belts and domains in the Fennoscandian terranes make up a number of 'thick-skinned' nappe packages, thrust towards the SSE and SE in the southern part of the Baltic-Belarus region, but towards the NE in Estonia and the area of Lake Ladoga (Fig. 1). Subsequently, these nappes were transected by sets of N N W - N W - and NNE-NE-trending post-collisional faults and the markedly east-west-striking Mesoproterozoic faults. The EUROBRIDGE profiles suggest that the formation of highvelocity layers in the crust was commonly associated with detachment, whereas lateral undulations may have been shaped by complementary deformation of the whole lithosphere (Figs 8, 10, 13 and 14). Almost all of the high-velocity layers are accompanied by distinct subhorizontal seismic reflectors and mark sharp compositional discontinuities in the crust. In many cases, mafic sheet intrusions were responsible or contributed.

EUROB RIDGE

617

ol

Density (g/cm 3)

], I,l~l,t,l,l,l,lll

9

d,

[.. ...............

ell ~.~~ ,~

, ,,,,~ = 9

0

III

N--

~

9

@

d

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0

4, M

t',,l

-

r

< ,,-,;I ~Z

I

r~

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2~

=

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.,..,

<m ~Z

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~
g

n

. ,.,.i

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==-

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d

= I

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6

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.

S. BOGDANOVAET AL.

618

VD

OMB

Osnitsk-Mikashevichi Igneous Bell PT

Pripyat Trough

~>

KP

[

PD

Volyn D o m a i n

]

Korosten Pluton

I

Podolian D o m a i n

r ~

o_

~

o_

o_

o_

E

o.

.tZ

50

100

150

200

250

300

Distance [ km ]

350

400

450

500

o

C~

3.2

-40

~3,0 .~==

-60

2.8

-80

VD

OMB

13.4

~20

1~

26

i

O

3

2.4

PD

2.2

PT

I KP Korosten Pluton

Pripyat Trough

Berdichev Zone r ~ LD

VinnitsaBlock r.0 r~ c0

E ,.c:

0

50

100

150

200

Archaean crust of 3.65-2.8 Ga age Palaeoproterozoic 2.2-2.1 Ga terranes: [!iii!iiiiii!iii!ii::!i]Upper and middle crust

l

300

350

Korosten Pluton (1.79-1.74 Ga): Various granitoids, incl. rapakivi-type E Layered body, a) anorthositic (?) layer

Lower crust

I

2.0-1.95 Ga batholiths

[ ~

Moho boundary

250

400

450

Distance [ km ]

N

Magma chamber ] Intrusions and conduits of ca. 1.8 Ga age

Mantle SSW-dipping reflector

~

500

V.E.

1.5:1

S

High velocity layer of ca. 1.8 Ga magmatic rocks High velocity Palaeoproterozoic upper mantle, partly eclogitized Devonian magmatic underplate? L_III ] Sedimentary cover

ones of high reflectivity in the mantle

Various faults, incl. Interlayer detachment faults

Fig. 14. The integrated tectonic interpretation of EUROBRIDGE'97 profile. Top: seismic model as in Figures 8 and 9. Middle: gravity-seismic model as in Figure 10b. Bottom: the tectonic model. V.E., vertical exaggeration.

Complex crustal and upper mantle structures characterize the Fennoscandia-Sarmatia junction area beneath the CBSZ and part of the BPG, where the more ancient and more rigid crust of Sarmatia has been particularly strongly deformed, laminated and also altered compositionally. As modelling of the gravity and magnetic data shows, a substantial proportion of the post-collisional and later post-orogenic faults have listric configurations (Garetsky et al. 2002). In a number of cases, their flat-lying deeper parts coincide with nearly horizontal reflectors (detachment zones) at the boundary between the upper and middle crust. However, some large faults (e.g. the Minsk Fault in the CBSZ) appear to extend to the Moho (Figs 2 and 13).

Geophysical images of accretionary and collisional tectonics Terranes related to the Fennoscandian crustal segment. The PolishLithuanian terrane with the West Lithuanian Granulite Domain (WLG) and the Mid-Lithuanian Suture Zone (MLSZ) resembles in many respects the crustal province of southeastern Sweden, situated to the south of the classical Svecofennian orogen. In that province, the rocks were formed during several orogenic events at c. 1.841.82, 1.81-1.78 and c. 1.75 Ga, tending to young towards the south. The structural trends are dominantly W N W - E S E to westeast, and there exist active-margin volcanic-arc and back-arc

EUROBRIDGE

type supracrustal belts (Sundblad et aL 1998; Beunk & Page 2001; Mansfeld et al. 2005). In southeastern Sweden, the primary, pre-collisional or pre-accretionary relationships of the different lithotectonic complexes are recognized fairly well in the seismic images. These indicate subduction towards the NNE (BABEL Working Group 1993; Abramovitz et al. 1997; Balling 2000). In Figure 13, steep dips towards the NW are indicated for the MLSZ and the adjoining area, whereas other studies, particularly those of metamorphism (see p. 603), have suggested that the WLG had been thrust towards the east, overriding the East Lithuanian Belt (EL). Skridlaite et al. (2003a) inferred that the widespread occurrence of high- to moderate-pressure granulites in the WLG indicated a subduction-collision tectonic regime between 1.84 and 1.80 Ga, whereas island-arc settings have been identified both in the MLSZ and the adjoining parts of the WLG (Rimsa et al. 2001; Skridlaite & Motuza 2001; Motuza 2005). The crustal thickness in the WLG ranges between 45 and 50 km, whereas the crust atop the uplifted Moho in the MLSZ has a thickness of only c. 40 km. Similar differences also characterize the lower crust, which measures 10-12 km in the WLG, but only 5 - 1 0 k i n in the WLSZ (see Kozlovskaya et al. 2001; Yliniemi et al. 2001; Grad et al. 2003b). The thicker lower crust beneath the WLG appears largely to be due to the presence of a basal crustal layer with densities as high as 3.0 g cm -3 and thus probably composed of mafic granulite (see Christensen & Mooney 1995). This layer has no equivalent in the MLSZ. The upper and middle levels of the crust have P-wave velocities of 6.1-6.4 and 6.5-6.8 km S - 1 , respectively, both in the WLG and the MLSZ. However, the upper crust in the WLG is largely granulitic, whereas that in the MLSZ is made up of various rocks in the granulite and amphibolite facies, such as blastomylonites, calc-alkaline metavolcanic rocks and gabbroic to tonalitic plutonic rocks. At depths of 1 2 - 1 5 k m (Kozlovskaya et al. 2001; Yliniemi et al. 2001), the crust of the WLG and, in part, that of the MLSZ have a low-velocity layer, which most probably consists of Mesoproterozoic granitic rocks created by the remelting of the upper crust. A large body of a c. 1.46 Ga monzogranitoid rock is present close to the northern end of the EB'95 profile (Cerys 2004; Motuza 2005). A prominent feature of the crust particularly in the northwestern part of the WLG is its multi-layered structure, built up of distinctly delimited, conformable, persistent individual layers. The seismic velocities at the base of the crust are high (Vp 8.258.35 km s-~). Although this might suggest a 'platformal' type of crust in the sense of Christensen & Mooney (1995), lithological and geophysical variations are substantial within the WLG. Thus, its northern part is largely made up of orthogneisses, whereas a mixture of metasedimentary granulites, charnockites, metavolcanic rocks and various granitic rocks dominates in the south. Major differences of rock composition and deep structure also exist between the eastern and western parts of the WLG (Kozlovskaya et al. 2001). Particularly worth noting is the recurrent granitoid magmatism both in the Palaeo- and Mesoproterozoic. The Lithuanian-Belarus terrane, including the EL and BPG, together with the Okolovo terrane forms a composite terrane where the crust is substantially thicker (up to c. 55 kin) and, as a whole, also denser than that in the Polish-Lithuanian terrane (Fig. 13). The principal mechanisms responsible for the development of this thick crust appear to have been collisional orogenic processes involving compression and folding, and the stacking of large piles of nappes in the junction zone between Fennoscandia and Sarmatia. Indications of tectonic thickening, thinning, folding and wedging-out of the rock units are common in the seismic profiles. With regard to the thickness of the upper and middle parts of the crust, the EL and BPG are not very different from the WLG, but no low-velocity layers appear to be present. Here, the seismic velocities vary substantially in accordance with lithological variation, but as most of the rocks are either mafic to

619

intermediate granulites or igneous rocks of similar compositions, the upper crustal P-wave velocities are mostly relatively high. They measure c. 6.25 km s -1 in the EL, 5.8-6.0 km s -~ in the BPG, and 6 . 1 - 6 . 2 k m s 1 in the largely metavolcanic Okolovo terrane. Substantially lower velocities are, naturally, found in the large, anastomosing systems of shear zones marked by blastomylonites and retrograde recrystallization of the granulites, and also the presence of metasediments. These occur particularly in the EL. Major west-dipping listric faults that could be traced to depths of 1 5 - 2 0 k m (see Aksamentova et al. 1994) have previously been found along the Grodno-Starobin seismic profile transecting the BPG and the Okolovo terrane in a westeast direction (see also Bogdanova et al. 2001b). In the middle crust, which has densities around 2.8 g cm -3 and P-wave velocities of 6.3-6.5 km s -1, granulites and TTG-type plutonic rocks appear to predominate. This part of the crust forms a 'trough' beneath the BPG. With regard to the lower crust, the BPG and EL are similar (Table 1, Fig. 13). Both have a c. 2 0 k m thick lower-crustal layer made up of mafic granulites with P-wave velocities of 6.8-7.1 km s -1 and densities of 2.9-3.1 g cm -3 (Fig. 10, Table 1). A remarkable feature is the southeasterly dips of this lower crustal layer, which appears to protrude into the upper mantle; its densitY3of 3.1-3.2 g cm -3 is substantially less than the 3.3-3.4 g c m - of the normal upper mantle in the region. Thus, this mantle offset-lower crustal protrusion may represent a 'fossilized' slab of subducted Palaeoproterozoic oceanic crust. The F e n n o s c a n d i a - S a r m a t i a junction. Within the CBSZ, the seismic and gravity characteristics of the crust change drastically across the major, west-dipping Minsk Fault (Figs 10a and 13). The latter extends to the Moho and, at the Earth's surface, separates two very different groups of tectonic units (Figs 1 and 2), the BPG and Okolovo terrane in the NW and the Vitebsk Domain (VG) and Borisov-Ivanovo Belt (B-I) in the SE. The P-wave velocities and rock densities in the upper crust are different on the two sides of the Minsk Fault, being 6.1-6.2 km s -1 and 2.7 g cm -3 in the NW and 5.8-6.0 km s- 1 and 2.60-2.67 g c m - 3 in the SE. This reflects the difference between the amphibolite- to granulitefacies mafic rocks of the Okolovo terrane and the granite-intruded metasedimentary gneisses and migmatites in the B - I . In the middle crust, the P-wave velocities are 6.3-6.5 km s -1 in the NW and 6.4-6.9 km s -1 in the SE, but there is apparently no corresponding difference in density values (Kozlovskaya et al. 2002). The best explanation appears to be that the higher P-wave velocities below the southeastern part of the CBSZ are due to a markedly laminated structure in a part of the crust where strongly deformed amphibolite- and granulite-facies rocks have been emplaced tectonically. In the CBSZ region, these diverse structural patterns in the upper to middle crust can be followed down to depths of 2 5 30 km, at which level the P-wave velocities reach 7.0 km s -1 and densities of 2.9-3.0 g cm -3 have been modelled. Farther down is a rather more uniform lower crustal high-velocity layer (HVLC in Figs 8, 10a and 13) with P-wave velocities of 7.27.4 km s -1 and densities of 3.0-3.1 kg m -3. These values correspond best to eclogitic granulites or garnet granulites. Similar high-velocity layers in the lower crust, with P-wave velocities between 7.0 and 7.7 km s - 1, appear to be common in Precambrian regions adjoining major tectonic sutures (Guggisberg & Berthelsen 1987; Korja etal. 1993; Korja & Heikkinen 1995; Funck etal. 2001; Hall et al. 2002). In some cases, they are associated with sizeable Moho offsets and are mostly explained as resulting from magmatic mantle underplating (Korsman et al. 1999; Funck et al. 2001 ; Hall et al. 2002). In the case of the CBSZ, the lower crustal high-velocity layer is a relatively young feature, as it appears to underlie adjacent terranes as well. These include the OMB, which adjoins the CBSZ in the SE, and, some 250 km farther SE, the part of the Volyn

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Domain that encloses the AMCG Korosten Pluton. In view of this spatial association, this high-velocity, high-density layer is most probably coeval with the c. 1.80-1.74 Ga Korosten body. Xenoliths from the lower crust (Markwick et al. 2001) and new data on the isotopic compositions of both OMB and Korosten rocks suggest that this high-velocity layer may be restitic, having been formed between 1.80 and 1.74 Ga by the successive removal of the AMCG melts from a 2.0 Ga lower crust similar to that beneath the OMB (Bogdanova et al. 2006). As indicated by the gravity-seismic modelling in Figures 10, 13 and 14, many crustal units in the CBSZ are wedge-shaped and imbricated, and are separated from each other by numerous distinct reflectors dipping in opposite directions. Structural patterns of this kind are characteristic of collisional-type crust (Meissner 1989; Cook et al. 1999) and therefore fit well with the location of the CBSZ region in the zone of collision between Fennoscandia and Sarmatia. A conspicuous major feature of the crust beneath the CBSZ and adjoining parts of the OMB and BPG is a large antiformal (domal) structure defined by convex high-velocity middle to lower crustal rock layers apparently related to the OMB (Fig. 13). As this structure coincides with the collisional zone between Sarmatia and Fennoscandia, it appears reasonable to assume that stacking and thickening of the crust with attendant metamorphism and gravitational instability must have been part of the early stages of its development (see Coney & Harms 1984). An important key to deciphering the formation of this antiform beneath the CBSZ and its evolution into a metamorphic core complex is the presence of a bulge of high-velocity material in the lower crustal core of the antiform (Fig. 13). In conjunction with other features, such as the association of the Korosten Pluton with the high-velocity lower crust, this suggests that c. 1.8 Ga magmatism could have been a major agent causing the doming and attendant metamorphism during the post-collisional stage. This evolution was associated with post-collisional extensional tectonics leading to c. 1.8 Ga AMCG magmatism in the OMB and the Lithuanian-Belarus terrane, local granulite metamorphism, intense mylonitization along the Minsk Fault, and, eventually, listric faulting and fast final uplift (Taran & Bogdanova 2003). Also, downfaulting of the edge of the OMB to expose the Borisov-Ivanovo and the Vitebsk tectonic units must have been part of this extension. The latter are connected with southwards dipping reflectors in the middle and lower crust of the OMB (Juhlin et al. 1996; Stephenson et al. 1996). The Archaean and Palaeoproterozoic crust of Sarmatia. The OMB was formed by voluminous magmatism between c. 2.0 and 1.95 Ga. In accordance with the OMB igneous mode of origin, the character of its c. 50 km thick crust is determined largely by the presence of numerous batholiths of granitic, granodioritic, dioritic or gabbroic composition. These obviously correspond to the seismic-velocity and density properties (Figs 12 and 13), which also suggest that the more felsic of the plutons dominate the upper crust, whereas the mafic ones prevail in its lower parts (Kozlovskaya et al. 2002, 2004; Yegorova et al. 2004). Markwick et al.'s (2001) study of deep crustal and mantle xenoliths indicates that the mafic plutonic rocks of the OMB have been partly transformed into eclogite-like, garnet-bearing granulites with P-wave velocities of 6.8-7.0 km s -1 and correspondingly high densities. In addition, the OMB contains younger, c. 1.8 Ga, mostly syenitic to quartz syenitic intrusions, which are associated with the coeval AMCG-type Korosten Pluton farther south and define a belt of marked, more or less isometric, magnetic anomalies (see Fig. 2). The distribution of these intrusions appears to have been controlled by major NE-trending, NW-dipping zones of faulting, which also follow some of the OMB boundaries. The upper and middle parts of the crust in the OMB in particular feature numerous major reflectors (Figs 8 and 13) that create

an overall multi-layered structure, presumably mostly caused by recurrent magmatism and tectonic deformation, especially in the vicinity of the Fennoscandia-Sarmatia junction. Some of the layering, however, must rather be due to the formation of the Devonian Pripyat-Dniepr-Donets Aulacogen (Fig. 1). As discussed by Stephenson et al. (1996), the system of listric faults in that structure coincides closely with Palaeoproterozoic wedge fabrics within the OMB, which indicates significant reactivation of Precambrian faults during Phanerozoic rifting. A well-preserved fine lamination of the crust also characterizes the OMB, presumably caused by deformation of rocks with contrasting elastic properties (Meissner & Rabbel 1999) and probably related to the latest major deformation event in the Devonian. In the lower crust, a high-velocity lower crustal layer with Vp of 7 . 2 - 7 . 4 k m s -1 exists also in the OMB, similar to that beneath the CBSZ, but substantially thinner and denser than in the latter. In the Volyn Domain (VD) with the large AMCG Korosten Pluton (KP), the crust is only 45 km thick; that is, substantially thinner than the 5 0 - 5 2 km crust in the neighbouring OMB and PD (Figs 8 - 1 0 and 14). The available seismic and gravity data suggest, however, that beyond the limits of the Korosten Pluton, the VD is similar to the OMB. All the crustal layers of the OMB appear to continue into and across that domain, extending southwards as far as the Berdichev Boundary Zone, which dips north and separates the Palaeoproterozoic VD from the Archaean interior parts of the Podolian Domain (PD). In the Berdichev Zone (BZ), the lower part of the Palaeoproterozoic crust appears to wedge out at depth, and the distinctly layered upper and middle parts are replaced southwards by seismically more uniform and less reflective Archaean crust. In this ancient crust there is a rather gradual increase of the P-wave seismic velocities with depth, from 6.1-6.2 to nearly 6.9 km s -1. The distribution of these velocities and the Vp/Vs ratios suggest a two-layered crust, but the S-wave data and the gravity-seismic modelling by Kozlovskaya et al. (2004) indicate the presence of three crustal layers. The relatively low Vp/Vs ratios of 1.69-1.74 indicate that, the Archaean crust of the Podolian Domain is much richer in quartz than the neighbouring Palaeoproterozoic crust (see also Yegorova et al. 2004). The crustal structures and boundaries in the Archaean part of the Podolian Domain dip gently towards the north, conforming to the inferred northerly dips of the Archaean-Proterozoic crustal boundary in the BZ. The grades of Palaeoproterozoic metamorphism increase southwards from amphibolite- and granulite-facies rocks to high-grade granulite-facies rocks. This suggests that the Pa|aeoproterozoic crust of the VD overlies the Archaean crust of the PD in a manner that may be a result of collisional tectonics at c. 2.1-2.05 Ga, followed by extension and the formation of a metamorphic core complex (Fig. 14). The greatest lateral and vertical variations of crustal composition and structure in western Sarmatia are associated with the 1.80-1.74 Ga multiphase Korosten Pluton. The influence of this intrusion is not restricted to its area of exposure, but extends for tens of kilometres in the surrounding region. All the crustal units have been updomed in a wide region (see Figs 8, 9 and 14). The gravity and magnetic modelling of the Korosten Pluton, employing also the data of the east-west-trending Geotraverse II seismic profile (Ilchenko & Bukharev 2001), indicates an extremely complex structure in the underlying crust (Bogdanova et al. 2004b). Whereas layered gabbro-anorthosite intrusions can be followed only to depths of less than 10 kin, granitoid rocks of various kinds, including rapakivi granites and monzonites, form flat-lying sheeted bodies at various levels of the upper and middle crust. The interlayering of these igneous sheets with earlier Palaeoproterozoic supracrustal and plutonic rocks is inferred to be responsible for the presence of low-velocity crustal layers with P-wave velocities of 6.1-6.5 km s -1 and densities varying between 2.6 and 2.8 g cm -3 (Fig. 14).

EUROBRIDGE At depths below 16 kin, mafic rocks form a single semicylindrical, lensoid body beneath the eastern part of the Korosten Pluton (see Fig. 14; note that the EB'97 profile crossed only the western half of the intrusion). This body measures c. 90 km across and extends to the high-velocity, high-density layer at the base of the crust, which underlies the VD, the OMB and even part of the PD (see above). This giant mafic body is considered to represent the feeding magma chamber of the Korosten Pluton (Bogdanova et al. 2004b). The underlying crust, with P-wave velocities of 7.4-7.8 km s -I, Vp/Vs of 1.77-1.79 and densities of 3.0-3.15 g cm -3, is therefore assumed to be mostly composed of mafic, ultramafic or eclogitic rocks, presumably representing a mixture between cumulates of the Korosten magma and restitic material (Bogdanova et al. 2006).

Aspects o f the mantle

With regard to Moho topography, the EUROBRIDGE seismic profiling and gravity modelling indicate depths varying between 40 and 55 km as well as a number of irregularities and offsets. Some of these were related to accretionary and/or collisional tectonics, or to superimposed late to post-collisional magmatism. The most obvious case of the former is the mantle irregularity beneath the Belarus-Podlasie Belt (VPG), which appears to connect with the Fennoscandia-Sarmatia collisional junction as defined by the Minsk Fault. Another, but so far less evident instance, may be the relationship between the Mid-Lithuanian Suture Zone (MLSZ) and a Moho offset beneath the West Lithuanian Domain (Fig. 13). An offset beneath the Volyn Domain (VD) may continue the Archaean-Proterozoic boundary in the Berdichev Zone, but coincides with and may have been masked by the root of the Korosten Pluton (Fig. 14). In addition, it is tempting to relate the Korosten magmatism to the extensive Moho uplift to 45 km beneath the Volyn Domain and the adjacent parts of the Osnitsk-Mikashevichi Belt (OMB) and Podolian Domain (PD). The upper mantle is rather inhomogeneous with regard to seismic velocities, the P-wave velocities ranging between 8.1 and 8.35 km s -1. This is probably due to lateral compositional variation from peridotite to eclogite. The latter composition is particularly characteristic of sites of collisional thickening in zones of deformation and fluidization of the lower crust and upper mantle (e.g. Austrheim et al. 1997). However, as mentioned above, a restitic origin of eclogites in the lowermost crust and upper mantle beneath the Korosten Pluton is also possible. Xenoliths in kimberlites of various ages and near-source alluvial placers have demonstrated that the upper mantle beneath the VD is made up of a 20 km thick layer of eclogites, underlain by garnet pyroxenites and peridotites (Tsymbal & Tsymbal 2003). According to Tsymbal & Tsymbal, the age of the mantle is Proterozoic beneath both the Volyn and the Podolian domains. A more eclogitic composition of the mantle beneath Sarmatia may be the reason why it has substantially higher Vp/V~ ratios than the mantle of Fennoscandia (1.83 and 1.72, respectively). Of particular interest in the tectonic interpretation of the lithosphere are reflectors in the upper mantle. Apart from some subhorizontal reflectors referrable to rheological and mineralogical changes with depth, the EUROBRIDGE transect also shows one distinct inclined reflector and more circumstantial evidence of several others. These may represent 'fossil' zones of subduction of oceanic as well as continental crust (see Balling 2000). The less distinct reflectors can to some extent be extrapolated on the basis of the Moho topography and compositional variation in the upper mantle, and also from lower crustal lenses of melting apparently related to post-collisional processes. The presence of a lens of lower-velocity and lower-density mantle beneath the edge of the Lithuanian-Belarus terrane where it faces Sarmatia thus suggests subduction into the mantle of a slab of Fennoscandian lower crust. Similar relationships, albeit less clearly

621

expressed, are also found in the upper mantle beneath the MLSZ and the adjoining parts of the WLG. Another zone of elevated mantle reflectivity is associated with the Berdichev Zone outlining the junction of the PD and VD. The distinct SSW-inclined mantle reflector beneath the OMB along the EB'97 profile (Figs 8 and 9) has previously been interpreted for purely geometrical reasons as the trace of a collisional boundary between Sarmatia and Volgo-Uralia (Thybo et al. 2003). An alternative interpretation (Aisberg & Starchik 2005) suggests that this reflector represents a detachment surface in the crust and upper mantle that was related to the formation of the Pripyat-Dniepr-Donets Aulacogen (PDDA) in the late Devonian. The latter interpretation accounts well for the part of the mantle with lower density and lower seismic velocity that overlies this reflector and is best explained as consisting of mafic and ultramafic igneous rocks. Support for this interpretation is provided by a Devonian (c. 380 Ma) age of lower crustal hornblendite xenoliths found in lamprophyric tuffites in southeastern Belarus (Markwick et al. 2001). Geochemically, these may represent remelting products of a garnetiferous mantle. Thus, it appears possible that the distinct flat-lying reflectors found at depths of c. 10 km below the undulating Moho can also be related to the Devonian event and mark the presence of mantle-derived melts (see Figs 8, 9 and 14). Similar conditions have also been observed along other seismic profiles running across the PDDA (Grad et al. 2003a; Maystrenko et al. 2003).

Conclusions The EUROBRIDGE traverse project has provided a new understanding of the structure and formation of the crust and upper mantle in the western part of the East European Craton. Although the results mostly concern the key region between the Baltic and Ukrainian shields, and the late Palaeoproterozoic collision of Fennoscandia and Sarmatia, they have relevance also for understanding the upper lithosphere in the entire EEC. The major conclusions are as follows. (1) The crust in the region between the Baltic and Ukrainian shields is Palaeoproterozoic and juvenile. It was formed between c. 2.0 and 1.8 Ga by accretionary plate-tectonic processes along the margins of the Archaean-earliest Palaeoproterozoic nuclei of Fennoscandia and Sarmatia. (2) Several Palaeoproterozoic terranes, related either to Fennoscandia or to Sarmatia, are recognized on the basis of their different ages, lithologies and tectonothermal evolution. They include various tectonic settings: juvenile island arcs, back-arcs and active continental margins. The Sarmatian terranes were formed between c. 2.2 and 1.95 Ga, whereas the Fennoscandian ones are, in general, younger, ranging between c. 2.0 and 1.8 Ga. Palaeomagnetic data indicate that the Fennoscandian and Sarmatian terranes belonged to different plates. (3) The complex, belt-shaped crustal structure and the fault zones that bound the various belts and domains mostly originated during the collision between the Sarmatian and Fennoscandian plates at some time between 1.85 and 1.80 Ga. However, the listric character of many faults and associated late to postcollisional magmatism, retrograde metamorphism and strong mylonitization along the inter-terrane boundaries as well as within their interiors all suggest that post-collisional extensional tectonics was of crucial importance for the following crustal development between c. 1.80 and 1.74 Ga, and even later at c. 1.71-1.67 Ga. Emplacement of the large AMCG plutons at 1.80-1.74 Ga in Sarmatia, and between c. 1.6 Ga and 1.50 Ga in Fennoscandia, substantially influenced the composition, petrophysical properties and geophysical structure of the crust and upper mantle. (4) The present major characteristics of the seismic profiles and potential fields in the Baltic-Belarus region were predetermined

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s. BOGDANOVAETAL.

by late Palaeo- and Mesoproterozoic accretionary, collisional and post-collisional geodynamics. The last, in particular, caused rearrangement of the lithospheric structure and shaped its present geophysical images. (5) Occasionally, the fault and suture zones between the Fennoscandian belts and domains can be traced throughout the entire crust (e.g. the Mid-Lithuanian Suture Zone). Displacements and offsets of the Moho boundary and various crustal layers along these zones, as well as the 'imbricate' character of the Palaeoproterozoic crust in southern Fennoscandia, allow comparison with 'thick-skinned' orogens. The offsets and irregularities of the M o h o boundary and lateral changes of petrophysical properties and compositions in the upper mantle may be interpreted as 'fossilized' Palaeoproterozoic zones of subduction and collision. This is particularly the case in the Central Belarus Suture Zone, between the Fennoscandian and Sarmatian terranes, where the crust is characterized by a pronounced tectonic layering and numerous reflectors. (6) The boundary between Fennoscandia and Sarmatia is defined by the major Minsk Fault, an extensional feature superimposed on the suture zone. Beneath the Minsk Fault, the crust was affected by doming of the collisionally stacked crustal layers, voluminous m a g m a t i s m at the base and the formation of a metamorphic core complex. (7) Subsequent rifting of the crust and the development of the Late Mesoproterozoic V o l y n - O r s h a Aulacogen was shallow and dispersed, roughly coinciding with the Central Belarus Suture Zone. Also, the Devonian rifting and the formation of the P r i p y a t - D n i e p r - D o n e t s Aulacogen did not cause substantial thinning of the c. 50 k m Palaeoproterozoic crust or its significant reworking. However, the underlying Palaeoproterozoic faults were reactivated and controlled the position of major listric faults (e.g. those outlining the Pripyat Trough). The SSW-dipping reflector beneath the northwestern margin of Sarmatia, thus, most probably represents a detachment surface bounding this aulacogen in the NE. The low-velocity upper mantle above this reflector is probably a Devonian mantle underplate. The EUROBRIDGE project (1994-2002) has been a highly successful WestEast co-operation enterprise. Despite the many economic difficulties in the East European countries involved, it produced a wealth of results and offered a unique experience to numerousjunior researchers. During its lifetime, geological and geophysical institutions, research councils and academies in 17 countries contributed financially. Particular thanks go to the Swedish Institute's Visby-Programme, the Royal Academy of Sciences in Stockholm and the INTAS organization (project 94-1664). Research and the workshops were always conducted in a warm and cordial atmosphere, with most participants feeling like members of a single family. In the above text, the section reporting the seismic results was compiled by M. Grad, A. Guterch and T. Janik, and E. Kozlovskaya authored the section on gravity-seismic modelling. L. Taran and G. Skridlaite produced most of the P - T - t data for the metamorphic rocks. G. Motuza, one of the founding fathers of the project, contributed invaluable material from Lithuania, and V. Starostenko did the same for the Ukraine. S. Bogdanova was the scientific leader of EUROBRIDGE, and R. Gorbatschev coordinated the INTAS effort. The authors thank D. Kurlovich from the Belarussian State University in Minsk for help with the preparation of the GIS-formated maps presented in this paper.

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SUNDBLAD, K. & CLAESSON, S. 2000. The Precambrian of Gotland. Geofizichesky Zhurnal, 22, 136. SUNDBLAD,K., MANSFELD, J., MOTUZA, G., AHL, M. & CLAESSON, S. C. 1994. Geology, geochemistry and age of a Cu-Mo-bearing granite at Kabeliai, Southern Lithuania. Mineralogy and Petrology, 50, 43-57. SUNDBLAD,K., GULLENCREUTZ,R. & FLOOgN, T. 1998. The Precambrian crust beneath the Baltic Sea. Geofizichesky Zhurnal, 20, 121-124. TARAN, L. N. & BOGDANOVA,S. V. 2001. The Fennoscandia-Sarmatia junction in Belarus: new inferences from a PT-study. Tectonophysics, 339, 193-214. TARAN, L. N. & BOGDANOVA,S. V. 2003. Metamorphism of the Palaeoproterozoic paragneisses of the Belarus-Podlasie Granulite Belt: a prograde-retrograde evolution. Petrology (Petrologija), 11, 444-461. THYBO, H., JANIK, T., OMELCHENKO, V. D., ET AL. 2003. Upper lithosphere seismic velocity structure across the Pripyat Trough and Ukrainian Shield along the EURUBRIDGE'97 profile. Tectonophysics, 371, 41-79. TSYMBAL, S. N. • TSYMBAL, Y. S. 2003. The upper mantle composition and diamond prospects in NW Ukrainian Shield. Mineralogichesky Zhurnal, 25, 40-56 [in Russian]. VALVERDE-VAQUERO, P., D(SRR, W., BELKA, Z., FRANKE, W., WISZNIEWSKA, J. & SCHASTOK, J. 2000. U - P b single-grain dating of detrital zircon in the Cambrian of central Poland: implications for Gondwana versus Baltica provenance studies. Earth and Planetary Science Letters, 184, 225-240. VERKHOGLIAD, V. M. 1995. Age stages of the Korosten pluton magmatism. Geokhimia i Metallogenia, 21, 34-47 [in Russian]. WOOLLARD, G. P. 1959. Crustal structure from gravity and seismic measurements. Journal of Geophysical Research, 64, 1521 - 1544. WYBRANIEC, S. 1999. Transformations and visualization of potential field data. Polish Geological Institute, Special Papers, 1. WYBRANIEC, S., ZHOU, S., THYBO, H., ETAL. 1998. New map compiled of Europe's gravity field. EOS Transactions, American Geophysical Union, 79, 437-442. YEGOROVA,T. P., STAROSTENKO,V. I., KOZLENKO,V. G. & YLINIEMI,J. 2004. Lithosphere structure of the Ukrainian Shield and Pripyat Trough in the region of EUROBRIDGE-97 (Ukraine and Belarus) from gravity modelling. Tectonophysics, 381, 29-59. YLINIEMI, J., TIIRA, T., LUOSTO, U., ET AL. 2001. EUROBRIDGE-95: deep seismic profiling within the East European Craton. Tectonophysics, 339, 153-176. Z1NCHENKO, 0. V., SKOBELEV,V. M., ESIPCHUK,K. E., SHEREMET,E. M. & VERKHOGLYAD, V. M. 1990. The Korosten complex. In: SHCHERBAKOV,I. B. (ed.) Petrology, Geochemistry and Metallogeny of Intrusive Granitoids of the Ukrainian Shield. Naukova Dumka, Kiev, 134-164 [in Russian].

The Archaean nucleus of the Fennoscandian (Baltic) Shield A. I. S L A B U N O V l, S. B. L O B A C H - Z H U C H E N K O 2, E. V. B I B I K O V A 3, P. S O R J O N E N - W A R D 4, V. V. B A L A G A N S K Y 5, O. I. V O L O D I C H E V 1, A. A. S H C H I P A N S K Y 6, S. A. SVETOV 1, V. P. C H E K U L A E V 2, N. A. A R E S T O V A 2 & V. S. S T E P A N O V 1

1Institute of Geology, Karelian Research Centre, RAS, Petrozavodsk, 185910, Russia 2Institute of Precambrian Geology and Geochronology, RAS, St. Petersburg, 199164, Russia 3Vernadsky Institute of Geochemistry & Analytical Chemistry, RAS, Moscow, 117975, Russia 4Geological Survey of Finland, Kuopio, 70211, Finland 5Geological Institute, Kola Science Centre, RAS, Apatity, 184209, Russia 6Geological Institute RAS, Moscow, 119017, Russia

Abstract: Archaean supracrustal complexes, known in the Fennoscandian (Baltic) Shield, are described and discussed by analysingthe time sections 3.1-2.9, 2.9-2.75 and 2.75-2.65 Ga. Data on granitoid complexes, interrelated in time and space, and evidence for Archaean metamorphic events are classified and presented briefly. Fragments of ophiolitic and eclogitic associations have been found in Archaean rocks in the Shield. The first evidence of continental crust in the Shield is from Meso-Archaean time (3.5-3.1 Ga); isolated microcontinents, such as Vodlozero, Iisalmi and North Finland, have been identified. New continental crust was mainly generated in the 2.9-2.65 Ga interval. The geodynamic settings in which the continental crust was formed in the Mesoand Neoarchaean included subduction (ensialic and ensimatic), accretion and collisional mechanisms. The continental and oceanic crust were affected by mantle plumes.

Archaean rocks form much of the eastern and northern parts of the Fennoscandian (Baltic) Shield, and can be divided into a number of discrete crustal provinces, each of which has a distinctive history of crustal formation and reworking. From SW to NE these are the Karelian, Belomorian, Kola and Murmansk Provinces, respectively (Fig. 1). Of these, the Murmansk Province has been little affected by younger events, whereas the Belomorian and Kola Provinces both record significant thermal and tectonic reworking and amalgamation related to the Palaeoproterozoic Lapland-Kola collisional orogeny (Daly et al. 2006). Further SW, the Karelian Province was blanketed by Palaeoproterozoic intracratonic basins; however, it shows significant thermal overprinting and tectonic reworking only along its southwestern margin, as a result of accretionary processes associated with the 1.9-1.8 Ga Svecofennian Orogeny. The Karelian and Murmansk Provinces form the cratonic nuclei of the Shield, and are therefore designated here as the Neoarchaean Karelian and Murmansk Cratons (Fig. 1). The Karelian Craton contains the typical range of granite-gneiss, greenstone, paragneiss and granulite complexes characteristic of the Archaean granite-greenstone association (Glebovitskii 2005; SorjonenWard & Luukkonen 2005). In contrast, the Murmansk Craton is composed dominantly of various granite-gneisses and granitoids, within which supracrustal rocks occur only as enclaves (Radchenko et al. 1994), usually metamorphosed to amphibolite grade; relics of granulite-facies mineral parageneses have also been described from the central part of the Craton (Petrov et al. 1990). The Belomorian Province is generally understood as a mobile belt, along the eastern and northeastern margin of the Karelian Craton (Fig. 1). Discriminating between Archaean and Palaeoproterozoic processes and events in the Belomorian mobile belt has been a source of controversy, but recent studies have clarified much of this and demonstrated that high-pressure (kyanite-facies) metamorphism (Volodichev 1990; Glebovitsky et al. 1996), with associated deformation, occurred in both in the Neoarchaean and the Palaeoproterozoic (Bibikova et al. 1996, 2001). Thus, although the Belomorian mobile belt contains rock units that are similar to those of the adjacent Karelian Craton in terms of age and composition, it has a distinctly different structural architecture, being composed of large-scale intensely folded nappe complexes

(Miller & Mil'kevich 1995). This generally recumbent structural development is also evident across the transition zone from the Belomorian mobile belt northeastwards into the Palaeoproterozoic Lapland-Kola orogen (Daly et al. 2001). The Kola Province consists of the Kola-Norwegian, Keivy and Sosnovska terranes and the Kolmozero-Voronya greenstone belt (Daly et al. 2006). Each of these terranes includes greenstone, paragneiss and granite-gneiss complexes, and the K o l a Norwegian terrane consists of a granulite-gneiss complex (e.g. Avakyan 1992). The entire Kola Province has been involved to a greater or lesser extent in the Palaeoproterozoic Lapland-Kola orogeny and it is separated from the Belomorian mobile belt by the Palaeoproterozoic (2.0-1.75 Ga) Lapland-Kola collisional suture (Daly et al. 2006) collisional sutures. The Inari and Tersky-Strel'na domains within the Kola suture zone comprise both Neoarchaean and Palaeoproterozoic rocks, the latter representing juvenile crustal protolith of the Lapland and Umba granulites (Glebovitsky et al. 2001; Daly et al. 2006). In the NE, the Kola Province borders against the Murmansk Craton. In summary, much of the Archaean of the Fennoscandian (Baltic) Shield is a typical granite-gneiss association (covering about 80% of the area), with various greenstone, paragneiss and granulitic complexes. However, the Belomorian mobile belt shows a distinct tectonic pattern, with large-scale thrusting and nappe complexes, and includes two tectonic and metamorphic rock associations that are rare in the Archaean: ophiolitic (the Central Belomorian greenstone belt and the Iringora complex) and eclogite-bearing complexes (Gridino zone and Salma area). This review commences with a description of the supracrustal greenstone and paragneiss complexes according to isotopically determined age groupings, followed by a presentation of intrusive and metamorphic rock associations.

Greenstone and paragneiss complexes Palaeoarchaean supracrustal rocks are represented by metakomatiites and spherulitic metabasalts in the Volotsk unit of the Vodlozero terrane in the southeastern part of the Karelian Craton (Kulikova 1993). Although they have not yet been studied in

From: GEE, D. G. & STEPHENSOY,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 627-644. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

627

628

A.I. SLABUNOV E T A L .

(a)

Archaean greenstone complexes (letters in square indicate major greenstone and schist belts: II = ilomantsi; KB = Khedozero-Bolsheozero, Ke = Keivy; Ks = Kostomuksha, KT = Kuhmo-Suomussalmi-Tipasj&rvi; KV = Kolmozero-Voron'ya; NK = North Karelian, OI = Olenegorsky, SK = Sumozero-Kenozero, SV = South Vygozero, T = Tulppio, V = Voche-Lambina; VS = Vedlozero-Segozero; Y = Yenb)

(b)

100Km White Sea

P IN , Murmansk Craton; LKO, Lapland-Kola Orogen; BP and KP, Belomorian and Kola Provinces, respectively; KC, Karetian Craton; SO, Svecofennian Orogen.

1

Archaean eclogitebearing complex (Gr = Gridino, Sa = Salma) Paragneiss complexes (2.7-2.78 Ga; N = Nurmes)

F~

CK

White Sea

~ , ~ --

"

2.75-2.68 Ga

o ~ m

2.9-2.85Ga

~

2.8-2.75 Ga (rhomb=lringora ophiolite complex)

E o ~ n

3.1.2.9 Ga

2.9-2.85 Ga ophiolite-like

complex of Central Belomorian Greenstone Belt 2.9-2.7 Ga granitoids, including Keivy alkaline granites, from Central Karelian (CK), Kola-Norwegian (KN), Keivy (Ke), Sosnovka (So) and North Sweden terranes

33 ~ E Lapland granulite belt (Lp; mainly 2.0-1.9 Ga), including Tanaelv melange

Caledonides Phanerozoic and Neoproterozoic platform cover []~-[

Neo- and Mesoproterozoic rocks Rapakivi granites

(1.65-1.54 Ga)

Granitoids (1.85-1.75 Ga) Supracrustal rocks (2.06-1.85 Ga)

r-q

Supracrustal rocks (2.5-2.06 Ga) Tectonic mixture of Neoarchaean and Palaeoproterozoic rocks (In and TS, Inari and Tersk-Strel'na terranes, respectively)

Sanukitoids and their analogues (2.74-2.72 Ga; Tv = Tavaj&rvi massif) Granulitic complexes (2.74-2.72 Ga; Vp = i~i!~i!i!i~!~i!;~i:~Varpaisj&rvi, Vk = Voknavolok, TI = Tulos, On = Onega, Nt = Notozero)

m

2.9-2.7 Ga granitoids from Belomodan (BMB) and Kianta (Ki) terranes 3.1-2.7 Ga granitoids from Ilomantsi-Voknavolok (IV) and Vodlozero (Vo) terranes 3.5-2.7 Ga granitoids from terranes: lisalmi (li), Pomokaira (P), Ranua (R), Vodlozero (Vo; core)

illmill ~

~

"~,',~b % ",, c

Major tectonic boundary Faults: thrust (a), normal (b) and strike-slip (c)

Fig. 1. (a) Tectonic units of the eastern Fennoscandian (Baltic) Shield. (b) Schematic representation of the major geological units and structures in the eastern Shield (based on the authors' data, and Kostinen et al. 2001; Sorjonen-Ward & Luukkonen 2005).

ARCHAEAN NUCLEUS, FENNOSCANDIANSHIELD detail, they have a S m - N d isochron whole-rock age of 3391 + 16 Ma (Puchtel et al. 1991). Some of the high-grade gneises and amphibolites I of the Vodlozero gneiss complex were perhaps derived from 3.3-3.55 Ga calc-alkaline volcanic rocks (Sergeyev et al. 1990; Lobach-Zhuchenko et al. 1993). The Meso- and Neoarchaean greenstone complexes of the Shield belong to at least four generations with ages of 3.1-2.9, 2.9-2.85, 2.85-2.75 and 2.75-2.65Ga, whereas paragneiss and schist complexes have been dated at 2.9-2.85 and 2.752.65 Ga. Each of these age groupings is considered separately below.

629

belt of the Kola Province. Contacts with younger complexes are typically tectonic. The Vedlozero-Segozero greenstone belt is located on the western margin of the Vodlozero terrane in the Karelian Craton (Fig. 1). It comprises two separate complexes with different ages: 3.05-2.94 Ga and c. 2.85 Ga (Fig. 2). The older of these includes two distinct associations formed between 3.05 and 2.94 Ga, as follows. ( 1) A basaltic - andesitic- dacitic association, which is most complete in the northern part of the Hautavaara structure (the Chalka zone), has a total thickness of 2.5 km and includes pillowed, amygdaloidal and fragmental lava flows, various volcaniclastic vents and dykes, metamorphosed to amphibolite and epidote-amphibolite grade. Ages of 2995 _+ 20 Ma have been obtained from subvolcanic rocks of andesitic composition (Sergeyev 1989) and 2945 _ 19 Ma for andesitic lava (Ovchinnikova et al. 1994). These rocks belong to a normally differentiated calc-alkaline series. The more primitive volcanic rocks are rich in Cr and Ni, whereas later differentiates are enriched in Co, Zr and Y. Their Sr/Y ratio < 12, Ce/Nb ratio <4.5 and Th/Nb ratio <0.72 are all

Greenstone belt associations f o r m e d between 3.1 and 2.9 Ga

Volcano-sedimentary complexes of this age are known in the Vedlozero-Segozero, South Vygozero, Sumozero-Kenozero and Kuhmo-Suomussalmi-Tipasj/irvi greenstone belts of the Karelian Craton, and also in the Kolmozero-Voron'ya greenstone

Proterozoic rocks: l

rapak~4granites (1.65-1.54 Ga)

~

supracrustalrocks (2.50-2.10 Ga)

Neoarchaean rocks: diorites, granodiorites, sanukitoids (2. 74 Ga)

N

T

gabbro.norites -~ mafic and ultramafic intrusive rocks Mesoarchaean rocks: /

~

granites (2.85-2.87 Ga) 2.86-2.85 Ga complexe

/ '%,

[

~

gabbro-diorites (2.85 Ga)

~

andesitic to dacitc volcanics, adakites and sediments (2.86-2.85 Ga)

/

3.05-2.9 Ga fiomplexe

/

high-Mg gabbro ~ ~

komatiitic-basalticassociation (lavas and tufts) (3.0-2.95 Ga) BADR-series volcanics, adakites (3.05-2.94 Ga)

~

amphibolites

~

gneissose-granites, migmatitegranites (3.15-2.95 Ga) - - ~ palaeovolcanic structures 1-5 - Hautavaara structure: 1 - Nyalmozero, 2 - Ignoila, 3 - Hautavaara, 4 - Maselga, 5 - Chalka, 6-9 - Koikary-Semch structure: 6- Yanish, 7- Korbozero, 8 - Elmus, 9 - Semch

~ km __J

faults

Fig. 2. Schematicgeologicalmap of the Vedlozero-Segozero greenstone belt (Svetov 2005).

630

A.I. SLABUNOVETAL.

typical of island-arc series. They also show enrichment in light rare earth elements (LREE) ((La/Sm)N = 1.7) with flatter heavy rare earth element (HREE) patterns ((Gd/Yb)N = 1.3). Subvolcanic rocks have adakite characteristics. All of these features suggest that the association can be described as a relict of the oldest island-arc complex preserved in the Shield (Svetov 2003). (2) A komatiitic-basaltic association, which is common in the Hautavaara, Koikary, Palaselga (or Palalamba) and Sovdozero structures (Svetov 1997 and references therein), has a S m - N d age (whole-rock isochron) of 2921 _+ 55 Ma at ~Nd(t) = +1.5 (Svetov et al. 2001). A minimum age constraint is provided by cross-cutting dacite dykes dated at 2935 _ 15 (Bibikova 1989) and 2860 _+ 15 Ma (A. Samsonov, pers. comm.) as well as gabbro-diorites, dated at 2890 ___40 Ma (Sergeyev et al. 1983). Sequences up to 2.7 km in thickness consist of various lava flows, including pillowed and spinifex-structured lavas; pyroclastic facies make up not more than 5% of the sequence (Svetova, 1988; Svetov et al. 2001). Pyroxenitic and basaltic komatiites and basalts predominate. Comagmatic Mg-rich gabbros and serpentinized ultrabasic rocks are also present. The rocks are metamorphosed to greenschist and amphibolite grade. Based on CaO/AI203 ~ 0.8, AlzO3/TiO2 ~ 22 and minor element and REE distributions (Gd/Yb)N = 0.99-1.10), the komatiites are classified as an Al-undepleted type. Unlike the lavas, komatiite tufts have low percentages of A1203 ( < 8 wt%), higher percentages of CaO ( 7 - 9 wt%) and low percentages of alkalies, although their REE distribution patterns are similar. The association is attributed to a spreading regime in a back-arc basin, initiated by the uplift of a mantle plume (Svetov 2005). In the South Vygozero greenstone belt, the Shilos belt is composed dominantly of basalts (including pillow basalts) with scarce komatiitic basalt and agglomerate (Sokolov 1981 and references therein). The rocks were metamorphosed at greenschist- to epidote-amphibolite-facies conditions and have been complexly folded (Glebovitsky 2005). Ages of 2915 _+ 30 Ma (LobachZhuchenko et al. 1999) and 3054 _+ 84 Ma (A. Samsonov, pers. comm.) have been reported for the basalt-komatiitic association. The basalts are cut by gabbros, tonalites and felsic dykes. The basalts belong to the tholeiitic series and are rich in Cr (250500ppm) and Ni (100-150ppm). Based on REE abundances, they can be subdivided into two groups: (1) LREE-depleted basalts; (2) basalts with only weakly fractionated REE patterns. The value ~Nd(t)----+1.6 suggests that the association was derived from a depleted mantle source and that there was no crustal contamination. Its formation has been attributed to the effects of a mantle plume (Lobach-Zhuchenko et al. 1999). In the Sumozero-Kenozero greenstone belt, two tectonically juxtaposed complexes are distinguished in Lake Kamennoe: an older komatiitic-basaltic complex dated by the S m - N d wholerock method at 2916 _+ 117 Ma (Sochevanov et al. 1991) and a younger (2875 Ma) differentiated complex. The komatiiticbasaltic complex is dominated by pillowed tholeiites and spinifexstructured komatiites, with intercalations of felsic volcanic rocks, tufts and graphite-bearing schists (Sokolov 1981). The complex is cut by mafic and felsic dykes and granitic bodies. The komatiites are also associated with C u - N i ore occurrences (Rybakov & Golubev 1999). Komatiites (MgO ---- 2 7 - 4 3 wt %) are of Al-undepleted character and, like the tholeiites, they have low LREE (La/SmN = 0.70 in komatiites and 0.68 in tholeiites) and high Nb contents, and their ~Nd(t)= +2.7. This association has been inferred to represent an oceanic plateau complex (Puchtel et al. 1999). The Kuhmo-Suomussalmi-Tipasj~irvi greenstone belt also records several distinct age groupings, with the majority of volcanic rocks being nearer 2.8 Ga in age, as discussed below. Older rocks are inferred from amphibolite enclaves within 2.84 Ga migmatites (Luukkonen 1988, 1992) and from a distinct intermediate rock association, in the so-called Luoma Group, which is restricted to the northern end of the Suomussalmi belt (Taipale 1988). The

Luoma Group also contains basalts and felsic volcanic rocks, the latter exhiditing strong REE fractionation. A zircon U - P b age of 2966 _+ 9 Ma (Hypp6nen 1983) has been obtained for the intermediate rocks, but it is possible that zircons are inherited, and the supposed andesites could in fact include reworked sediments (Vaasjoki et al. 1999). The Kolmozero-Voron' ya greenstone belt of the Kola Province consists predominantly of 2.9-2.85 Ga supracrustal complexes. Associated with the belt, however, are strongly tectonized and internally differentiated gabbro-anorthosite bodies, such as the Patchemvarek massif. These gabbroic massifs are isolated within granitoids of the Murmansk Craton in close proximity to the greenstone belt (Mints et al. 1996; Pozhilenko et al. 2002), and because of their age (2925 _+ 6 Ma, Kudryashov et al. 2001) they are considered to mark the initiation of greenstone belt evolution in the province. Greenstone and paragneiss sequences formed b e t w e e n 2.9 a n d 2.85 G a

Rocks of this age are known in the belts surrounding the core of the Vodlozero terrane of the Karelian Craton (VedlozeroSegozero, Sumozero-Kenozero and South Vygozero); they are also common within the Belomorian mobile belt and in the greenstone belts of the Kola Province (Kolmozero-Voron'ya, Keivy). In the Vedlozero-Segozero greenstone belt, dacitic to andesitic volcanism was characterized by subaerial pyroclastic eruptions, with inferred palaeovolcanic complexes up to 1.5 km thick (Svetova 1988). The most complete section is present in the Koikary-Semch structure (Fig. 2), where the Yanis palaeovolcanic edifice comprises lava breccias, lavas and blocky agglomeratic deposits; the feeder channel consists of subvolcanic dacites. Siliceous chemical precipitates and graded beds deposited as craterlake facies are also present. The distal part of the complex contains generally finer-grained volcaniclastic deposits and siliceous chemical sediments. Subvolcanic intrusions are dacitic and rhyolitic. The felsic volcanic rocks of 2860 _+ 15 Ma (A. Samsonov, pers. comm.) and rhyolite dykes (2862 _+ 45 Ma, Ovchinnikova et al. 1994) correlate well with the Hautavaara structure dacites (2854 + 14 Ma, Sergeyev 1989). These felsic volcanic rocks are rather sodic: KzO/Na20 ----0.39 and SiO2 varies from 52 to 76 wt % (Svetov 2005). REE ratios are as follows: (La/Sm)N = 3.0, (Gd/Yb)N = 2.3, (Ce/Yb)N = 5.8, with the finer tufts showing more differentiated REE patterns: (La/Sm)N-~3.5-4.1, ( C e / Y b ) N = 2 2 - 2 6 . The andesitic and dacitic lavas have similar abundances of Th (5.2-6.2 ppm), La (13-29ppm), Hf (2.6-5.4ppm) and Nb ( < 1 0 p p m ) and in La/Yb ratios (7-9), Ti/V (60-70), Hf/Yb (3-4), Ti/Zr (16-37) to those of volcanic rocks from the Andean-type active continental margins. This follows also from the adakitic nature of the Semch volcanic rocks (Ba = 270-500 ppm, Sr = 200-320 ppm, Nb = 3.0-3.8 ppm, Ti = 3600-3800 ppm, low HREE). In the Sumozero-Kenozero greenstone belt, the Lake Kamennoe structure contains, in addition to the old (c. 2916Ma) komatiitic basalts, another unit comprising basaltic-andesiticdacitic-rhyolitic lavas and tufts, with carbonaceous and carbonate schists and quartzites, metamorphosed to greenschist grade (Sokolov 1981; Bogatikov 1988), and subvolcanic adakites. The age of the rhyolites and adakites is 2875 _ 2 Ma (Puchtel et al. 1999). This unit also contains distinctive metasomatic lithologies, including carbonate rocks, listwanites and carbonate-quartz veins (Kuleshevich 1992). Volcanic rocks of this highly differentiated series tend to be poor in Mg, Nb and Ti but enriched in LREE ((La/Sm)N = 1.4 in basalts and andesites and 3.3 in rhyolites); adakitic rhyolites exhibit marked REE fractionation (La/ SmN = 3.8-5.1, Gd/YbN = 2.8-4.5; there is no Eu anomaly; Puchtel et al. 1999). Such differentiated series and adakites are considered to be formed in subduction zones upon melting of

ARCHAEANNUCLEUS,FENNOSCANDIANSHIELD the mantle wedge and the slab of a subducting plate, respectively (Drummond et al. 1996). The North Karelian greenstone belt in the Belomorian mobile belt consists of the Tikshozero and Keret belts. Its supracrustal rocks are represented by metavolcanic rocks and metasediments of the Lake Keret (2.88-2.82 Ga) and Hisovaara (2.8-2.77 Ga) series. The Lake Keret series makes up a large part of the Keret belt and consists of three tectonostratigraphic associations: a komatiitetholeiitic association, an intermediate-felsic volcanic association and an andesitic basalt to basaltic association (Slabunov 1993). The petrogeochemical characteristics and geochronology of these rocks have been discussed by Stepanov & Slabunov (1989), Slabunov (1993, 2005) and Bibikova et al. (1999). The metabasalts of the komatiite-tholeiitic association represent a rather Na-rich type of tholeiitic trend, ultramafic basalts are classified as LREE-enriched and pyroxene-bearing, and basaltic komatiites belong to the Al-undepleted type (AlzO3/TiO2 ~ 20, CaO/ A1203 = 0.64-0.9; Zr/Y = 2-3). Intermediate to felsic volaniclastic deposits and lavas (ranging from andesitic basalts to rhyolites, with predominant andesites and dacites) of 2877 _+ 45 Ma (Bibikova et al. 1999) together with subvolcanic bodies of 2829 _+ 30 Ma, make up the bulk of the greenstone complex. Their petrogeochemical characteristics are comparable with those of volcanic rocks from more mature island arcs. The andesites also have similar ~ d = 2.80 Ga (eNd(t) = +2.8), indicating that the magma was not contaminated significantly with material of long crustal residence. The metavolcanic rocks of the andesitic basalt-basalt association contain metasedimentary intercalations containing large concentrations of Cr (up to 570 ppm) and Ni (up to 130 ppm), which could have been derived from weathering of intermediate to felsic volcanic rocks, basalts and komatiites. Greywackes with such compositions are also typical in oceanic island-arc settings, and the overall association is considered to mark an early (2.9-2.8 Ga) subduction stage in the crustal evolution of the Belomorian mobile belt (e.g. Slabunov & Bibikova 2001; Slabunov 2005). The Central Belomorian greenstone belt (CBGB) (Slabunov 2005) is primarily identified as a mafic zone (Stepanov & Slabunov 1989; Lobach-Zhuchenko et al. 1998; Bibikova et al. 1999). It is a 0.5-3.0 km wide structure that can be traced for 150-160 km N W - S E along the axial line of the Belomorian mobile belt; it presumably extends even further SE. There is a distinctly discordant relationship between the CBGB and the other structures of the Belomorian belt (Fig. 1). Four fragments (Seryak, Nigrozero, Lake Louhi-Pizemsky and Nizhemsky) have been recognized, each consisting predominantly of amphibolites, with some ultramafic rocks of 2.88-2.86 Ga (Bibikova et al. 1999). The Seryak fragment is the best exposed and best preserved structural element, and has been traced for over 70 km. This unit contains the largest (c. 300 m thick) deformed ultramafic body within the amphibolites. The ultramafic rocks include serpentinites, melanocratic amphibolites and chlorite-serpentineamphibole schists. They locally retain primary magmatic structures and minerals such as olivine (Fa84-86), orthopyroxene (Enss-s6) and spinel (ferrialumochromite with up to 29% Cr203), and can thus be identified reliably as metamorphosed harzburgites, dunites and pyroxenites (Stepanov et al. 2003). The ultramafic rocks are typcially poor in LREE ((La/Yb)N = 0.52), but some varieties showing a U-shaped REE distribution have also been reported (Slabunov 2005). Metabasalts (amphibolites) from the CBGB correspond to tholeiites in chemical composition and are comparable in many respects with mid-ocean ridge basalt (MORB) and ocean-island basalt (OIB) (Stepanov & Slabunov, 1989; Lobach-Zhuchenko et al. 1998; Bibikova et al. 1999). The mafic-ultramafic rocks show petrological and geochemical features similar to those of Phanerozoic ophiolite complexes (Lobach-Zhuchenko et al.

631

1998; Bibikova et al. 1999; Slabunov & Bibikova 2001; Stepanov et al. 2003; Slabunov 2005). The minimum age of the CBGB is 2878 +_ 13 Ma (Bibikova et al. 1999), which is the age of magmatic zircons from high-Fe trondhjemitic gneisses resembling those associated with tholeiitic series. Because these gneisses have ~ d _ 2840 Ma and eNd(t) = +2.7, there was clearly no significant Palaeoarchean material in their protoliths. The Seryak mafic-ultramafic unit is also intruded by diorites dated at (U-Pb) 2850 _+ 10 Ma (Borisova et al. 1997). The Chupa paragneiss complex is composed of migmatized kyanite-garnet-biotite and biotite gneisses, among which garnet-biotite gneisses (the relics of the least altered primary rocks) occur as small lenticular bodies. The latter are generally understood as metasediments, although an alternative viewpoint is that they include a metavolcanic component (Volodichev 1990). The enrichment in Ni, V, Co and Cr in some gneisses, for example, is also consistent with a greywacke protolith composition (Myskova et al. 2003 and references therein), derived from weathering of intermediate to felsic volcanic rocks and mafic and ultramafic rocks in a fore-arc basin environment. Scarce amphibolite bodies, occurring amongst the paragneisses, are chemically equivalent to tholeiites. The Sm-Nd model age estimates for these rocks ( ~ d = 3.01-2.83 Ga; Timmerman & Daly 1995; Bibikova et al. 2001, 2004; Myskova et al. 2003) defines the maximum depositional age for the paragneiss protoliths, which is consistent with the U-Pb detrital zircon datings of 3.2-2.9 Ga. In contrast, the earliest metamorphogenic zircons have U-Pb ages of 2.85-2.80 Ga (Bibikova et al. 2004). Therefore, sedimentary precursors to the paragneisses could have been deposited between 2.9-2.82 Ga. This assumption is supported by an age of 2883 + 22 Ma, obtained by U-Pb dating of magmatic zircons from garnet-biotite gneisses (Levchenkov et al. 2001). This further demonstrates that sedimentation and volcanism were coeval, irrespective of whether the gneisses represent metasediments or metavolcanic rocks. The Kola paragneiss complex of the Kola-Norwegian domain is made up of garnet-biotite gneisses (often with sillimanite and cordierite, occasionally with kyanite and staurolite). Sporadic intercalations of two-mica gneisses, biotite, amphibole and amphibole-biotite gneisses, as well as amphibolites, amphibole-pyroxene schists, amphibole-magnetite schists and pyroxene-magnetite schists and iron formation are also present (Radchenko et al. 1994, and references therein). The rocks were metamorphosed during several events, and twice under granulitefacies conditions (see below). Compositionally, they correspond to sediments (dominated by greywackes and pelites) and various volcanic rocks. Isotopic data (Bibikova 1989; Avakyan 1992; Balashov et al. 1992; Timmerman & Daly 1995; Kudryashov et al. 2001, and references therein) suggest that the protoliths to the paragneisses were derived from juvenile Neoarchaean material (tN~ is usually not older than 3.0 Ga). Because the first metamorphic event occurred at 2.80-2.85 Ga (see below), these metasediments were probably deposited at 3.0-2.85 Ga, which is effectively simultaneous with the metagreywackes from the Chupa paragneiss complex of the Belomorian mobile belt. The Kolmozero-Voron'ya greenstone belt extends as a narrow linear belt delineating the boundary zone between the Kola Province and Murmansk Craton. It consists of komatiitic-tholeiitic, basaltic-andesitic-dacitic, and terrigenous tectonostratigraphic associations (Vrevsky 1989; Radchenko et al. 1994, and references therein); all contacts between these units are tectonic and record overthrusting during Neoarchaean collision (Mints et al. 1996; Pozhilenko et al. 2002). Komatiites of the komatiitic-tholeiitic association commonly exhibit spinifex textures, and lava breccias are also observed (Vrevsky 1989; Smolkin et al. 2000). Both komatiitic series rock types and Mg-rich and Fe-rich basaltic compositions are present (Ruzh'eva et al. 2002). All these rocks are poor in REE (REE abundances in the komatiites are 0.5-1.2 times that of

632

A.I. SLABUNOVETAL.

chondrites, whereas those of tholeiites are 4 - 6 times greater) and show a 'flat' distribution pattern (Vrevsky 1989). The mantle source for the komatiitic melts was also poor in LREE. The komatiites were formed at 2826 _ 60 Ma (Sm-Nd whole-rock method, Ruzh'eva et al. 2002). These dates are close to the time of extraction of komatiitic melts from a depleted mantle source (2870 _ 90 Ma, ENd(t) = +2.5 + 0.3, Vrevsky 2002). The basaltic-andesitic-dacitic stratotectonic association consists of calc-alkaline series metavolcanic rocks. A quartz-porphyry from the Voron'ya tundra, belonging to this association, has been dated by the U - P b method at 2828 + 8 Ma (Kudryashov et al. 2001). One terrigenous association is made up of homogeneous metasediments (garnet-biotite gneisses), whereas the other has more variable compositions ranging from metagreywacke (alumina gneiss) to polymictic conglomerates. The formation of this greenstone complex is either attributed to the subduction of the intervening crust between the Murmansk Craton and the Kola Province (Mints et al. 1996), or is related to the development of a mantle plume (Vrevsky 2002). The Keivy schist belt, located in the Kola-Norwegian domain of the Kola Province, consists dominantly of felsic metavolcanic rocks (Lebyazh'ya sequence) and structurally overlying, highly differentiated terrigenous metasediments represented by garnetiferous, kyanite, staurolite-kyanite and sillimanite schists, carbonaceous shale, metasandstone and quartzite (Radchenko et al. 1994, and references therein). The metavolcanic rocks often occur as hastingsite- and microcline-bearing gneisses that are inferred to represent metavolcanic rocks that underwent alkaline metasomatism. The age of volcanism is estimated at 2871-t-15 Ma (Belyaev et al. 2001). Isotopic S m - N d data and the ages of detrital zircons show that both ortho- and paragneisses were derived from Neoarchaean juvenile material (Mints et al. 1996; Bridgwater et al. 2001). The internal sequence within the belt has generally been regarded as stratigraphical and concordant, but there is evidence (Bridgwater et al. 2001) to suggest that it consists of several tectonically imbricated slices. G r e e n s t o n e belt s e q u e n c e s f o r m e d between 2.85 and 2.75 Ga

Greenstone associations formed during this period include the Kostomuksha greenstone belt and large parts of the KuhmoSuomussalmi-Tipasj~irvi greenstone belt in the Karelian Craton, the North Karelian and the Yena greenstone belts of the Belomorian mobile belt, and the Olenegorsky greenstone belt of the Kola Province. The Kuhmo- Suomussalmi- Tipasjiirvi greenstone belt includes the following associations: komatiite-basalt, (andesite)-daciterhyolite, Fe-basalt and sedimentary rocks, metamorphosed under epidote- amphibolite-facies conditions (Piirainen 1988; Sorj onenWard et al. 1997; Sorjonen-Ward & Luukkonen 2005). A critical constraint on the evolution of the belt is provided by the differentiated Moisiovaara mafic sill, which intrudes the older deformed migmatite-amphibolite complex and has a U - P b zircon age of 2790 _+ 18 Ma (Luukkonen 1988). The best preserved section is in the southern part of the Kuhmo greenstone structure, where a stratigraphic sequence has been recognized, commencing with at least 1000 m of massive and pillowed tholeiitic flows and intercalated magnetite-grtineritequartz banded iron formations (BIF). This Pahakangas phase of volcanism was followed by the Sivikko komatiitic volcanism, which contains, in addition to numerous spinifex-textured flows, a significant proportion of serpentinites derived from olivine accumulates. The KellojS.rvi cumulate complex has yielded a zircon U - P b age of 2757 _+ 20 Ma, which is currently one of the few constraints from within the greenstone sequence (Tulenheimo 1999). Although the komatiites are mostly LREE-poor varieties,

LREE-enriched flows low in the sequence are attributed to assimilation of felsic material; partially assimilated granitic enclaves are also found in serpentinite cumulates (Tulenheimo 1999). The basalts typically belong to the tholeiitic series but compositions also extend into Mg-rich and komatiitic basalt fields. In addition, some highly unusual Cr-rich basalts (Cr = 13004500 ppm) have also been recognized and attributed to a relatively low oxygen fugacity in the source region (Halkoaho et al. 2000); these rocks also display weakly fractionated REE (La/Smy = 0.3-1.2), and Eu anomalies are absent (Piirainen 1988). The komatiitic and Cr-basaltic sequence is overlain discordantly by a sequence of coarse and graded mafic to felsic pyroclastic and epiclastic deposits (Taipale 1988; Nieminen 1998). Isolated occurrences of the (andesite)-dacite-rhyolite are not well constrained stratigraphically in the Kuhmo greenstone structure, but by analogy with the sequence further south in the TipasjSxvi structure, these felsic volcanic rocks and related sediments appear to underlie the basaltic and komatiitic volcanic rocks. This is consistent with the few age determinations from Kuhmo felsic rocks, namely 2798 + 15 (Hypp6nen 1983) and 2810 _+ 48 (Luukkonen 1992), given that the zircon age of 2791 __ 8 Ma from Tipasj/irvi is considered to be a reliable analysis (Vaasjoki et al. 1999). The felsic volcanic rocks at Tipa@irvi also host the Taivalj~irvi stratiform A g - Z n - P b deposit (Kopperoinen & Tuokko 1988). The Kostomuksha greenstone belt hosts one of the largest iron deposits, and comprises rocks of the basalt-komatiite, rhyolitedacite and sedimentary associations, with the last containing banded iron formations (Rayevskaya et al. 1992). The S m - N d isochron whole-rock ages obtained for the basalts and komatiites are 2843 _+ 43 Ma (Puchtel et al. 1998) and 2 8 0 8 _ 95 Ma (Lobach-Zhuchenko et al. 2000a). The basalts include massive, pillowed and variolitic lavas with sporadic iron formations and graphitic schist intercalations. The basalts belong to the tholeiitic series, have variable alumina contents, and have low LREE ((La/Sm)y = 0.66) and Th contents, and a 'flat' HREE pattern ((Gd/Yb)y = 1). A chondritic Ti/Zr ratio and aNd(t) value of +2.8 were published by Puchtel et al. (1998), although negative (to -3.44) aNd(t) values have been reported for two basalt samples by Lobach-Zhuchenko et al. (2000a). Komatiites are represented by spinifex-structured massive and pillow lavas, breccias, tufts and subvolcanic bodies (Bogatikov 1988). They are poor in both LREE and HREE ((La/ Sm)N = 0.48, (Gd/Yb)N = 1.2), and A1 (A1203/TiOz)N = 0.76; Puchtel et al. 1998). The overlying rhyolitic-dacitic association is represented by various tufts and tuffites with carbonaceous shale and iron formation streaks. The ages obtained for the rocks of the association are 2795 + 10 Ma (Lobach-Zhuchenko et al. 2000a). Their aNd(t) values vary from --1.3 to --6.2 (Puchtel et al. 1998). They are similar petrochemically to felsic volcanic rocks of FII type (Condie 1983). The sedimentary association consists of conglomerates, iron formation and greywackes. The basal conglomerates contain metarhyodacite (60%), amphibole schist and amphibolite (30-40%) clasts (Rayevskaya et al. 1992). Terrigenous sediments are homogeneous, immature and are similar to greywackes from Neoarchaean greenstone belts, such that their source area would have comprised approximately equal amounts of mafic and felsic rocks (Mil'kevich & Myskova 1998). In the North Karelian greenstone belt (NK), the entire Tikshozero belt and the northem part of the Keret belt consist of supracrustal rocks of the Hisovaara series (2.8-2.77 Ga), which have been described in detail by numerous workers (Kozhevnikov 1992, 2000; Shchipansky et al. 1999, 2001, 2004; Thurston & Kozhevnikov 2000; Miller et al. 2002; Kozhevnikov et al. 2005). The following tectonostratigraphic associations have been recognized: supra-subduction ophiolites, volcano-sedimentary rocks, coarse clastic sedimentary rocks and basalts. This

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complex is considered to be related to 2.8-2.75 Ga subduction during the later evolution of the Belomorian mobile belt. A supra-subduction ophiolitic fragment makes up the northern part of the Hisovaara structure (Fig. 3a); from the base upwards, it consists of: (1) ultramafic rocks (peridotitic cumulates); (2) tholeiite-series metabasalts; (3) Mg-rich basalts with a 0.5-1 m thick metaboninite unit; (4) Ti-rich ferrobasalts; (5) homogeneous and amygdaloidal andesites (calc-alkaline and tholeiite series). The age of the last is not less than 2777 _ 5 Ma (Bibikova et al. 2003a). This sequence is cut by subvolcanic rhyodacites dated at 2799 4- 67 Ma (Kozhevnikov 1992) and trondhjemites, dated at 2804 _ 27 Ma (Bibikova et al. 2003a). The sedimentary-volcanic association consists of andesites with intercalations of quartz arenite and conglomerates with granule- to pebble-sized clasts at the base, thinly laminated (kyanite)-biotite-muscovite pelitic schists, as well as graphitic schists and distinctive quartz-kyanite rocks. The quartz-kyanite rocks are attributed to hydrothermal alteration related to 2.77 Ga volcanism (e.g. Bibikova et al. 2003a). The intermediate to felsic volcanic rocks are composed of calc-alkaline andesitic, dacitic and rhyodacitic lava and tufts; a marine eruptive environment is indicated by alternating horizons of carbonaceous, chemically precipitated sediments and iron formations. The U - P b age of dacitic volcanism ranges from 2778 + 21 Ma (at Hisovaara) to 2782 + 9 Ma (at Iringora). Subvolcanic dacite, rhyodacite and rhyolite within the last two associations are comparable in

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Fig. 3. Geologicalmaps of (a) the Hisovaara area and (b) the Iringora area. The map of Hisovaara structure is modified after Thurston & Kozhevnikov (2000); the map of Iringora structure is modified after Shchipanskyet al. (2004). The abbreviationsin the inset map, showing the eastern Fennoscandian (Baltic) Shield, are as in Figure 1.

composition with Phanerozoic adakites (Bibikova et al. 2003a). Some workers have noted, however, that they are poorer in Mg and are less strongly enriched in Fe group elements; this suggests that interaction between the melts that were derived from the subducting plate and the overlying mantle wedge was insignificant. This has been attributed to relatively gently dipping subduction, as a result of the greater thickness and hence negative buoyancy of the Archaean oceanic crust. The Hisovaara upper basaltic association includes Mg-rich basalts (MgO = 7-12%,) basalts and basaltic andesites (SiO2 = 49-54%), enriched in LREE in varying degrees, with L a / YbN = 0 . 9 - 4 (Kozhevnikov 1992). The Iringora ophiolitic sequence, compared with the Hisovaara sequence, is a domain of relatively low strain (Fig. 3a and b). The Iringora ophiolitic sequence (Fig. 3b) occurs within an imbricate thrust stack that dips gently to the NNE. This stack preserves several slices of c. 2.8 Ga mafic volcanic rocks, of both boninitic and MORB-like tholeiitic affinity, overthrust onto a tectonically juxtaposed complex of arc-derived felsic volcanic rocks and siliciclastic turbidites (Fig. 3b). Although the entire Iringora sequence has been intensely deformed and metamorphosed, primary igneous, volcanic and sedimentary features are locally well preserved. The best preserved fragment of the ophiolitic sequence occurs along the northern shore of Lake Irinozero, where not only gabbroic and lava units, but also remnants of a sheeted-dyke complex, including half-dykes and dyke cusps,

634

A.I. SLABUNOVETAL.

are exposed, representing the transition to the overlying lava unit. Moreover, at the base of the ophiolitic nappe there is a tectonic m61ange, metamorphosed in the mode of a metamorphic sole. This complex consists of internally fragmented rock bodies containing a mixture of locally derived rocks of the boninite series and exotic F e - T i basalt blocks in a lithologically diverse matrix composed of coarse-grained garnet-, staurolite- and kyanitebearing biotite-amphibole schists with lenses of subarkosic meta-arenites and graphitic schists. Although the thickness of the lava units along the north shore of Lake Irinozero is no more than 100-150 m, the preservation of primary volcanic textures is striking. The lava unit consists predominantly of pillowed and massive flows, but also contains some pillow breccias indicative of hyaloclastic processes. Medium- to coarse-grained, compositionally layered, mafic rocks of the Iringora sequence are interpreted as the gabbro unit, which appears to be the deepest exposed level of the ophiolite. This unit is strongly foliated, so that primary gabbro textures such as modally graded layering are preserved only in a few lower-strain locations. The transition from the gabbro unit to the sheeted-dyke complex is abrupt and appears to be faulted. However, it is important to note that several dykes have been found intruding the gabbro unit. Uniformly layered medium- to fine-grained rocks occur between the lavas and the gabbro units, and are interpreted as a dense swarm of mafic and intermediate dykes; that is, a sheeted dyke complex. This differs from the foliated gabbro in terms of both dyke width (average 50-60 cm) and texture, locally displaying relict screens of earlier-formed crust. A single outcrop within the dyke complex displays well-preserved chilled margins, demonstrating the development of dyke-within-dyke relationships. The transition from sheeted dykes to volcanic rocks was preserved only in a few localities where the sheeted dykes have intruded the overlying volcanic rocks at a nearly perpendicular angle. Sheeted dykes represent one of the best diagnostic criteria for defining extensional magmatic systems, such as ophiolites. Although sheeted dykes may be developed during rifting of volcanic edifices built on sialic crustal basement, those of the Iringora sequence, combined with their distinctive geochemical features, strongly support interpretation as a supra-subduction zone ophiolite. Indeed, the major and trace element chemistry of the Iringora gabbro, dyke and lava units suggests a coherent series with a continuous compositional variation from more primitive mafic to strictly boninitic melts. Moreover, rocks of boninitic composition are found both in the lavas and in the dyke units (Shchipansky et al. 2001, 2004). The volumetrically dominant low-Ti tholeiites define a chemically primitive (Mg-number = Mg/(Mg + Fe 2+) = 0.72-0.78) group of mantle melts, strongly depleted in incompatible elements. The group identified as strictly boninitic, including high-Mg hypersthene-normative dacites, is characterized by volatile-free recalculated SIO2=53.4-66.2%, M G O = 8 . 4 13.1%, Mg-number = 0.65-0.74, and TiO2 = 0.34-0.48%. The relatively high CaO (up to 12.3%) contents and the high CaO/ A1203 values (mostly 0.8-1.1) suggest that these rocks may belong to the high-Ca boninite group (Crawford et al. 1989). Co, Ni and Cr abundances are high and vary from 51 to 79 ppm Co, 187-452ppm Ni and 324-1680ppm Cr, respectively. Mantle-normalized plots show a marked depletion in HFSE and REE compared with MORB. There are also variable negative Nb anomalies (average Nb*/Nb = 0.59), but dominantly positive Zr and Hf peaks (average Z r * / Z r = 1.14 and average Hf*/ Hf = 1.13), which are also distinctive features of many recent boninite series (Hickey & Frey 1982). Although the chemistry of the NK greenstone belt boninite series has many of the characteristic attributes of their Phanerozoic high-Ca counterparts, it is surprising that the major and trace element abundances of the Iringora ophiolite-like units are practically indistinguishable from groups I and II of the Upper Pillow Lavas of the Troodos ophiolite (Cameron 1985).

In summary, the Iringora ophiolite complex is assigned to the 2.80-2.77 Ga subduction episode, recorded in the Belomorian mobile belt (e.g. Shchipansky et al. 2001, 2004). The Olenegorsk greenstone belt, located in the KolaNorwegian terrane of the Kola Province, is composed of metavolcanic rocks that range in composition from basalts to rhyodacites and metasediments represented by garnet-biotite (sometimes with staurolite or sillimanite) gneisses and economic, banded iron formations (Radchenko et al. 1994, and references therein). Sulphide-magnetite quartzites also host gold-tellurium ore occurrences (Ivanyuk et al. 1999). The age of the felsic metavolcanic rocks determined by the U - P b zircon method is 2760+7 Ma (Kudryashov et al. 2001, and references therein). The belt was deformed and metamorphosed to amphibolite grade shortly after the deposition of the supracrustal rocks, as shown by the ages of 2738 _+ 6Ma, obtained for postdeformational gabbro-norite dykes (Kudryashov et al. 2001, and references therein). G r e e n s t o n e belt s e q u e n c e s f o r m e d b e t w e e n 2.75 a n d 2.65 G a

Rocks of this age are known in the Ilomantsi greenstone belt in Finland, where they have been thoroughly studied because of their demonstrated potential for gold mineralization (Nurmi & Sorjonen-Ward 1993), and in the Khedozero-Bolsheozero greenstone belt, where the age of the rhyolites is 2730+6 Ma (Samsonov et al. 2001). In the northwestern part of the Belomorian mobile belt, high-grade supracrustal rocks of the SuomujS.rvi complex also have a maximum estimated age of deposition of 2731 + 8 Ma, and detrital zircons extracted from quartzites have ages up to 3413 • 5 Ma (Evins et al. 2002). In the Ilomantsi greenstone belt, the Hattu structure is composed dominantly of sedimentary-volcanic rocks and volcanic rocks are scarce (Nurmi & Sorjonen-Ward 1993; Sorjonen-Ward & Luukkonen 2005; Sorjonen-Ward et al. 1997); therefore, it is often referred to as the Hattu schist belt. Despite high strain, it has been possible to reconstruct lithofacies relationships with confidence, and to recognize two proximal to distal, partly overlapping felsic volcanic complexes, and a number of thin mafic to ultramafic horizons. The age of a basaltic to andesitic unit in the stratigraphically lowest part of the sequence is 2754_+6 Ma; however, felsic rock clasts from overlying conglomerates yield an age of 2727 _+ 14 Ma and syntectonic granitoids, which demonstrably intrude the latter, range from 2748___6 to 2724+ 5 Ma (Nurmi & Sorjonen-Ward 1993). Thus, despite the relatively precise isotopic data, there are some age anomalies that need to be reconciled with the geological constraints. The sediments contain detrital zircons with an age of up to 2.86 Ga, and their ~Nd(t) varies from -- 0.6 to § 1.2, which points to a contribution of detritus from basement rocks (Nurmi & SorjonenWard 1993). Sedimentary rocks are represented dominantly by feldspathic greywackes and correspond in composition to Cr-, Ni- and Venriched intermediate to felsic volcanic rocks. This is reasonable in view of the textural immaturity of the sediments and lateral and vertical facies transitions. A smaller contribution from granitoids and mafic volcanic rocks is also demonstrable, with a polymicitic conglomerate component. Thin sulphide-facies iron formations are closely associated with mafic volcanic horizons, and some more extensive magnetite-grfinerite-quartz BIF occur in more distal turbidite facies. Thin basaltic flows and pyroclastic units occur sporadically, with the most significant unit, known as the Pampalo Formation, occurring towards the top of the exposed stratigraphic section. As well as basalts, coarsegrained gabbros may represent sills, and a single komatiitic unit is present, which is distinctive in being a clastic breccia, rather than a coherent lava flow.

ARCHAEAN NUCLEUS, FENNOSCANDIANSHIELD Among the volcanic rocks, komatiitic, tholeiitic and calc-alkaline series have been distinguished. Because of this, as well as the effects of gold-related hydrothermal alteration, chemical data should be treated with caution, but the komatiites appear to be enriched in LREE. The tholeiites fall into two groups: low-Ti, LREE-enriched ( ( L a / S m ) N ~ 1.1-1.9) and varieties with a flat REE pattern. The tholeiites are similar to basalts from island arcs or active continental margins. The andesites, dacites and calc-alkaline basalts are rich in LREE and poor in Ta, Nb and Ti, which is typical of island-arc volcanic rocks (Nurmi & Sorjonen-Ward 1993). The Nurmes paragneiss complex consists of migmatized biotite gneisses with intervals of garnet-biotite gneiss containing distinctive graphite and sulfides (Kontinen 1991; Sorjonen-Ward & Luukkonen 2005). Their geochemical character is consistent with derivation from weathering of felsic and mafic rocks (SiO2 content 6 7 - 6 8 wt%, with Cr, Ni and V enrichment). They are thus similar to Hattu metasediments and it is reasonable to expect that they will yield evidence for deposition within the age range 2750-2730 Ma.

Granitoids As with the greenstone complexes, the various Archaean granitegneiss associations in the Shield can also be classified in terms of several distinct age groupings of granitoids and are described below accordingly, from oldest to youngest.

Granitoids f o r m e d b e t w e e n 3.5 and 3.1 Ga

The oldest known granitoids occur in four widely separated parts of the Karelian Craton, and are represented by tonalitic-trondhjemitic-granodioritic (TTG) rocks. Recent secondary ionization mass spectrometry (SIMS) studies of a trondhjemitic gneiss from Siuru, within the Ranua terrane, approximately midway between the Tojottomaselk~i and Iisalmi localities, have demonstrated a population of 3.5 Ga zircons, with the oldest zircon core giving an age of 3.73 Ga (Mutanen & Huhma 2003). In North Finland, TTG form a small (1.4 x 2.6 km) dome known as the Tojottamanselk~i dome. Tonalitic magmatism is dated at 3.11 Ga, and inherited zircons with ages of 3.16-3.25 Ga and average eNd(t) ~ -- 3.7 suggest a Palaeoarchaean age of the source material (Krrner & Compston 1990, and references therein). TTG rocks from the Iisalmi terrane have a similar age of 3136 _+ 20 Ma, and also originated from a Palaeoarchaean source ( t ~ -- 3.23.4 Ga; Paavola 1986). Inherited zircons with ages up to 3.18 Ga were also found in the Silvevaara granodiorite intrusion, which has an age of 2757 +_ 4 Ma and eNd(t) = --0.4 tO --2.1, and thus the Hattu schist belt indicates the presence of older crust at depth (Sorjonen-Ward & Claour-Long 1993). Mesoarchaean granitoids are also common in the Vodlozero terrane (Fig. 1), which was deformed and metamorphosed repeatedly (of 3.11, 2.86, 2.7-2.65 and 2.5 Ga). In this terrane there are 3166 +_ 14 Ma and 3138 _+ 63 Ma TTG rocks, as well as a leucosome from migmatized amphibolites of 3210_+ 12 Ma (Lobach-Zhuchenko et al. 1993). All the Vodlozero TTG are poor in Rb, Y, Zr, Nb and Ba, have negative Ti, Nb and Ta anomalies, and exhibit fractionated REE patterns (La/YbN = 30-50); Eu anomalies have been reported only from tonalites from northern Finland. eNd(t) values in these TTG are both positive (0.3-1) and negative ( - 4 . 3 to -0.7), with ~ d = 3.2--3.5 Ga and, in gneisses and amphibolites from the Vodlozero complex, eNd(t) ranges from -1.2 to +3, and Nd tbM = 3.0-3.2 Ga (Chekulaev et al. 1997; Lobach-Zhuchenko et al. 2000b).

635

Granitoids a n d gneisses f o r m e d a n d r e w o r k e d b e t w e e n 3.0 a n d 2.8 Ga

Widespread granite generation in the Vodlozero terrane was coeval with the formation of the mafic complexes in the Vedlozero-Segozero greenstone belt. These are the trondhjemites of the Chebino pluton (2985 __+ 10Ma), tonalites of the Lake Chernoe area (2957 ___ 23 Ma) and diorites and granodiorites of the River Kalya area (2971 ___ 11 and 2908 _ 12 Ma; Lobach-Zhuchenko et al. 1999). All these granitoids exhibit posiNd around 3.0 Ga (Chekulaev tive eNd(t) values (2.3-4.2) for tom et al. 1997; Lobach-Zhuchenko et al. 1999, 2000b). In this part of the Karelian Craton, younger TTG rocks also occur as small plutons, such as the pre-2884 Ma Lizhmorechensky pluton and the 2859 +_ 24 Ma Shilos pluton, which intruded older greenstone complexes and TTG gneisses (Lobach-Zhuchenko et al. 1999). In addition, they are represented by migmatites in the Kianta terrane (Sorjonen-Ward & Luukkonen 2005) of the Karelian Craton (2843 +__ 18 Ma; Luukkonen 1988) and in the Voknavolok-Ilomantsi terrane, where trondhjemitic gneisses near Lake Verkhneye Kuito have an age of 2887 +__24 Ma (Samsonov et al. 2001). The Vodlozero terrane also contains the oldest known twofeldspar granites in the Shield, dated at 2876 +__21 Ma (Kovalenko & Rizvanova 2000). These granites have an 1-type character and have poorly fractionated REE patterns ( L a / S m N - - 3 . 5 - 5 ; G d / L u N = 1.6-2) and negative Eu anomalies (Eu/Eu* = 0.350.60). Variations in eNd(t) from - 1 . 5 to +1.8 are caused by the heterogeneity of the source (Lobach-Zhuchenko et al. 2000b). In the western part of the Belomorian mobile belt, granitoids of this age group include the 2803 Ma quartz metadiorites and trondhjemites from the Lake Keret area and 2826 __ 18 Ma dioritesgranodiorites from the Hisovaara area (Bibikova et al. 1999, 2003a). Tonalite gneisses, dated at 2.81 Ga, are present along the northernmost margin of the Belomorian mobile belt (Mitrofanov & Pozhilenko 1991) and as clasts in basal conglomerates in the Voche-Lambina greenstone belt (Kislitsyn 2001). In the Tuntsa area in the northwestern part of the Belomorian mobile belt, granites are dated at 2896+_8 Ma, and granites and tonalites at 2805 __ 4 Ma; syenitic intrusions into the Tulppio greenstone belt are also dated at 2.8 Ga (Juopperi & Vaasjoki 2001). In addition, the SuomujSxvi complex further south is dominated by 2823-2810 Ma tonalites and granodiorites (Evins et al. 2002). In the Murmansk Craton and the Kola-Norwegian domain of the Kola Province, 2.9-2.8 Ga granitoids are widespread and are represented by enderbites, with a compositional range from quartz diorites and tonalites to trondhjemites (Radchenko et al. 1994). The enderbites and associated complementary rocks have low to normal alkaline abundances and tend to be sodic rather than potassic. Age estimations for enderbites from the V e z h e Tundra are of the order of 2830 _+ 70 Ma (Bibikova 1989) and isotopic characteristics preclude derivation from Palaeoarchaean protoliths (Balashov et al. 1992; Timmerman & Daly 1995). The oldest granitoids are tonalite-gneisses from the Kola Superdeep Drillhole in the Kola-Norwegian domain (2.93 Ga, Bibikova et al. 1993). Zircon U - P b ages of 2902 + 9, 2813 -t- 6 and 2803 _ 15 Ma have been obtained from TTG in the northwestern part of the Kola-Norwegian domain in the Kola Province. Detailed studies by Levchenkov et al. (1995) and Vetrin et al. (1995, 1999b) have revealed the presence of two types of TTG. Type I is richer in A1203 (17-22%), poor in Fe, Mg and Ti, rich in Sr, Rb and Ba, and exhibits a fractionated REE pattern in tonalites (La/ YbN = 44-112), whereas in trondhjemites La/YbN is 48-392. Type II TTGs are poorer in A1203 and rich in Fe, Ti, Mg, Mn, P, Zr, Y, Co and total REE, with La/YbN = 7 - 2 5 ; trondhjemites are richer in HREE ( L a / Y b N = 13-41) and show positive Eu anomalies. A similar TTG, dated at 2.83 Ga ( ~ a = 2.85-2.95 Ga;

636

A.I. SLABUNOV ETAL.

aN,i(0 = 0.5--2.5), occurs in the Archaean unit of the Kola Superdeep Drillhole (Vetrin et al. 2002). The granitoids are assumed to have been formed at varying degrees of partial melting of mafic lower crust, with pressure varying over a broad range (Vetrin et al. 1999b).

Granitoids f o r m e d b e t w e e n 2.75 a n d 2.50 Ga

Granitoids of this age are widespread throughout the Shield and fall into four groups. (1) TTG rocks, diorites and enderbites: in the Karelian Craton, these occur predominantly as late tectonic intrusions. In the Belomorian mobile belt, many of the granitic gneisses, which tend to be tonalitic and trondhjemitic in composition, with occasional leucogranites, were intruded and migmatized during the interval 2.78-2.70 Ga (Bogdanova & Bibikova 1993; Bibikova et al. 1995, 1996, 1999; Kaulina & Bogdanova 1999). In contrast, the period from 2.73 to 2.66 Ga included intrusion of hypersthene diorite massifs near Pongoma (2728 • 21 Ma, Levchenkov et al. 1996) and Chupa (2728 ___4 Ma, Glebovitsky et al. 2000) and migmatites and calc-alkaline and tholeiitic-series enderbitic and charnockitic intrusions near Lake Notozero (Volodichev 1990; Glebovitsky 2005) and Lake Kovdozero (Bibikova et al. 1995). The Tulppio belt in the northwestern part of the Belomorian mobile belt also contains abundant tonalites, dated at 2744__5 and 2702_+5 Ma, and granites dated at 2721 __ 15 Ma (Juopperi & Vaasjoki 2001). Small tonalitic, trondhjemitic and dioritic veins in the northwestern Belomorian mobile belt were intruded even later, c. 2.68-2.64 Ga (Bogdanova & Bibikova 1993; Kudryashov 1996). In the southern Belomorian mobile belt, 2.67 Ga subalkaline granitic massifs and related dyke swarms have been recognized (Chekulaev et al. 1997; Glebovitsky 2005). In the western part of the Belomorian mobile belt, near Lake Kichany, rocks of the North Karelian greenstone belt are cut by post-kinematic granites dated at 2674___4 Ma (Drugova et al. 1995). In the Kolvitsa structure (northeastern Belomorian mobile belt) metatonalites have an age of 2708 _+ 10 Ma and Nd tDM = 2.82--2.83 Ga at ~Na(t) = 0.1-0.7 (Balagansky 2002). In the Kola-Norwegian terrane of the Kola Province, monzodiorites are dated at 2720 __+3 Ma, quartz diorites at 2679 __+ 18 Ma (Radchenko et al. 1994), and enderbites at 2656 __ 14 Ma (Pozhilenko et al. 2002, and references therein). In the Tersky-Strel'na domain (Fig. 1), tonalitic gneisses have yielded ages of 2722 __+ 18 and 2692 _+ 19 Ma (Daly et al. 2006). The tDM Nd of all granitoids in the Belomorian mobile belt and Kola Province fall within the time interval 2.93-2.72 Ga (Timmerman & Daly 1995; Bibikova et al. 1999; Lobach-Zhuchenko et al. 2000b; Daly et al. 2001) and almost coincide with the ages of magmatic crystallization. Therefore, these intrusions have practically no older crustal components in their respective source regions. (2) Sanukitoids (Mg-rich diorites) form post-tectonic plutons, dated at 2.75-2.70 Ga (Chekulaev et al., 1997; Bibikova et al. 1997, 2005; Halla 2005; Lobach-Zhuchenko et al. 2005; Samsonov et al. 2005; K~ipyaho et al. 2006) in the Central Karelian and Voknavolok-Ilomantsi terranes of the Karelian Craton. The compositions of the plutons vary from monzodiorite to granodiorite to quartz monzonite and tonalite; homblendite, pyroxenite, monzogabbro are also present. The sanukitoids have high Cr and Ni concentrations and Mg-numbers of 0.45-0.70; they are enriched in Sr, Ba and LREE, show strongly differentiated LREE patterns (La/ Yby > 20) and have no Eu anomalies. They are also commonly associated with lamprophyre dykes. (3) Subalkaline rocks occur as typically post-tectonic syenitic massifs, as at Khizhjarvi in the Karelian Craton, and vary in composition from monzodiorite to leucosyenite. They resemble sanukitoids, but are richer in alkalis and poorer in Mg.

In the Kola Province, in the southeastern part of the Kola-Norwegian terrane, this group seems to include quartz monzonites, quartz syenites and latites with an age of 2657 _ 9 Ma (Balashov et al. 1992) and aNd(t) = --0.8 (Timmerman & Daly 1995). The Porosozero polyphase monzodioritic-granitic massif, dated at 2733 + 6 Ma (Kudryashov et al. 2001), intrudes supracrustal rocks of the Kolmozero-Voron'ya belt. A very distinctive suite of granitoids in the Kola Province is represented by the alkaline granites of the Keivy terrane, which typically consist of aegirine- arfvedsonite-bearing and lepidomelanehastingsite-bearing granites (Radchenko et al. 1994). These rocks form sheet-like bodies and dykes that have been dated by the U - P b zircon method at 2630 _+ 31 and 2654 _+ 5 Ma (Belye Tundry massif, Bayanova et al. 1999; Mitrofanov et al. 2000), and at 2751 • 41 Ma (Ponoy massif, Vetrin et al. 1999a). Syenitic granites from the West Keivy massif have an age of 2674 _ 6 Ma (Mitrofanov et al. 2000). Ponoy alkaline granites are rich in Fe, poor in P and Sr, very poor in Cu, Ni, V, Cr and Co, and rich in Li (up to 1000 ppm), Zr, Nb, Y, U, Th and REE (Vetrin et al. 1999a). Their aNd(t) values are 0.1-2.9 (~;~ = 2.64-2.91 Ga), although one sample has aNd(t) = --5.9 ( t ~ = 3.62 Ga). These granites are defined as anorogenic. A final notable feature of Neoarchaean evolution in the Shield was the emplacement of the Siilinj/irvi carbonatite into the Iisalmi terrane, near the western margin of the Karelian Craton, at 2.58-2.61 Ga (Patchett et al. 1981). (4) Two-feldspar granites are ubiquitous in this age group, occurring as post-tectonic plutons and gently dipping sheets dated at 2680-2710 Ma. They have both I- and A-type affinities and aNd(t) values depend on the age of the terrane, ranging from - 0 . 4 to - 4 . 9 for plutons in the old Vodlozero terrane to 0.1 to -1.2 in the Voknavolok-Ilomantsi terrane; in the relatively young Central Karelian terrane the values are as high as +0.8 to +2.2 (Lobach-Zhuchenko et al. 2000b). In the Kola Province similar types of granitoids are represented by the monazite-bearing granites (2634 ___ 12 Ma) in the southeastern Kola-Norwegian domain (Balashov et al. 1992).

Characteristics of the Archaean metamorphism of supracrustal sequences, granulitic and eclogite-bearing complexes A r c h a e a n m e t a m o r p h i s m in the K a r e l i a n Craton

The oldest rocks of the Karelian Craton have been repeatedly and variously metamorphosed, including under high-temperature amphibolite-granulite-facies conditions. For example, 3.2-3.1 Ga granulite metamorphism has been recognized in the 3.54 Ga rocks of the Vodlozero complex, which were subsequently metamorphosed at amphibolite (2.86 Ga) and epidote-amphibolite (2.7 Ga) grade (Glebovitsky 2005, and references therein). In addition, a Neoarchaean (2.65 Ga) granulite metamorphic event was identified within this terrane (Glebovitsky 2005, and references therein). The main stages of metamorphism of rocks in the Karelian Craton coincided with the main stages of formation of greenstone, paragneiss and granitiod complexes. Supracrustal and granitegneiss complexes are usually polymetamorphic. Their P - T - t evolutionary trends typically include the following two stages: (1) early (3.0-2.75 Ga) low-pressure metamorphism; (2) late (2.72-2.65Ga) high-pressure dynamothermal metamorphism (Fig. 4a). The earlier stage is recorded throughout all greenstone belts, regardless of age, and is manifest by greenschist- to lowpressure amphibolite-facies assemblages with either andalusite or sillimanite. The later stage is more typically developed in discrete domains, particularly in zones of transpression. The metamorphic grade tends to be higher-pressure amphibolite facies with kyanite-sillimanite-bearing assemblages. However, even

ARCHAEAN

NUCLEUS,

(a)

FENNOSCANDIAN

SHIELD

637

(b) 16 P. Kbar

P. Kbar

14 12

E

~

//

~""

.2[/" /-

10 .7 Ga

~

//"

S~171-

2.73 Ga

4

2(J0

400

600

800

T, ~

2(t(/

J 4~t

4102~3~4

I 600

....

J 8o0

~1, T, ~

/?5,/6

P-T diagram for the Archaean metamorphic events in the rocks of (a) the Karelian Craton and (b) the Belomorian mobile belt (Volodichev 1990; Lobach-Zhuchenko et al. 1995). Numbers with 'Ga' show the age of the metamorphic processes. In (a), the colours indicate the evolution of different terranes of the Craton. In (b): 1, earliest granulitic metamorphism; 2, Neoarchaean granulitic metamorphism; 3, Neoarchaean eclogitic metamorphism; 4, Neoarchaean collisional metamorphism; 5, evolution of metamorphism in the western part of the Belomorian mobile belt; 6, evolution of the Neoarchaean metamorphism in the eastern part of the Belomorian mobile belt. Fields of metamorphic facies: Gr, greenstone; A, amphibolitic; G, granulitic; E, eclogitic.

F i g . 4.

within single shear zone, pressures may vary greatly, from 5 - 7 to 10 kbar. At least two stages of transpressive tectonics and metamorphism can be distinguished in the Karelian Craton. Metamorphism occurred simultaneously with development of an early regional system of shear zones (transpressional and transtensional), with granitoid (including sanukitoid) intrusions formed at depth in the crust and pull-apart basins near the surface. The evolution of second-generation shear zones was accompanied by metamorphism and the formation of subalkaline granitoids (Volodichev et al. 2002). These Neoarchaean metamorphic and tectonic processes in the Karelian Craton were accompanied by hydrothermal and metasomatic alteration processes and gold mineralization (Nurmi & Sorjonen-Ward 1993), and are presumably a consequence of the coeval Neoarchaean (2.72-2.65 Ga) collisional tectonics in the Belomorian mobile belt. Several granulite (or granulite-enderbite-charnockite) complexes are known in the Karelian Craton (VarpaisjSxvi, Tulosozero, Voknavolok and Onega complexes), in the Belomorian mobile belt (Notozero complex) and in the Kola Province (Fig. 1). They all have some features in common, in that they are dominated by enderbites of dioritic to tonalitic composition; supracrustal rocks are less abundant but include mafic and intermediate granulites, and aluminous garnet-cordierite-sillimanite gneisses. Charnockites do not occur in all granulitic complexes. The Varpaisjarvi complex, which has been studied most thoroughly, is used as an example to discuss their structure and evolution. The Varpaisj~irvi complex occurs among migmatized TTG granitoids of the Iisalmi terrane, dated at c. 3.1 Ga (Paavola 1986) and includes enderbites (with sporadic anorthositic layers) and mafic, intermediate and aluminous granulites. Dioritic to tonalitic enderbites predominate and are transitional to surrounding granitoids, or locally intrude them. Magmatic crystallization of the enderbites took place at 2.72-2.70 Ga ago; they were then metamorphosed at P = 9-11 kbar and T = 800-900 ~ probably at around 2.63 Ga, followed by decompression and cooling at 7 kbar and 700 ~ (H61tt~i et al. 2000; M/intt~iri & H61tt~i 2002). The Proterozoic Svecofennian thermal and tectonic overprint on the VarpaisjSxvi complex is probably relatively minor, whereas the adjacent Rautavaara supracrustal gneiss complex may record granulite- to amphibolite-facies exhumation and re-equilibration during Proterozoic time.

Two-pyroxene and hornblende granulites (often with garnet) occur as small lenticular bodies within the TTG rocks and enderbites. No primary protolith textures are preserved, but geochemically they can be classified into two distinct groups of basaltic and andesitic composition: (1) tholeiitic-series basalts with a chondritic Ti/Zr ratio (c. 110), flat REE patterns (sometimes poor in LREE), with tDM Na = 3.1--3.2 Ga; (2) basalts and andesites with Ti/Zr ratios < 100, enriched in LREE, and with Nd model ages of 2.7-2.9 Ga. It should be noted that the latter group is more common in the southeastern part of the complex (Hrltt~i & Paavola 2000; Hrltt~i et al. 2000). The U - P b zircon ages of the protoliths of granulite from group 1 are 3.05-3.2 Ga, whereas those from group 2 are 2.65-2.68 Ga. The enderbites and granulites of both groups were metamorphosed to granulite grade at 2.62-2.70 Ga (M~intt~iri & Htltt~i 2002). Aluminous granulites are related spatially to the Group 2 mafic granulites and have similar Nd model ages (2.7-2.8 Ga). Zircon cores usually give ages of 2.70-2.81 Ga, whereas newly-formed metamorphic rims have ages of 2.60-2.68 Ga (MS_ntt~iri & H61tt~i 2002). The S m - N d isochron ages of garnet from aluminous and marie granulites, corresponding to cooling to T ~, 600 ~ are 2.59 + 0.01 and 2.52 + 0.05 Ga, respectively (M~intt~iri & Hgltt~i 2002). The Varpaisj~irvi complex is assumed to have been formed upon accretion of two terranes differing in composition and age (M~intt~iri & H61tt~i 2002).

Neoarchaean metamorphism of Belomorian mobile belt rocks The most distinctive feature of the Belomorian mobile belt is the two superimposed Neoarchaean and Palaeoproterozoic high-pressure (including eclogite-facies) metamorphic events (Volodichev 1990; Glebovitsky et al. 1996). The Neoarchaean metamorphism occurred in different ways in the eastern and western domains of the Belomorian mobile belt (Fig. 4b). The eastern domain records a clockwise P - T - t evolutionary trend, which is particularly well expressed in rocks from the Gridino zone (Fig. 1) and the Pongoma Bay area, where prograde metamorphism culminated at eclogite-facies conditions ( P = 1 4 - 1 7 . 5 k b a r , T = 7 4 0 - 8 6 5 ~ followed by

638

A.I. SLABUNOVETAL.

multistage retrogression. This included near-isothermal decompression from 14.0kbar to 6.5 kbar and cooling from 770 to 650 ~ (Fig. 4b). Thus, the conditions correspond to high-pressure granulite-facies conditions for mafic garnet granulite assemblages, transitional to amphibolite-facies conditions (Cloos 1993). The granulite-facies conditions (P = 6 - 7 kbar, T = 750-800 ~ associated with migmatitic and intrusive enderbites, and moderate-pressure amphibolite facies were maintained at 2.72 Ga (Levchenkov et al. 1996) (Fig. 4b). At 2691 + 5 M a (Levchenkov et al. 2001) a return to higher-pressure conditions (P = 9-11 kbar, T = 650-700 ~ occurred. This metamorphic event is assumed to have been caused by transpressive tectonics during the Belomorian collisional orogeny. In contrast, the western domain (e.g. in the Lake NotozeroLake Kovdozero area) is characterized by an 'anticlockwise' P - T - t trend (Fig. 4b). The earliest event, dated at 2820 _+ 5 Ma (Bibikova et al. 2004), was moderate-pressure granulite-facies metamorphism (P = 5.5-6.5 kbar, T > 700 ~ reported from the Chupa paragneisses and from mafic granulites (LobachZhuchenko et al. 1993). A granulite complex, consisting of mafic granulites, enderbites and charnockites of calc-alkaline and Fe-tholeiitic series (Volodichev 1990), was formed later, at c. 2.72 Ga (Glebovitsky 2005). High-pressure metamorphism, up to 9-11 kbar at 650-700 ~ was accompanied by intense migmatization and occurred somewhat later, around 2.7 Ga (Volodichev 1990; Glebovitsky et al. 1996, 2000). The latter episode is common to both western and eastern domains and reflects a response to collisional processes (Searle & Rex 1989). A similar P - T evolution characterizes the northern part of the Belomorian mobile belt, where the first stage of the Neoarchaean metamorphism was at 620-680 ~ and 7.6-8.8 kbar whereas the later event occurred at 665-695 ~ and 9.7-10.6 kbar (Belyaev & Petrov 2002). An eclogite-bearing complex is present in the Gridino zone (Volodichev et al. 2004) and Salma area (Shchipansky et al. 2005) in the eastern and northern parts of the Belomorian mobile belt (Fig. 1). It can be followed as a zone of intensely migmatized mrlange from NW to SE for about 50 km and has a width of up to 10 km (Fig. 5a). The granitoid components within the eclogite complex have been transformed to granite gneisses by multiple deformational and metamorphic events and are of tonalitic-trondhjemitic composition. Relict enderbites, indicative of granulite-facies metamorphism, are also present. The mrlange comprises a chaotic mixture of diverse lithologies (Fig. 5b), including eclogites, gamet, garnet-clinopyroxene and feldspathic amphibolites, meta-ultramafic rocks, metagabbroids and zoisite rocks (meta-anorthosites), aluminous and amphibole-bearing gneisses, and marbles. They are very heterogeneous in terms of composition, the primary nature of the protolith and the metamorphic history. Archaean eclogites, consisting of omphacite (28-33 to 40% jadeite) and garnet (22-30% pyrope and 22-30% grossular; Fig. 5c), are preserved as relics among symplectitic eclogites and garnet-clinopyroxene amphibolites derived from the ecologites during retrograde decompression. The eclogites were formed at pressures of 14.0-17.5 kbar and temperatures of 740865 ~ corresponding to depths to 60-65 km. The U - P b age of zircons from the eclogites and symplectite eclogites is 2720.7-+- 5.8 Ma (Bibikova et al. 2003b; Volodichev et al. 2004). The morphology of the analysed zircon is characteristic of high-pressure granulites and eclogites (Bibikova 1989). The mrlange zone is cut by post-tectonic trondhjemite veins dated at 2701.3 • 8.1 Ma (Bibikova et al. 2003b) and gabbro-norite dykes (2.4 Ga). The eclogites correspond petrogeochemically to tholeiitic-series mafic rocks (47-51% SiO2, 1.38-4.3% Na20 + K20) (FeO*/ MgO = 0.5-2.5). REE abundances are 2 - 1 2 times that of chondrites, and they typically show flat or poorly fractionated REE patterns (La/Smy = 0.99-1.8; Ga/YbN = 0.77-1.17). Compositions

correlate with the metabasalts (amphibolites) from the ophiolite complexes, including from the Central Belomorian greenstone belt (Slabunov 2005). The recognition of Archaean crustal eclogites in the Belomorian mobile belt clearly provides a strong argument in favour of possible deep subduction during the Neoarchaean, to explain such high pressures under a relatively low geotherm (see Yardley 1989).

M e t a m o r p h i s m in the K o l a P r o v i n c e a n d in the M u r m a n s k Craton

The available data (Avakyan 1992; Belyaev & Petrov 2000) show that in the granulitic zone of the Kola-Norwegian terrane (the Kola Province) maximum pressure estimates for cordierite-free parageneses from sillimanite gneisses are 6.2 _+ 1.2kbar at T = 700 ~ whereas in the zone of transition to amphibolite facies, peak pressures were 5.2 • 0.9 kbar. Cordierite-bearing paragneses show lower pressures: 4.5 _+ 0.6 kbar at 700 ~ For rocks with kyanite-sillimanite assemblages in the transition zone, P = 5.3 kbar and T = 580 + 20 ~ Multiple (retrograde) metamorphism in the granulit,~-facies zone took place under amphibolite-facies conditions at P : 3.5 + 0.5 kbar and T = 590 ~ R b - S r isotopic data (Avakyan 1992) suggest that the first thermal event within the Kola granulite-facies paragneisses, prior to the granulite metamorphism, occurred at 2880 + 45 Ma (Isr : 0.7005 _+ 0.0004). The age of the earliest granulite metamorphism is estimated at 2.83 Ga (Bibikova 1989; Balashov et al. 1992), whereas the age of metamorphic zircon from later shear zones, also formed under granulite-facies conditions, is 2648 ___ 18 Ma (Balashov et al. 1992). Metamorphic zircons from the granulite-facies paragneisses, sampled near Lake Pulozero, suggest that yet another metamorphic event occurred at 2724 +_ 49 Ma, followed by retrogression under amphibolitefacies conditions at 2640 + 20 Ma (Pozhilenko et al. 2002). The last recorded Archaean tectonometamorphic event is related to emplacement of discordant pegmatite bodies at 2556 ___27 Ma (Balashov et al. 1992). At least three metamorphic events can be distinguished (Belyaev & Petrov 2000) in the Kolmozero-Voronya greenstone belt. Early metamorphism (2.83-2.76Ga) took place under epidoteamphibolite- and amphibolite-facies conditions at low temperature and low pressure (T = 460-560 ~ P = 2.5-4.3 kbar). A second event (2.76-2.68 Ga) occurred at the same temperatures, but at higher pressures (T = 470-530 ~ P = 3.9-5.8 kbar). During the final stages of Archaean evolution (2.68-2.52Ga), even greater temperatures and pressures were attained in late shear zones (T = 530-640 ~ and P = 6.0-8.5 kbar). In rocks of the Murmansk Craton, the earliest metamorphic events occurred under granulite-facies conditions, at temperatures up to 750 ~ and pressures of 4 - 6 kbar (Petrov et al. 1990). Subsequent metamorphic reworking took place under amphibolitefacies conditions and was accompanied by migmatization.

Summary of the main stages of crust formation The earliest 'sialic crustal nuclei' were the 3.5-3.1 Ga Vodlozero terrane in the southeastern Karelian Craton and two small blocks in central and northern Finland, at the western margin of the Karelian Craton (Figs 1 and 6). In the period 3.1-2.95 Ga, new crust accreted around the Vodlozero terrane and, to a lesser extent, in the northwestern part of the Karelian Craton. This stage is reflected most completely in the Vedlozero-Segozero greenstone belt of the Vodlozero terrane (Fig. 6), which provides the oldest evidence of orogeny in the Shield (Fig. 6). An ensialic (mature) volcanic arc and a

ARCHAEAN NUCLEUS, FENNOSCANDIAN

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Fig. 5. (a) Location of the Gridino eclogite-bearing complex. (b) Schematic geological map of the southeastern part of Stolbikha Island (by A. I. Slabunov & O. S. Sibelev, in co-operation with O. I. Volodichev). (c) Photomicrograph showing a thin section of eclogite with homogeneous non-zoned garnet crystals (Grt; Prp26 = pyrope content of Grt) and omphacite content (Omp32; number shows jadeite content). Secondary alteration is represented by formation of plagioclase (P122; number shows anorthite content), diopside (Di16; number shows jadeite content) and pargasitic hornblende (Prg-Hbl).

deep-water back-arc basin evolved there between 3.05 and 3.0 Ga (Svetov 2005). Mantle plumes gave rise to oceanic-plateau type basalts and komatiites, which were obducted later onto the continent (the Sumozero-Kenozero greenstone belt), and a subcontinental-plateau basalt type in the South Vygozero greenstone belt (Arestova et al. 2003). The first record of crustal growth in the Voknavolok-Ilomantsi terrane also dates from this time ( ~ not more than 3.1 Ga). In the period 2.95-2.85 Ga, the continental crust grew mainly in the southeastern part of the present Karelian Craton, adjacent to the Belomorian belt and in the Kola Province. In the western Vodlozero terrane, subduction produced a volcanic arc with calc-alkaline series volcanic rocks (the Vedlozero-Segozero

greenstone belt) The formation of the oldest two-feldspar granites records the final stages of evolution of this accretionary event. However, the major sialic crustal formation event related to subduction took place somewhat later (2.88-2.82 Ga), as recorded in the Sumozero-Kenozero greenstone belt at the margin of the Vodlozero terrane and in the North Karelian greenstone belt in the Belomorian belt. There are fragments of c. 2.88 Ga oceanic crust in the Belomorian mobile belt. Continental crust was formed simultaneously in the central part of the Karelian Craton, in the Murmansk Craton and Kola Province (~dM is commonly < 2 . 9 Ga). In the period 2.82-2.75 Ga, a new island-arc system was formed and began to evolve along the northeastern boundary of

640

A.I. SLABUNOV ET AL .

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Fig. 6. Correlation of Archaean greenstone, paragneiss, magmatic complexes and metamorphic events in the eastern Fennoscandian (Baltic) Shield (terranes: CK, Central Karelian; Ii, Iisalmi; IV, Ilomantsi-Voknavolok; Ke, Keivy; Ki, Kianta; KN, Kola-Norwegian; P, Pomokaira; R, Ranua; Vo (VS) Vodlozero (Vedlozero-Segozero greenstone belt); Vo (Core), core of the Vodlozero.

the Karelian Craton with the Belomorian mobile belt. It has survived as fragments of various volcanic rocks (including boninites and adakites) and supra-subduction ophiolites in some of the greenstone belts in both the Karelian Craton and the Belomorian mobile belt. In the western part of the Karelian Craton, supracrustal complexes were generated (the Kuhmo-Suomussalmi-Tipasj/irvi and Kostomuksha belts). Some workers (Luukkonen 1988; Lobach-Zhuchenko et al. 2000a) have interpreted them as riftogenic, whereas others (Piirainen 1988; Puchtel et al. 1998; Kozhevnikov 2000; Samsonov et al. 2001) have considered them in terms of accretionary-collisional and mantle plume tectonic processes. Much of the present Karelian Craton had thus been formed by 2.75 Ga, when this new system accreted to the older amalgamation. Between 2.75 and 2.65 Ga, the growth of the crust continued in the Belomorian mobile belt and culminated in the collision of microcontinents, which gave rise to an orogen with thick continental crust (Glebovitsky 2005; Slabunov 2005). During the final stage of subduction and the start of collision (c. 2.72 Ga) eclogites of the Belomorian mobile belt were exhumed. The intrusion of the 2.73-2.70 Ga Notozero intrusive complexes and their analogues, which correspond geochemically to active continental margin rocks, occurred in the zone between the Karelian Craton and the Belomorian mobile belt, and in the central Karelian Craton, pullapart basins evolved. They were filled with sediments, and felsic and intermediate volcanic rocks, and are now represented by the Ilomantsi, Khedozero-Bolshozero and Gimoly belts. At 2.742 . 7 0 G a subalkaline and sanukitoid granitoids were intruded during closure of these basins. The Kola Province became a collage of exotic terranes (such as the Kola-Norwegian and Keivy terranes) by the end of the Neoarchaean.

In summary, the formation of the Archaean sialic crust of the Fennoscandian (Baltic) Shield can be understood in terms of the subduction and collision of lithospheric plates (Gafil & Gorbatschev 1987; Mints 1998), with an additional influence from mantle plumes (Lobach-Zhuchenko et al. 1999; Puchtel et al. 1998; Arestova et al. 2003). The architecture of the Archaean continental crust can be attributed to accretionary-collisional processes, with at least four major phases of accretion and one collisional event being recognized (Fig. 6). The SVEKALAPKO Project of the EUROPROBE programme initiated this work. This paper is a contribution to programmes O N Z - 6 "Geodynamics and mechanisms of lithosphere deformation" and O N Z - 8 "Isotope systems and isotope fractionation in natural processes". We acknowledge financial support from the Russian Foundation for Basic Research (RFBR) (grants 00-05-64295, 00-05-64701, 03-05-64010, 03-05-64501 and 06-05-64876).

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Ilomantsi, eastern Finland. In: NURMI, P. & SORJONEN-WARD, P. (eds) Geological development, gold mineralization and exploration methods in the late Archean Hattu schist belt, Ilomantsi, Eastern Finland. Geological Survey of Finland, Special Papers, 18, 25-29. SORJONEN-WARD, P. & LUUKKONEN, E. 2005. Archean rocks. In: LEHTINEN, M., NURMI, P. A. & RAM0, O. T. (eds) The Precambrian Geology of Finland--Key to the Evolution of the Fennoscandian Shield. Elsevier, Amsterdam, 19-99. SORJONEN-WARD,P., NmONEN, M. & LUUKKONEN,E. 1997. Greenstone associations in Finland. In: DE WIT, M. J. & ASHWAL,L. (eds) Greenstone Belts. Oxford Monographs on Geology and Geophysics, 35, 677-698. STEPANOV, V. S. & SLABUNOV,A. I. 1989. Precambrian amphibolites and early basic-ultrabasic rocks in northern Karelia. Nauka, Leningrad [In Russian]. STEPANOV, V. S., SLABUNOV, A. I. & STEPANOVA, A. V. 2003. Rock-forming and accessory minerals of Late Archaean peridotites from the Lake Seryak area, Belomorian mobile belt, Fennoscandian Shield. Geology and commercial minerals of Karelia, Issue 6. KarRC RAS, Petrozavodsk, 17-25 [in Russian]. SVETOV, S. A. 1997. Komatiitic-tholeiitic associations of the VedlozeroSegozero greenstone belt, Central Karelia. Karelian Research Centre, Petrozavodsk [in Russian]. SVETOV, S. A. 2003. New data on geochemistry of the oldest 2.953.05 Ga andesite association in Eastern Fennoscandia. Doklady Earth Sciences, Moscow, 389(2), 195-198. SVETOV, S. A. 2005. Magmatic systems in the ocean-continent transition zone in the Archean of the eastern Fennoscandian Shield. Karelian Research Centre, Petrozavodsk [in Russian]. SVETOV, S. A., SVETOVA,A. I. & HUHMA, H. 2001. Geochemistry of the komatiite-tholeiite rock association in the Vedlozero-Segozero Archean greenstone belt, Central Karelia. Geochemistry International, Moscow, 39 (Supplement 1), $24-$38. SVETOVA, A. I. 1988. Archean Volcanism in the Vedlozero-Segozero Greenstone Belt of Karelia. Karelian Research Centre, Petrozavodsk [in Russian]. TAIPALE, K. 1988. Volcanism in the Archean Kuhmo greenstone-granite terrane in the Tipasj~irvi area, Eastern Finland. In: MARTTILA, E. (ed.) Archaean geology of the Fennoscandian Shield. Geological Survey of Finland, Special Papers, 4, 151-160. THURSTON, H. C. & KOZHEVNIKOV, V. N. 2000. An Archean quartz arenite-andesite association in the eastern Baltic Shield, Russia: implications for assemblage types and Shield history. Precambrian Research, 101, 313-340. TIMMERMAN, M. J. & DALY, J. S. 1995. S m - N d evidence for late Archaean crust formation in the Lapland-Kola Mobile Belt,

Kola Peninsula, Russia and Norway. Precambrian Research, 72, 97-107. TULENr~EIMO, T. 1999. The Kellojgirvi layered ultramafic complex in the Kuhmo greenstone belt. MSc thesis, University of Turku [in Finnish]. VAASJOKI, M., TAIPALE, K. & TUOKKO, I. 1999. Radiometric ages and other isotopic data bearing on the evolution of Archaean crust and ores in the Kuhmo-Suomussalmi area, eastern Finland. Geological Survey of Finland Bulletin, 71, 155-176. VETR~N, V, NORDCULEN, 0 , COBBINO, J., STURT, B. & DOBRZHINETSKAYA,L. 1995. The pyrozene-bearing tonalite-granodiorite-monzonite series of the northern Baltic Shield: correlation and petrology. In: ROBERTS, D. & NORDGULEN, 0. (eds) Geology of the Eastern Finnmark-Western Kola Peninsula Region. Norges Geologiske UndersCkelse, Special Publications, 7, 65-74. VETRIN, V. R., KAMENSKY, I. L., BAYANOVA,T. B., TIMMERMAN, M., BELYATSKY, B. V., LEVSKY,L. K. & BALASHOV,Yu. A. 1999a. Melanocratic nodules in alkaline granites of the Ponoiskii massif, Kola Peninsula: a clue to petrogenesis. Geochemistry International, Moscow, 37(11), 1061-1072. VETRIN, V., TURKINA,O. & NORDGULEN,0. 1999b. Surface analogues of 'grey gneiss' among the Archaean rocks in the Kola Superdeep Borehole. Kola Science Centre, Apatity. VETRIN, V. R., TURKINA, O. M., LUDDEN,J. & DELENITSIN, A. A. 2002. Correlation and petrology of the basement rocks of the Pechenga palaeorift. In: MITROFANOV,F. P. (ed.) Geology and mineral resources of the Kola Peninsula. 2. Mineral resources, mineralogy, petrology, geophysics. Kola Science Centre, Apatity, 208-230 [in Russian]. VOLODICHEV, O. I. 1990. The Belomorian Complex of Karelia (Geology and Petrology). Nauka, Leningrad [in Russian]. VOLOI3ICIqEV, O. I., KUZENKO, T. I. & KOZLOV, S. S. 2002. On the structural-metamorphic study of Kontokki Series volcanics from the Kostomuksha structure. Geology and commercial minerals of Karelia, Issue 5. Karelian Research Centre, Petrozavodsk, 15-26 [in Russian]. VOLODICHEV, O. I., SLABUNOV,A. I., BIBIKOVA,E. V., KON1LOV,A. N. & KUZENKO,T. 2004. Archaean eclogites in the Belomorian mobile belt, Baltic Shield. Petrology, Moscow, 2, 540-560. VREVSKY, A. B. 1989. Petrology and geodynamic regimes of the evolution of the Archaean lithosphere. Nauka, Leningrad [in Russian]. VREVSK',', A. B. 2002. Petrology and geochemistry of komatiites and isotope evolution of their mantle sources: implications for the mantle plumes evolution beneath the Archaean lithosphere of the Fennoscandian Shield. In: GRACHEV, A. F. (ed.) Mantle plumes and metallogeny, Extended Abstracts. Probel-2000, Moscow, 500-504. YARDLEY, B. W. D. 1989. An Introduction to Metamorphic Petrology. Longman, Harlow.

Archaean terranes, Palaeoproterozoic reworking and accretion in the Ukrainian Shield, East European Craton S. CLAESSON 1, E. BIBIKOVA 2, S. BOGDANOVA 3 & V. SKOBELEV 4

1Swedish Museum of Natural History, Box 50 007, SE-I04 05 Stockholm, Sweden (e-mail: [email protected]) 2Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, Kosygin str. 19, Moscow 119991, Russia 3Department of Geology, Srlvegatan 12, SE-22362 Lurid, Sweden 4Institute of Geochemistry, Mineralogy and Ore Formation of the National Academy of Sciences, Ukraine, Palladin Ave., 34, 03680 Kiev-142, Ukraine

Abstract: The UkrainianShield is a large coherentregion of exposed Archaeanand Proterozoic crust in the southwestern, Sarmatian part of the East EuropeanCraton. It is traditionallydividedinto blocks, or domains, separatedby major suture zones. The Azov Domain in the east and the PodolianDomain in the SW are Archaean complexes that have been highly reworked in the Palaeoproterozoic; in contrast, the Archaean (3.2-3.0 Ga) granite-greenstone terrane dominated Middle Dniepr Domain, in the central part of the Shield, is virtually untouched by Proterozoic processes. Palaeoproterozoicrocks dominate the Kirovograd domain in the central Shield. We review previous and recent geochronologicalresults and demonstrate that the Volyn Domain and adjacent parts of the Ros-Tikich Domain in the NW are largely juvenile, c. 2.2-2.0 Ga segments of Palaeoproterozoiccrust accreted to the Palaeo- to Mesoarchaean crust in the Podolian Domain. The Podolian Domain includes 3.65 Ga granitoids, with traces of 3.75 Ga material. It has been reworked, at 2.8 Ga and c. 2.0 Ga. Its temporal evolution is thus similar to that of the Azov Domain in the eastern part of the Shield. However, in view of the complex terrane pattern of Sarmatia, this does not necessarilymean that the Podolian and Azov domains were parts of the same continentin the Archaean.

Critical to all analysis of Archaean and Proterozoic terranes is the establishment of their age relationships by accurate isotope dating. The interpretation of the EUROPROBE transect EUROBRIDGE (extending from Sweden to the Ukraine) and other geophysical data requires knowledge of the precise timing of the major tectonothermal events for reconstructions of the lithosphere geodynamics. Among the three major crustal segments composing the East European Craton (EEC), Sarmatia differs from Fennoscandia and Volgo-Uralia particularly in the timing of major crust-forming events (Bogdanova et al. 1996, 2005) There is no temporal correlation of the crustal histories in Sarmatia and Fennoscandia in the Archaean and the Palaeoproterozoic, and the evolution of Volgo-Uralia appears distinct until the earliest Palaeoproterozoic. Sarmatia occupies most of the southern half of the EEC (Fig. 1). It comprises the Ukrainian Shield and the Voronezh Massif, which are separated from each other by the Palaeozoic Dniepr-Donets Aulacogen with a deeply subsided Precambrian basement. The structural pattern of the Voronezh Massif closely resembles that of the Ukrainian Shield (e.g. Shchipansky & Bogdanova, 1996). Both feature Archaean and Palaeoproterozoic terranes and a distinct pattern of magnetic anomalies. The Precambrian crust of Sarmatia was formed between c. 3.7 Ga and 1.7 Ga. Differences in tectonic pattern, lithology and geochronology allow subdivision into Archaean granulite-gneiss and granite-greenstone terranes, and Palaeoproterozoic mobile belts. The aim of this paper is to give a general picture of the Archaean evolution and Palaeoproterozoic reworking of the oldest parts of the Ukrainian Shield, and present the evidence for the Palaeoproterozoic accretion of new continental crust, based on previously published and recent isotope geochronological results. The main emphasis is on the western parts of the Shield.

Structural outline of Ukrainian Shield Much of our present knowledge of the crust-forming events in Sarmatia derives from the Ukrainian Shield, which is situated at the southern boundary of the EEC (Fig. 1). Geochronological

studies of most major provinces and terranes in the Shield have a long history, starting in the 1950s. During this period much geochronological information, including several books (e.g. Shcherbak et al. 1981; Shcherbak & Bartnitsky 1995), has been published, much of it in Ukrainian and Russian journals, which are not widely accessible in the West. In recent decades, results based on S m - N d and U - P b ion-microprobe investigations have been presented. These new results have demonstrated a need for revision of previous ideas on the geodynamic evolution of the Precambrian in the Ukrainian Shield, particularly in relation to geophysical studies of the crust and upper mantle. The EUROBRIDGE seismic profiling, complemented by other geophysical surveys, has provided new insights into the structure and composition of the lithosphere in the western part of the Shield (Thybo et al. 2003). The Ukrainian Shield is commonly described as several blocks, or domains, separated by suture zones. The Archaean high-grade Azov Domain in the east and the Podolian Domain in the SW were highly reworked in the Palaeoproterozoic (Fig. 1). In contrast, the granite-greenstone Middle Dniepr Domain, in the central part of the Shield, was virtually untouched by Palaeoproterozoic orogenic processes. Palaeoproterozoic rocks compose most of the Kirovograd Domain in the central part of the Shield, and also the R o s - T i k i c h and Volyn domains in its northwestern part. The Orekhov-Pavlograd (OSZ), Krivoy Rog (KSZ) and Golovanevsk (GSZ) suture zones (Fig. 1) have complex tectonic fabrics including strong shearing of rocks from the adjacent domains on both sides of the suture zones. In the OSZ, rocks of the Azov Domain appear to prevail; juxtaposition of rocks from both the Middle Dniepr Domain and the Palaeoproterozoic Krivoy Rog formations characterize the KSZ, whereas both Podolian Archaean rocks and Palaeoproterozoic rocks of the Kirovograd Domain occur in the GSZ. There are two major concepts concerning the Archaean and Palaeoproterozoic tectonic evolution of the Ukrainian Shield. The first envisages a continuous process of reworking of a single Palaeoarchaean continent of 'mafic-tonalitic-trondhjemitic composition' (e.g. Artemenko 1995; Shcherbak & Ponomarenko 2000). It is thought to have been rifted in the Mesoarchaean

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society,London,Memoirs, 32, 645-654. 0435-4052/06/$15.00 9 The GeologicalSociety of London 2006.

645

646

S. CLAESSONET AL.

Bait

25 ~

e Shield

(/I]

33 ~

52 ~ Ukraiaian

Shield

39" ................................... ......................5 0 " L.

'

(b)

......

48 ~

25~ 0

200

km

~ 29 ~

(3.2-2.8 Ga) with the formation of the Middle Dniepr granitegneiss-greenstone terrane, and then again in the Neoarchaean (2.8-2.6 Ga), when the Ros-Tikich granite-gneiss terrane was built up. Subsequent, mainly Palaeoproterozoic rifting (2.6-1.7 Ga) of the Archaean crust caused the development of large grabens in the Kirovograd and Volyn domains, and the interdomainal suture zones. According to the proponents of the alternative theory, involving a plate-tectonics approach to the reconstruction of the geodynamic evolution in the Ukrainian Shield, the tectonic structure of the Ukrainian basement is a collage of Archaean and Palaeoproterozoic terranes (e.g. Kalyaev 1976; Kalyaev et al. 1984; Shchipansky & Bogdanova 1996; Glevassky & Kalyaev 2000; Glevassky & Glevasska 2002). These could have evolved separately, and been amalgamated by accretion to the Palaeoarchaean cores at different times, both in the Archaean and in the Palaeoproterozoic.

The oldest Archaean components The oldest components of the Ukrainian Shield appear in the Azov and Podolian domains (Fig. 1). The oldest identified rocks in the east belong to the Novopavlovsk complex, which appears within the Orekhov-Pavlograd Suture Zone (OSZ), separating the Azov and the Middle Dniepr domains. These have been discovered by deep drilling and consist mainly of ultramafic rocks and tonalites, metamorphosed to amphibolite to granulite facies. The tonalites have been dated at 3.65-3.6 Ga by both ion-microprobe and conventional U - T h - P b dating of zircon (Shcherbak et al. 1984; Bibikova & Williams 1990), and by S m - N d isochron dating (Bibikova & Baadsgaard 1986). A second phase of tonalite emplacement has been dated to 3.4 Ga. The Podolian (Dniestr-Bug) Domain (Fig. 1) is bordered by major deep fault zones in the NE and by the Golovanevsk (GSZ) suture zone in the SE. It is mainly composed of high-grade Archaean and Palaeoproterozoic igneous rocks. The oldest Palaeoarchaean rocks, which belong to the Dniestr-Bug formation, have been found in outcrops along the banks of the

46~ 39 ~

'

Fig. 1. (a) Geologicalsketch map of the East European Craton (EEC), showingits division into the three segments Fennoscandia, Sarmatia and Volgo-Uralia. The Baltic and Ukrainian Shields, with Proterozoic and Archaean rocks that crop out, and the Voronezh Massif, north of the Ukrainian Shield, are also shown. (b) Geologicalsketch map of the Ukrainian Shield, with its subdivisioninto domains and major shear zones indicated. The Paleoproterozoic Osnitsk-Mikashevichi Igneous Belt in the extreme NW is also shown. A, Azov Domain; K, Kirovograd Domain; MD, Middle Dniepr Domain; P, Podolian Domain; RT, Ros-Tikich Domain;V, Volyn Domain; GSZ, GolovanevskSuture Zone; KSZ, Krivoy Rog Suture Zone; OSZ, Orekhov-Pavlograd Suture Zone; OMB, Osnitsk-MikashevichiIgneous Belt; TESZ, Trans-European Suture Zone. The sampling localitiesfor the U-Pb zircon studies in the Palaeoarchaean Dniestr-Bug formation of the Podolian Domain, and near Novograd-Volynskin the Volyn Domain, which are discussedin the text, are shown as rectangles. DB, Dniestr-Bug; NV, Novograd-Volynsk.

southern Bug River and in nearby quarries. They are represented by granulite-facies granitoids, mainly of tonalitic composition (enderbites), and subordinate mafic and supracrustal rocks. The supracrustal rocks are dominated by two-pyroxene, amphibolepyroxene, garnet-pyroxene, garnet-sillimanite and garnet-cordierite granulites. A sequence of granulite-facies metamorphic and structural events has been described from these rocks (Lesnaya 1988; Lesnaya et al. 1995). The emplacement of these granulites into the upper crust was probably accompanied by movements in the deep Tal'nov fault zone (Lesnaya et al. 1995). Multigrain U - P b dating of zircon from the Dniestr-Bug highgrade rocks has given upper intercept ages for discordant zircon fractions of up to 3.1-3.4 Ga (Bibikova 1984; Lesnaya et al. 1995), but did not provide precise results. This was due to the complex structure of the zircons, which commonly show several periods of growth of individual crystals. The old ages, which were regarded as minimum age estimates, are corroborated by whole-rock S m - N d model ages of 3.6-3.9 Ga (Fig. 3; Stepanyuk et al. 1998; Dovbush et al. 2000). We have recently reported new results (Claesson et al. 2003) from U - P b spot analyses of zircons from high-metamorphic enderbites of the Dniestr-Bug region. A full presentation of these results is given here. The analyses were carried out on the Cameca 1270 secondary ion mass spectrometer at the NORDSIM laboratory, Swedish Museum of Natural History, Stockholm. This method permits U - P b isotopic analysis of selected spots in single crystals, and thus has the capacity to resolve the ages of cores and overgrowths in complex zircon crystals. Analytical results and a reference to analytical methods are given in Table 1. Zircons separated from three samples of enderbite collected near the small town of Zavalje, in the Kazachii Yahr and Odessa quarries on the opposite banks of the southern Bug River, have been analysed. These crystals have complex internal structures, usually containing cores with at least one discernible generation of overgrowth. Two main zircon generations were identified. The earliest consists of subprismatic high-U

ACCRETION AND REWORKING OF THE UKRAINIAN SHIELD zircon with zoning of magmatic appearance. The colour, which varies from dark brown to grey and pinkish, appears to reflect different degrees of recrystallization. The magmatic crystals are commonly overgrown by zircon of a second generation. There also occur new transparent, isometric crystals. These younger zircons are discussed further below. Results from the U - P b isotope analyses of these highmetamorphic enderbites from the Podolian Domain are shown in concordia diagrams in Figure 2. The zircon generations differ in isotopic age. Most analyses give discordant U - P b ages, but a few concordant and nearly concordant analyses of the older zircon type show ages of about 3.65 Ga. This can be taken as the age of magmatic zircon generation. The new U - P b isotopic results confirm the previously reported evidence (Bibikova 1984; Lesnaya et al. 1995) that this region includes rocks that have had a complex Archaean evolution, and are among the oldest in the Shield. The morphology of the zircons suggests that the enderbites originally were plutonic rocks rather than metavolcanic rocks, which had been suggested previously. Our interpretation is that the c. 3.65 Ga age dates a crust-forming event. Three analyses in the unzoned core of crystal 28 from enderbite sample 186 gave a nearly concordant age of c. 3.75 Ga (Table 1, Fig. 2). This cannot be used to reconstruct any detailed pre-3.65 Ga history. However, it indicates that zircon-bearing material, probably derived from differentiated continental crust of this age, was available when the enderbite-generating magmas formed c. 3.65 Ga ago. The Nd model ages for rocks from the Podolian Domain (Stepanyuk et al. 1998; Dovbush et al. 2000), which are summarized in Figure 3, must be interpreted with some caution because the strong polymetamorphic reworking may also have disturbed the S m - N d isotope systematics. However, the age pattern in Figure 3, with a dominance of Archaean ages and a maximum at 3.6 Ga, corroborates the U - P b results and strongly indicates that the Podolian Domain is dominated by Archaean crust.

Archaean granite-greenstone terranes In contrast to the highly reworked granulite-gneiss terranes in the Azov and Podolian domains, the Archaean Middle Dniepr Domain in the central part of the Ukrainian Shield (Fig. 1) is a typical granite-greenstone terrane with narrow greenstone belts occurring between large areas of pre-, syn- and post-tectonic granitoids (Orsa 1988). Numerous geochronological results have shown that volcanites of the Konka, Verkhovtsevo and other greenstone belts, and syntectonic TTG granitoids of the Sura complex were emplaced in the same time interval between 3180 and 3000 Ma (Shcherbak & Ponomarenko 2000). Orthogneisses of the Auly series, which have been regarded as the basement for the greenstone belts, have been investigated in detail by both conventional (Shcherbak et al. 1984) and ion microprobe U - P b (Samsonov et al. 1996) dating of zircon. No ages older than 3.2 Ga have been obtained in these studies. These results are supported by S m - N d data (Artemenko 2000). Archaean (3.1-3.0 Ga) TTG gneisses and granitoids, similar to those in the Sura Complex, have been identified in the Kirovograd (Ingul) Domain as the basement to the Palaeoproterozoic supracrustal rocks. Several greenstone belts containing metasediments, metabasalts and metakomatiites have also been described from the Azov Domain (Artemenko 1995). These appear to be approximately coeval with the greenstone belts in the Middle Dniepr Domain, but are more strongly reworked.

647

analogous rocks of the Middle Dniepr Domain. S m - N d model ages as old as 4.1-4.2 Ga have been presented by Artemenko (2000). However, these ages have not been confirmed by other methods and should be interpreted with caution. Our recent, unpublished results show Archaean U - P b ion microprobe ages in the range 2.90-2.93 Ga for post-tectonic intrusions varying in composition from gabbro to granodiorite and even trondhjemite. Two stages of high-grade metamorphism, overprinting the oldest rocks in the Azov Domain, occurred at 2.8 and c. 2.0 Ga (Bibikova & Williams, 1990). High-Mg, post-tectonic granitoids in the Middle Dniepr Domain have recently been shown by us (unpublished results) using ion microprobe U - P b dating to be as old as 2975 Ma. The posttectonic magmatic activity in the region, which is represented by more potassic anatectic granitoids, continued until c. 2.7 Ga (Shcherbak & Ponomarenko 2000). The metamorphism of the rocks in the Middle Dniepr Domain reached amphibolite grade in the Archaean, at least locally. It has long been known that the Palaeoproterozoic overprint in the Middle Dniepr Domain, if present, was mild. This is also supported by Archaean K - A r ages, which were obtained from this region in the early geochronological investigations of the Ukrainian Shield (Shcherbak & Bartnitsky 1995). The K - A r isotope system is easily disturbed by later metamorphic overprint, but these ages have been shown to be relatively reliable. The Archaean evolution of the Podolian Domain, which we infer (above) originated during a crust-forming event at c. 3.65 Ga, is difficult to unravel in detail because of the strong metamorphic reworking. We have identified a zircon generation in the Dniestr-Bug formation enderbites with distinct morphology and optical appearance. It is younger than the previously described 3.65 Ga population and commonly appears as light-coloured, transparent, commonly low-U overgrowths on magmatic zircons. There also occur new low-U, isometric light-coloured crystals with high Th/U ratios and bright lustre. The strong metamorphic reworking has affected the U - P b isotope systematics and makes it difficult to date this zircon-forming event precisely. However, many of the data points for this zircon generation plot near 2.8 Ga (Fig. 2). Our preferred interpretation is that it is 2.8 Ga, and the characteristics of the zircon suggests it formed during granulite-facies metamorphic conditions. The Nd model age distribution for rocks from the Podolian Domain presented in Figure 3 indicates maxima at 3.2 and 2.8 Ga, which corroborates this conclusion. The Ros-Tikich Domain (Fig. 1) is characterized by mostly NNW-trending structures. It is bordered by Palaeoproterozoic north-south-trending zones of strong faulting. New Nd isotope data (Fig. 3; Stepanyuk et al. 1998; Dovbush et al. 2000) suggest a subdivision of this domain into two parts, the southern being dominantly of Archaean and the northern of Palaeoproterozoic origin. The southern (Tikich) part is composed of amphibolites, amphibole- and subordinate garnet-bearing gneisses and migmatites with Nd model ages of 3.2-3.3 Ga (Fig. 3; Stepanyuk et al. 1998; Dovbush et al. 2000), similar to model ages for Archaean rocks in the Podolian Domain (Fig. 3). These have been intruded by Archaean, c. 2.7-2.6 Ga granitoids. A gradual transition of metamorphic grade from amphibolite-facies Tikich rocks to the Dniestr-Bug granulites in the Podolian Domain in the south has been described by Shcherbakov (1975). This suggests that this region may represent a less eroded, more shallow crustal section of the Archaean crust in the Podolian Domain, in agreement with the tectonic interpretation of the EUROBRIDGE-97 seismic profile (Thybo et al. 2003; Bogdanova et al. 2006).

Younger Archaean evolution

Palaeoproterozoic reworking of Archaean crust

Most of the crust in the Azov Domain is occupied by the West Azov and Central Azov formations. Greenschist rock assemblages from these formations are c. 3.0-3.2 Ga, i.e. similar in age to

Palaeoproterozoic processes have reworked and added new material to the Archaean crust in the Azov Domain. The Palaeoproterozoic rocks are dominated by intrusive granitoids,

S. CLAESSON ETAL.

648

Table 1. Ion microprobe U- Th-Pb analytical data for zircons from enderbites, Dniestr-Bug formation, Ukrainian Shield Analysis number

Concentration (ppm) U

Th

Th/U

f2o6" (%)

Pb

Age (Ma) 2ovpb/2O6pb

Isotopic ratios 2~

-+ %

2~

Disc. (%)

-+ %

Sample 46 04a 04b 05a 06a 07a 08a 08b 08c 09a 09b 10a lla llb 14a 18a 18b 19a 19b 21a 22a 23a 23b 25a

271 191 670 2142 1251 502 31 617 852 1005 789 1326 553 198 329 646 1290 1278 38 372 227 48 146

125 75 84 855 62 181 40 253 534 38 38 147 232 74 127 235 486 844 8 40 32 30 285

256 117 543 1731 769 537 17 610 860 959 551 962 521 157 275 671 836 1248 17 236 168 39 130

0.460 0.390 0.130 0.400 0.050 0.36 1.290 0.410 0.630 0.640 0.050 0.110 0.420 0.370 0.390 0.360 0.380 0.660 0.202 0.110 0.140 0.620 1.950

0.00 0.10 0.00 0.10 0.00 0.10 1.00 0.00 0.30 0.10 0.10 0.00 0.60 0.00 0.04 0.10 0.00 0.00 0.10 0.20 0.00 0.20 0.10

0.6574 0.6768 0.6234 0.5857 0.4913 0.7420 0.3625 0.6787 0.6858 0.6422 0.5489 0.5593 0.6633 0.5744 0.5993 0.7233 0.4797 0.6548 0.3776 0.5053 0.5916 0.5945 0.5183

3.15 5.75 3.13 3.11 3.11 3.14 3.13 5.75 5.75 5.75 3.13 3.11 5.75 6.85 6.97 6.86 6.86 6.88 2.12 2.13 2.12 2.18 2,13

27.738 25.213 23.217 20.910 16.106 34.852 6.125 32.451 29.038 27.731 19.594 20.382 28.424 22.359 24.291 33.819 16.204 27.980 7.997 15.724 17.767 17.394 14.582

3.17 5.86 3.15 3.12 3.12 3.15 3.88 5.75 5.76 5.75 3.13 3.13 5.75 6.86 6.99 6.87 6.87 6.89 2.56 2.17 2.14 2.29 2.23

3501 3307 3305 3240 3105 3666 1994 3692 3506 3537 3240 3272 3524 3375 3439 3659 3152 3520 2386 3022 2964 2922 2858

+5 +7 +3 +3 +5 + 40 +3 +4 +2 +3 +5 +3 -+6 -+9 +4 -+4 -+4 + 24 +6 -+5 -+ 11 -+ 11

3.0 1.0 1.0 5.0 16.0 0.0 0.0 0.0 2.0 2.0 10.0 10.0 5.0 5.0 3.0 0.0 13.0 0.0 9.0 11.0 0.0 0.0 2.0

473 214 124 1934 409 584 10 143 3092 908 893 2494 1694 620 705 662 322 449 594 360 205

276 64 49 121 129 180 48 87 601 100 388 1009 440 172 573 348 78 34 51 85 90

443 160 109 1049 415 611 13 118 2952 692 690 1638 1005 424 560 530 190 275 463 319 153

0.584 0.297 0.400 0.063 0.316 0.307 4.617 0.606 0.194 0.111 0.434 0.405 0.260 0.278 0.813 0.525 0.243 0.076 0.086 0.238 0.440

0.14 1.47 0.02 0.00 0.01 0.00 0.37 0.07 0.00 0.02 0.00 0.00 0.01 0.08 0.06 0.14 0.44 0.06 0.08 0.01 0.02

0.6356 0.5514 0.6291 0.4517 0.7194 0.7390 0.4793 0.5732 0.7080 0.5983 0.5713 0.5054 0.4719 0.5289 0.5420 0.5712 0.4686 0.5014 0.6023 0.6498 0.5551

0.97 0.94 0.94 0.94 0.94 0.93 1.18 5.75 0.93 0.93 0.95 0.93 0.94 3.11 3.15 3.30 3.14 5.75 5.75 5.75 5.75

26.883 21.130 25.094 11.3779 32.436 34.347 12.497 22.208 28.992 20.042 17.991 12.787 11.809 15.919 16.661 22.217 12.585 13.984 22.665 27.847 17.021

1.03 1.05 1.12 1.00 0.96 0.96 2.24 5.84 0.95 0.95 1.03 0.95 0.97 3.23 3.18 3.42 3.19 5.76 5.76 5.78 5.76

3505 3351 3413 2677 3603 3650 2734 3368 3455 3139 3041 2685 2666 2968 3001 3374 2783 2845 3323 3525 2998

-+6 -+7 -+9 -+6 _+ 4 _+3 _+ 31 -+ 16 -+3 +3 -+6 ___3 -+4 -+5 -+6 -+ 14 _+9 +7 -+6 _+9 _+ 14

12.0 19.1 9.9 12.3 3.9 3.0 9.3 16.5 0.1 4.6 5.2 2.2 7.9 9.5 8.6 17.0 13.2 9.6 10.7 10.7 6.3

369 281 354 1772 574 1082 517 196 247 170 2683 1362 460 497 432 1378 2129

4 4 5 511 129 50 735 35 169 146 285 136 245 227 119 223 269

345 241 348 1465 488 645 590 160 280 133 2324 1051 340 343 341 1203 1846

0.01 0.01 0.01 0.29 0.22 0.05 1.40 0.18 0.68 0.86 0.11 0.10 0.53 0.46 0.27 0.16 0.13

0.00 0.00 0.00 0.00 0.00 0.00 0.13 0.10 0.00 0.00 0.00 0.00 0.13 0.15 0.20 0.00 0.00

0.6962 0.6470 0.7482 0.6152 0.6339 0.4910 0.6712 0.6043 0.7414 0.5381 0.6521 0.5921 0.5302 0.5009 0.5761 0.6476 0.6529

0.98 1.18 5.89 1.10 1.00 0.93 0.94 0.98 5.76 0.94 0.94 0.98 0.94 5.78 5.79 5.77 5.76

32.830 28.276 32.062 22.570 24.605 13.579 29.621 25.127 32.488 15.743 26.954 22.637 18.106 17.274 23.603 27.315 26.474

1.24 1.87 5.79 1.39 1.06 0.99 1.00 1.32 5.76 1.17 0.98 1.05 0.98 5.75 5.75 5.76 5.75

3671 3556 3525 3283 3371 2831 3572 3478 3559 2922 3470 3347 3170 3185 3455 3500 3439

-+ 11.5 ___ 22 + 16 -+ 13.2 -+ 5.4 _+ 5.3 + 5.3 -+ 13 _+5 + 11 -+ 4.0 -+6 + 4.3 _+9 -+ 11 _+5 +6

___

15

Sample 108 01a 02b 02c 03a 04a 04b 04c 04d 05a 06a 08a 09a 09b 10a lla llb 12a 14a 15a 16a 18a

Sample 186 01a 01b 01c 04a 04b 05a 06a 06b 06c 07a 08a 08b 10a lla 12a 13a 13b

4 -1 -2 4 6 9 7 12 -1 -6 9 10.9 15 12 9 10 7

(Continued)

ACCRETION AND REWORKING OF THE UKRAINIAN SHIELD

649

Table 1. Continued Analysis number

Concentration (ppm) U

14a 15a 16a 17a 18a 21a 21b 22a 25a 25b 25c 26a 26b 26c 27a 28a 28b 28c 28d 30a 30b 31a

352 1321 2784 655 983 580 3500 64 112 71 182 48 41 67 666 491 330 424 125 141 203 366

Th

120 236 228 328 76 566 165 123 132 139 285 105 141 280 184 262 160 231 95 125 185 46

Th/U

f206" (%)

Pb

319 1095 2564 664 379 678 2435 63 90 65 149 32 52 79 647 567 390 492 96 107 153 282

Isotopic ratios 2~

0.34 0.18 0.08 0.50 0.08 0.976 0.047 1.924 1.172 1.958 1.568 2.179 3.400 4.180 0.276 0.533 0.484 0.546 0.766 0.888 0.913 0.127

0.00 0.00 0.30 0.15 0.00 0.01 0.01 0.49 0.60 0.58 0.56 0.24 0.09 2.56 0.22 0.25 0.03 0.03 0.01 0.03 0.55 0.02

0.6571 0.6308 0.7091 0.6994 0.3299 0.7454 0.5733 0.5757 0.5260 0.5137 0.5272 0.3879 0.5727 0.5536 0.7093 0.7838 0.7936 0.7769 0.5387 0.5307 0.5370 0.5982

_ %

5.78 5.76 5,76 6.28 1.72 1.70 1.71 1.70 2.83 2.00 1.81 2.83 2.10 1.75 1.70 1.70 2.86 2.20 3.12 2.82 1.99 3.00

2~

25.740 23.301 26.935 30.542 6.540 33.302 15.904 16.012 14.581 14.326 14.636 7.585 16.240 14.905 29.400 38.352 39.537 39.039 14.991 14.172 14.583 21.088

Age (Ma) 2~176

Disc. (%)

• %

5.75 5.75 5.83 6.22 1.81 1.79 1.73 1.96 2.89 2.31 2.26 3.00 2.33 3.00 1.76 1.87 2.89 2.22 3.79 3.00 2.87 3.00

3385 3294 3337 3554 2273 3589 2836 2840 2835 2845 2837 2249 2872 2787 3473 3728 3755 3768 2841 2774 2801 3220

• • • • + • • • • • + • • + • • + • • + • •

9 6 4 12 15 9 4 16 9 16 22 20 16 45 7 12 6 6 35 22 33 18

5.0 5 4 5 12 -0.1 -3.8 -4.0 4.8 7.4 4.6 7.1 - 2.0 - 2.4 0.6 - 0.1 - 0.4 2.2 2.7 1.3 1.3 7.7

U-Th-Pb analyses were performed on a Cameca IMS1270 ion microprobe at the Swedish Museum of Natural History, Stockholm (NORDSIM facility) using a method identical to that described by Whitehouse et al. (1999, and references therein). Errors on ratios and ages are quoted at the lo- level. The results are presented in concordia diagrams in Figure 2, with the exclusion of analyses with uncertainty in 2~ larger than c. 3%. This high uncertainty is due to an imprecise U/Pb calibration during one analytical session. Analysis number represents grain number and analysed spot on SIMS mount. Th/U ratios presented are those calculated from measured Th and U signals, calibrated to the 91500 zircon reference. Disc., degree of discordance in per cent; positive numbers are normally discordant. *Percentage of 2~ that is common Pb, calculated from the 2~ signal assuming a present-day Stacey & Kramers (1975) model terrestrial Pb isotope composition.

and some supracrustal rocks are probably present, particularly in the eastern part of the domain. The magmatic rocks include c. 2.1 Ga alkaline intrusions and high-Mg granodiorites, and anatectic c. 2.0 Ga granites (Esipchuk et al. 2004). Whereas the Middle Dniepr Domain is virtually untouched by Palaeoproterozoic processes, the Kirovograd Domain, further to the west (Fig. 1), is dominated by Palaeoproterozoic rocks, which in parts are thought to be deposited on an Archaean basement. These metasedimentary rocks include the famous banded iron formations (BIF) within the Krivoy Rog suture zone, which separates the two domains. The Krivoy Rog BIF were formed between c. 2.6 and 2.3 Ga ago, probably in a craton-margin environment, and the sedimentary processes lasted until 2.1 Ga (Bibikova et al. 1963). Supracrustal rocks of the Kirovograd Domain include flyschoid metasediments of the I n g u l o - I n g u l e t s k formation, which are metamorphosed to granulite facies in the west, and to amphibolite facies in the east. Calc-alkaline metavolcanic rocks occur in the eastern part of the domain, while carbonates and other sedimentary rocks prevail in the western parts. The Proterozoic magmatic activity continued with the emplacement of the Novo-Ukrainian polyphase intrusion of monzonitic rocks at c. 2 . 0 3 - 2 . 0 0 Ga, followed by the formation of the anatectic Kirovograd granites at c. 2.0 Ga. The Bug formation in the southern and central parts of the Podolian Domain is composed of stratified metasedimentary sequences including high-aluminous gneisses and quartzites, calciphyres, and graphite- and m a g n e t i t e - p y r o x e n e - b e a r i n g gneisses which are preserved in several narrow synforms. Based on lithological similarities and palaeogeographical reconstructions, the Bug formation has been correlated by most workers with Palaeoproterozoic metasediments of the Kirovograd and Volyn domains (e.g. Ryabenko 1993). However, a U - P b zircon age determination of c. 2.6 Ga for a mafic dyke cutting the Bug metasediments

(Stepanyuk 1997) indicates that the latter may be Archaean. Nevertheless, considering the sedimentary, terrigenous origin of the host rocks it cannot be excluded that detrital Archaean zircons have been entrapped by the mafic dyke m a g m a and influenced the age determination. The Bug formation is, like much of the Archaean D n i e s t r - B u g formation, metamorphosed to granulite facies. The rocks are intruded by Palaeoproterozoic chamockites, by garnet granitoids of 2.08-2.02 Ga age belonging to the Berdichev complex of anatectic granites, and by younger (c. 1.98-1.92 Ga) pegmatites and granitic dykes (Shcherbak et al. 1989). The youngest Nd model ages from the Podolian Domain, 2.5 Ga and less (Fig. 3), also indicate that the Palaeoproterozoic reworking involved the introduction of juvenile mantle-derived material into the continental crust. Evidence for a c. 2.0 Ga granulitic metamorphic overprint in the Podolian Domain includes results from conventional U - P b zircon analysis of a mafic dyke that intersects the Archaean enderbites (Lesnaya 1988). W e see little evidence of this metamorphism in our new ion probe U - P b results on zircon from the enderbites (Fig. 2), probably because all the Zr already was contained in zircon at that time. However, one low-uranium overgrowth (sample 46, analysis 08b) with a concordant age of 2.0 Ga indicates that these rocks were also affected by a Palaeoproterozoic metamorphic event at that time.

Accretion of n e w Palaeoproterozoic crust Much of the crust in the Volyn and R o s - T i k i c h domains has traditionally been considered to be Archaean (Shcherbak & Ponomarenko 2000). A large region in the Volyn D o m a i n dominated by supracrustal rocks (Teterev metasediments and other

650

S. CLAESSON E T A L .

0,85 108+46, enderbRe 0,75

0,65

4.0

oo r

3.5

3.0

0j

I . I . I . I,.~,/,,.~,1,~,H,pl,., 2.5 2.0

1.5 Ga

Nd model ages from the western Ukrainian Shield (Ga)

c~ 0,55 n

r ~ =PodolianDomain

[]=Ros-Tik~ch Domain

[ ] =VolynDomain

o

Fig. 3. Histogram showing Nd isotope depleted mantle model ages for rocks from the Podolian, Ros-Tikich and Volyn domains, western Ukrainian Shield (Stepanyuk et al. 1998; Dovbush et al. 2000). The model ages for the Podolian Domain vary over a wide range, reflecting the protracted evolution of this domain. The distinct peak at 3.6 Ga is in accordance with the U-Pb age dating results, and supports that this was an important period of crustal formation. The smaller peak at 2.8 Ga is also supported by independent U-Pb ages. The Volyn Domain ages are mainly for metasediments, but also include one sample from a dacitic dyke and one from the anatectic Berdichev granites. These ages form a distinct peak at 2.3 Ga, demonstrating that this domain is essentiallyjuvenile Palaeoproterozoic.It is 0.1-0.2 Ga older than independently determined ages for intruding granitoid rocks. This may be explainedby a minor Archaean component in the detrital material that formed the sediments. The model ages for the Ros-Tikich Domain form two groups. The Archaean ages are for rocks from the southern part, and the Palaeoproterozoicages are for rocks from the north. This demonstrates that the southern part of that domain has characteristics resembling the Podolian Domain, whereas the northern part is similar to the Volyn Domain.

0,45

0,35 l 800/ f,

I

'

I

'

I

'

I

0,25 10

20 207pb/235 U

30

40

t,0 186, ender bite

0,8 340O ~8 t-~ 0,6

22~

0,4

0,2

r

I

0

10

J

I

20

'

.....

I

30

'

I

40

'

50

207pb1235U

Fig. 2. U-Pb concordia diagrams showing results and photomicrographs of analysed very old and younger zircon generations from three enderbite samples collected in quarries on opposite sides of the Southern Bug river in the Dniestr-Bug formation (Fig. 1). Results from samples 46 and 108, which were collected in the same quarry, are shown together in a single diagram. Analytical methods and results are presented in Table 1. For clarity, only analyses with an analytical precision for 2~ better than c. 3% are shown. The results illustrate that the studied rocks are Palaeoarchaean and have complex polymetamorphicArchaean and Palaeoproterozoicevolutions, as discussed in the text. For sample 186, three analyses in the clear, transparent core of crystal 28 (Table 1) yield concordant ages of c. 3.75 Ga. A c. 3.65 Ga age defined by concordant and near-concordant zircon cores with magmatic appearances is interpreted as the age of a crust-forming event. A metamorphic event at c. 2.8 Ga, indicated by analyses of metamorphic overgrowths and new, low-U, isometric light-coloured crystals, is clearly seen in sample 186. A second period of metamorphic zircon growth at c. 2.0 Ga is indicated by one concordant analysis in crystal 8, sample 46 (Table 1).

formations) was thought to be a segment of reworked Archaean crust, which had been intruded by voluminous Palaeoproterozoic granitoids. A m i n i m u m age for Teterev metasediments was provided by a U - P b zircon date of 2.43 Ga for a dacitic dyke (Skobelev 1987) near N o v o g r a d - V o l y n s k . However, new Nd model ages of 2 . 2 5 - 2 . 4 5 Ga for metapelitic gneiss, including the wall rock of this dated dyke, demonstrate that the detritus in the Teterev metasediments cannot be entirely Archaean in age. A Nd model age for the dacite is similar to those for the

metasediments (Fig. 3; Stepanyuk et al. 1998; Dovbush et al. 2000). Furthermore, our unpublished ion microprobe U - P b results for detrital zircons separated from a metagreywacke sample next to the dacite demonstrate that it is dominated by Palaeoproterozoic, < 2.2 Ga crystals, with some input of Archaean material. Microanalysis of zircon from the dyke, as well as from leucosome material in the partly migmatitic metagreywacke, shows that these are 2.05 Ga old and include some inherited older zircon material. The new results, combined with lithostratigraphical correlations, indicate that the Volyn Domain is a largely juvenile segment of Palaeoproterozoic continental crust. The dominance of < 2 . 2 Ga detrital zircons in the investigated Teterev metagreywackes near N o v o g r a d - V o l y n s k (Fig. 1), combined with the 2.05 Ga age for the dacitic dyke and for leucosome material in the metagreywacke, provides close constraints on the age of deposition of the sediments in this area. Various granitoid types have been distinguished in the Volyn Domain, the dominant ones being the Zhitomir granites in the north and, farther to the south, the Berdichev granites. The ages of these granites are consistently in the range 2 . 0 - 2 . 2 Ga, as demonstrated by Ukrainian and Russian workers (Shcherbak 1993). The Berdichev granites are widespread in the southern part of the Volyn Domain, and also in the northern parts of the Podolian Domain. A S m - N d model age for a Berdichev granite sample is similar to those for Teterev metasediments (Fig. 3; Stepanyuk et al. 1998; Dovbush et al. 2000), demonstrating that it has a dominantly Proterozoic provenance. The northern part of the R o s - T i k i c h Domain (Fig. 1) has many features similar to those of the Volyn Domain. S m - N d ages of 2 . 3 - 2 . 5 Ga for amphibolites and gneisses from the Belaya Tserkov' region strongly indicate that the crust is Palaeoproterozoic (Fig. 3, Stepanyuk et al. 1998; Dovbush et al. 2000). A m o n g intrusive rocks, the Fastov and Zvenigorodka TTG intrusions are dated to between 2.14 and 2.04 Ga (Shcherbakov 2005). Many magmatic rocks in this region have island arc-type geochemical signatures. The crust in the Ros region also includes carbonate and magnetite-rich rocks, similar to those in the Bug complex of the Podolian Domain, which are well defined by magnetic anomalies.

ACCRETION AND REWORKINGOF THE UKRAINIAN SHIELD In the extreme NW of the Shield (Fig. 1), the OsnitskMikachevichi Igneous Belt consists of major granodioriticgranitic c. 2.0 Ga batholiths with subordinate gabbros and diorites. This huge NNE-striking structure follows the northwestern margin of the Sarmatian continent. Geologically and geophysically, the crust of the Volyn Domain and the Osnitsk-Mikashevichi Belt can be traced northwestwards up to the Minsk fault, within the Central Belarus Suture Zone, where Sarmatia meets Fennoscandian terranes (Bogdanova et al. 2006). The seismic structure of the crust to the south of the Berdichev zone shows that the Archaean crust of the Podolian Domain dips to the north and underlies that zone. The Teterev belt is difficult to identify by geophysical methods because of the strong reworking of the crust and upper mantle during the Korosten magmatic event and later, in the Devonian, when the PripyatDniepr-Donets palaeorift was formed. There is, to the best of our knowledge, no evidence that the Palaeoproterozoic rocks in the Volyn and northern Ros-Tikich domains were deposited on a significantly older basement. This does not preclude the existence of Archaean microcontinents, but the crust in this part of the Ukrainian Shield is best understood as a southwards extension of the accretional-collisional type of geological terranes that has been demonstrated along the northwestern margin of Sarmatia and further to the NW, under the sedimentary cover of the East European Craton (Claesson et al. 2001). Much of the Palaeoproterozoic crust in the Kirovograd Domain may also be included in this accretion-collision scenario. The c. 2 Ga metamorphism and magmatic activity in the Volyn Domain, as well as the 2 Ga metamorphic overprint in the Podolian Domain further south, may be related to major collisional events during these crust-forming processes.

651

Summary and conclusions General

The temporal evolution of the various parts of the Ukrainian Shield is summarized in Figure 4. Crust dated at 3.65 Ga is found in the Azov and Podolian domains, which also include younger Archaean components. The granite-greenstone terranes that dominate the crust in the Middle Dniepr Domain are c. 3.2-3.0 Ga old, whereas the crust in the Kirovograd and Volyn domains is dominated by Palaeoproterozoic metasediments and various 2.2-2.0 Ga granitoid rocks. The new isotopic ages for the key rock complexes in the western part of the Ukrainian Shield strongly suggest that some of the tectonostratigraphic units either have ages different from those previously assumed (e.g. the 'Novograd-Volyn series'), or are more heterogeneous than previously recognized (e.g. the 'Ros-Tikich series'). In consequence, the traditional tectonic subdivisions of the crust must be revised. A 2.2-2.1 Ga belt of juvenile crust extends from the Volyn Domain towards the SE, occupying the northern ('Ros') part of the Ros-Tikich Domain and extending into the Kirovograd Domain. Figure 5 shows a sketch map of the Ukrainian Shield, based on these results. The Shield is divided into strongly reworked Archaean crust with Palaeoarchaean components in the Azov and Podolian (including the southern part of Ros-Tikich) domains, a well-preserved Mesoarchaean region composed of granite-greenstone rocks in the Middle Dniepr Domain and a wide region of largely juvenile Palaeoproterozoic crust extending from t h e Volyn Domain over the northern part of the Ros-Tikich Domain into the Kirovograd Domain. The oldest A r c h a e a n regions in the Ukrainian

Post-tectonic Palaeoproterozoic evolution

Shield and a worldwide comparison

At c. 1.9 Ga, the tectonic and magmatic activity in the Ukrainian Shield ceased. During the following cratonically more stable period, sandstones and shales were deposited on the newly developed crust in the northwesternmost part of the Shield; remnants of these undeformed platform sediments being preserved locally. The major Korosten igneous suite of dominantly gabbro-anorthositerapakivi granite (Fig. 1; Amelin et al. 1994) was emplaced at 1.81.73 Ga in the eastern part of the Volyn Domain. At about the same time (c. 1.75 Ga), the large anorthosite-mangerite-charnockite-granite (AMCG) Korsun-Novomirgorod suite (Fig. 1) was emplaced into the Kirovograd Domain (Shcherbak & Bartnitsky 1995), and syenites and granites (c. 1.8 Ga) formed several small intrusions in the easternmost part of the Azov Domain (Esipchuk et al. 2004). The youngest rocks in the Volyn Domain are volcanic and sedimentary rocks deposited on top of the northern parts of the Korosten suite.

The oldest known crust in the Ukrainian Shield has been identified in the Novopavlovsk complex in the westernmost part of the Azov Domain (Shcherbak et al. 1984; Bibikova & Williams 1990) and in the Dniestr-Bug formation in the Podolian Domain. In both regions magmatic rocks have been dated to c. 3.65 Ga, and even older (c. 3.75 Ga) provenance is indicated. These Ukrainian rocks are among the oldest on Earth. The most ancient (>3.6 Ga) rocks that have been identified worldwide are dominated by so-called grey gneisses; that is, highgrade orthogneisses of mainly tonalitic composition (Windley 1976). The oldest U - P b zircon age from a magmatic rock has been obtained from the Acasta tonalitic gneisses, Slave Province, Canada. These have been dated to 4.03 Ga (Bowring et al. 1989; Bowring & Williams 1999), but some authors have suggested that these zircons are xenocrysts and thus do not date the host rock.

VOLYN DOMAIN PODOLIAN DOMAIN

B

KIROVOGRAD DOMAIN MIDDLE DNIEPR DOMAIN AZOV DOMAIN

EOARCHAEAN! PALAEOARCHAEAN! MESOARCHAEAN NEOARCHAEA Greenstone belts

PALAEOPROTEROZOIC

TTG (tonalite-trondhjemite-granodiorite) rocks

Intracratonic dfting (Banded Iron Formations)

mmmmmmmmmmi Late/post tectonic magmatism

Metasediment:s

~,

~

AMCG (anorthosite-mangedte-charnockite-granite) magmatism

high-grade metamorphism and deformation

Fig. 4. Compilationof rock ages and periods of high-grade metamorphismin the various domains of the Ukrainian Shield, based on previouslypublishedresults and our own partly unpublishedresults from the Podolian, Volyn and Ros-Tikich domains referred to in the text.

652

S. CLAESSON ETAL.

r Shield

Balt

25 ~ 52 ~

33 ~

, . _

Ukrainian

Shield

39 ~

50 ~ :! .....

i,

....

i

(b)

48 ~

2s~ 0

200

km

~p~

46040 '

29 ~

39 ~

3.2-3,0 Ga Middle Dnicpr domain

Undivided Archacan and l}alaeoproterozoic rocks

3.7- 2,8 Ga Azov and Podolian domains re~orked in the Palacoprotcrozoic

Palaeoproterozoic suture zones

2,2~2.1 G~ Tctercv - Ros orogenic bell

1,8-1.7 (ia gabbro-anorthosite-granitic plutons

"'"

"

Major boundaries or'domains anti belts

The most studied and best preserved area of ancient rocks occurs in West Greenland, where the oldest Am~tsoq tonalitic gneisses have been dated using several isotopic systems ( U - P b , S m - N d , R b - S r , P b - P b , L u - H f ) to 3.6-3.9 Ga (e.g. Black et al. 1971; Baadsgaard 1973; Nutman et al. 1996; Whitehouse et al. 1999). Even older ages, >_3.8 Ga, have been obtained for volcanic rocks from the Isua greenstone belt in Greenland (Baadsgaard et al. 1984; Nutman et al. 1997). Similar ages have been obtained for the Saglek gneisses in Labrador (Schi6tte et al. 1989; Bridgwater & Schi6tte 1991), for granulites from the Napier complex in Antarctica (Black et al. 1986), for tonalitic gneisses and greenstones from Swaziland in South Africa (Compston & Kr6ner 1988; Kr6ner et al. 1996), for gneisses from NE China (Liu et al. 1992), and for gneisses and anorthosite inclusions from western Australia (Kinny et al. 1988). The Palaeoarchaean Ukrainian rocks have a strong metamorphic overprint and complex deformational histories, which make more detailed interpretation difficult. In both the Podolian and the Azov Domain, the 3.65 Ga rocks appear to have been affected by high-grade metamorphism, in the late Archaean at c. 2.8 Ga, and again in the Palaeoproterozoic at c. 2.0 Ga. However, the Azov and Podolian domains are separated by the Middle Dniepr and Kirovograd domains, where the oldest known rocks are dated at 3.1-3.2 Ga, and the Palaeoproterozoic overprint on the Middle Dniepr Domain was very mild. There is no evidence that the temporal similarities in the evolution of the Dniestr-Bug formation and the Novopavlovsk complex reflect a common history. On the contrary, kinematic indicators in the fault systems of the Ukrainian Shield (Gintov 2004) show that the Archaean tectonic

Fig. 5. (a) Geological sketch map of the East European Craton (EEC), as in Figure 1. (b) Schematic geological map of the Ukrainian Shield, indicating its subdivision into highly reworked Palaeoarchaean core regions in the Podolian and Azov domains, a major Mesoarchaean granite-greenstone terrane that is largely unaffected by younger metamorphism or deformation in the Middle Dniepr Domain, and a large region dominated by juvenile crust, accreted in the Palaeoproterozoic, which includes the Volyn Domain, the northern part of the Ros-Tikish Domain and the Kirovograd Domain. Regions of major post-tectonic magmatism in the Volyn, Kirovograd and Azov domains, and Palaeoproterozoic suture zones, are also shown. The abbreviations are the same as in Figure 1.

evolutions to the east and to the west of the Kirovograd Domain were independent. Despite the striking age similarities between the Palaeoarchaean and younger rock complexes in the western and eastern parts of Sarmatia, tectonic correlation must therefore be conducted cautiously. The two parts of the shield were juxtaposed not earlier than 2.1-2.0 Ga, probably concomitantly with the collision between Sarmatia and Volgo-Uralia. In view of the lack of evidence for a common history, we consider that the temporal similarities between the Azov and Podolian domains indicate that 3.65, 2.8 and 2.0 Ga events were periods of tectonic activity in Sarmatia in a more general sense.

Palaeoproterozoic accretion

Our results emphasize the important role of Palaeoproterozoic geodynamics for the assembly of the Archaean components and the formation of the dominant structure in western Sarmatia. This applies even to its oldest, Archaean parts in the Podolian Domain. Palaeoproterozoic processes have also been important in the evolution of the Azov Domain in the east. The only parts that have not been significantly affected by these Palaeoproterozoic processes are the granite-greenstone terranes in the Middle Dniepr Domain. The multiple magmatism and granulite-facies metamorphism in the Podolian Domain may reflect active-margin tectonic settings both in the Neoarchaean at c. 2.8 Ga and in the Palaeoproterozoic between 2.1 and 2.0 Ga. The eastern part of the Ukrainian Shield, in particular the Azov Domain, should be studied further, not only to give more

ACCRETION AND REWORKING OF THE UKRAINIAN SHIELD

information about the oldest crust, but also to correlate Sarmatia and Volgo-Uralia, and to clarify the relationships with the Palaeoproterozoic in the western parts of the Shield and the role of tectonic processes along the S a r m a t i a - V o l g o - U r a l i a margin. This paper is a result of collaboration within the framework of the EUROBRIDGE project (EUROPROBE/ILP/ESF). Funding by INTAS, the Swedish Institute and the Royal Swedish Academy of Science helped E.B. to visit the Laboratory for Isotope Geologyin Stockholm to carry out isotope work. We thank the laboratory staff for analytical help and support, and R. Gorbatschev and D. Gee for constructive comments on the manuscript. Grants from the Swedish Research Council to S.B. are acknowledged. This is NORDSIM Contribution 133.

References

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Index Figures are indicated in italic and tables in b o l d Aar massif 131,134 Acadian deformation 293, 297, 300 Acadian event 324, 356 accretion 323-329, 551 Proterozoic 12, 510 accretionary complex 69, 70, 75, 265 Anatolia 396-400 Magnitogorsk arc 408, 410, 411,413 accretionary margin, East European Craton 291-306 accretionary prism 348, 349, 351 Oman 236-238, 240-242, 250 acritarch, Vendian 508 Adana Basin 270 Adria terrane 85, 93 Adriatic indenter 192, 193 Adriatic Moho 133, 134, 135, 138, 141 Adriatic plate 130, 136, 139, 192, 193 Aegean back-arc basin 171, 177, 272 ~ g i r Sea 86, 87-88 Aeolian Islands volcanism 174, 175, 179 African plate 358 convergence 223-225, 263, 265, 267-269 plate boundary 272 age see chronology, isotope and zircon Agnostus [trilobite] 513 Akershus Graben 365, 366 Albanian ophiolites 242-244 Alboran, volcanic province 172-176, 227 Alcapa terrane 193-195, 198, 199-202 alkaline magmatism 167, 172, 198, 362 volcanism 147, 150, 152-155, 158 Alleghanian orogeny 57, 59 ALP seismic experiment 143 Alpide orogeny 389, 390 Alpine deformation 1, 3-4, 129-143, 227, 281 thermal-mechanical controls 113-123 Alpine deformation front 98, 100, 102, 356 Alpine Orogen 51,277 Alpine subduction 171 - 172 Alpine Tethys 50-51, 59-62, 65, 75 Alpine-Mediterranean geodynamics 180-182 Alps 29-30, 129, 130, 131 geophysical data 21, 29-30 lithosphere thickness 24 Amazonia 326, 328, 568, 570 Amerasia Basin, opening 507, 508 amphibolite 243, 327, 351,395 East European Craton 293, 296, 300 Anatolia 64, 73, 246, 271,389-396 accretionary complex 396-400 chronostratigraphy 391, 397 seismic profile 271 volcanism 154, 177 Anatolian Fault 171 Andean-type collision 357 Anglo-Brabant Deformation Belt 301 anorogenic magma see alkaline magmatism anorogenic volcanic province 147, 151 Antalya Basin 73, 252, 268, 271 Apenninic arc 171 Apulia terrane 85, 130, 153 Arabian plate convergence 264 Arabian Platform 389 Archaean crust 600-601 Archaean Fennoscandian Shield 627-640 accretion 627 correlation 640 granitoids 635-639 greenstone complexes 627-634 metamorphic rocks 629-634, 640 metamorphism 636-639 provinces 628

archaeocyathids 89 Arctic Caledonides 509 Arctic, magnetic map 514 Armorica 89, 93 Armorica Massif, Quaternary folding 118 Armorican Archipelago 297 Armorican microplate (Franconia) 335 Armorican Terrane 85, 323, 325, 333 deformation 327-328 Armorican Terrane Assemblage 300 Asturian Phase 356 Atlantic opening 50, 223,224 Aubrac Cenozoic volcanism 152 Austroalpine nappes 135, 199 Autun Basin 370, 371 Avalonia 3, 57, 293, 296, 323 East European Craton, 324-329 emplacement mechanism 326-327 reconstruction 83-85, 89, 333, 460 soft collision 324 transect 301, 302, 305 triple plate collision 20 Avalonian suture 291 Thor-Tornquist 46

BABEL upper mantle project 545, 546-548, 554, 565 participants 555-556 seismic line 580 working group 15,299, 313, 315 back-arc basin 4, 5, 20, 192, 207 Aegean, 171, 177, 272 Avalonia 327, 328 extension 45, 48, 52, 170, 176 Guevgueli 248, 389, 400 Pannonian 191,200 Rheno-Hercynian 357, 339, 357 rift 30, 57, 62, 70, 238 topography 123 Baer-Bassit ophiolite 244, 247 Baikalian orogeny 510 Balkan orogen 59-60, 328, 399-400 Balkan suture 68 Balonia 293 Baltic Basin 458, 460 Baltic Shield see Fennoscandian Shield Baltica 84, 85-86, 89 East European Craton 294 Eurasia 507-509, 514-515, 521-536 mantle 2, 5, 24 rotation 447 transect 301, 303 Barentsian Caledonides 507, 509 basanite intraplate 147, 150, 152, 154, 161, 162 subduction-related 172, 176, 177 basement, East European Craton 482, 484-486, 510 Bashkirian Anticline 411,412, 413 BASIN seismic profile 375, 378 basins, east Mediterranean 263-273 bathymetry 264, 268, 269 Bay of Biscay, subduction 50-51, 75 B6k6s Basin 197, 199 Belomorian terrane 528 Belomorides, Archaean crust 546 Benioff zone 175, 177 Betic Cordillera 154, 225, 226, 227 Bey~ehir ophiolite 245,246 BIRPS seismic profile 29 bituminous shale 370 Black Forest, magmatism 148, 151,338 Black Sea orogeny 399-400

blueschist 237 Cenozoic 199, 390 Cretaceous 65, 398 Palaeozoic 397, 408 Triassic 72 Bohemian Massif 296, 323, 328, 333, 338, 339 Cenozoic 50, 97, 100, 107, 117, 123 volcanic activity 51, 148, 149, 153, 154 Bohemian Terrane 298 boninitic lava 243, 409, 410 Bouguer anomaly 227-230, 551 Brabant Massif 324 Brenner fault 201 Bresse Graben 97, 99, 103, 120 Brianqonnais domain 61, 62, 130 Bruno-Silesian Promontory 323, 325, 327 Bruno-Silesian Terrane 296 Budva domain 75 Burgundy Transfer Zone 51, 52, 120 Cadomian crust 455, 457, 460 Cadomian orogeny 3, 86, 323, 389, 516 Calabrian arc 265, 272 calc-alkaline magmatism 333 Balkan Terrane 328 Cenozoic 147, 152, 154, 168, 169, 176 Betic-Rif province 172-174, 178 Magnitogorsk arc 408, 410 Variscides 356, 357, 361-363, 370, 380 Calcareous Alps 59-60, 72 Caledonian Deformation Front 24, 301 Caledonian orogeny 85, 89, 93, 327, 453 Eurasian Arctic 507-516 Caledonian suture 380, 508 Caledonides 1-5, 19, 20, 25 crustal domain 43, 45-47 geophysical data 21 lithosphere thickness 24 ~amlik granodiorite 395, 400 isotope ratios 396 Campania, volcanism 175, 179 Cantabrian Mountains, Permian 351,372, 373 Cantal, Cenozoic volcanism 152 Cappadocia ignimbrite 177 carbonate platforms and ophiolites 240-242, 245, 249 carbonates 245, 251,400 Lower Palaeozoic 454-455,457, 512 carbonatite 147, 151, 174, 181 Carmel Fault 267 Carnic Alps, Variscan 334, 336, 337 Carpathian arc 178, 191, 194, 210-213, 216 lithosphere thickness 21, 23, 29 Pannonian region 173, 176, 180 Western domain 59-60, 170 Catalan Coastal Ranges 224, 226, 227 CELEBRATION project 6, 200, 313, 314 Cenozoic basins east Mediterranean 247, 268-272 Cenozoic magmatism 103-104, 149, 155, 172 volcanism 152 Central European Rift System 20, 21, 23, 30-31 Central Russia Rift System 19, 20, 24 Channel Basin 50 Chios accretionary complex 397 chronology 496 Anatolia accretionary complexes 397 Archaean greenstone 640 Carboniferous-Jurassic 74 Cenozoic Pannonian Basin 195 Cenozoic volcanism 150, 172 Cretaceous, Oman ophiolite 236 Dramala ophiolite 242

From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 655-662. 0435-4052/06/$15.00 9 The Geological Society of London 2006.

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656

chronology (Continued) Evia ophiolite 241 Greater Caucasus 285 Lycian ophiolite 245 Mersin ophiolite 245 Palaeozoic, East European Craton 452 Permian, Germany 363 Pontic terranes 391 Pyrenees 373 Cimmerian cycle 5, 57, 59, 277, 278, 358 tectonics 400, 494-498 Ciudad Roderigo Basin 226, 227 coal 359, 369, 372, 373, 392 Carboniferous 335-336, 338 collision structure 138-139, 140, 240 collision tectonics 4, 5, 6, 20, 25, 51 arc-continent 357, 407-410 Avalonia 327, 328 Baltica 407-409, 416 continent-continent 43, 99, 167, 356 Gondwana-Laurussia 355, 357, 358 iberia 345-351 Plio-Quaternary 253 Proterozoic 579-597 Pyrenees 224 soft 85, 170 Svecofennia 568-569 Variscides 99, 345-351 compression 498 Greater Caucasus 281-282, 283 Iberian Peninsula 224-227 conductivity 544-547 Conrad discontinuity 134 controlled source seismology 130 convergence 390, 410, 443, 454-457 convergence rate 50, 51, 167 Cordilleran-type ophiolite 237, 247-253 Crete Basin 272 Crimea, Triassic to Jurassic 496 cross-sections see transects crust 24, 130-136 Archaean 639-640 properties 542-546, 549-550 Proterozoic 548, 606-608, 613, 614 Russia 521-536, 585-589 shortening values 100 strength 116-117, 1 2 3 - 1 2 4 crustal evolution 43-53, 365,380, 481-482 crustal lamination 103-104 crustal structure, East European Craton 291,300 transects Avalonia- Rhenohercynian 302 Caledonide 304 Variscan 303, 305 see also delamination crustal thickness 18, 116, 131,200, 209 map 45, 46, 47, 48, 49 Southern Permian Basin 102-104 crustal wedge 195, 351,348 crystalline basement 391,474, 529, 531,533 East European Craton 292, 298 Cycladic domain 64, 73 Cyclopyge [trilobite] 328 Cyprian arc 266, 268, 272 Danish graben 104, 365, 376 Danube Basin 198, 199, 326 Dead Sea Rift 264, 268 debris flow 251,409 Balkans 240, 242-244, 246, 247, 249 DEKORP basin research group East European Craton 299, 300, 301, 304, 305 DEKORP seismic profile 24, 103, 314, 371,375 delamination 273, 285, 351,379, 561 Alps-Mediterranean 167, 169, 170, 174, 175, 179, 181 rifting 101, 102, 154 mantle 97, 574

INDEX

Denmark, TOR project 314 density variation 18, 18, 32 density, Baltic Shield 32 diamond 100, 337, 408 diatremes 151 Dinaride-Hellenide passive margin 59 Dinarides 75 Dinarides, volcanism 173, 176-177 Dnestr Basin 449, 457, 460 Dnieper-Donets-Pripyat rift 21, 22, 27 -28 lithosphere thickness 24 Dniepr-Donets Basin 380, 463-468 DOBRE reflection profile 28, 104, 465-468, 474, 475 Franconian Platform 104 Dobrogea Platform 326 Dobrogea terrane 298 Dobrudzha Trough 72 Dowsing Lineament 297 Dramala ophiolite see Pindos ophiolite Drinova Thrust 326 Dublin graben 45 Duero Basin 226-228, 231 dunite 444, 446 dyke swarm 4, 18, 361 Earth model IASP 138 earthquake 120, 124, 284 hypocentres 213 tomography 139 TOR project 314, 315, 319 East Barents Basin 470-472 East European Craton 1- 8, 11, 277, 389 accretion 293-299, 323-329 break-up 299 correlation 451, 452 geodynamic evolution 599-621 geophysical interpretation 616-621 gravity map 295 gravity model 609-615 magnetic map 294 named basins 450 passive margin 327, 328 Proterozoic crust 407, 422, 606-609 rift basins 463-476 southeast margin 481-498 subsidence curves 453 sutures 299-300 tectonic map 459, 464, 600, 601 terrane provenance 291-293 terranes 602-606 western accretionary margin 291-306 terranes and sutures map 292 East European Platform 11, 17-19, 20 geophysical data 21, 27-28, 117 lithosphere thickness 24 Ebro Basin 51,226, 227, 231 eclogite 103, 327, 351,638 Anatolia 390, 397 Bohemia 298 Caspian 470 Uralides 422, 515 Variscan 335, 337, 338 ECORS, Alps seismic profile 24, 129-132, 135, 139 lithosphere slab geometry 141-143 ECRIS see European Cenozoic Rift System Ediacaran fauna 86 EEC see East European Craton effective elastic thickness 213, 215, 219 Eger Graben 97, 149, 152 EGT seismic experiment 305 Eifel plume project 313,320 Eifel volcanic field 150, 155 elastic-plastic plate model 119 Elbe Fault 358 Elbe Line 314, 317 Elbe Lineament 291,301

electromagnetic conductivity 542-547, 554, 555 techniques 543-544 see also EUROPROBE enriched mantle component 181 Entomozoe [ostracode] environment change, Eocene-Oligocene 283-284 Eocene volcanic activity 148 Eocimmerian 59, 69, 70, 498, 278-280 Eo-Variscan orogeny see Caledonian Eratosthenes Seamount, continental crust 266, 268, 397 ERCEUGT group 303 EREGT working group 29 erosion 155, 216, 217, 337, 339 Erzgebirge diamonds 337 ESRU 408, 414, 416, 421,422 data processing 431-433 seismic profile 24, 26 seismic reflection data 427-440 Etna 153, 174 EUGENO working group 313,314, 315 East European Craton 299, 300 Eurasian Arctic 507-516 EUROBRIDGE 17 EUROBRIDGE profiles 599-621 European Cenozoic Rift System 51, 97-108 lithosphere 113-115, 117, 123-124 map of faults 98, 100 reflection lines, location 100 sedimentary basins 98, 102 European Geotraverse 11, 26, 117, 374 refraction profile 117, 129, 130, 131, 133, 134 EUROPROBE 11, 27-28, 129, 192, 203, 207 Dniepr-Donets Basin 421,463 station map 543 SVEKALAPKO 521-536 tomography experiment 549-554 Urals 407 evaporite 94, 227, 231,529 Messinian 264, 268, 270 Palaeozoic 365, 370, 375, 379, 467, 471 Evia ophiolite 240, 241 extension 4, 99, 333,568 in compressional setting 191,195 Greater Caucasus 280-281 Iberian Peninsula 223-224, 230 post-collision 170 Variscan 103, 333 extinction event, Permian 83 extrusion tectonics 191, 194, 202, 217 Fallotaspsis [trilobite] 513 far-field effects 381,458, 460 faults 46, 47 FENNOLORA seismic profile 313, 319, 541 Fennoscandia 541-556 tomography experiment 543, 549-554 Fennoscandian Shield 114, 117, 542, 582 crustal structure 522-529 geophysical studies 21, 23, 32, 541-556 lithosphere 24, 579-593 map 7 provinces 628 seismic data 12-15, 20 tectonic map 580, 628 terranes 602-605 Finike Basin 271 fish fauna, Old Red Sandstone 512 flexural downloading 210-212, 283, 285 flood basalt 94 flower structure 140 flysch 424 Devonian-Carboniferous 335-338, 348, 350 fold-and-thrust-belt 347, 348-350 folding and lithosphere strength 117-119 folding, Greater Caucasus 281-282 folds 217, 218, 282, 283 fore-arc basin, Palaeotethys 69-70 forebulge, Caledonides 455, 457 foreland basin 45, 59, 449

INDEX

fossils, Gondwanan affinity 292, 297 Fourier' s law 117, 318 Franconian Platform 107, 335 gabbro, Silurian 443-447 Galahetes [trilobite] 513 Galatia volcanic province 154 garnet, age 335, 348 geochemistry 157-160, 173 Betic-Rif province 172-176, 178 Fennoscandia 585-589, 627-636 primitive magma 158-160 subduction-related magma 168-170, 181 GEON center 523, 531 geophysical data 21-22 Alps 28-29 Fennoscandia 12-15, 20, 21, 23 lithosphere properties 19-24 lithosphere thickness 24 geophysical interpretation 11 - 33 Baltica 521-536 East European Craton 313- 320, 616- 621 Fennoscandia 541-556, 581-583 Permian basins 374-376 upper mantle 156-157 Trans-European Suture Zone 18, 19 GEORIFT 27-28 geotherms 16, 19, 26, 31,410 Germany 314 Cenozoic volcanism 155, 156 Variscan orogen 333-340 Germany, north, transect 301 Gfrhl Suture 300 Giessen Ocean 99, 327 Giessen ophiolite 297 Gissen-Harz Basin 45, 102 Giudicarie belt 136, 139, 140, 143 glacial deposits, Ordovician 89, 329 glaciation 86, 92, 94 Gloria Fault 223 Gltickst~idler Trough 363 Gondwana 5, 57, 83, 99, 355 accretion 323-329 break-up 328, 391 collision with Bohemia 336, 337 convergence 356, 379, 381 map showing named units 58-59 reconstruction 84, 87-91 tiffing 293 Variscan 333 Goniatites [ammonoid] 338 graben 97, 151,266, 267, 270, 363 granite 18, 65, 393 Variscides 70, 338, 339 granitoid plutonism, Pontides 395-396 granitoid, Fennoscandia 545,635-639 granitoid, Uralian 411,414 granulite 103, 243, 298, 337, 338 Lapland Belt 528, 579, 581,582, 585-589 Lapland and Umba terrane 591 Lithuanian 602-603 Pontides 395, 398 Urals 425 graphite as conductor 542, 544, 545 gravitational collapse 171, 174, 203, 217 Fennoscandia 569-570, 573, 592-593 gravity 18, 27, 32, 253 Danube Basin 198 East European Craton 600-615, 616 maps of Europe 3, 7, 18, 295 NW Russia 523 Poland 375-376 Timan Range 531 Uralides 26-27 Greater Caucasus chronology of tectonic events 285 compression (mid Jurassic) 281-282, 283 compression, Cenozoic 283 Eo-Cimmerian 278-280 extension (early Jurassic) 280-281

inversion, Cenozoic 283-286 rate of shortening 284 post-rift succession 282-283 structure 278 topography 278 and tectonics 284 greenschist 243, 327, 395, 398 Fennoscandia 627-639 Neoproterozoic 510 Grenville orogeny 3,532, 564 Guadalquivir Basin 226, 227 Guidicaria Fault 193, 201 Gulf of Lyons, opening 225 gypsum 282, 511, 513 Halloporina [bryozoa] 513 Hanseatic terrane 57 Hatay ophiolite 244, 247, 25 l Hawasina Complex 241,246 Haybi Complex 241, heat flow 12, 19, 25, 26, 27, 208, 214 modelling 375 values 15, 17, 19 Hebediscus [trilobite] 513 Hegau volcanic field 151, 159 Hellenic arc 85, 93, 171, 176-177, 191 Helvetic nappes 129 Hercynian suture zone 30 Hessian Graben 97, 114 Himalayan-type collision 43, 99, 167, 356 H6d-Mak6 Basin 197, 199 Holstein-Horn Graben 363 Holycross Mountains 304 Palaeozoic sequence 295-296 Hun superterrane 57, 58 Hungary Cenozoic volcanism 155 Hyblean Platform 269, 270 hydrocarbons 6, 219, 362, 377 hydrothermal vent communities 93 Iapetus Ocean 87-89, 460 opening 293 subduction 296 Iapetus Suture 20, 45, 83, 85 East European Craton 299, 304 Iberia lithosphere thickness 229-230 Iberian basins 372-374, 376 Iberian Massif 345-351 structure 346, 349 Iberian Peninsula 223-231 compression, Cenozoic 224-227 mantle plume 351 Mesozoic extensional basin 223-224, 230 Neogene basins 227 inversion 230-231 tectonic map 226 topography 223,227-231 Variscan basement 228 Iberian pyrite belt 349, 351 IBERSEIS seismic profile 20, 346-348, 351 Iceland plume 50 ILIHA seismic profile 20, 25 imbricate thrust 227 Indonesian archipelago 570, 574 inselberg pattern 195-196, 196 Insubric Line 131,135, 171 intermontane basins 227, 231,373 intraplate lithosphere 113-119 Intrapontide Ocean 67 inversion 122, 208, 209, 281 Cenozoic 283-286 Iberian 230-231 Laramide 155 Moscow Basin 460 Pannonian Basin 195-196, 208, 214, 216, 219 Polish Basin 378-379 Tethyan 50 Tornquist 319 Ionian Basin 171,263, 264-265 IPL-ALCAPA project 207

657

iron and seismic velocity 12, 32 iron formation, Archaean 632, 634 Iskenderun Basin 270 island arc 407, 421 Mid-German High 333 isotope age 443 East European Craton 293 Magnitogorsk intrusives 409 Maksyutov Complex 422 Sakarya Zone 396 see also zircon isotope chemistry 158-160 isotope measurement 444 isotope ratio 169, 175, 178, 180 Kazda~ metamorphic rocks 396 Isparta Angle 250, 251 Israel, graben and horst 266 Istanbul Block 326-327, 391-394, 400 deformation 394 succession, Mesozoic 394 Italy, volcanism 174-176, 178 Cenozoic 153, 155 Ivrea Zone 131,133, 135, 138, 139, 141 Izanca 64-66 Izhma Domain 532, 533 Izmir-Ankara ocean 67, 73, 75 accretionary complex 397, 400 rift 72 suture 64, 171, 177, 238, 398 Jura Mountains, compression 120 Kachkanar massif 445 Kaiserstuhl, volcanic activity 151, 155 kamafugite 181 Kanin peninsula 531 Kara magnetic anomaly 513-514, 515, 516 Kara Terrane 86 Karaburun accretionary complex 397- 398 Karakaya Basin 72 Karakaya Complex 238 Karelia Domain 12, 541,545, 546 Karelian Craton 579 Kasimlar basin 59 Kazakhstan 422 Kazakhstan arc 58 Kazakhstan plate 26, 409-411,416 Kazakhstania 84, 86, 87, 94 Kazbek volcano 284 Kazda~ metamorphic rocks 395, 396 Kempersai massif 443 Khanty-Mansi ocean 58, 71 Khoreyver Domain 533 Kimzha graben, seismic profile 530 Kipchak arc 58 Kir~ehir Massif 389 Kola deep borehole 524 Kolva deep well 471 Kujandaspis [trilobite] 513 Kuloy graben, seismic profile 530 Kumba gabbro, age 446 Ktire 62, 63, 64, 71 Ktire complex 397-398 Kttre ophiolite 238 Kytlym dunite 446-447 lacustrine deposits Iberia 226 Oligocene 226 lamproite 174, 176, 177, 178 lamprophyre 153 Lapland-Kola orogen 579-593 cooling age 592 Laramide basin inversion 155 Larnaca Basin 270 Latakia Basin 270 Laurasia 399, 400

658

Laurentia 5, 24, 293, 507, 508, 509 palaeoequator 89 reconstructions 83, 84, 85, 86 Laurussia 57, 58, 93, 99, 392, 400 collision 355, 356, 379 palaeogeography 394 passive margin 327, 328 reconstructions 85, 86 Leshukona-Pinega Rift 531 leucities 150, 176 Levant margin, seismic profile 267 Levantine Basin 266-268, 270 Ligeran Phase 356 Limagne Graben 97, 103, 114 Linosa, magmatism 153 Linosa, seismic profile 269 listric fault 199 LITHOPROBE 15, 19 lithosphere 472, 474, 475 conductivity 545 folding 117-119, 124 profiles 23 strength 117-119 lithosphere structure, Svecofennia 561-562, 571-573 lithosphere thickness 24, 31, 97 East European Platform 27-28 Iberia 229-230 Pannonian Basin 201,202 south Europe 25-26, 27-28, 31 lithosphere, Baltica 459-460, 542-546 lithosphere, East European Craton 291 lithosphere, Fennoscandia 579-593 lithospheric mantle composition 31-32, 117 Lizard 45, 99, 333 MORB 71 ophiolite 394 peridotite 293, 297, 300, 327 Lizard-Rhenish suture 358 London-Brabant Massif 45, 117 strength of lithosphere 123 Lower Rhine Graben 97 Lower Rhine Lineament 297 Lviv slope, carbonate 456 Lycian domain 63 ocean 66, 67, 75 ophiolote 245, 246, 247 Lyciam nappes 64 lydites 70 Lysogory terrane 85 magma geochemical characteristic 157-160 subduction-related 158, 167-203 magmatic fields Permo-Carboniferous 76, 48-49, 339 Oligocene 49 magmatic underplating 547, 548, 551 magmatism, age 172 Carboniferous 99 Carboniferous-Permian 97, 101-102, 104 Cenozoic 4, 31, 114 Devonian-Carboniferous 6, 348, 349 Permo-Triassic 5 magmatism, Armorican Terrane 327 magmatism, Cenozoic intraplate 147-162 age 150-155 and basement uplift 155-157 geochemistry 157-160 geodynamic setting 147-150 source 160-161 magmatism, Greater Caucasus 280-282 magmatism, Variscan 359-362, 380-381 distribution map 360 foreland 359 Germany 361 Iberia and west Mediterranean 362 Massif Central 362 Scotland Midland Valley 361

INDEX

Variscan foreland (externides) 362-367 Variscan internides 367-374 Whin sill 361 magnetic anomaly 532, 533 Massif Central 48 Paris Basin 48 Saar-Nahe 103 magnetic map Arctic 514 East European Craton 601 Europe 2 NW Europe 294 NW Russia 522 magnetite 634 magnetotelluric (MT) studies 29 Magnitogorsk arc 407-411,413 Magnitogorsk block 27 Magnitogorsk-Tagil island arc 422, 424-425 Main Caucasian Thrust 278, 279 Main Uralian fault 27 seismic interpretation 436-437 suture 407, 414, 421,443 Maksutovo Complex, radiometric age 409 Maksyutov Complex, subduction 421-422 Maladiodella [trilobite] 513 Maliac 62, 63, 64 back-arc basin 70 ocean 73 Malopolska Terrane 85, 293, 295 Malta Trough 269 seismic profile 270 Mamonia Complex 252 mantle conductivity 545 mantle convection 156, 162 mantle diaper 147, 162 mantle discontinuity 553 mantle events, Svecofennia 566 mantle experiments 542 mantle model, East European Craton 611 mantle peridotites 552 mantle plume 30, 46, 51 Archaean 632 Cenozoic 115, 122, 123, 157 rift-related 97, 101, 103 subduction-related 175, 181, 182 Mesozoic 360, 380, 584 Palaeozoic, East European Craton 473 mantle plume dynamics 345-355 mantle properties 542, 546-547 mantle reflector 15, 566 mantle structure 11-33, 549, 551 mantle temperature 12-14 mantle wedge 410 marginal ocean sequences 71-75 Massif Central 101,336-338, 367-370 geophysical data 21 lithosphere thickness 24 lithosphere strength 123 Permo-Carbioniferous trough 107 uplift 149 Variscan basement 152, 155 volcanic activity 31, 51, 148, 152 mechanical strength of lithosphere 115-119 Mediterranean basins 52 Mediterranean Ridge deformation front 263, 264, 265, 265 accretionary wedge 264-265 Mediterranean, geophysical model 28-29 mrlange 327, 410 east Mediterranean 239, 243,245,246 collapse of platform successions 240, 242 Tanaelv 592 Meliata 62, 63, 64 Meliata-Hallstatt ocean 72 melilite 151, 152, 161, 174, 181 melting curves 160 Menderes Massif 390 Mersin ophiolite 245, 247 Mesozoic basins, east Mediterranean 263-268 Messinian salt 264, 268, 270

metamorphic sole/ophiolite 239, 243, 244, 246 metamorphism 61, 63, 65 Armorican Terrane 327, 329 Fennoscandian 584, 636-639 Maksutovo Complex 408-409 Pontides 395-396 Variscan 57, 328, 333,337-339 meta-sediments, Archaean 631-634 metasomatism 170 meta-volcanics and intrusions Archaean 629-634, 640 Mezen Basin 522, 523-525, 528-530, 531 Mid-German Crystalline High, Variscan 297, 333, 335 cooling age 370 Mid-German Crystalline Rise, Cenozoic 99, 100 Midland Valley, Scotland 45 Midlands Microcraton 296, 301 mid-ocean ridge basalt 158, 173 trace element 169 mid-ocean ridge basalt, Palaeotethys 69, 70 Mid-Polish Trough 375, 376 Milankovi6 cyclicity 196 mineralization 6 Miocene volcanic activity 149, 150 Mobil Search, seismic survey 15 Moesia terrane 85, 298-299 Moesian Platform 298, 326, 391,392 Moho depth 348, 566, 609 Adriatic 133, 134, 135, 138, 141 depth controls 43-45 East Barent Sea 471,474 Fennoscandia 547, 548 Liguria 133, 138 NE Baltica 523, 528, 531,533 Pannonian Basin 199, 200 petrology 103 post-Variscan 97, 98, 101,375, 376 seismic 104 temperature 15, 102 topography 131 - 134 Uralides 416 Moho offset 131,133, 138, 139, 141 Tornquist Zone 300 Moho reflector, TOR data 318- 319 Mohorovirid discontinuity see Moho molasse 455 Carboniferous 335, 336 Cenozoic 193 Permian 400 Moldanubian ocean 336, 337, 338, 339 Moldanubian Terrane 100, 293,298, 305 Moldova carbonate platform 455, 456 MONA LISA Working Group 19, 20, 25 East European Craton 297, 299-301 Mont Dore, Cenozoic volcanism 152 Monte Vulture, magmatism 153, 174 Moravian Suture 300 Moravicum nappe 326 Moscow Basin 449, 454, 457 inversion 460 isopachs 455 Mugodzhar-Khanty-Mansi ocean 71 Mut Basin 270 Nagorskaya drilling project 508 nappes Alpine and Carpathian 51 Austroalpine 135, 199 Helvetic 129 Lyciam 64 Moravicum 326 Pindos 68 Semail 69 NARS seismic profile 20 Neogene basins, Iberian Peninsula 223, 226, 227 Neogene uplift 113, 114 Neoproterozoic accretion 484-486 Neoproterozoic subduction 392

INDEX

Neotethys 5, 59, 104, 130, 399 active margin 249, 251,252 closure 283 evolution 70-71 mrlange 244, 245 nomenclature 66-69 opening 358, 359 ophiolite emplacement 237, 238, 248 reconstructions 61-68 rift 50 rotation 253 nephelinite 150, 151, 152, 161 Nevadella [trilobite] 513 New Red Sandstone 94 NFP transect, Alps 134, 136, 139, 143 seismic profiles 130, 132, 135, 142, 131 Nicholsonella [trepostom] 513 Nordaustlandet Terrane 507-508 Normannian complex 300 North Aegean Trough 272 North African plate margin 247 North Anatolian Fault 272 North Atlantic 120 rift system 59 opening 51-52 uplift 123 North Danish Basin 47-50 North Danish-Polish Trough 117 North German Basin 356, 360, 378 North Kara Terrane 508, 510-513, 515 sedimentary succession 512 suture 511 North Sea Rift System 29, 50, 117, 123 North Sea subsidence 52 Northern Permian Basin 49, 355, 363, 365 Northumberland graben 45-46 Norwegian-Greenland Sea 508 Novaya Zemlya 509-510 nuclear explosion, Murmansk-Kizil 523 obduction Apulian plate 153 Beja-Acebuches 348, 351 Lizard complex 300 Saxo-Thuringian ocean 338 Urals 423, 424 Occam model 547 Ocean Island Basalt 157, 172, 174, 177, 179, 181 oceanic spreading 487-490 octupole component 86 Old Red Sandstone facies 512 Scandian orogeny (Silurian) 455 Oligocene volcanic activity 149 Oligocene-Miocene continental deposits 226, 227 olistolith 71 olistostrome 72, 227, 283, 398 see also debris flow deposit olivine conductivity 547 Oman-type ophiolite 235-238, 240, 244, 253 sedimentary cover 244 ophiolite 4, 71, 72, 398 Anatolides 63 Arabian 69 Archaean 565, 627, 632-634, 640 Baer-Bassit 271 Balkan 238-244 Bay of Islands 235 Croatian 73 Guevgueli 248 Palaeozoic 70, 99, 286, 356 Uralides 515 ophiolite obduction Apulian plate 153 Beja-Acebuches 348, 351 Dinarides 59, 75 ophiolite protolith, Bohemia 298 ophiolite, east Mediterranean 235-254 active margin (Cordilleran-type) 237, 247-253 collision-trench 235-237

passive margin (Oman-type) 235-237, 238, 240, 244 volcaniclastic sediments 249, 250 orogenic chains 224 orogenic magma 168, 178, 181 orogenic wedge 191,203 Oslo Graben 47-48 Oslo Rift 365-367 geophysics 376 Palaeozoic 27-28 post-Variscan reactivation 380 rheological model 377, 378 sedimentary fill 363 volcanism 360-361 PACE network 300 Pagetiellus [trilobite] 513 Palaeocene volcanic activity 148 palaeogeography Baltic Basin, Silurian 458-459 Devonian 464 East European Craton depocentres 451 Laurussia 394 Oman ophiolite 239 Palaeotethys 398, 399 Permian 355-359 Tethys 399 Palaeolenus [trilobite] 513 palaeomagnetism 193 Neotethys rotation 253 plate reconstruction 83-94 Palaeoproterozoic basement 389 Palaeotethys 5, 57, 58, 94, 99 accretionary complex 75 closing 358, 359 forearc sequence 69-70 mid-ocean ridge basalt 69 nomenclature 66-69 opening 333, 337, 339, 359, 357 ophiolite emplacement 237-238, 239 palaeogeography 398, 399 reconstructions 60-68, 91-93 sequence in Iran 69 subduction 59, 281 Palaeozoic orogens 19-28 Panafrican deformation see Cadomian orogeny PANCARDI project 192, 207 Pangaea 57, 86, 104, 355,359, 416 break-up 5, 50, 75, 266, 363 formation 293 reconstruction 92-93, 94 Pannonian Basin 4, 29-30, 52, 117 crustal thickness 200 extrusion tectonics 202 geophysical data 21, 23 lithosphere strength 123-124 lithospheric structure 200-202 lithosphere thickness 24, 201, 202 magmatism 179, 181, 195, 196-198 MORB diagram 173 pre-Neogene basement 194 seismic reflection section 197, 199 stratigraphy 195, 196-198 tectonic framework 192-193, 195 topography 215 volcanic deposits 195 volcanism, Cenozoic 152, 154 Pannonian, basin evolution 207-208 back-arc extension 170 deformation 198-200, 214-216 depositional environment 195, 196 formation 191-203 inversion 195-196, 208, 214, 216, 219 rifting 208-209 Pannonian-Carpathian system lithosphere strength 213-215 subsidence 208, 209, 211,212 stretching 209, 210, 214, 216-217

659

thermomechanical modelling 207-219 Pannotia supercontinent 296, 328 break-up 299, 303 Pantelleria, magmatism 153, 174 Panthalassic Ocean 86, 87, 359 Paphlagonian Ocean 71 Paradoxides [trilobite] 326, 513 Paratethys stages, 195 Paris Basin 50, 104, 119, 376 magnetic anomaly 49 subsidence curve 106, 107 Parnassus block 239 partial melting 158 and conductivity 542 and velocity 156 Partnach Basin 72 passive margin 6, 120, 123,359 Arabia 237 Caledonian collision 320 East European Craton 327, 328, 475-476 east Mediterranean 267, 272 Eurasian Arctic 509 inversion 122 Neoproterozoic 313 ophiolites 235-237, 238, 240, 244 Ordovician 486 peri-Tornquist 449-454, 457 passive rift 27-28, 30 inversion 208, 209 Peaceful Nuclear Explosion Profile Quartz 17, 26 Pechora Basin 470-472, 522-525, 531-534 igneous rocks 533 sedimentary cover 533 Pelagian block 268, 269, 270, 272 delamination 273 Pelagonian carbonate platform 240, 241 Pelgonian terrane 59, 390 Peltura [trilobite] 513 Penninic front 368 Periadriatic Line 200 magmatism 171, 172, 177-179 Periadriatic Lineament 140 Peri-Caspian Basin 451-452, 456-457, 468 -470 peridotites 30, 158, 160 mantle temperature 15 peri-Uralian basins 449-451,456-457 Permian basin evolution 355-381 Permian basins basin fill 355, 362 chronology 363 development 376-380 France 367-370 Iberian succession 373, 374 isopach maps 364 magmatism 358 marine deposits 373 modelling and basin history 376-379 palaeogeography 355-359 rifting 379-380 syn-rift sequence 71 Switzerland 367, 368 Variscan foreland (externides) 360, 362-367 Variscan internides 367-374 Permian orogens 5 Permian peri-glacial fauna 71 Peronopsis [trilobite] 326 Perunica (Bohemia) 85, 89, 293 petrogenesis, Cenozoic magma 178-182 Phlegrean Fields, volcanism 175 phonolite 152 Piedmont-Ligurian ocean 129 Pindos 63-67, 70 domain 73 nappe 68 ophiolite 238, 239, 242, 243 Pindos Zone 390 plate margin, Anatolia 389

660

plate reconstruction 57, 498 Armorican Terrane Assemblage 325 Cambrian 84, 87-93 Carboniferous 91, 92 Cretaceous 225 Gondwana 84, 87-91 Neotethys 60-68, 83-94 Ordovician 88-89 Palaeotethys 60-68, 91-94 Pangaea 92-93 Permian 5, 92, 93 Permian-Triassic 359 Precambrian 485 Silurian-Devonian 90, 91 Tethys 6 0 - 6 7 Urals 416 Vendian 84 plate tectonic evolution 208 Precambrian 15 Svecofennia 571-573 Urals 424 Variscides 334, 336, 351 plateau basalt 152 platform carbonates Palaeozoic 450 and ophiolites 240-242, 245, 249 platinum-bearing belt, Urals 443-447 playa lake 363 Plinian eruption 150 Pliocene volcanic activity 149 plume activity 379 plume induced magmatism 328 plume-related structures 20 geophysical data 21 Po basin 29 Po Plain 131, 153 Podlasie basin 449 Poland, seismic profile 375-376 transect 303 POLAR seismic profile 583 Polish Basin 364-365 cross-section 365 inversion 378-379 Polish Trough 49-50 POLONAISE seismic survey East European Craton crustal model 304, 304 Fennoscandia 6, 26, 313, 314 Trans-European Suture Zone 318, 319 interpretation 375,608-615 Pontes Basin 226 Pontide volcanic arc 170 Pontides 63, 93, 389-390, 391-396 molasse, end-Variscan 400 suture 393 Pontides, Karakaya Complex 238, 249, 250 pop-up structure 226 potential field data 521-522 Precambrian 422 lithosphere 11-17 see also Archaean, Neoproterozoic and Proterozoic Precordillera Terrane 87 Proterozoic crust, East European Craton 606-608 Protoatlantic 99 pull-apart basin 101, 195 pull-apart structure 361 Cornwall 380 Pyrenees 65 Carboniferous-Permian rocks 373 fold and thrust belt 225 geophysical data 22 lithosphere thickness 24 rift 75 volcanism 372 pyroclastic deposits 153, 154 Permian 361 Racha-Lechkhumy Fault 278, 279, 283 radiolarites 238, 240, 243 Ran Ocean 86

INDEX

Rayleigh wave 313, 315, 317, 552, 553 Rechnitz window 193 reconstruction see plate reconstruction Rheic Ocean 57, 58, 293, 327, 333, 349 suture 46, 83, 88-89, 327, 359 Rheic Suture 297, 299, 300 Rhenish Massif 97, 336 volcanic centre 51 Rheno-Hercynian Belt 45 Rheno-Hercynian Ocean 58, 71,334, 375 subduction 336, 338 Rheno-Hercynian Shelf 45 Rheno-Hercynian suture 48, 103 Rheno-Hercynian terrane 85, 90, 93, 94 rheology 116, 123,213-215,218, 319 Rhine Graben 30, 150, 154 lithosphere thickness 24 Rhine Rift System 117 Rhinish Massif, volcanism 148, 150, 155 Rhodes Basin, seisnfic profile 272 Rhodope Massif 398, 399 rhyolite dome 362 rift basin 49-59 Norwegian 120, 121 Permo-Triassic 5, 7 rifting 104-105, 333, 488-490 Caucasus 278, 280-282 Cenozoic 99 and magmatism 149-150, 154 Cretaceous-Palaeocene 50-51 East European Platform 18 Fennoscandia 583-585 Levantine Basin 266-267, 272 and lithosphere strength 121 maps 47-48 Mesozoic 266 Neotethyan 264, 270, 271 Palaeozoic 27-28, 463-476 Pannonian Basin 208 Strait of Sicily 268-270, 272 Triassic 223, 400 Urals 450-451,457 Variscan 358, 363, 376-380 Ringkcbing-Fyn High 313, 314, 319 Roccamonfina, mantle-source contamination 179 Rockall-Faeroe Bank 51 Rockall-Faeroe Trough 50 Rodinia 3, 7, 296 break-up 299, 303, 449-454, 457 Roer Valley Graben 120 roll back 35,273 Carpathians 210 Pannonian Basin 191, 193, 202, 203 Roman province 174 Romanian Terrane 298 Rondonian event 323, 326, 328 Rondonian-type crust 296, 305 rotation Baltica 447 East European Craton 293 Neotethys 253 Troodos 251 Rotliegend 363, 365, 366 depth to base (Upper) 365 volcanism 376 Russia, northwest geological provinces 525, 528 geophysical provinces 522 geophysical studies 541 state geophysical company 524 Saalian unconformity 362 Saar-Nahe Basin 49, 370, 371 Saar-Nahe Trough 103 sabkha 365 Sakarya domain 72 Sakarya Zone 328, 389-391, 395-397, 400 Sakmara arc 447 salt see evaporites

Sardinia magmatism calc-alkaline magmatism 178, 179 intraplate 153, 155 Sarmatian terrane 605-606, 617 map 7 Saros Trough 272 Saxo-Thuringian Basin 100, 335 ocean closure 336, 338, 339 Saxo-Thuringian Suture 300 Saxo-Thurngian terranes 297, 305 Scandian event 293 Scandian orogeny 455, 460, 509 Scandinavian basin 453-454, 455 Scandinavian Caledonides 516 schist as conductor 544 Schmidtiellius 296 Scythian Platform 3, 5, 7, 277, 279, 391,400 sea-floor spreading 584-585,593 east Mediterranean 272 sedimentary basins 470-472 East European Craton 449-460 subsidence curves 450, 453, 454 Iberia 225, 226-227 sedimentary sequence 493, 490-493 Eurasian Arctic 508, 511-513 Permian basins 362-374 Permo-Triassic 491 Proterozoic 529, 530, 531,533 sedimentary succession, post-rift Greater Caucasus 282-283 seismic activity, Vrancea 201,213, 216 seismic data/studies 6 East European Craton 15-17, 608-615 Fennoscandian Shield 547-554 Iberia 346-348, 351 NW Russia 523-525 Strait of Sicily 269-270 Urals 27 Variscides 25 seismic profile Antalya Basin 271 Baltic Basin 456 Fennoscandia 582 Ionian Basin 265 Levant margin 267 Linosa Trough 269 Malta Trough 270 Pannonian Basin 197, 199 Rhodes Basin 272 Skagerrak Graben 367 Urals foreland 429 seismic velocity and temperature 115 seismic, deep sounding Caucasus 278 Semail nappe 69 Semail Ocean, subduction 71 Semail ophiolite 235,244 serpentinite 123,243, 251 Urals 407, 423 Severnaya Zemlya archipelago 511, 512, 513, 515 shear wave velocity 6 Shelvian deformation 297 soft collision 301 Shemshak basin Iran 59 shield volcano 150 shoshonite 99, 181 Siberia 507, 508 Siberian terrane 86 Siberian traps 94 Sibumasu Terrane 87 Sicily intraplate magmatism 153, 155 transform fault 270 Silurian collision 328 Sitia microcontinent 73 Skagerrak Graben 363, 365, 366, 367 slab detachment 28, 29, 47-49, 52 Alpine arc 139, 143 Alpine 157, 167, 176, 181 Cenozoic rift 97, 101

INDEX

Iberia 51 magmatism 169, 175 Pannonian 52, 191,192, 196, 202 Permian 379 Rheno-Hercynian 99, 103 seismicity 170 Variscides 351,358 slab extrusion, Urals 423 slab roll-back 57, 59, 60, 62, 167, 171 see also roll-back slope stability 219 soft collision 291,293, 297, 301 Solenopleura [trilobite] 513 Sorgenfrei Line 47 Sorgenfrei-Tornquist Zone 291,301,300 South Hewett Lineament 297 South Portuguese Zone 351 Southern Pennine Basin 49, 355, 359, 363, 376 crustal thinning 102-104 Spain, Cenozoic volcanism 153, 155 spinel lherzolite (xenolith) 161 Sporades Trough 272 Srednogorie arc 61, 62-63 Strandja Massif 389, 391,395, 397 accretionary complex 398 stratovolcano 152, 177 strength and deformation 115-119 in rifts 121 strength lithosphere Pannonian-Carpathian 213-215 strength map 118, 208 strength models 116 stress field 123,214, 216, 218-219, 284 map 115 strike-slip faults 284 Pyrenees 370 Stromboli, volcanism 174, 175, 179 structure, Greater Caucasus 278 structure, Pannonian Basin 194 subduction 358 Cenozoic 29-30 Carboniferous 489 Palaeotethyan 69 subduction and conductive material 546 subduction and ophiolite emplacement 235, 236 active margin (Cordilleran-type) 247-253 mid-ocean ridge 237, 238, 240, 243 suprasubduction zone 237, 239-240, 245 subduction polarity 167,201,203,327, 349 Bohemia 298 Western and Eastern Alps 139-140, 143 subduction processes 167-203 active zones 170, 175 zones and migration 170 subduction relicts 553 subduction zones Avalonia 329 Calabrian-Hellenic arcs 264 Cyprus 247 European plate 134 Fennoscandia 585-589 Kazakhstan 26 Palaeoproterozoic 12, 15 Uralides 421-424 Variscides 338, 339 subduction zone, properties of 19 subduction-related magmas 178 subsidence curve 105, 106-107 Norwegian margin 122 Pannonian- Carpathian 211, 212 Pechora Basin 450, 453 subsidence modelling 105 Sudetian Phase 356 sulphide deposits 349, 351,634 as conductor 542, 545 suture 45, 99, 459 Belarus 605 Bohemian-Moldanubian 45 east Mediterranean 237, 238 Iberia 348, 351

Uralian 1,414, 415, 443 Variscides 333, 334 suture zones 356, 357 suture, East European Craton 299-300 Svecofennia, geological map 563 Svecofennian orogen 561-574 Svecofennian Province geophysical data 21-22 lithosphere thickness 24 SVEKALAPKO 313, 320, 521-536, 579 participants 555 Sweden, TOR project 314 Switzerland, Permo-Carboniferous basin 367 Tagil arc 422, 424-425 arc-continent collision 407-408, 410 crustal structure 414-415 platinum-bearing belt 421 seismic 437 Taimyr Terrane 510, 515, 516 geology map 511 Tajo Basin 226-227, 228, 231 Tauern Window 136, 140, 193, 200, 201 Tauric-Anatolian plate 64 Tauride carbonate platform 245, 246 Tauride ophiolite 247, 248, 249, 252 Tavas Nappe 69 tectonic evolution 11-33,488 tectonic map Anatolia 392, 393 east Mediterranean 236, 237, 266, 390 Europe 1, 12 Fennoscandia 580-581 Greater Caucasus 280 Urals 422 tectonic units 483 tectono stratigraphy Evia ophiolite 240, 241 Lycian ophiolite 245, 246, 247 Mersin ophiolite 245, 247 Oman ophiolite 236 Teisseyre-Tornquist Zone 27, 33,314 East European Craton 291,300 Tekirova ophiolite 251 temperature, mantle 12-14 terrane analysis 293-294 terrane defined 389 terrane reconstruction see plate reconstruction terranes 84, 498 Caledonide 1 East European Craton 291-293,602-606 Fennoscsandian Shield 564-565,568, 569 terranes, timing of break-up Vendian-Permian 84 TESZ see Trans-European Suture Zone Tethyan collision zone 154 Tethyan inversion 50 Tethyan Ocean 223,396-399 closure 400 evolution 65-75 Tethyan oceans 359 age of 398 reconstructions 6 0 - 6 7 Tethyan ophiolite 239, 248 Tethys rift system 59 Tethys suture 68, 277 thermal age 208, 215 thermal anomaly 156, 376 thermal data 16, 17 Uralides 26 thermal decay curve 99 thermal destabilization 101, 103, 107 thermal evolution and tiffing 120 thermal model 15, 16, 551 lithosphere thickness 24, 26, 27-28, 31 southern Europe 30 Variscides 25 thermal regime Bohemian Massif 337 Palaeoproterozoic terranes 603

661

thermal sag basin 49-50, 107 thermal structure 116 thermal subsidence intra-cratonic basins 44-45, 50 Pannonian Basin 195 Rotliegend 363 thermal subsidence and rifting 104-105 thermal thickness 30 thermal thinning 124 thermal-mechanical controls on Alpine deformation 113-123 thermo-mechanical model Pannonian- Carpathian system 207 - 219 Permian basins 377 thermotectonic age 118-119 thickness of lithosphere 12, 16, 20, 24, 156 Variscides 25 thickness, Variscan crust 103, 104 thinning of crust 44, 49-50 tholeiite 150, 153, 162, 174, 243 Archaean 630-634 tholeiitic basalt 172, 421 Thor Suture 313, 314, 317, 319, 320 East European Craton 297, 299-300, 301,304 Thor-Tornquist suture 45 thrust faults 284 thrust systems, Iberia 224, 226 Thulean flood basalt province 122 Timan Range 523, 524, 525, 530-531 seismic profile 533 Timanian Ocean 449 suture 452 Timanian Orogeny 3-4, 86, 532 Timanide Orogen 470, 507-516, 530 Timanides 1, 4, 5, 6 Tisza-Dacia terrane 193, 194, 200, 202 extrusion tectonics 202 tomography 6, 11, 13, 16 Fennoscandia experiment 549-554 high-resolution teleseismic 129-143 lower lithosphere structure 137-138 mantle model 136, 137 thermo-mechanical controls 119 see also TOR project topographic map East European Craton 483 Europe 114, 148 southeast 483 Greater Caucasus 278 Iberian Peninsula 228 Netherlands 120 topography and tectonics 337 Alpine deformation 113, 114, 123 Greater Caucasus 284 Iberian Peninsula 223, 227-231 Pannonian Basin 215, 216 Variscan 338 TOR project 313-319, 320 earthquakes 314, 315, 319 P-wave travel time residuals 316 intraplate magmatism 156 lithosphere thickness 24, 31 seismic experiment 551, 552 seismic, Pannonian Basin 200-202 seismograph location 314 subduction and magmatism 171, 176, 177, 180 tomography experiment 18, 32-33 Tornquist Ocean 83, 88, 89 Tornquist Sea 293,296 closure 299, 301 Tornquist Zone 325, 358 inversion 319 subduction 326-327 TOR data 314, 316, 317, 318, 320 Tornquist-Sorgenfrei Line 49 Tornquist-Teisseyre Zone 377, 379 trace element ratio 158, 160 trace elements, data sources 157

662

TRANSALP transect, Alps 117, 131, 136, 138-143 seismic profile 130, 132, 135 working group 29, 200, 201 transects Adriatic microplate 136 Alps 131, 132, 141, 142 Avalonia- Rhenohercynian 302 Caledonide 304 Carpathian-Libyan 74 Danube Basin 198 Dniepr-Donets Basin 466, 468 East European Craton 301, 302, 303 Fennoscandian Shield 567, 573, 582, 617, 618 German Variscides 305, 335 Greater Caucasus 279, 280 Iberian Massif 346, 347, 350 Lod~ve Basin 371 Massif Central 369 Montagne Noire 369 North Sea 304 Pechora Basin 472, 510, 534 Peri-Caspian Basin 469 Polish Trough 365 Sarmatia 617 southern France, Permian Basin 371 Svecofennia 567 Timan Range 534 Urals 410, 412, 413, 415, 422 Variscan 303, 305 Trans-European Suture Zone Baltica 516 East European Craton 291 Trans-European Suture Zone 11, 83, 364 defined 1-2 location 5, 7, 12 dispersal, peri-Tethyan 299 gravity 17 seismic velocity 18 crustal thickness 18, 22 crystalline crust 375 lithosphere properties 18-24 lithosphere, deep 313,319, 362 Palaeozoic accretion 323 transform fault, Strait of Sicily 270 transpression 6 Variscan 345-351 Transylvanian Basin 202, 215- 216 Cenozoic volcanism 152 trapdoor basin 49, 101 trench-passive margin collision 247 Triassic basin, isopach map 105 trilete spores 93 trilobite provinciality 89, 298 trilobites 328, 513 Troodos ophiolites 235, 247, 252, 253 rotation 251 Tulcea Terrane 298 Turkey 389-410 Tyrrhenian Basin 272 Ukrainian Shield 11,463 ultramafic complex 443-447 unconformity Bathonian 281 Kan'on River 513 Oligo-Miocene 283 Pechora Basin 450 Vendian 515 underplating 45, 574 magmatism 584 Upper Rhine Graben 51, 52, 97, 103, 107 Cenozoic tectonism 114, 120 Urach province 151, 159 Uralian Ocean 515, 516 subduction 509 Uralian orogeny 86, 93, 476

INDEX

Uralide orogen 444, 507 Uralides 4, 6, 24, 26-27 Maksyutov Complex 421-422 subduction and collision 421-425 Uralides tectonism 409-417 arc-continent collision 407-409 crustal features 434-436 folds and thrust belt 411-413,416-417 granitoid emplacement 411,413-415 ocean closure 414-416, 443-447 seismic profile 439 strike-slip faulting 411,413-416 subduction 409-411 Urals 26 geophysical data 21 lithosphere thickness 24 Urals foreland basin 533 Urals seismic reflection survey 427-440 crustal root 439 crustal structure 433-438 mantle reflectors 438-439 URSEIS seismic survey 407, 408, 411,421,422 interpretation 414-416 seismic profile 20, 26 Valais ocean (Alpine Tethys) 129 Valence graben 114 Valencia Trough 226, 227, 230, 231 Vardar 64-66 Vardar Ocean 170, 248, 253 subduction 59, 61-63, 65, 67, 75 Mesozoic-Cenozoic 176 volcanism 247, 248 Vardar zone volcanism 247-253 variation diagrams 158-160, 161 Carpathian-Pannonian region 173, 180 Variscan cycle 43, 45, 57 Variscan Deformation Front 98, 100, 356, 363, 375, 376 Variscan foreland 305 Variscan orogen 97, 333-340 Cenozoic Rift System 97-101, 104 Balkans and Black Sea 399-400 Iberia 345-351 Variscan orogeny, 277, 325, 327, 355-359 East European Craton 297 compression rate 357 final stage 372 fracture system 359 lithosphere 379 magmatism 359-362 Massif Central 367 molasse 335, 336, 400 plate configuration 359 reactivation 103 suture 357 tectonic map 356 unroofing 101, 102 Variscan plate model 334, 335, 339 Variscan terranes boundaries 102 collision chronology 348-350 cross-section 350 depositional history 348 Variscan thrusts, Germany 303 Variscides (Hercynian) 20, 93, 333, 334, 337 development 3-6 geophysical data 21 Himalayan-type and Andean-type 101 location map 1 lithosphere thickness 24, 19-28 seismic model 25 thermal model 26 Variscides orogen 97 terrane 100 VARNET seismic profile 24 Velay, Cenozoic volcanism 152 velocity 531-532 TOR data 313-319, 320

velocity indicating partial melting 156 velocity model EUROBRIDGE 612 Fennoscandia 549-550, 553, 554, 555 velocity variation, P-wave 137, 138, 140, 142 Veneto province 153, 154 vent, volcanic 154 Veporic nappe 75 Vestfold Graben 365, 366 Vesuvius, volcanism 175 Viking Graben 30 micro seismic activity 379 Vivarais, Cenozoic volcanism 152 Vogelsberg volcanic complex 150, 155, 162 volcanic arc 168 volcanic fields 150-155 in Europe 148, 149, 151, 168 Greater Caucasus 280-282, 285 Iberia 230, 231, 360, 372 Mid-German Crystalline High 327, 329 Norwegian margin 122 Oman 240 Polish Basin 364 Vardar zone 247 volcanicity 349-351 Cenozoic 154-155 chronology 150 Permian 365, 376 Tremadoc 513 Permo-Carboniferous 102 Variscan 335, 336, 360-361,362 volcanoes, active 153, 174, 175,280 Volgo-Uralia 7 Vosges magmatism 148, 151,338 Vosges-Black Forest arch 52, 53, 97 Voykar massif 443,447 Vrancea slab 196, 202 Vulcano, volcanism 174 Vulsini, mantle source contamination 179 wedge structures, lower crust 134-135 wedges, orogenic 62 Wessex Basin 50 West Ny Friesland Terrane 507 Westerwald volcanic activity 150, 155 Whin Sill Complex 361,381 Wilson cycle 277, 584 wrench faults 268, 380, 416 Carboniferous-Permian 45, 48 Cenozoic Rift System 97, 99, 100-102 map of Europe 48, 49 xenolith 12, 13, 15, 381,590 anorthosite 302 Cenozoic Rift System 103 crustal 361,380 garnet 196 lherzolite 25, 161,553 mantle 31, 156, 361,551,554 ultramafic 172, 176 Xystridura [trilobite] 513 Zagrab Fault 193, line 201 Zechstein salt 365, 375, 379 Zechstein Sea 59, 94 zircon age 326, 327, 333,336-338, 362 Anatolia 390, 392, 395, 396 Archaean 564, 632, 635 Dnestr 457 East European Craton 292, 293, 296, 297 Fennoscandia 583, 586, 587, 588 Proterozoic 532 see also isotope age zircon analysis, U-Pb 443-447 zircon, detrital (Palaeozoic) 513 zircon, U-Pb isotope data 445, 446 zircons, inherited 380