The Neoproterozoic Timanide Orogen Of Eastern Baltica (Geological Society Memoirs No. 30)

The Neoproterozoic Timanide Orogen of Eastern Baltica

Geological Society Special Publications Society Book Editors R. J. PANKHURST (CHIEF EDITOR) P. DOYLE F. J. GREGORY J. S. GRIFFITHS A. J. HARTLEY R. E. HOLDSWORTH

J. A. HOWE P. T. LEAT A. C. MORTON N. S. ROBINS J. P. TURNER

Society books reviewing 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 has a team of Book Editors (listed above) who 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. Geological Society books are included in the ISI Index of Scientific Book Contents, but they do not have an impact factor, the latter being applicable only to journals. More information about submitting a proposal and producing a Society Publication can be found on the Society's 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. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30. BECKHOLMEN, M. & GLODNY, J. 2004. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125–134.

GEOLOGICAL SOCIETY MEMOIRS NO. 30

The Neoproterozoic Timanide Orogen of Eastern Baltica EDITED BY

DAVID G. GEE Uppsala University, Sweden and

VICTORIA PEASE Stockholm University, Sweden

2004 Published by The Geological Society London

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Contents Acknowledgments

vi

Introduction GEE, D. G. & PEASE, V. The Neoproterozoic Timanide Orogen of eastern Baltica: introduction

1

Pre-orogenic successions and foreland basins ROBERTS, D., SIEDLECKA, A. & OLOVYANISHNIKOV, V. G. Neoproterozoic, passive-margin, sedimentary systems of the Kanin Peninsula, and northern and central Timan, NW Russia MASLOV, A. V. Riphean and Vendian sedimentary sequences of the Timanides and Uralides, the eastern periphery of the East European Craton GRAZHDANKIN, D. Late Neoproterozoic sedimentation in the Timan foreland Timanide fold-and-thrust belt ROBERTS, D. & OLOVYANISHNIKOV, V. Structural and tectonic development of the Timanide orogen LORENZ, H., PYSTIN, A. M., OLOVYANISHNIKOV, V. G. & GEE, D. G. Neoproterozoic high-grade metamorphism of the Kanin Peninsula, Timanide Orogen, northern Russia LARIONOV, A. N., ANDREICHEV, V. A. & GEE, D. G. The Vendian alkaline igneous suite of northern Timan: ion microprobe U-Pb zircon ages of gabbros and syenite Timanide hinterland PEASE, V., DOVZHIKOVA, E., BELIAKOVA, L. & GEE, D. G. Late Neoproterozoic granitoid magmatism in the basement to the Pechora Basin, NW Russia: geochemical constraints indicate westward subduction beneath NE Baltica GLODNY, J., PEASE, V., MONTERO, P., AUSTRHEIM, H. & RUSIN, A. I. Protolith ages of eclogites, Marun-Keu Complex, Polar Urals, Russia: implications for the pre- and early Uralian evolution of the northeastern European continental margin REMIZOV, D. & PEASE, V. The Dzela complex, Polar Urals, Russia: a Neoproterozoic island arc BECKHOLMEN, M. & GLODNY, J. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia Post-Timanian Palaeozoic platform successions KORAGO, E. A., KOVALEVA, G. N., LOPATIN, B. G. & ORGO, V. V. The Precambrian rocks of Novaya Zemlya BOGOLEPOVA, O. K. & GEE, D. G. Early Palaeozoic unconformity across the Timanides, NW Russia MOCZYDLOWSKA, M., STOCKFORS, M. & POPOV, L. Late Cambrian relative age constraints by acritarchs on the post-Timanian deposition on Kolguev Island, Arctic Russia Regional relationships and correlations SIEDLECKA, A., ROBERTS, D., NYSTUEN, J. P. & OLOVYANISHNIKOV, V. G. Northeastern and northwestern margins of Baltica in Neoproterozoic time: evidence from the Timanian and Caledonian Orogens GEE, D. G. & TEBEN'KOV, A. M. Svalbard: a fragment of the Laurentian margin JOHANSSON, A., LARIONOV, A. N., GEE, D. G., OHTA, Y., TEBENKOV, A. M. & SANDELIN, S. Grenvillian and Caledonian tectono-magmatic activity in northeasternmost Svalbard VERNIKOVSKY, V. A., VERNIKOVSKAYA, A. E., PEASE, V. & GEE, D. G. Neoproterozoic Orogeny along th margins of Siberia Index

5 19 37 47 59 69

75 87 107 125

135 145 159

169 191 207 233 249

Acknowledgments Much of the research reported in this Memoir began during the IUGS-IUGG Inter-Union Commission on the Lithopshere's EUROPROBE programme in the context of the URALIDES and TIMPEBAR (Timan–Pechora–Barents Sea) projects. For ten years (1992-2001) EUROPROBE received much appreciated sponsorship from the European Science Foundation member organizations. Our Timanide research has been promoted by the Brussels-based INTAS programme and two INTAS projects—HALE 96-1941 (High Arctic Lithosphere of Europe) and NEMLOR 01-0762 (Northern European Margin and Lomonosov Ridge) have provided support for Russian colleagues and joint fieldwork. Grants from the Swedish Royal Academy of Sciences and financing of expeditions by the Swedish Polar Research Secretariat have, likewise, promoted our international collaboration. It is a pleasure to acknowledge the importance of these contributions to this synthesis of eastern European geodynamics. During our work on the Timanides, we have only had access via translation to a small part of the relevant Soviet literature. However, our Russian colleagues have participated and, in many cases, guided EUROPROBE studies and drawn our attention to the main controversies. We particularly appreciate their collaboration and friendship. In addition, we would like to express our appreciation for help with production of the Memoir, particularly to Dr Olga K. Bogolepova in Uppsala, the staff of the Geological Society Publishing House in Bath, and the many reviewers (below) who have improved the quality of the various chapters with thorough evaluations of scientific merit and, often, with help improving the linguistics. Reviewers: P. G. Andreasson, A. Andresen, H. Austrhiem, S. Bogdanova, D. Brown, F. Corfu, I. Dalziel, P. F. Friend, R. Gabrielsen, R. Gorbatschev, P. Gromet, N. Henriksen, A. K. Higgins, R. Ingersoll, F. Kalsbeek, M. Lindstrom, K. Ludwig, M. Moczydlowska-Vidal, A. Maslov, V. Melezhik, V. Puchkov, D. Roberts, J. Scarrow, F. Schaffer, R. Scott, A. Siedlecka, S. Sindern, R. Strachan, M. Tichomirova, A. Willner, M. Whitehouse.

The Neoproterozoic Timanide Orogen of eastern Baltica: introduction 1

D. G. GEE1 & V. PEASE2 Uppsala University, Department of Geosciences, Villavagen 16, 752 36, Uppsala, Sweden 2 Department of Geology and Geochemistry, Stockholm University, SE-106 91, Sweden

This volume was conceived during EUROPROBE's investigations into the dynamic evolution of the Palaeozoic Uralide Orogen and relationships northwards into the Eurasian high Arctic. During these European Science Foundation studies, the preservation of Neoproterozoic deformation over large regions of northern Europe became increasingly apparent. This mainly Vendian tectonic event is referred to as the Timanian Orogeny and became the focus of many recent and on-going investigations. Much progress has been made in understanding Timanian Orogeny and a Memoir synthesizing our current knowledge is not only timely, but also relevant to Neoproterozoic global tectonic reconstructions. The type area for the Timanide Orogen is located in the Timan Range of northwestern Russia, which separates the East European Craton from the Pechora Basin and Polar Urals. The orogen extends over a distance of at least 3000 km, from the southern Ural Mountains of Kazakhstan to the Varanger Peninsula of northernmost Norway, flanking the eastern margin of the older craton (Fig. 1). From the Timan Range, it reaches northeastwards below the thick Phanerozoic successions of the Pechora Basin and Barents Shelf (O'Leary et al. 2004), and reappears in the Polar Ural Mountains and northwards through Pai Khoi to Novaya Zemlya. Timanian orogeny thus influenced a vast region of northwestern Russia. The Phanerozoic cover, Arctic shelf areas and, further east, Uralian deformation, obscure the importance of this orogenic event for the geodynamic evolution of Europe. The Timanide Orogen has been referred to by various other names, most frequently as the 'Baikalides'. The term 'Baikalian Orogeny', with a type area along the southern margin of the Siberian Craton, was introduced by Edelstein (1923) and promoted by Shatsky (1963), and suggested a tectonic event that started in the Late Precambrian and finished in the Early Palaeozoic. Other authors prefer to restrict 'Baikalian' events to those that took place in the Neoproterozoic time interval of 850-650 Ma (e.g. Khomentovsky 2002). The term 'Baikalian' has also been used to designate a late Precambrian stratigraphic system in Siberia, corresponding to the Cryogenian of the IUGS International Stratigraphic Chart (2000). To avoid ambiguity, we advocate the use of the term Timanian' Orogeny to describe the late Neoproterozoic tectonic events documented along the eastern margin of the East European Craton, best exemplified in the Timan–Pechora region, and restrict the use of the term Baikalian to tectonic events associated with Siberia. For much of the last century, the dominating hypothesis for the evolution of northwestern Europe has explained Timanian tectono-thermal activity in terms of rift basin (aulacogen) inversion. Thick Neoproterozoic and partly Mesoproterozoic sedimentary successions were described and interpreted to separate blocks of older Precambrian crust that previously had been a part of the Archaean and Palaeoproterozoic core of Europe. Thus, Stille (1958) inferred that the Timanides were a result of deformation between the Fennoscandian Craton and an outboard continent, which he called Barentsia. Subsequent geophysical studies, particularly potential field, but also seismic, suggested a more

complex crustal evolution. Deep drilling (up to c. 5 km) of the Pechora Basin provided convincing evidence (Belyakova & Stepanenko 1991) that a broad belt of calc-alkaline igneous rocks flanked terrigenous slope-to-basin deposits of the Timan Range. Late Neoproterozoic granites carry Grenville-age zircon xenocrysts and complexes of this age were shown to exist further towards the hinterland within the Palaeozoic allochthons of the Subarctic Urals. Late Neoproterozoic ophiolites, albeit fragmented, were described from the Polar Urals (Dushin 1997). Thus, despite powerful resistance (e.g. Ivanov & Rusin 2000), an alternative hypothesis has emerged that favours the existence of a Timanian accretionary orogen, on the eroded roots of which were deposited the early to mid-Palaeozoic rifted and passive margin successions which flanked the Uralian ocean. Continent–ocean collision played an important role in the orogenic process and some authors (e.g. Sengor et al. 1993) have speculated on the possible continuity between the Timanian' and Uralian oceans; however, the question remains unresolved. The studies of the Timanides included in this Memoir are structured to provide a comprehensive overview of the orogen. The first three contributions treat the pre-Timanian rifted margin of the East European Craton. Roberts et al. describe the Neoproterozoic passive margin sedimentary successions of the Kanin Peninsula, and northern and central parts of the Timan Range. Maslov provides comprehensive descriptions of the Mesoproterozoic and Neoproterozoic (Riphean–Vendian) stratigraphy preserved within the Uralian foreland and western flank of the Ural Mountains, making regional correlations to the Timan–Pechora area. Grazhdankin follows with an overview of the late Neoproterozoic differential subsidence patterns of the East European Craton in the Mezin Basin SW of the Timan Range, significantly relating this to development of a Timanian foreland basin in the late Vendian. The magmatic, metamorphic and structural evolution of the Timanide Orogen is described regionally, divided into the Timan Range, the Pechora Basin, and Ural Mountains. Roberts & Olovyanishnikov present the structural and tectonic development of the Timanide Orogen in the Timan region. Larionov et al. present U–Pb ages on an alkaline igneous suite in northern Timan which provides constraints on the beginning of Timanian Orogeny. Using Neoproterozoic high-grade metamorphic rocks from the Kanin Peninsula, Lorenz et al. document P / T conditions associated with Timanian orogenesis. In the Pechora Basin region, drillcore samples of pre-Palaeozoic basement provide the foundations for our understanding of the pre-Palaeozoic events. Belyakova & Stepanenko's paper (1991) documenting the different structural and metamorphic zones within the basement to the Pechora Basin, is particularly important. New geochemical evidence from Dovzhikova et al. (2004) suggests that the Precambrian mafic complexes in the Pechora zone represent Neoproterozoic oceanic crust, probably accreted during Timanian orogenesis. Pease et al. provide geochemical evidence for the calc-alkaline affinity of Vendian granitoid rocks which are interpreted to indicate late-orogenic westward subduction beneath northeastern Baltica at about 560 Ma.

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 1-3. 0435-4052/04/$ 15 © The Geological Society of London 2004.

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Fig. 1. Geological map of the eastern margin of Baltica, showing the extent of the Timanides from the southern Urals to Novaya Zemlya and the Varanger Peninsula (VP).

THE NEOPROTEROZOIC TIMANIDE OROGEN OF EASTERN BALTICA: INTRODUCTION

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ute significantly to a clearer understanding of its role in the tectonic evolution of Baltica. Several different geological timescales are in routine use within the scientific community at present. Though the use of the IUGS International Stratigraphic Chart (2000) has been encouraged, the older International Stratigraphic Chart of Plumb (1991), as well as the Russian timescale (Keller & Chumakov 1983) in which Riphean and Vendian are used to subdivide parts of the Precambrian, have also been used. For the convenience of the reader, we provide a cross-reference for these three timescales (Fig. 2). Additionally, the International Commission of Stratigraphy has recently revised the Precambrian timescale (Gradstein et al. 2004), but it has not yet received wide usage. Regarding nomenclature and especially the translation of Russian Stratigraphic terms, a few more years are needed to achieve consensus on these matters. References Fig. 2. Comparison of Meso- and Neoproterozoic timescales used in this volume.

Within the Ural Mountains, the evidence for Timanian orogeny is fragmentary and the contributions are geographically restricted to the Polar and Northern Urals. Glodny et al. report Timanian protolith ages within the eclogitized Marun-Keu complex and discuss their implications for the pre- and early Uralian evolution of the northeastern European continental margin. Remizov & Pease present geochemical and U – Pb age data from the Dzela complex which indicate Neoproterozoic island arc magmatism. Beckholmen & Glodny follow with a description of, and age constraints for, blueschist metamorphism in the pre-Ordovician basement to the Kvarkush anticline, also interpreted within a Timanian tectonic framework. The sections on Timanian Orogeny are followed by descriptions of post-Timanian platform successions, which are important for interpreting the timing of orogeny and the post-Timanian return to a passive margin setting. These include assessment of the regional Early Palaeozoic unconformity across the Timanides (Bogolepova & Gee), as well as Late Cambrian age constraints from acritarchs of Kolguev Island on post-Timanian deposition (Moczydlowska et al.). Finally, regional correlations are explored in which it is concluded that Timanian Orogeny does not extend to Svalbard (Gee & Tebenkov; Johansson et al. ), but is present on Novaya Zemlaya (Korago et al.). Work in progress also suggests it influences Franz Josef land basement (Pease et al. 2001). Comparison is made between the Neoproterozoic passive margin of western Baltica, in the Scandinavian Caledonides, and contemporaneous orogeny in the Timanides (Siedlecka et al.). Similarities in the Neoproterozoic tectonic evolution of Baltica and Siberia are also explored (Vernikovsky et al. ). Syntheses of Timanian orogenic evolution have been provided by several authors (e.g. Sengor et al. 1993; Roberts & Siedlecka 2002; Dovzhikova et al. 2004; Gee 2004). The contributions presented in this Memoir will promote further elaboration. In pursuing research on the Timanides, critical aspects related to this orogeny have been identified for future work. The nature of the hinterland beneath the Pechora Basin and as it is exposed in the Ural Mountains needs more investigation. Determining the role and extent of subduction along the orogen and characterization of the arc-related magmatic rocks remain a critical point. Future collaborative studies with Russian partners which seek to understand Timanian orogenesis better will undoubtedly contrib-

BELYAKOVA, L. T. & STEPANENKO, V. Ya. 1991. Magmatism and geodynamics of the Baikalide Basement of the Pechora Syneclise. Doklady Akademii nauk SSSR (geologiya), 106–117 [in Russian]. DOVZHIKOVA, E., PEASE, V. & REMIZOV, D. 2004. Neoproterozoic island arc magmatism beneath the Pechora Basin, NW Russia. GFF, 126, 353-362. DUSHIN, V. A. 1997. Magmatism and Geodynamics of the Palaeocontinental Sector of the Northern part of the Urals. Nedra, Moscow, 211 pp [in Russian]. EDELSTEIN, Y. 1923. Tectonics and ore deposits of Siberia. Izv. Geol. Kommittee, 42, 23–50 [in Russian]. GEE, D. G. 2004. Timanides of northern Russia. In: Selley, R. C., Cocks, R. & Plimer, I. R. (eds), Encyclopedia of Geology. Elsevier, Amsterdam. GRADSTEIN, F., OGG, J., SMITH, A., BLEEKER, W. & LOURENS, L. 2004. A new geologic time scale with special reference to Precambrian and Neogene. Episodes, 27, 83–100. IUGS International Commission on Stratigraphy, 2000. International Stratigraphic chart, REMANE, J., CITA, M. B., DERCOURT, J., BOUYSSE, P., REPETTO, F. L. & Faure-MURET, A. (eds). Division of Earth Sciences, UNESCO. IVANOV, S. N. & RUSIN, A. I. 2000. Late Vendian tectonic evolution of the Urals. Geotektonika, 3, 21–32 [in Russian]. KELLER, B. M. & CHUMAKOV, N. M. 1983. Stratotype of Riphean Stratigraphy and Geochronology. Nauka, Moscow, 184 pp [in Russian]. KHOMENTOVSKY V. V. 2002. Baikalian of Siberia (850-650 Ma). Russian Geology and Geophysics, 43, 313–333. O'LEARY, N., WHITE, N., TULI, S., BASHILOV, V., KUPRIN, V., NATAPOV, L. & MACDONALD, D. 2004. Evolution of the Timan–Pechora and South Barents Sea basins. Geological Magazine, 141, 141–160. PEASE, V., GEE, D. & LOPATIN, B. 2001. Is Franz Joseph Land affected by Caledonian deformation? European Union of Geosciences Abstracts, 5, 757. PLUMB, K. A. 1991. New Precambrian time scale. Episodes, 14, 139–140. ROBERTS, D. & SIEDLECKA, A. 2002. Timanian orogenic deformation along the north eastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian–Cadomian connections. Tectonophysics, 352, 169-184. SENGOR, A. M. C., NATAL' IN, B. A. & BURTMAN, V. S. 1993. Evolution of the Altaid tectonic collage and Palaeozoic crustal growth in Eurasia. Nature, 364, 299-307. SHATSKY, N. S. 1963. On Cambrian-Proterozoic relations and Baikalian orogeny. Izbran. Trudi. M., Izd-vo Acad. Nauk SSSR 1, 581–587 [in Russian]. STILLE, H. 1958. Die assyntische Tektonik im geologischen Erdbild. Beihefte zum Geologischen Jahrbuch, 22, 255 pp.

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Neoproterozoic, passive-margin, sedimentary systems of the Kanin Peninsula, and northern and central Timan, NW Russia DAVID ROBERTS1, A. SIEDLECKA1 & V. G. OLOVYANISHNIKOV2 Geological Survey of Norway, N-7491 Trondheim, Norway (e-mail: [email protected]) 2 Institute of Geology, Komi Research Centre, Ural Division RAS, Pervomayskaya 54, Syktyvkar, 167000 Russia 1

Abstract: Neoproterozoic, slope-to-basin, lithostratigraphic successions are discontinuously exposed within the Timan Range, in NW Russia, NE of a faulted basinal margin that marks the outer edge of a former, fluvial to shallow-marine, pericratonic domain. The Mid to Late Riphean, deep-water depositional systems of the Kanin Peninsula, and northern and central Timan attain considerable thicknesses, up to 10 000 m in the case of Kanin Peninsula. Basements to these successions are nowhere exposed. Although the successions accumulated along a comparatively stable, passive margin of Baltica, there are notable differences in sedimentary facies from area to area. Whereas the successions in northern and central Timan preserve a record of relative stability, with sedimentation keeping pace with subsidence, the nearby Kanin succession shows evidence of repeated faulting. This may reflect a non-contemporaneity of the diverse successions or a segmentation of the basin margin. Comparisons are also made with deep-water, turbidite-fan systems in northwestern parts of the Timan–Varanger Belt, on the Rybachi and Varanger Peninsulas. The lateral differences in sedimentary facies in these areas, seen in relation to the situation in Timan and Kanin, do, in fact, suggest that the 1800 km long Timanian Basin margin may have been segmented, and possibly into sub-basinal domains.

Deep-water sedimentary facies, in particular turbidite systems, are good indicators of the plate-tectonic setting, development and morphology of a basin margin and slope, and the composition and topography of the hinterland. Such facies have been studied in varying detail in five separate areas of the extensive Timan– Varanger Belt (TVB) bordering Baltica along its northeastern and northern margins. In this account, we describe and interpret the Neoproterozoic slope-to-basin systems of central and northern Timan and compare them with the better known turbidite systems of the Rybachi and Varanger Peninsulas in the northwestern parts of the TVB. This approach illustrates the similarities and differences between the various successions and shows how they reflect variations in the tectonic development of the southwestern margin of the Timanian Basin. Regional setting The TVB consists of pericratonic and basinal regimes of sedimentation that together constitute the southwestern marginal part of the Timanian Orogen, the central outboard parts of which occur beneath the Palaeozoic and younger cover of the Pechora Basin and its continuation beneath the southern Barents Sea (Getsen 1991; Bogatsky et al 1996; Olovyanishnikov et al 2000; Roberts & Siedlecka 1999, 2002; Siedlecka et al. 2004) (Fig. 1). Common for all the exposed basinal systems along the TVB is that they represent either the oldest or the only exposed sedimentary facies association, they contain turbidites, and their substratum is unknown. Although their age constraints are poor, their Neoproterozoic (?Mid and Late Riphean), pre-Late Vendian age is fairly certain, based on biostratigraphic data, isotopic ages, the Vendian Timanian Orogeny and, in the NW, evidence of Varangerian glacial deposits and the Caledonian Orogeny (e.g. Vidal & Siedlecka 1983; Roberts 1995; Gee et al. 2000). The degree of metamorphism of the successions is low, mostly in lower greenschist facies or anchizone grade (Getsen 1987; Rice & Roberts 1995; Siedlecka & Roberts 1995; Roberts & Siedlecka 1999, 2002), with the exception of northern Timan and Kanin Peninsula where metamorphic grade reaches amphibolite facies (Lorenz et al 2004). Palaeocurrents, where known, are generally towards NE or ENE, i.e. from the continent Baltica and into the Timanian Basin. Any detailed chronostratigraphic correlation between the separate turbidite and associated sedimentary

facies systems within the broad Neoproterozoic interval is difficult. Criteria for definition of submarine turbidite systems Turbidite successions that result from deposition by repeatedly occurring turbidity currents commonly consist of 'Bouma sequences' (Bouma 1962). Loss of shear strength and resulting instability, eventually leading to redeposition, are favoured by high slope gradients and high rates of sedimentation. Typically, in all known modern and ancient turbidites, the model Bouma sequences are incomplete; mostly, the top and/or bottom intervals are missing. The lack of a bottom interval (the graded A interval) is most critical, as this provides the essential criterion for defining a turbidite. Several depositional sequences, therefore, have to be examined in order to establish if a sequence actually represents, or is a part of a turbidite system. Turbidites do not occur exclusively in submarine turbidite systems (STS). They may be present in other settings where turbidity currents are generated, e.g. in crevasse splays in rivers and deltas. In defining a STS, it is therefore necessary to establish the presence of a continuous succession of turbidites up to hundreds to thousands of metres in thickness. Such successions are most likely to accumulate at the foot of a major slope. Additional criteria include the presence of interbedded debrisflow deposits (debrites) and olistostromes, the latter occurring in the most proximal parts of a STS. Other characteristic features are load casts resulting from soft-sediment deformation, and slump folds and intraformational slump breccias, testifying to high sedimentation rates and accumulations on a slope. Deposition on a slope is also indicated by synsedimentary folds, rotated blocks and slump scars. Farther out, pelagic fallout and planktonic organisms/fossils may be present. In the case of the Neoproterozoic deposits the fossils are represented by acritarchs. Petrography of coarse-grained turbidites is a helpful tool in defining both the composition of the hinterland and the passive versus active nature of the basin margin and slope. In ancient, deformed and metamorphosed successions, purely sedimentological criteria may be partly or even largely obliterated. There is, thus, no single criterion for defining a STS; in order to establish its presence and a palaeocontinent constellation, several criteria and indications are required. With this in mind, we describe and propose interpretations for the Kanin-Timan basinal successions.

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 5-17. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. The Neoproterozoic Timan-Varanger Belt (TVB), with palaeocurrent directions indicated. VP, Varanger Peninsula; RP, Rybachi Peninsula; KP, Kanin Peninsula; NT, Northern Timan; VR, Vymskaya Ridge; TKFZ, Trollfjorden-Komagelva Fault Zone; SRFZ, Sredni-Rybachi Fault Zone; WTF, West Timan Fault; CTFZ, Central Timan Fault Zone; ETF, East Timan Fault; PK, Pechora-Kozhva Fault. Modified from Olovyanishnikov et al. (2000).

Kanin Peninsula Meso- and Neoproterozoic rocks occur in a NW-SE-elongated ridge, the Kanin Kamen Ridge, extending from Kanin Nos Cape in the NW to Cape Mikulkin in the SE (Figs 1 & 2). Southeastern coastal sections of the Kanin Kamen Ridge expose, in an anticline, an apparently continuous succession of interbedded clayey, silty and sandy rocks subdivided by Getsen (1975) into the Mikulkinskaya (oldest) and Tarkhanovskaya Series, now termed Groups. The youngest unit, the Tabuyevskaya Group, is poorly exposed along the Barents Sea coast in the northwestern part of the peninsula. The substratum of the c. 10 000 m thick succession comprising the three above-mentioned groups, here named the Kanin Kamen Supergroup, is unexposed and the top is erosional. The succession exhibits a decreasing degree of metamorphism

from amphibolite to lower greenschist facies, from bottom to top, and its main characteristics are summarized in Fig. 3. The description that follows is based largely on Olovyanishnikov (1998, 2000, 2001). In its lower part, the Mikulkinskaya Group (1500m) is composed mainly of massive, thick-bedded, fine-grained sandstone and siltstone (now represented by gneiss and mica schist). Cross-bedding and slump folds are present locally. The middle part is dominated by schistose, clayey-silty beds alternating with lenticular sandy-silty bodies. Higher up, the number of sandstone beds increases and there is a distinct alternation of thin- to medium-bedded sandstone, siltstone and claystone (schists). Upward-fining sequences c. 1.5 m thick and larger fining-up motifs 30-40 m in thickness (the flyschoid and transgressive cyclicity of Olovyanishnikov 2000) have also been observed. The amount of carbonate cement and the number of

Fig. 2. Geological map of Kanin Peninsula, showing the principal lithostratigraphic groups and formations of the Neoproterozoic succession. The map is simplified from Olovyanishnikov (2000).

NEOPROTEROZOIC, PASSIVE-MARGIN TURBIDITES

Kanin Kamen Supergroup >10 000 m

Apparently continuous succession of alternating sandstones, claystones and clayey shales with stromatolitic carbonate beds at the top. Decreasing degree of metamorphism towards the top.

7

Tabuyevskaya Group 3700 - >4000 m Gnilskaya Formation Interbedded terrigenous and algal-laminated carbonate beds Yaneyskaya Formation Quartzitic sandstone, pale gey, thickbedded with claystone-siltstone interbeds in upper part Bolvansky Formation Parallel-laminated claystone-siltstone Cross-bedded sandstones Graded sandstone and siltstone Quartzitic and feldspathic sandstones Tarkhanovskaya Group c. 5000 m Upper formation Quartzitic and feldspathic sandstones and silty and clayey beds Middle formation Silty and clayey beds, lenticular dolomites Lower formation Interbedded sandstones, mudstones and claystones Mikulkinskaya Group 1500m Carbonate skarnoid unit Thin-bedded sandstone-siltstone Clayey beds and lenticular sandy-silty beds Massive sandstones, in places cross-bedded

concretions increase towards the top of the section, where lenticular bedding is dominant. Metamorphosed carbonate beds and concretions (scarnoids) occur at the top of the Mikulkinskaya succession. The Tarkhanovskaya Group (c. 5000 m) is divided into three informal formations (Fig. 3). The lowest formation consists of medium-grained quartzitic sandstones and muddy-clayey beds (now schist or phyllite) interbedded in various proportions. In the middle formation, thin schistose, silty-clayey beds and lenses predominate. Lenticular, dolomitized limestone bodies (boudinage) occur and probably represent diagenetic features. The upper formation consists primarily of clayey-silty metasediment (now low-grade schist or phyllite), commonly with lenticular bedding. In the highest part, there are interbedded grey sandstones, siltstones and cleaved claystones showing graded bedding, and also pale-grey quartzitic and feldspathic sandstones in lenticular beds, in places cross-bedded and with small slump folds. The Tabuyevskaya Group (c. 4000 m thick) is believed to lie conformably on the Tarkhanovskaya Group, though the contact is only locally exposed; and in many parts of the area it has been described as tectonic (Olovyanishnikov 1998, p. 93). More recently, however, the same author has indicated an unconformity between these two groups (Olovyanishnikov 2001, Fig. 5). Exposure of the Tabuyevskaya Group is generally poor and incomplete, except in incised river sections. The group is divided into three formations: from bottom to top, the Bolvansky, Yaneyskaya and Gnilskaya Formations (Figs 2 & 3). The Bolvansky Formation consists, in its lower part, of fine- to medium-grained quartzitic and feldspathic sandstones with siltyclayey interbeds. Upwards, c. 300-350 m from the base, there is

Fig. 3. Lithostratigraphy of the Neoproterozoic succession of Kanin Peninsula; based on the subdivisions of Olovyanishnikov (1998, 2000) but modified to accord with international stratigraphic nomenclature.

a transition into a 600 m thick interval of graded sandstonesiltstone beds (Fig. 4a) with clearly erosional bottom surfaces. In addition to the upward-fining depositional sequences (14-40 cm), a 15-25 m scale cyclicity has been observed. This interval is overlain by a unit with small-scale cross-bedding, including herring-bone type, and finally by a parallel-laminated, clayey-silty succession. On top, there are c. 2000 m of thinbedded claystone-siltstone (now mostly phyllite) with sporadic, thin limestone intercalations and a mafic tuff layer. The overlying Yaneyskaya Formation is a c. 200-400 m thick unit of pale-grey, pink or greenish-grey, medium- to thickbedded, quartzitic sandstones with intercalations of cleaved claystone-siltstone. In the upper parts of the formation there are thinner-bedded silty sandstones with laminated mudstones showing graded bedding and local channelling (Fig. 4b). The Gnilskaya Formation (900 m thick) is a variable succession of alternating, blue-grey, quartzitic sandstone and black, cleaved mudstone, as well as green-grey, thin-bedded, tuffogenic graded siltstone and claystone, with at least one horizon of basaltic lava (Olovyanishnikov 2000). In its uppermost parts there is a terrigenous-carbonate unit, locally with algal lamination and stromatolites. There are few, positive, diagnostic indications in the Kanin succession of deposition in a submarine turbidite system by turbidity currents. There is only one c. 600 m interval consisting of turbidites with observed A-intervals, in the Bolvansky Formation. Other, sporadically recorded, sedimentary structures, e.g. crossbedding or minor slump folds, may occur in diverse environments of deposition and, on their own, are not diagnostic of any particular environment. The uppermost part of the Tabuyevskaya Group (Gnilskaya Formation) contains carbonate beds with algal structures, testifying to shallow-water deposition. In parts of the

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D. ROBERTS ETAL.

Fig. 4. (a) Interbanded, strongly cleaved, silty sandstone and dark grey mudstone, showing graded bedding; looking WNW. Bolvansky Formation; upper part of the Bolshoy Pidertselkha stream section, NW Kanin Peninsula. (Photo: D. Roberts.) (b) Thin-bedded, silty sandstone and intercalated laminated mudstone, with an erosional channel developed a few centimetres below the pencil. Upper part of the Yaneyskaya Formation, lower reaches of the Bolshoy Pidertselkha river. Photo taken looking down on a near-horizontal surface; north is approximately at the top of the picture. (Photo: D. Roberts.)

Mikulkinskaya Group there are fining-up, massive, sandstoneshale units on a scale of metres, but these are not laterally persistent. This is suggestive of accumulation in channels, that might have been developing in a submarine turbidite system, but another depositional environment cannot be excluded. As noted above, the unconformity between the Tarkhanovskaya and Tabuyevskaya formations is not well exposed. However, the Kanin Kamen Supergroup (c. 10 km thick) shows no record of subaerial exposure. This indicates that deposition occurred on a subsiding slope, where variations in subsidence rate and the supply of terrigenous material were the main factors influencing the water depth and mechanisms of deposition. Both bottom currents and fallout from suspension were operating, probably most of the time, while the generation of turbidity currents was largely subordinate. Sedimentation might thus have occurred in a basinal setting and on a slope, with an inferred STS to prodelta sedimentation and repeatedly occurring shallow-water environments, perhaps at a sandy delta front, with embayments in which sporadic carbonates accumulated. The succession, on the whole, is complex and exposed discontinuously, and because of these limitations we cannot propose a more detailed interpretation. We, therefore, tentatively conclude that the Kanin succession, is not strictly representative of a STS, although some parts of the succession might have accumulated in this particular environment.

Northern Timan From the coastal areas of northern Timan, Getsen (1975) and Olovyanishnikov (1998, 2000) described an entirely terrigenous, Barminskaya Group (c. 3600m thick), composed of the Yambozerskaya (800-900 m), Malochernoretskaya (2000m) and Rumyanichnaya (c. 300-700 m) Formations (Fig. 5). The basement to the succession is not exposed. The group is unconformably overlain by either low-grade, Silurian sedimentary rocks or basaltic rocks of Devonian age. The most extensive research on the Barminskaya Group has been carried out over the last forty years by Getsen and published in Russian with little, if any, English translation. His most comprehensive papers are listed in Olovyanishnikov (1998). Descriptions of sections provided by Olovyanishnikov (1998, 2000) refer to rhythms or cycles on a scale ranging from centimetres to tens of metres.

Fig. 5. Simplified geological map of the coastal district of northern Timan (modified from Olovyanishnikov 2000).

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NEOPROTEROZOIC, PASSIVE-MARGIN TURBIDITES

Barminskaya Group up to 3600 m ?Continuous terrigenous succession Degree of metamorphism decreasing upwards Upward-diminishing number of mafic dykes

Yambozerskaya Formation 800 - 900 m Shallow-water sandstones, siltstones and clayey slates Numerous fining-up and coarsening-up cycles Sporadic mafic dykes Malochernoretskaya Formation up to 2000 m Slates or phyllites with subordinate quartzitic sandstones. Numerous fining-up cycles and megacycles. Abundant mafic dykes Rumyanichnaya Formation 300 - 700 m Cleaved mudstone/phyllite interbedded with metasandstones. Fining-up cycles and megacycles Swarms of mafic dykes

The first-, second- and third-order rhythms in his flyschoid succession start with sandy intervals and terminate with muddy or clayey sediment. We summarize briefly the existing lithostratigraphical data published by Getsen (1975) and Olovyanishnikov (1998, 2000) in Fig. 6. Additional field observations by one of us (DR) and examination of abundant photographic material and thin-sections (DR & AS) suggest that all three formations in northern Timan consist of turbidites comprising mainly fine sand, silt and clay. The fine-grained sandstones and siltstones are moderately to well sorted and quartz dominated, with very few feldspar grains. Quartz overgrowths are the main component of the cement and sericite is subordinate. Diagenetic concentrations of ferruginous carbonate are common. Usually, there is a good grain-size separation between the silty/sandy and clayey laminae and, in places, clay-dominated laminae display a weakly pronounced graded bedding. Intervals A, A-B, A-C and A-D of typical Bouma sequences were recorded, with particularly well-developed rippled intervals in the coastal section south of Cape Rumyanichny (Rumyanichnaya Formation) (Fig. 5). Various current ripple forms are represented, ranging from starved to climbing ripples, testifying to variations in current velocity and sand and silt supply (Figs 7a, b). Most of the current ripples indicate palaeocurrents directed towards the NE. In addition to the diverse Bouma sequences, sporadic load casts (Fig. 7c), flame structures, minor slumps and debrites are present, all testifying to rapid deposition on a slope. In addition, in the middle reaches of the Chernaya River (?Malochernoretskaya Fm.), packages of thin-bedded to laminated, fine-grained sandstone and clay stone show evidence of widespread penecontemporaneous slump-folding or slump scars with rotated blocks on the fault surfaces. The minor 'rhythms' probably represent single turbidites, whereas the larger 'cycles' represent fining-up packages or motifs, which we believe may reflect tectonic activity along the marginal fault(s) of the basin. Thus, with the exception of the striking abundance of rippled intervals in the assumed oldest, Rumyanichnaya Formation, there are no major differences in character of the three formations. A transitional stratigraphic contact between the Rumyanichnaya and Malochernoretskaya Formations has been reported from a part of the Chernaya river (Olovyanishnikov 2001). The contact between the Malochernoretskaya and Yambozerskaya Formations is not exposed; thus although the formations are arranged by Getsen

Fig. 6. Lithostratigraphy of the Neoproterozoic succession of northern Timan (based on Getsen 1975; Olovyanishnikov 2000).

(1975) in a stratigraphic order (Fig. 6), the precise stratigraphic relationship of the two higher formations is uncertain. Assuming that the three formations constitute one continuous succession, we propose that they represent a thick accumulation of fine-grained turbidites. The presence of A-intervals and finegrained sand, along with various penecontemporaneous deformational structures, suggests that there was a lack of supply of coarser-grained material and not that the turbidites were particularly distal. However, this is a tentative interpretation based on just one summarized section. Previous descriptions show that the turbidites are arranged in units with fining-up motifs on various scales ('transgressive cycles' of Olovyanishnikov 2000, 2001), reflecting either tectonic and/or eustatic events. The Barminskaya Group turbidites probably accumulated on a gently sloping basin margin. The presence of dolerite dykes and sheetlike gabbroic rocks suggests that the extensional regime also extended into the period immediately following sedimentation, and supports the idea that tectonic events were primarily responsible for the cyclicity along this extensional margin.

Central Timan: Vymskaya Ridge The Dimtemyol and Pokju rivers cut through the Vymskaya Ridge and expose the upper part of the Bystrynskaya Group and the bulk of the Vymskaya Group (Figs 1, 8 & 9). Both sections were examined in 1995 during the expedition organized by Vsevolod Olovyanishnikov (Olovyanishnikov 1995; Siedlecka & Roberts 1995). Olovyanishnikov's description of these sections was published in 1998. The upper portion of the Paunskaya Formation of the Bystrynskaya Group (c. 200 m thick), exposed only in the section of the Pokju river (Fig. 8), consists of massive siltstones, some with parallel lamination in the upper parts of the beds. There is an upward-coarsening motif in the uppermost part and a transition into the 3500 m thick Pokjuskaya Formation of the Vymskaya Group. Its lower, more than 300 m, interval consists predominantly of sandstones which exhibit several features diagnostic of turbidites. The sandstones vary from thin- to thickbedded and occur in c. 2-4 m thick (in places 10m thick) packages separated by 1 -2 m thick shaly units. The sandstones ar mostly fine-grained and massive, less commonly graded-bedded,

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D. ROBERTS ET AL

Fig. 7. (a) Fine sandy and muddy turbidites, mostly with just laminated and rippled intervals developed. A variability in sand supply is indicated by starved ripples and beds composed of climbing ripples. An east-dipping spaced cleavage is well developed here. This particular photo is of the surface of a dislodged, wave-washed block. Rumyanichnaya Formation, c. 750 m south of Cape Bolshoy Rumyanichny, northern Timan. (Photo: D. Roberts.) (b) A c. 30 cm thick, sandy, rippled bed between sandy-muddy, thin-laminated intervals and subordinate horizons of starved ripples; looking NNW (the palaeocurrent transport direction here is towards 050-055°). Rumyanichnaya Formation, c. 750 south of Cape Bolshoy Rumyanichny, northern Timan. (Photo: D. Roberts.) (c) Contact between a massive, sandy interval, probably a channel fill, and a muddy-sandy package beneath, with faintly laminated and rippled intervals. Note load casts, resulting from soft-sediment deformation. Rumyanichnaya Formation, c. 1 km south of Cape Bolshoy Rumyanichny, northern Timan; looking NNE. (Photo: D. Roberts.) (d) Medium-thick, graded sandstone beds interbedded with packages of massive or indistinctly graded and laminated, thin, sandy-muddy beds; looking NW. Lower part of the Pokjuskaya Formation, Vymskaya Ridge, Pokju river section, Central Timan. (Photo: A. Siedlecka.)

with laminated and rippled intervals at the top (A-B, B, A-C intervals of the Bouma sequence, Figs 7d & 9). The sandstones exhibit erosional bottom surfaces and some beds contain mud chips derived from subjacent beds. Cross-bedding was observed in one place in a thin lenticular bed of sandstone. Soft-sediment deformation and clastic dykelets were also recorded. The middle part of the Pokjuskaya Formation, more than 1000 m thick, consists predominantly of grey-black claystone, that typically exhibits an extremely fine parallel lamination, and of ripple cross-stratified siltstone. Subordinate thin beds of sandstone increase in thickness and amount upwards in the section and exhibit graded bedding. Only the lower part of the upper

Pokjuskaya Formation (c. 400 m thick) is exposed; and in its development it resembles the lower part of the formation. It is characterized by interbedded grey-black claystone and grey siltstone-sandstone, and exhibits an increase in grain size of the clastic material and in the thickness of the individual beds upwards in the section, resulting in a clear, coarsening-up motif. Sedimentary structures include parallel lamination with subordinate ripple-cross stratification on a centimetre scale in the clay-dominated beds, that are mostly just a few centimetres thick. The silty-sandy beds are c. 15-30 cm thick and are mostly lenticularly bedded throughout, and in some beds ripplecross stratification is preserved. There are also a few massive beds.

NEOPROTEROZOIC, PASSIVE-MARGIN TURBIDITES

11

Fig. 8. Simplified geological map of Vymsky Ridge, Central Timan (modified from Olovyanishnikov 1995).

Stratigraphically above a c. 300 m break in exposure in the Pokju river section, that may conceal a possible tectonic contact, the lower part of the Lunvozhskaya Formation is discontinuously exposed. The remainder of this 3500 m thick unit crops out in the section of the Dimtemyol river (Fig. 8). The lower part of the formation consists of dark grey, massive or parallel-laminated mudstone to very fine-grained sandstone. The middle part of the Lunvozhskaya Formation is a monotonous unit of greenish-grey, parallel- to lenticular-laminated mudstone with some greyblack, clayey shale beds. Upwards, there are subordinate sandstone beds with a clear grain-size separation from the predominant, laminated mudstone. This tendency continues into the upper part of the formation. The uppermost part is c. 200 m thick and composed of black, muddy shale with very few, thin lenticular or parallel-laminated sandstones. With the exception of this uppermost unit, in the Lunvozhskaya Formation there are two 1000m thick, slightly coarsening-up motifs forming the lower and middle parts of the succession. The base of the Kikvozhkaya Formation is represented by a c. 10m thick, fine- to coarse-grained, white quartzitic to feldspathic sandstone, that contains dark shaly clasts and dark thin shaly intervals. This bed contrasts markedly in composition, grain size and colour with the substratum, and the contact is sharp. We have not observed any erosional relief along this interface. This quartzitic sandstone unit is clearly folded and, although the exposure is poor, it seems that the same bed is tectonically repeated. Olovyanishnikov (1998) interpreted this bed as resulting from a regressive and erosive event, and considered it to lie unconformably upon the Lunvozhskaya Formation (Fig. 9). Alternatively, the white quartzite could represent a storm or channel deposit in a continuous succession. The poor exposure in this area makes the interpretation uncertain. There is a rapid upward transition into a grey-black, finegrained interval of c. 700 m exposed thickness. Carbonaceous, clayey beds containing > 1% of organic matter predominate. The thin, lenticular beds of sandstone with ripple cross-stratification to

flaser-bedding gradually increase in number upwards, and eventually predominate in the c. 150 m thick upper part of the section. The sandstones are fine-grained, or medium- to finegrained in the graded intervals, and there is a gradual transition from sand to silt. Thin 'sandy' beds are, in fact, mostly siltstones. Sandstones and siltstones are quartz-dominated with a few recognizable plagioclase grains in some beds. Quartz overgrowths constitute the predominant cement, with sericite forming a subordinate component. Sandstones are moderately to well sorted. Some grains exhibit clastic outlines showing that they were subrounded. Diagenetic concentrations of ferruginous carbonate several millimetres in size are common. The clearest indications favouring deposition of the Vymskaya Group in a STS are present in the lower part of the succession, including the exposed upper part of the Bystrynskaya Group. This part appears to be deposited from low-density turbidity currents partly reworked by bottom currents. Higher up, there was deposition from suspension with uncommon bottom currents and a deficiency of silt and sand. In the upper exposed part of the Pokjuskaya Formation there is evidence of an increasing activity of turbidity currents that produced graded, rippled and laminated beds, but not in any clearly defined parts of the Bouma sequence. The Kikvozhskaya Formation shows a distinct upward-coarsening motif, from an extremely low rate of sedimentation from suspension to one characterized by an increasing influence of bottom currents and finally some turbidity currents. The origin of the c. 10 m thick, well-sorted, fine- to coarsegrained sandstone is uncertain. If it represents a new period of sedimentation separated from its substratum by an unconformity, then a rapid deepening has to be assumed, returning to a regime similar to that in which the subjacent beds accumulated. If, on the other hand, it represents a storm or channel deposit, and there is no unconformity, then the upper part of the Bystrynskaya Group and the bulk of the Vymskaya Group would represent a continuous marine succession that accumulated on a gentle slope by turbidity or bottom currents, and probably representing

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D. ROBERTS ETAL

Vymskaya Group c. 7500 m Terrigenous succession of sandstones, mudstones and claystones.

Bystrynskaya Group, Top part, c. 200 m

Kikvozhskaya Formation c. 700 m Black-grey claystone, lenticular thin-bedded sandstone with parallel lamination and ripple-cross stratification. Pale grey quartzitic sandstone Unco nfo rmity Lunvozhskaya Formation c. 3500 m Black-grey mudstone and fine sandstone with parallel and lenticular lamination and ripple-cross stratification. Massive and graded-bedded sandstones are subordinate. Coarse sand and fine gravel/gritstone ?Un c o nfo rmity Pokjuskaya Formation c. 3500 m Lower and upper parts dominated by thick sandstones, massive or A-B, B, A-C. Clastic dykelets and soft-sediment deformation structures are common

Paunskaya Formation Siltstones, massive to parallel-laminated

first STS and finally prodelta conditions in a generally upwardshallowing depositional regime. Along the entire Timanian belt there are stromatolitic dolomites of Late Riphean age along the outer shelf margin, above the footwall of the Central Timan Fault (Olovyanishnikov et al. 2000). This fact also supports the notion of an overall shallowing of the marginal part of the basin, either continuously or in two episodes separated by an emergence and unconformity. A more detailed, definitive interpretation is not possible in view of the poverty of exposure and scarcity of reliable sedimentological criteria.

Comparison between the successions of Kanin Peninsula, northern Timan and the Vymskaya Ridge Even a brief review of the successions in the three separate areas reveals significant differences. What they have in common is their location NE of major faults defining the marginal part of the Timanian Basin; also, they are several kilometres thick and show no signs of subaerial exposure, although unconformities may be present. None of the three successions contains coarsegrained sediments and there are no clear indications of major channelling. Their substratum is nowhere exposed. Each succession has been affected by the Timanian deformation and lowgrade metamorphism, then uplifted, eroded and overlain above an angular unconformity by Lower Palaeozoic sedimentary rocks. The rock successions are Neoproterozoic in age, covering a time range of more than 400 million years. The presence of Vendian strata is open to discussion, as there are no positively diagnostic microfossils. Olovyanishnikov (1998, p. 119) assumed the possible presence of Lower Vendian beds in Central Timan and on Kanin Peninsula. We failed to obtain microfossils from the Vymskaya Group (Vidal 1996). The stromatolites of the Kanin

Fig. 9. Lithostratigraphy of the Neoproterozoic succession of central Timan, Vymskaya Ridge area (based on Olovyanishnikov 1995; Siedlecka & Roberts 1995).

Peninsula (Ludovatovskaya Formation) are of Late Riphean age (Raaben et al. 1995). The thin limestones and dolostones of the Gnilskaya Formation are interpreted as terminating the thick, terrigenous basinal succession (Fig. 3) (Olovyanishnikov 2000, 2001); and they are considered to be approximate stratigraphic equivalents of the Ludovatovskaya Formation. Numerous K-Ar and Rb-Sr isotopic dates have been reported in the Russian literature going back to the 1950s (see e.g. Olovyanishnikov et al. 2000; Gee et al. 2000), but these are not wholly reliable in terms of depositional ages. The sparse information available on the ages of the successions thus means that time correlation is not possible. Those referred to above illustrate that the ages of these successions are positively Neoproterozoic, with the exception of the Upper Vendian, and we believe that the presence of Lower Vendian strata is doubtful. This leaves an extensive Late Riphean time interval during which the discussed successions accumulated, though not necessarily contemporaneously. Olovyanishnikov (1995, 1998) proposed lithostratigraphic correlations based on lithological grounds and taking into account the presence of unconformities. In this contribution, comparisons are restricted to the overall sedimentological interpretation, combined with regional geological aspects, an approach that assumes that the events controlling the accumulation of sediments may or may not have been contemporaneous. The successions exposed on Kanin Peninsula and in northern Timan occur immediately NE of the East Timan Fault and are only about 50 km apart (Fig. 1). The Vymskaya Ridge is adjacent to the Central Timan Fault and the distance between central and northern Timan is roughly 500 km. There are considerable differences embracing interpretation of sedimentary environments, thicknesses and metamorphism even between the two successions in the north, in northern Timan and on Kanin Peninsula. The succession of northern Timan, according to our interpretation, represents the deposits of a STS. In contrast, only a minor part

NEOPROTEROZOIC, PASSIVE-MARGIN TURBIDITES

of the Kanin Kamen Supergroup has probably been deposited in a submarine turbidite system, most of it representing prodelta and shallow-water deposits, terminating with the uppermost stromatolitic dolostones. The origin of the 'scarnoid' carbonates is not clear, but they may represent metamorphosed equivalents of elongate carbonate concretions, or perhaps carbonate-bearing beds (Olovyanishnikov 1998, Fig. 44). Such concretionary beds and lenses are common in, for example, the Rumyanichnaya Formation of nearby northern Timan. The succession of the Vymskaya Ridge represents deposits of a STS, grading up into a prodelta environment. The fine-grained character of most of the sediments may have formed on a gentle and stable slope or perhaps a distal portion of the depositional system. Thus, from a sedimentological point of view, the three successions show more differences than similarities. The different facies of a large system reflect differences in the characteristics of the slope of the Timanian Basin margin, and also the likely presence of faults trending perpendicular to the margin (e.g. Roberts & Siedlecka 2002); in addition, although they are Neoproterozoic in age, they are not necessarily contemporaneous.

Comparisons with the basinal successions of the Rybachi and Varanger Peninsulas The basinal Barents Sea Group of northeastern Varanger Peninsula is a shallowing-up succession of c. 9000 m thickness. The substratum of the succession is not exposed. Its lower and middle portions (c. 6000 m thick), consist of a STS and prodelta accumulations (Siedlecka 1972; Siedlecka & Edwards 1980; Pickering 1981, 1982, 1983, 1985; Siedlecka et al 1989; Drinkwater et al 1996). The STS of the oldest Kongsfjord Formation was described by Siedlecka (1972) as flysch and later interpreted by Pickering (1981) as a submarine fan with inter-, middle-, and outer-fan facies associations that developed along a passive continental margin. We refer to the Kongsfjord turbidites as the Kongsfjord Turbidite System (KTS) (Fig. lOa). The oldest part of this system is dominated by coarse-graded sandstones and also comprises channelled conglomerates, including debrisflow deposits (Fig. lOb). Also present, however, are packages of mud-dominated sediment with Bouma C-D intervals. Higher up, the succession is dominated by fine-grained, thin-bedded packages of sediment with A-C, B-C or only C intervals and, in places, there is an abundance of diverse, erosional (Fig. lOc) and softsediment deformation structures. The sandstones are texturally immature, quartz-dominated greywackes, derived from a relatively stable source area (Siedlecka 1972; Pickering 1981). The KTS grades upwards into the mud-dominated lower Basnaering Formation, interpreted as an upper slope-prodelta deposit (Siedlecka & Edwards 1980; Pickering 1982; Siedlecka et al 1989). Palaeocurrents are consistently towards the NE and ENE. The Rybachi Turbidite System (RTS) on the Rybachi Peninsula, only 60 km from the KTS, is about 4000 m thick, including the uppermost 200 m that are interpreted as prodelta accumulations (Siedlecka 1985; Siedlecka et al 19950). The RTS is considerably coarser grained than the KTS, its lowermost exposed portion being an olistostrome. This is overlain by thick, very coarse-grained, sandstone turbidites interbedded with conglomerates and breccias, mainly matrix-supported (Fig. 11 a). Boulders and cobbles of granite and gneissic, sedimentary and volcanic rocks are identified as having been derived from the older Precambrian of the Kola Peninsula. There are abundant soft-sediment deformation structures and synsedimentary folds (Fig. 1 Ib), which, along with th olistostrome, testify to the presence of a fairly steep slope. The sediment gradually fines upwards, though still dominated by sand, and there are sandy packages with A-B intervals in every bed, interbedded with mud-dominated packages with A-C or B-C intervals. The bottom surfaces of beds are erosional. The

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uppermost c. 200 m consist mainly of thin, parallel-laminated beds characterized by slump scars and abundant synsedimentary faults (Siedlecka et al 19950). The sandstones are immature and vary from feldspathic greywackes (coarse sands) to quartz-dominated greywackes (finer sands). Palaeocurrents are consistently towards the NE and ENE (Siedlecka et al 19950; Drinkwater et al 1996). A comparison between the KTS and the RTS shows both similarities and differences. The similarities include thickness, finingand shallowing-up trends, palaeocurrent directions and location immediately NE of a major fault zone, the TrollfjordenKomagelva Fault Zone (TKFZ) on Varanger, and the SredniRybachi Fault Zone (SRFZ) on the Sredni-Rybachi Peninsula (Fig. 1). The ages of these successions are not constrained satisfactorily. Late Riphean microfossils were described from a formation resting conformably above the KTS (Vidal & Siedlecka 1983). In contrast, Lyubtsov et al (1989) had argued, on biostratigraphic grounds, that the uppermost RTS may possibly be of Vendian age, but Samuelsson (1997) maintained that there was no evidence for the presence of Vendian strata. Collectively, the various data suggest that both turbidite systems are Late Riphean in age. The main difference between these two turbidite systems lies in the coarseness of the sediments. This may have been caused by the steepness of the submarine slope and the topography of the hinterland (Siedlecka et al 19950; Drinkwater et al 1996). In addition, the two successions may not be exactly time equivalent and, therefore, may represent two diachronous phases of the turbidite systems developing along the faulted margin of the Timanian Basin in these areas. A comparison between the Varanger and Rybachi turbidite systems and the successions of the Kanin Peninsula and northern and central Timan shows considerable differences, even greater than those between the Kanin and Timan deep-water systems. This is not unexpected, however, taking into account the reasons considered above. The distance between Varanger-Rybachi and the Kanin Peninsula is more than 500 km, and between the Kanin Peninsula and central Timan there is an additional 500-600 km. The topography of the faulted margin might have ranged from an escarpment to a gentle slope, and the character and magnitude of the fault could have been different from one area to another both in time and in space. Some geophysical and other evidence suggests the presence of transverse, NE-SWtrending faults (Bogatsky et al 1996), now largely concealed beneath the Palaeozoic cover. This implies that the elongate basin and its bordering faults may have been segmented (Roberts & Siedlecka 2002) and resulted in the development of sub-basins, which, in turn, would help to explain the lateral facies differences in the deep-water systems. Palaeocurrents in the KTS and RTS show an easterly swing from the general northeasterly direction seen across the faulted basinal margin. This divergence towards a more longitudinal transport was previously explained by a 'failed rift' or aulacogen model related to the opening of the palaeo-Uralian ocean (Provodnikov 1970; Siedlecka 1975; Drinkwater et al 1996). There is now mounting evidence, from several studies of drillcores penetrating the Pechora Basin, of the importance of both primitive and later, evolved arc development in the oceanic realm to the NE (Getsen 1987, 1991; Bogatsky et al 1996; Olovyanishnikov et al 1996; Gee et al 2000; Dovzhikova et al 2004). The Timan Range thus represents the marginal part of a major oceanic basin that developed adjacent to northeastern Baltica in latest Mesoproterozoic to Neoproterozoic time. These deposits currently extend about 300 km to the NE of the Central Timan Fault Zone and are preserved beneath the Palaeozoic cover of the Pechora Basin (Olovyanishnikov 1998). During the early stages of this time interval, a passive margin prevailed along northeastern Baltica, but this was subsequently converted to an active, compressional, and locally transpressional margin in latest Neoproterozoic time (Roberts & Siedlecka 2002; Roberts & Olovyanishnikov 2004).

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D. ROBERTS ETAL.

Fig. 10. (a) Sandy, fine-grained turbidites. Kongsfjord Formation, Veineset in Kongsfjord, Varanger Peninsula. (Photo: A. Siedlecka.) (b) Thick bed of matrix-rich debnte. Kongsfjord Formation, coastal section NE of Kongsfjord, Varanger Peninsula. (Photo: A. Siedlecka.) (c) Exceptionally large groove-casts on the bottom surface of a sandstone bed (upper half of photo), probably representing a channel fill. Kongsfjord Formation, Veineset in Kongsfjord, Varanger Peninsula (Photo: A. Siedlecka.)

The large-scale regional factors that could have influenced, simultaneously, the development of the basinal depositional systems from Varanger to Central Timan are eustatic sea-level changes, tectonic subsidence following rifting, and climate. The first two factors could have produced both relief and lowering of the erosional base — both factors triggering an input of terrigenous material, perhaps bypassing the shelf areas and being redeposited downslope from the faulted basin margin. Intensive weathering and erosion would have been promoted by a wet and warm climate in adjacent land areas that were free of vegetation in the Neoproterozoic era. Palaeoclimatic studies suggest that, in Late

Riphean time, Baltica was close to the palaeo-equator and the climate was warm (Torsvik et al 1995). Traces of evaporites are present in both pericratonic and basinal realms (e.g. Siedlecka & Roberts 1992; Siedlecka et al. 19956, 1998). Factors that influenced the variability of facies and facies associations along and across the faulted margin were: (1) slope gradient and amplitude; (2) topography and lithological composition of the hinterland that provided terrigenous debris; and (3) topography and extent of the pericratonic, shelf areas. The great thicknesses, up to several kilometres, of the discussed systems suggest that they have a considerable lateral and

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Fig. 11. (a) Thick-bedded, graded, sandy turbidites with an overlying conglomerate. Perevalnaya Formation, eastern coastal section of Rybachi Peninsula. (Photo: A. Siedlecka.) (b) Synsedimentary fold in a package of fine-grained turbidites. Perevalnaya Formation, eastern coastal section of Rybachi Peninsula. (Photo: A. Siedlecka.)

downslope extent, perhaps over hundreds of kilometres. The KTS, for example, in its deformed state, can be observed across strike over a distance of 70-80 km, and the RTS over a distance of 50-60 km, which is probably close to its original extent since shortening by folding is only moderate. Both systems exhibit channelling and are coarse grained, which suggests, especially for the RTS, deposition on a high-gradient slope of approximately 10° (Stow et al. 1996). A narrow shelf, initially erosional, and a

mountainous hinterland have been suggested for this system, whereas a broader shelf and progradation of a large delta from a hilly hinterland formed the setting for the KTS (Siedlecka et al. 19950). There was one main faulting episode with subsequent infilling of the basin. A strongly contrasting picture emerges from an examination of the successions of the Kanin Peninsula and the Timans. The Kanin Kamen Supergroup reflects a moderately to gently inclined slope

16

D. ROBERTS ETAL

with several episodes of faulting (synsedimentary tectonics), and the presence of fine, sandy-muddy sediment provided from shelf areas. The succession of northern Timan reflects one episode of downfaulting, followed by subsidence that kept pace with sedimentation, and the fine sandy-silty sediment was transported by dilute turbidity currents and bottom currents across a probably moderately steep slope from a sandy-silty shelf. The fairly short distance involved between Kanin and northern Timan, combined with the great thicknesses and contrasting lithologies of the two lithostratigraphical successions, suggest that they may be of different age. Finally, the mud-dominated system of the Vymskaya Ridge reflects deposition on a low-gradient slope (possibly 1-2°) with little or no tectonic activity in the hinterland, recorded only by the assumed pre-Kikhvozskaya unconformity. Redeposition from a flat continental area and a broad muddy shelf, and perhaps from a large river (the pericratonic realm), may be envisaged for this slope-to-basin system. If the main Timanian Basin faulted margin was segmented by transverse faults, some stretches of the basin-bordering faults may have been more active than others (Roberts & Siedlecka 2002). Tectonic segmentation of the elongate basin into sub-basinal domains might also explain some of the differences in sedimentary facies in different parts of the TVB.

Conclusions 1. The deep-water depositional systems of the Kanin Peninsula and the Timans are similar in having developed NE of a major fault system bordering the Timanian Basin in the SW. They have considerable thicknesses and their basement is unknown. 2. The described and interpreted Neoproterozoic successions show considerable differences in their sedimentary facies development. 3. The close location between the Kanin Peninsula and northern Timan successions, combined with differing sedimentary facies development, suggest that they are not contemporaneous. They are, however, both believed to have accumulated along a comparatively stable basin margin. The Kanin part of the margin was repeatedly affected by faulting, whereas that of northern Timan was comparatively stable with less faulting, such that sedimentation kept pace with subsidence. 4. The Vymskaya Ridge succession represents a mud-rich, slopeto-basin system reflecting accumulation on a gentle and fairly stable slope. 5. Differences in sedimentary facies between the Kanin-Timan deep-water systems and those from the Varanger and Rybachi Peninsulas are striking, suggesting a steeper slope for the latter which, once created, was comparatively stable during sedimentation. We conclude that the deep-water systems exposed along the Timan-Varanger Belt show that there was a considerable variability in topography and development of the passive margin of the Timanian Basin during Mid to Late Riphean time. Fieldwork in the central Timans, for D. Roberts and A. Siedlecka, was made possible via collaboration between the Geological Survey of Norway and the Geological Institute of the Komi Branch of the Russian Academy of Sciences, Syktyvkar. Here, we are particularly indebted to the director of the Geological Institute, Professor N.Yushkin; and also to Dr A. Pystin for assistance, support and discussions during fieldwork in 1995. Fieldwork in 2000, for D. Roberts and V. Olovyanishnikov, in northern Timan and on Kanin, was a part of the EUROPROBE Timpebar' project, supported largely by the Swedish Polar Research Secretariat. We are grateful to the referees, Professor R. Ingersoll and Dr A. Maslov, for their constructive and helpful reviews of the manuscript; and to the guest editor, Dr V. Pease, for diverse pertinent comments and suggestions. Thanks also go to Irene Lundquist for drafting most of the figures.

References BOGATSKY, V. I., BOGDANOV, N. A., KOSTYUCHENKO, S. L., SENIN, B. V.,

SOBOLEV, S. F., SHIPILOV, E. V. & KHAIN, V. E. 1996. Tectonic map of the Barents Sea and the northern part of the European Russia: explanatory notes. Institute of the Lithosphere, Russian Academy of Sciences, Moscow, 101 pp. BOUMA, A. H. 1962. Sedimentology of some flysch deposits. Elsevier, Amsterdam, 168 pp. DOVZHIKOVA, E., PEASE, V. & REMIZOV, D. 2004. Neoproterozoic island arc magmatism beneath the Pechora Basin, NW Russia. Geologiska Foreningen i Stockholm Forhandlingar, 126, 353-362. DRINKWATER, N. J., PICKERING, K. T. & SIEDLECKA, A. 1996. Deepwater fault-controlled sedimentation, Arctic Norway and Russia: response to Late Proterozoic rifting and the opening of the lapetus Ocean. Journal of the Geological Society, London, 153, 427-436. GEE, D. G., BEYLAKOVA, L. T., PEASE, V., DOVSHIKOVA, E. & LARIONOV, A. 2000. Vendian intrusions in the basement beneath the Pechora Basin, northeastern Baltica. Polarforschung, 68, 161-170. GETSEN, V. G. 1975. [The basement structure of the northern Timan and Kanin Peninsula]. Nauka, Leningrad, 144 pp [in Russian]. GETSEN, V. G. 1987. [Tectonics of Timan.} Nauka, Leningrad, 170 pp [in Russian]. GETSEN, V. G. 1991. Geodynamic reconstruction of the northeastern European part of the USSR in the Proterozoic. Geotectonics, 25 (5), 391-400. LORENZ, H., PYSTIN, A. M., OLOVYANISHNIKOV, V. G. & GEE, D. G. 2004. Neoproterozoic high-grade metamorphism of the Kanin Peninsula, Timanide orogen, northern Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 59-68. LYUBTSOV, V. V., MIKHAILOVA, N. S. & PREDOVSKY, A. A., 1989. [Lithostratigraphy and microfossils of the Late Precambrian of the Kola Peninsula.] Kola Science Centre of the USSR Academy of Sciences, Apatity, 129 pp [in Russian]. OLOVYANISHNIKOV, V. G. 1995. Guide for study of stratotypic section of Vymskaya series of Upper Precambrian of Timan (unpublished field-guide). Institute of Geology, Russian Academy of Sciences, Syktyvkar, 20 pp. OLOVYANISHNIKOV, V. G. 1998. [Upper Precambrian of Timan and Kanin Peninsula.] Russian Academy of Sciences, Ekaterinburg, 163 pp [in Russian]. OLOVYANISHNIKOV, V. G. 2000. Neoproterozoic of the north Timan and Kanin Peninsula. Svedarctic Timan International Expedition 2000. Europrobe-Timpebar and Intas-Hale Projects, Syktyvkar (Unpublished guide-book). OLOVYANISHNIKOV, V. G. 2001. Neoproterozoic of the north Timan and Kanin Peninsula. (A report). Institute of the Komi Science Centre RAS, Syktyvkar, 45 pp. OLOVYANISHNIKOV, V. G., BUSHUEV, A. S. & DOKHSAN'YANTS, E. P. 1996. The structure of the conjugation zone of the Russian and Pechora plates from geological and geophysical data. Transactions of the Russian Academy of Sciences, Earth Science Section, 351, 1228-1232. OLOVYANISHNIKOV, V. G., ROBERTS, D. & SIEDLECKA, A. 2000. Tectonics and sedimentation of the Meso- to Neoproterozoic Timan-Varanger Belt along the northeastern margin of Baltica. Polarforschung, 68, 267-274. PICKERING, K. T. 1981. The Kongsfjord Formation—a Late Precambrian submarine fan in north-east Finnmark, North Norway. Norges geologiske unders0kelse, 367, 77-104. PICKERING, K. T. 1982. A Precambrian upper basin slope and prodelta in northeast Finnmark, North Norway—a possible ancient upper continental slope. Journal of Sedimentary Petrology, 523, 171-186. PICKERING, K. T. 1983. Transitional submarine fan deposits from the late Precambrian Kongsfjord Formation submarine fan, NE Finnmark, N. Norway. Sedimentology, 30, 181-199. PICKERING, K. T. 1985. Kongsfjord Turbidite System, Norway. In: BOUMA, A. H., NORMARK, W. R. & BARNES, N. E. (eds) Submarine fans and related turbidite systems. Springer, New York, 237-244.

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PROVODNIKOV, L. Y. 1970. Basement of the Timan-Pechora Region. Doklady an SSSR, 191, 40-43. RAABEN, M. E., LYUBTSOV, V. V. & PREDOVSKY, A. A. 1995. Correlation of stromatolite formations of northern Norway (Finnmark) and northwestern Russia (Kildin Island and Kanin Peninsula). Norges geologiske unders0kelse Special Publication, 7, 233-246. RICE, A. H. N. & ROBERTS, D. 1995. Very low-grade metamorphism of Upper Proterozoic rocks of the Sredni and Rybachi Peninsulas and Kildin Island, NW Kola Region, Russia. Norges geologiske unders0kelse Special Publication, 7, 259-270. ROBERTS, D. 1995. Principal features of the structural geology of the Rybachi and Sredni Peninsulas, Northwest Russia, and some comparisons with Varanger Peninsula, North Norway. Norges geologiske unders0kelse Special Publication, 7, 247-258. ROBERTS, D. & OLOVYANISHNIKOV, V. G. 2004. Structural and tectonic development of the Timanide orogen. In: Gee, D. G. & Pease, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 47-57. ROBERTS, D. & SIEDLECKA, A. 1999. Baikalian/Cadomian deformation and metamorphism along the northern margin of Baltica, northern Russia and northern Norway. Extended abstract volume, International meeting on Cadomian orogens, Badajoz, Spain, 223-228. ROBERTS, D. & SIEDLECKA, A. 2002. Timanian orogenic deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian-Cadomian connections. Tectonophysics,352, 169-184. SAMUELSSON, J. 1997. Biostratigraphy and palaeobiology of Early Neoproterozoic strata of the Kola Peninsula, Northwest Russia. Norsk Geologisk Tidsskrift, 77, 1-28. SIEDLECKA, A. 1972. Kongsfjord Formation—a Late Precambrian flysch sequence from the Varanger Peninsula, Finnmark. Norges geologiske unders0kelse, 278, 41-80. SIEDLECKA, A. 1975. Late Precambrian stratigraphy and structure of the north-eastern margin of the Fennoscandian Shield (East Finnmark-Timan Region). Norges geologiske unders0kelse, 316, 313-348. SIEDLECKA, A. 1985. Development of the Upper Proterozoic sedimentary basins of the Varanger Pennisula, East Finnmark, North Norway. Bulletin of the Geological Survey of Finland, 331, 175-185. SIEDLECKA, A. & EDWARDS, M.B. 1980. Lithostratigraphy and sedimentation of the Riphean Basnaering Formation, Varanger Peninsula, North Norway. Norges geologiske unders0kelse, 355, 27-44. SIEDLECKA, A. & ROBERTS, D. 1992. The bedrock geology of Varanger Peninsula, Finnmark, North Norway: an excursion guide. Norges geologiske unders0kelse Special Publications, 5, 45 pp.

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SIEDLECKA, A. & ROBERTS, D. 1995. Report from a visit to the Komi Branch of the Russian Academy of Sciences in Syktyvkar, Russia, and from fieldwork in the Central Timans, August 1995. Norges geologiske unders0kelse Report, 95.149, 24 pp. SIEDLECKA, A., PICKERING, K. T. & EDWARDS, M. B. 1989. Upper Proterozoic passive margin complex, Finnmark, North Norway. In: Whateley, M. K. G. & Pickering, K. T. (eds) Deltas: Sites and Traps for Fossil Fuels. Geological Society, London, Special Publications, 41, 205-219. SIEDLECKA, A., NEGRUTSA, V. Z. & PICKERING, K.T. 1995a. Upper Proterozoic turbidite system of the Rybachi Peninsula, northern Russia—a possible stratigraphic counterpart of the Kongsfjord submarine fan of the Varanger Peninsula, northern Norway. Norges geologiske unders0kelse Special Publication, 7, 201216. SIEDLECKA, A., LYUBTSOV, V. V. & NEGRUTSA, V. Z. I995b. Correlation between the Upper Proterozoic successions in the TanafjordenVarangerfjorden Region of Varanger Peninsula, northern Norway, and on Sredni Peninsula and Kildin Island in the northern coastal area of Kola Peninsula in Russia. Norges geologiske unders0kelse Special Publication, 7, 217-232. SIEDLECKA, A., ROBERTS, D. & OLSEN, L. 1998. Geologi pa Varangerhalv0ya: En oversikt med ekskursjonsforslag. Norges geologiske unders0kelse, Grdsteinen, 3, 122 pp. SIEDLECKA, A., ROBERTS, D., NYSTUEN, J. P. & OLOVYANISHNIKOV, V. G. 2004. Northeastern and northwestern margins of Baltica in Neoproterozoic time: evidence from the Timanian and Caledonian orogens. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 169-190. STOW, D. A. V., READING, H. G. & COLLINSON, J. D. 1996. Deep seas. In: READING, H. G. (ed.) Sedimentary environments and fades 3rd edn Blackwell, Oxford, 395-453. TORSVIK, T. H., ROBERTS, D. & SIEDLECKA, A. 1995. Palaeomagnetic data from sedimentary rocks and dolerite dykes, Kildin Island, Rybachi, Sredni and Varanger Peninsulas, NW Russia and NE Norway: a review. Norges geologiske unders0kelse Special Publication, 1, 315-326. VIDAL, G. 1996. Examination of samples of Neoproterozoic rocks from Central Timans for microfossil content. Norges geologiske unders0kelse Report, 96.115, 9 pp. VIDAL, G. & SIEDLECKA, A. 1983. Planktonic, acid-resistant microfossil from the Upper Proterozoic strata of the Barents Sea Region of Varanger Peninsula, East Finnmark, northern Norway. Norges geologiske unders0kelse, 382, 145-179.

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Riphean and Vendian sedimentary sequences of the Timanides and Uralides, the eastern periphery of the East European Craton ANDREY V. MASLOV 7, Pochtovy per., 620151, Ekaterinburg, Russia, Institute of Geology and Geochemistry, Urals Branch of RAS (e-mail: [email protected])

Abstract: The northeastern and eastern margin of the East European Craton exposes numerous Riphean and Vendian (Meso- and Neoproterozoic) sedimentary successions that were deposited in alluvial and shallow-marine environments in intra- and pericratonic basins. A review is presented of the lithostratigraphy, sedimentary environments and architectural style of these sedimentary sequences in the Southern, Middle, Subarctic and Polar Urals, on the Poludov Range, and in the Volga-Urals and Timan-Pechora regions. The Riphean sequences are subdivided into three major sedimentary units: Lower, Middle and Upper, based on type areas in the Bashkirian Anticlinorium. During the Early and Middle Riphean, in the Southern Urals there were several short episodes of 'diffuse' and linear rifting and long intervals of more stable development in intracratonic sedimentary basins. During the Late Riphean, in the territory under review, larger shallow marine basin developed. Two laterally connected zones existed along the eastern periphery of the East European Craton: one in the Southern and Middle Urals, with a predominance of shallow marine arkosic deposits, and the other, with moderately deep marine (continental slope and rise) turbidites in the Timan-Pechora region. Subsequent Vendian successions were largely shallow marine and deposited in epicratonic basins; they generally give way upwards in the Late Vendian into non-marine clastic formations, derived from the east.

There are numerous Riphean and Vendian (Meso- and Neoproterozoic) sedimentary successions along the eastern margin of the East European Craton that were deposited in diverse fluvial and shallowmarine environments in intra- and pericratonic basins. The most complete and best known of these sequences is located in the Bashkirian and Kvarkush-Kamennogorsk anticlinoria (Shatsky 1945; Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Semikhatov et al. 1991). The aim of this paper is to review the lithostratigraphy, sedimentary environments and architectural style of the Riphean and Vendian sedimentary sequences in their main areas of development: the Southern, Middle and Subarctic Urals, in the Poludov Range, and in the Volga-Urals and Timan-Pechora regions (Fig. 1). The concept of the Riphean (derived from the Roman name for the Ural mountains) was established by Shatsky (1945). He included four main sedimentary units in his Riphean Group: the Burzyan, Yurmatau, Karatau and Asha suites. The last of these was subsequently moved into the Vendian. Shatsky (1945) considered that the succession of the pre-Ordovician siliciclastic and carbonate rocks of the Southern Urals was similar to that found for the Hercynian and Alpine tectonic cycles. '... This succession was thought to have resulted from a single Baikalian tectonic cycle that corresponded stratigraphically to the Riphean era ...' (Semikhatov 1991, p. 10). Isotope age constraints on the Riphean and Vendian successions provide a time frame (Fig. 2) for this interval of Precambrian history from c. 1650 Ma to the base of the Cambrian (c. 545 Ma). Although the correlation of stratigraphic boundaries and isotope age data, as shown in Figure 2, is accepted by many authors (Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Ancygin etal. 1994), Semikhatov (2000) favours a Middle-Late Riphean boundary at 1030 Ma, the base of the Vendian at 600 Ma and a subdivision of the Vendian into Early and Late at 570 Ma. The Riphean and Vendian successions are described below, starting in the best preserved and documented type areas in the southernmost parts of the orogen. Shatsky's suites are upgraded here to groups and the Riphean and Vendian are treated as major stratigraphic units. Riphean and Vendian successions Southern Urals Of the Riphean sequences along the western front of the Urals, those of the Bashkirian Anticlinorium in the Southern Urals are

the most complete. They are composed of siliciclastic and carbonate deposits and divided into three parts: the Burzyan, Yurmatau and Karatau Groups (Fig. 3), these being the standard units of the Lower (1.650-1.350 Ma), Middle (1.350-1.000 Ma) and Upper Riphean (1.000-650 Ma) of northern Eurasia, within the borders of the former USSR (Resolution 1979; Keller 1979; Keller & Chumakov 1983; Sokolov 1990; Semikhatov 1991; Ancygin et al. 1994). Each of these groups is separated by a hiatus and, in places, also by an angular unconformity. The total thickness of the Riphean deposits in the Bashkirian Anticlinorium is 12-15 km. Vendian deposits of the Southern Urals are represented mainly by siliciclastic strata of the Asha Group (Keller & Chumakov 1983; Sokolov & Fedonkin 1990). In the Bashkirian Anticlinorium, the deposits of this group, with total thickness of the c. 2200-2500 m, occur in three zones and span the entire Vendian (650-545 Ma). Lower Riphean. The Lower Riphean (Burzyan Group) of the Bashkirian Anticlinorium combines three main subdivisions (in ascending stratigraphic order): the Ai, Satka and Bakal formations. The type sections of these formations are situated in the northeastern part of the Bashkirian Anticlinorium (see IIa on Fig. 1); in its eastern areas (Yamantau Anticline), the Bolshoi Inzer, the Suran and the Yusha formations are correlated with the Ai, Satka and Bakal formations (Keller & Chumakov 1983; Semikhatov 1989). The total thickness of the Lower Riphean deposits in the northeastern part of the Bashkirian Anticlinorium is c. 5.500-6.000 m. The Ai Formation is represented by conglomerates, sandstones, siltstones and black shales, with trachybasalts in the lower part. The formation can be subdivided into two parts: a lower member (thickness up to 2000-2500 m), that is dominantly coarse-grained and an upper member (up to 1000 m), represented mainly by dark fine-grained siliciclastic deposits. The lower member consists mainly of breccias and conglomerates, poorly sorted arkoses and other polymict sandstones and subordinate interbeds of siltstones and shales, largely of non-marine origin (Semikhatov 1989; Parnachev et al. 1990; Maslov 1993; Maslov et al. 1997); subordinate basalts and other volcanic rocks are also present. The upper member is composed of monotonous black shales (Corg contents up to 1-2%), siltstones and finegrained sandstones, with subordinate gravelites and smallpebbled conglomerates; they are thought to have been deposited

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 19-35. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 2. Comparison of the Riphean-Vendian (Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Ancygin et al. 1994) and Proterozoic subdivisions (Remane et al. 1996) of the Precambrian timescale.

Fig. 1. Sketch map of the eastern periphery of the East European Craton (from Becker 1996), showing the regions as they are treated in this paper. Ia, VerkhneKamsk depression of the Volga-Urals region; Ib, Kamsk-Belsk depression (aulacogen) of the Volga-Urals region (axial zone of the Permian Pre-Uralian foredeep); IIa, western, central and northeastern parts of the Bashkirian Anticlinorium; IIb, eastern part of the Bashkirian Anticlinorium; ina, western zone of the Kvarkush-Kamennogorsk Anticlinorium; inb, eastern zone of the Kvarkush Kamennogorsk Anticlinorium; IV, Poludov Range; V, Timan-Pechora region (Va, Mezen-Vychegodsk zone; V b , Obdyr-Nivshera subzone; V c , ChetlasDzhezhimparma subzone; Vd, Tsil'ma-Ropchino zone; Ve, Vymsk-Volsk zone; V f , Kanin-Pechora zone); VI, Izhma-Pechora depression; VII, Khoreiver depression.

in moderately deep basins, on an open shelf, in which stagnant environments prevailed (Maslov 1997). The Ai Formation is underlain by high grade metamorphic rocks of the Taratash Complex. Granitic and other gneisses of the latter have yielded U-Pb and U-Th-Pb zircon ages c. 1950 Ma (Keller & Krasnobaev 1983; Semikhatov 1989). The U-Pb isotopic ages of zircons from the youngest granites and diabase dykes cutting the gneisses of the Taratash Complex vary from 1610 to 1570 Ma (Tugarinov et al. 1970; Keller & Chumakov 1983). U-Pb zircon isotopic ages of 1635 ± 30 Ma were obtained from trachybasalts in the lower part of the Ai Formation (Semikhatov 1989). In the fine-grained siliciclastic rocks of the upper member of the Ai Formation, microfossils with a wide Riphean range have been recognized (Weiss et al. 1990). Correlatives of the Ai Formation are considered to occur in the central parts of the Bashkirian Anticlinorium (in the Yamantau

Anticline) where the Bolshoi Inzer Formation (thickness more than 2200m) is exposed. This unit is represented mostly by fine- and medium-grained quartz and feldspathic sandstones with subordinate interlayers of shales, siltstones, limestones and dolostones. In the lower and middle parts of the formation, moderately deep-water deposits predominate, whereas the formation's upper part consists of shallow-marine siliciclastic and carbonate sediments (Maslov 1993, 1997). The overlying Satka Formation (thickness from 1700-3500 m) consists mainly of massive or thin-bedded dolostones with minor stromatolitic dolostones, shales and fine-grained siltstones. It is subdivided into five members, the middle one of which (the Polovinka Member) comprises mainly fine-grained siliciclastic rocks (Garan 1969; Keller & Chumakov 1983; Semikhatov 1989). The Satka Formation was deposited mostly in near-shore and shallow marine environments (Maslov 1993, 1997). The formation is cut by rapakivi granites and syenites of the Berdyaush massif that have yielded an isotopic age (U-Pb in zircons) of 1348 ±16 Ma (Krasnobaev 1986). In the lower and upper parts of the Satka Formation, the following stromatolites have been described Paniscollenia satka Kom., Conophyton punctatus Kom., Crateri melodia Kom., Kussiella kussiensis Kryl., Gongulina differenciata Kom., Stratifera omachtella Semikh. and Conophyton garganicus Kom. (Krylov 1963; Keller & Chumakov 1983; Semikhatov 1991). There are also spheromorphites of very simple structure, such as Leiosphaeridia crassa (Naum.), L incrassata (Naum.), Nucellosphaeridium minutum Tim. and Protosphaeridium densum Tim. (Weiss et al. 1990; Semikhatov 1991). Chert nodules from the dolostones contain the silicified microfossils Eomycetopsis sibiriensis Lo, Gunflintia minuta Bargh., Gloeodiniopsis uralicus Kryl. et Serg., Oscillatoriopsis sp. and Eosynechococcus amadeus Knoll et Golub. etc. (Krylov & Sergeyev 1986; Semikhatov 1991). The upper unit of the Burzyan Group, the Bakal Formation (1200-1600 m) includes black shales, siltstones and fine-grained

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21

Fig. 3. Stratigraphic correlation chart for the Riphean and Vendian deposits in the Bashkirian Anticlinorium and the Volga-Urals region (from Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Maslov 1997; Maslov et al 1997; Maslov & Isherskaya 1998; Romanov & Isherskaya 1998, 1999, 2001).

sandstones, and also dolostones and limestones, containing the stromatolites Conophyton cylindricum Masl., C. lituum Masl., Gaia irkuskanica Kryl. and Jacutophyton sp. The Bakal Formation consists of two members. The lower one (500-650 m) comprises mainly thin-bedded or massive black shales deposited in stagnant basins of a shallow shelf. The upper member (thickness of 800-1000m) consists of siliciclastic and carbonate deposits, which were formed in near-shore and shallow-marine environments (Krupenin 1983, 1999; Maslov 1997). Black shales of the Bakal Formation contain the microfossils Leiosphaeridia crassa (Naum.), L. incrassata (Naum.), L. bicrura Jank., Protosphaeridium densum Tim., Leiominuscula minuta Naum., Germinosphaera todasii A. Weiss, etc. (Weiss et al. 1990). Limestones yielded an isotopic age (Pb-Pb method) of 1430 ± 3 0 Ma (Kuznetsov et al 2001). The terrigenous and carbonate deposits of the Bakal Formation are intruded by gabbro-diabase with an isotopic age (magmatic biotite, Rb-Sr method) of 1360 ± 35 Ma (Ellmies et al 2000). In central parts of the Bashkirian Anticlinorium, a Satka Formation correlative occurs in the Suran Formation. Its thickness ranges from 1000 m in the west to 2800 m in the eastern parts of the Yamantau Anticline. The formation is divided into five members, the first and the fifth of which are dominated by limestones and dolomites. The second, third and fourth

members consist mainly of shales and siltstones. In dolomites of the upper, Lapyshta Member, the stromatolites Kussiella kussiensis Kryl. and Chimaera metabole Vlass. were described by Radchenko & Fedonkin (1974) and provide a basis for correlation with the Satka Formation. The siliciclastic and carbonate deposits of the Suran Formation accumulated in shallow and moderately deep-marine environments, which were periodically anoxic. The Yusha Formation (650-1000m) of the Yamantau Anticline, is represented exclusively by siliciclastic rocks: finegrained sandstones, siltstones and shales. These are mostly shallow-marine deposits (Maslov 1993) and are thought to be correlatives of the Bakal Formation. Middle Riphean. The Middle Riphean Yurmatau Group overlies the rocks of the Burzyan Group with angular unconformity. The Yurmatau Group has a total thickness of 5-6 km and consists of four formations, comprising volcanic, terrigenous and carbonate deposits (Keller & Chumakov 1983; Ancygin et al 1994) and referred to (from base upwards) as the Mashak, Zigalga, Zigazino-Komarovo and Avzyan formations. The Mashak Formation is developed only in the eastern regions of the Bashkirian Anticlinorium, where thicknesses range from 2000 m in the eastern parts of the Yamantau Anticline to

22

A. V. MASLOV

3500 m in its western parts. It is composed of basalts and rhyolites as well as siliciclastic rocks: fine- and medium-grained quartzitic sandstones, tuffaceous sandstones, siltstones, shales and conglomerates (Maslov et al. 1997). In the conglomerates of the lower levels of the Mashak Formation, rounded pebbles of sandstones and quartzitic sandstones of the underlying Lower Riphean Yusha Formation are found (Rotaru 1983). Mainly near-shore deposits characterize the lower and middle levels of the Mashak Formation, whereas the upper ones consist of the shallowmarine sediments (Maslov 1993). Rhyolites and liparite-dacites of the Mashak Formation have yielded a Rb-Sr whole rock isochron age of 1341 ± 41 Ma and a U-Pb zircons age of 1348 ± 30 Ma (Krasnobaev 1986; Semikhatov 1989). The overlying Zigalga Formation (thickness up to 550m) mainly comprises fine- and medium-grained quartzitic sandstones and siltstones. In the northeastern sections of Bashkirian Anticlinorium, interbeds and lenses of conglomerates are present in the lower part of the Zigalga Formation (Garan 1969; Keller & Chumakov 1983; Maslov & Krupenin 1991). In the central and eastern parts of the Bashkirian Anticlinorium, there are dark shales and fine-grained sandstones in the middle part of the formation (Keller & Chumakov 1983; Maslov & Krupenin 1991). The Zigalga Formation mainly comprises sediments deposited in near-shore and shallow-marine environments (Maslov 1993). The Zigazino-Komarovo Formation (1000-1300 m) consists of siliciclastic rocks: dark shales and siltstones with thin interbeds of sandstones, limestones and dolomites (Maslov 1991, 1993). It is subdivided into three members, differentiated mainly by their colour. The structual peculiarities of the rocks and their geochemistry (Maslov & Krupenin 1996) indicate that the sediments of the Zigazino-Komarovo Formation were formed in near-shore, often desiccating environments. The presence, locally, in the lower part of this stratigraphic level of nodular pyrite concretions, lenticular interbeds of diagenetic siderite and black shales with an organic carbon content up to 3% suggests that reducing environments were established in certain parts of the basin. Fine-grained siliciclastic rocks of the Zigazino-Komarovo Formation contain abundant simply-organized microfossils, namely Leiosphaeridia crassa (Naum.), L. bicrura Jank., L. incrassata (Naum.), Leiominuscula minuta Naum., Satka favosa Jank. and S. elongata Jank. (Weiss et al 1990; Semikhatov 1991). The Avzyan Formation (800-2000 m) is represented by siliciclastic and carbonate rocks and subdivided into six members (Kozlov et al. 1990; Ancygin et al. 1994). Of these, the first, third and fifth (from the bottom) consist of limestones and dolomites, often with the stromatolites Baicalia aborigena Schap., Svetliella avzianica Kom., Colleniella evoluta Schap., Conophyton metula Kir., Baicalia nova Kryl. et Schap., Strati/era flexurata Kom., and Cryptophyton convolutum Kom. (Krylov 1975; Keller & Chumakov 1983). Among the carbonate rocks of these members, there are thin interbeds of black and greenish-grey shales, siltstones and intraclastic carbonate breccias. The second and fourth members of the Avzyan Formation comprise mainly grey, greenish-grey and black shales, siltstones and, more seldom, fine-grained quartz sandstones. The uppermost Tulmen Member is composed of shales with interbeds of sandstones, siltstones and dolostones. From the greenish-grey and dark-grey shales of the Avzyan Formation, microfossils similar to those in the Lower Riphean and also several new, exclusively Middle Riphean species have been described, namely Leiosphaeridia crassa (Naum.), L. incrassata (Naum.), L. bicrura Jank., Nucellesphaeridium minutum Tim. Leiominuscula minuta Naum., Protosphaeridium ternatum (Tim.), Leiosphaeridia minutissima (Naum.), and L. jacutica (Tim.) (Weiss et al 1990; Semikhatov 1991). There are also abundant Eomycetopsis robusta Schopf, Eoentophysalis belcherensis Hoffmann, Polybessurus bipartites Fairch., Gloeodiniopsis lamellosa Schopf and large Leiosphaeridia in chert nodules from the first and fifth members (Sergeyev 1992). The Avzyan Formation combines a wide spectrum of

shallow-marine, near-shore and lagoonal siliciclastic and carbonate deposits (Maslov 1997). Glauconites from the Avzyan Formation have yielded an isotopic age (K-Ar method) of c. 1220 Ma (Keller & Chumakov 1983). The Middle Riphean deposits of the Bashkirian Anticlinorium are intruded by gabbro-diabase dykes with K-Ar whole rock isotopic ages ranging from 1090 ± 20 to 1080 ± 30 Ma (Garris 1977; Keller & Chumakov 1983). Upper Riphean. The Upper Riphean, Karatau Group, in the western part of the Bashkirian Anticlinorium comprises (from bottom to top) the Zilmerdak, Katav, Inzer, Minyar and Uk formations. In the southeastern limb of the anticlinorium, an additional uppermost siliclastic unit, the Krivaya Luka Formation, is also present (Keller & Chumakov 1983; Sokolov 1990; Ancygin et al 1994). The Zilmerdak Formation comprises, in its lower part (Biryan Member), red and pink, coarse and medium-grained arkosic sandstones, with subordinate beds of gravelites, siltstones and conglomerates. The maximum thickness of these deposits (c. 2500-3000 m) occurs in the westernmost part of the Bashkirian Anticlinorium (Olli 1948; Keller & Chumakov 1983; Maslov & Krupenin 1991). This succession overlies, with angular unconformity, the Middle Riphean Avzyan Formation (Keller & Chumakov 1983; Semikhatov 1991; Maslov et al 1997) and is composed of four units (from base upwards): the Biryan, Nugush, Lemeza and Bederysh members. Sandstones of the Biryan Member contain detrital zircons with isotopic ages (U-Pb multigrain method) as young as 1100 Ma (Keller & Chumakov 1983; Krasnobaev 1986). Higher in the Zilmerdak Formation there are green and greenish-grey sandstones and siltstones with shales. The Nugush Member (200-350 m) consists of grey, dark-grey and greenish-grey thin-bedded siltstones, shales and argillites. The Lemeza Member (150-300 m) is composed mostly of by light medium- and fine-grained quartz sandstones with interbeds of siltstones. The Bederysh Member (250-400 m) comprises sandstones, siltstones and argillites, with interbeds of limestones and dolomites in its middle part. Shales of the Bederysh Member contain a diverse and rich assemblage of microfossils: Leiosphaeridia crassa, L. incrassata, L. bicrura, L. jacutica, Protosphaeridium densum Tim., Leiominuscula minuta Naum., Nucellosphaeridium minutum Tim., N. nordium trichoides typicus Hermann, large and gigantic Chuaria, Brevitrichoides bashkiricus Jank., B. karatavicus Jank., Eomycetopsis psilata Maihy et Schukla, E. rugosa Schopf et Blacic, Palaeolyngbya minor Schopf, P. zilimica Jank. and other forms (Weiss et al 1990; Semikhatov 1991). Within the overlying middle levels of the Upper Riphean, carbonate rocks (mainly red limestones and marls) and sandstones with glauconite, siltstones and shales occur. The Katav Formation (thickness 200-300 m) contains mainly red and pink, thin-bedded clayey limestones and marls, with thin interbeds of red argillites and carbonate breccias in the lower part. In southeastern areas of the Bashkirian Anticlinorium at this stratigraphic level, grey and greenish-grey limestones predominate. Carbonate rocks of the Katav Formation contain the stromatolites Inzeria tjomusi Kryl., Jurusania cylindrica Kryl., /. nisvensis Raab., Malginella malgica Kom. and M. zipandica Kom. (Krylov 1963, 1975; Komar 1978; Keller & Chumakov 1983). The isotopic age (KAr method) of glauconite from the upper part of the formation is 970-938 Ma (Garris 1977; Keller & Chumakov 1983). The overlying Inzer Formation (100-1000m) is made up of siliciclastic and subordinate carbonate deposits. The latter contain characteristic Upper Riphean stromatolites: Gymnosolen ramsayi Steinm., Katavia karatavica Kryl., and Gonophyton garganicus Kor. (Komar 1978; Kozlov 1982; Keller & Chumakov 1983). In the westernmost parts of the Bashkirian Anticlinorium, there are two siliciclastic and two carbonate subdivisions in the Inzer Formation, the lower siliciclastic unit being known as the Todinzer Beds'. In the southeastern limb of the anticlinorium

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23

(near the Belaya river) there are only siliciclastic deposits in the Inzer Formation. In the shales of the Todinzer Beds' abundant, simple sphaeromorphites with new species (Leiofusidium dubium Jank., Pterospermella simica Tank, and some others) have been reported (Semikhatov 1991). Isotopic ages (K-Ar method) of glauconite, sampled from sandstones and siltstones in the upper part of the Inzer Formation, range between 800 and 900 Ma (Garris et al 1964; Garris 1977; Keller & Chumakov 1983). Glauconites from near the boundary between the Katav and Inzer formations have yielded an isotopic age (K-Ar method) of approximately 940 Ma (Keller & Chumakov 1983). Rb-Sr isochrons for Fe-illites of the Inzer Formation indicate an age of 820 ±15 Ma for early diagenetic processes (Gorokhov et al. 1995). Limestones in the Todinzer Beds' have yielded a Pb-Pb isotopic age of 836 ± 25 Ma (Ovchinnikova et al. 1998). The upper part of the Upper Riphean (the Minyar and Uk formations) includes mainly stromatolitic and microphytolitic dolostones and limestones with minor shales, siltstones and glauconitic sandstones. The Minyar Formation (500-800 m) is represented mainly by dolostones (in the upper part, extensively silicified), with subordinate limestones. Microfossils, typical of the Upper Riphean (Sergeyev 1992) occur in the cherts of the Minyar Formation. The dolostones of the Minyar Formation contain numerous stromatolites: Minjaria uralica Kryl., Gymnosolen ramsayi Steinm. and Katavia karatavica Kryl. (Raaben 1975; Komar 1978; Keller & Chumakov 1983). Shales from the middle part of the Minyar Formation in the vicinity of Yuruzan contain melanocyrriliums (Maslov et al. 1994), which are characteristic of Riphean deposits younger than 850 Ma, and the microfossils Leiosphaeridia crassa, L. incrassata, L. kulgunica, Protosphaeridium densum Tim. and Myxococcoides grandis Horodysku & Donaldson (Weiss et al. 1990). Silicified microfossils, similar to the forms present in the Uralian Middle and Lower Riphean, are known from the numerous chert nodules in the dolostones (Nyberg & Schopf 1984; Sergeyev & Krylov 1986; Semikhatov 1991). Glauconites from the lower part of the Minyar Formation have K-Ar isotopic ages in the range of 740 to 710 Ma (Garris 1977; Keller & Chumakov 1983). According to Ovchinnikova & Gorokhov (2000), the average weighted isotopic age (Pb-Pb method) of the Minyar dolostones is 780 ± 85 Ma. The Uk Formation (160-300 m) overlies, the Minyar Formation, with minor angular unconformity. It comprises glauconitic sandstones, siltstones, argillites, and limestones with the stromatolites Linella ukka Kryl. in the lower member and L. simica Kryl. in the upper one (Krylov 1975; Keller & Chumakov 1983; Semikhatov 1991). Fine-grained siliciclastic deposits of this level also yield abundant microfossils Eomycetopsis psilata Maithy et Schukla and Palaeolyngbya zilimica Jank. (Semikhatov 1991). Rb-Sr isochron ages of glauconite from the Uk Formation range from 687 ± 29 Ma (the maximum ages are up to 700713 Ma, Gorozhanin 1995) to 664 ± 11 Ma (Zaitseva et al. 2000). In the southeastern limb of the Bashkirian Anticlinorium (IIb on Fig. 1), above the Uk Formation, another Upper Riphean unit has been described, the Krivaya Luka Formation (thickness 400500 m). It combines predominantly quartzitic sandstones, shales (partly phyllitic) and siltstones with thin beds of limestones. Gabbro-diabases, intruding the Krivaya Luka Formation, have a Rb-Sr whole-rock isotopic age of c. 660 Ma (Gorozhanin 1995). The rocks of the Krivaya Luka Formation are overlain unconformably by the Lower Vendian Kurgashlya Formation (Chumakov 1998).

The Tolparovo Formation (thickness up to 600 m) consists mainly of massive coarse-grained, grey and yellow-grey, feldspathic sandstones, matrix-supported breccias and conglomerates with minor intercalations of argillites. This unit is overlain by the Suirovo Formation (~300m), which comprises diamictites (tillite-like conglomerates) with the boulders of dolostone, sandstone, granite and diabase, and also beds of siltstones, argillites and thin sandstone. According to Gorozhanin (1995), argillite clasts (fraction <0.1 |xm) from the tillite-like conglomerates have a model Rb-Sr isotopic age 638 + 70 Ma. The Kurgashlya Formation (total thickness 250 m) is represented by matrix-supported breccias and mottled shales with rare limestones. The lower part of the formation has been correlated with the Tolparovo Formation and the upper part with the Bakeevo Formation (Ancygin et al 1994). The Arsha Formation is correlated with the Lower Vendian deposits of the western limb of the Bashkirian Anticlinorium (Ancygin et al. 1994). It comprises shales of variable composition, siltstones, chlorite-epidote-albite schists (metamorphosed volcanic rocks), diamictites and thin carbonate beds. The total thickness of the Arsha Formation is 1000-1800 m. The Rb-Sr whole-rock isotopic age of the volcanic rocks is 677 + 31 Ma (Gorozhanin 1995). The relationship between underlying and overlying formations has not been clearly established.

Lower Vendian. Lower Vendian deposits are known only from isolated regions in the Bashkirian Anticlinorium: in the western limb of the anticlinorium in the vicinity of Tolparovo (the Tolparovo and Suirovo formations), in the southeastern part of the anticlinorium around Bainazarovo (the Kurgashlya Formation), and 35-40 km to the north of Beloretsk (the Arsha Formation).

In the Volga-Urals region (see Ia and Ib on Fig. 1) siliciclastic and carbonate deposits of the Lower, Middle and Upper Riphean and Upper Vendian have been described (Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Kozlov et al 1999; Romanov & Isherskaya 1999; Maslov & Isherskaya 1998; Lozin 1999; Resolution 1999).

Upper Vendian. In the western part of the Bashkirian Anticlinorium, the Upper Vendian deposits contain siliciclastic rocks of the Bakeevo, Uruk, Basa, Kuk-Karauk and Zigan formations (Becker 1988; Sokolov & Fedonkin 1990; Maslov et al 1996). The Bakeevo Formation (up to 140 m) consists of finegrained hematitic sandstones, glauconitic sandstones and siltstones and layers of hematite ores. The overlying Uruk Formation (200-300 m) includes arkose and other feldspathic cross-bedded sandstones, conglomerates and gritstones. The Basa Formation (thickness up to 1000 m) is represented mainly by subgreywackes and feldspathic sandstones and siltstones, interbedded with mottled siltstones and shales. Higher up in the section, the KukKarauk Formation comprises predominantly massive and crossbedded conglomerates with minor sandstones and siltstones. The thickness of these deposits is no more than 200-250 m. The highest unit, the Zigan Formation (thickness 500-600 m) is represented by thin intercalations of grey and greenish-grey sandstones, siltstones and shales. These Upper Vendian deposits of the Bashkirian Anticlinorium were formed mainly in fluvial, deltaic and shallow-marine environments. According to Becker (1968, 1988) these deposits are very similar to molasse. In the southeastern part of the Bashkirian Anticlinorium, the deposits of the Kuk-Karauk and Zigan formations are absent, either due to non-deposition or erosion, below the Lower Ordovician unconformity. Glauconites from sandstones of the Bakeevo Formation have yielded a Rb-Sr isochron age of 618 + 13 Ma (Gorozhanin 1995). Glauconites from the Uruk Formation provided minimum K-Ar ages ranging from 582 to 569 Ma. In the sandstones and siltstones of the Basa Formation, glauconites that have yielded K-Ar ages ranging from 600 to 558 Ma (Keller & Chumakov 1983; Raaben 1994; Gorozhanin 1995).

The Volga-Urals region

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A. V. MASLOV

Lower Riphean. The Lower Riphean is represented by the Kyrpin Group. It includes the Prikamsk, Kaltasa and Nadezhdino formations (Lisovski et al. 1981; Isherskaya & Romanov 1993; Romanov & Isherskaya 1994; K'ozlov et al 1999). The most complete sections of the Kyrpin Group have been described from numerous deep boreholes in the western part of the KamskBelsk Aulacogen in the vicinity of Arlan (see Ia on Fig. 1) (Romanov & Isherskaya 1998; Maslov & Isherskaya 1998). The Prikamsk Formation was deposited directly on the crystalline basement and is composed exclusively of siliciclastic rocks: medium- and coarse-grained grey and red sandstones, gravellites, siltstones and in its upper part, fine-grained strata, locally with a carbonate cement.The thickness of the Prikamsk Formation varies from 100 to 1800 m. The pebbly-gravelly and sandy deposits of the lower part of the Prikamsk Formation were formed in alluvial and alluvial-delta environments, whereas the sandy-siltstone sequences of the upper part accummulated in near-shore zones of the shelf basin (Maslov & Isherskaya 1998). Glauconite from sandstones in the Prikamsk Formation has yielded K-Ar ages ranging from 1430 to 1330 Ma (Gorozhanin 1983). Gabbro-diabase dykes cutting the Prikamsk Formation have yielded K-Ar isotopic ages of 1300-1316 Ma (Keller & Chumakov 1983). Isotopic ages from the less than 0.001 mm fraction from argillites of the Prikamsk Formation range from 1482 ±15 to 1408 ± 14 Ma (Rb-Sr method, Isherskaya & Romanov 1993). The overlying Kaltasa Formation (230-1700m) is dominated by various types of carbonate rocks, with siliciclastic and carbonate deposits in its middle part (the Arlan Member). The base of the formation is concordant and gradational with underlying strata of the Prikamsk Formation. The greatest thickness (3000-3500 m) of the Kaltasa Formation is found in the central part of the Kamsk-Belsk Aulacogen, whereas along its northern periphery it is only a few tens of metres thick (Frolovich 1988; Isherskaya & Romanov 1993). Investigations of the sedimentary structures characterizing the carbonate rocks of the Kaltasa Formation have shown that they were mostly deposited in shallow-marine and moderately deep-water environments (i.e. below storm wave base). In the Kaltasa dolostones, according to Semikhatov and Panova (in Romanov & Isherskaya 2001) the stromatolite Stratifera omachtella Kom. is present, a form that is characteristic of Lower Riphean deposits in other regions of northern Eurasia. K-Ar isotopic dating of authigenic glauconite from siltstones of the Arlan Member yielded an age of c. 1470-1490 Ma (Kazakov et al 1967; Keller & Chumakov 1983). The Nadezhdino Formation (180-780 m) consists of red and mottled fine-, medium- and coarse-grained sandstones, siltstones and argillites with interbeds of gravellites and small-pebble conglomerates; in its upper part, there are fine-grained, partly calcareous siliciclastic deposits. Accumulation of these sediments occurred mainly in littoral and sublittoral environments. The rocks of the Nadezhdino Formation are cut by dykes of gabbrodiabase, which have a K-Ar whole-rock isotopic age of c. 1370 Ma (Keller & Chumakov 1983). Authigenic glauconite from the sandstones of this stratigraphic level has given a K-Ar isotopic age of 1366 ± 6 Ma (Maslov & Isherskaya 1998). In the Kamsk-Belsk depression (aulacogen), the Lower Riphean successions are unconformably overlain by Palaeozoic strata, which transgress various underlying formations. Middle Riphean. The Middle Riphean deposits (the Serafimovo Group) in the Volga-Urals region were deposited transgressively upon either the Lower Riphean strata or on the crystalline basement, and are represented by the Tukaevo and Olkhovo formations (Isherskaya & Romanov 1993; Romanov & Isherskaya 1994; Maslov & Isherskaya 1998). The most complete sections of the Serafimovo Group and the overlying deposits of the Upper Riphean, Abdulino Group (Romanov & Isherskaya 1998), occur in the axial zone of the Pre-Uralian foredeep (see IIb on Fig. 1).

The Tukaevo Formation (thickness up to 630 m) consists mainly of mottled arkoses and sub-arkosic sandstones with subordinate interbeds of shales and siltstones. The Tukaevo Formation was deposited predominantly on the inner shelf, and subjected to the influence of bottom currents (Maslov & Isherskaya 1998). Authigenic glauconite from the Tukaevo sandstones has yielded K-Ar isotopic ages ranging from 1250 to 1336 Ma (Garris et al 1964; Isherskaya & Romanov 1993) and sericite from the same stratigraphic level has an isotope age of c. 1270 Ma (Zaides 1973), by the same method. The overlying Olkhovo Formation (340-840 m) is represented in its lower part (the Akberdino Horizon) by dark siltstones and shales (Andreev et al 1981; Keller & Chumakov 1983) and in its upper part by red, green and pink-grey argillites, marls, siltstones and dolostones. The deposits of the Akberdino Horizon were formed under littoral or sublittoral conditions, whereas higher levels of the Olkhovo Formation comprise mostly shallow-marine sediments (Maslov & Isherskaya 1998; Maslov 2000a). K-Ar whole rock isotopic ages of gabbro-diabase dykes, cutting the Olkhovo Formation range from 1338 to 1120 Ma (Keller & Chumakov 1983; Romanov & Isherskaya 1998). Upper Riphean. The Upper Riphean, Abdulino Group, of the Volga-Urals region (the Usa, Leonidovo, Prijutovo and Shikhan formations) overlies, with a stratigraphic break, the Middle and the Lower Riphean sedimentary sequences and, in a number of places, oversteps onto the crystalline basement. Its upper boundary is controlled almost everywhere by pre-Vendian erosion, as shown by its preservation in central and eastern areas of the Kamsk-Belsk Aulacogen, in the Pre-Uralian foredeep and in the northern flank of the Sernovodsk-Abdulino Aulacogen. The basal Usa Formation (thickness usually up to 400 m) consists mainly of mottled and grey feldspathic sandstones and arkoses with thin interbeds of siltstones and argillites. These siliciclastic deposits were formed in alluvial, alluvial-deltaic and near-shore environments, above normal wave base (Maslov & Isherskaya 1998). The overlying Leonidovo Formation (thickness at least 1300 m) is represented predominantly by fine and mediumgrained mottled quartz sandstones with a kaolinitic cement. Siliciclastic deposits of this stratigraphic level of the Abdulino Group were formed under shallow-marine and near-shore conditions. The Prijutovo Formation (thickness at least 670m) comprises mottled and red shales and argillites, siltstones, marls, dolostones and very subordinate sandstones. Accumulation of these sediments occurred in near-shore and shallow-marine environments. According to Weiss (Kozlov et al 1999), microfossils found in the fine-grained siliciclastic rocks of the Prijutovo Formation are similar to those known from the Bederysh Member of the Upper Riphean Zilmerdak Formation of the Bashkirian Antic linorium. K-Ar isotopic ages of glauconite from sandstones of the Prijutovo Formation range from 843 to 896 Ma (Keller & Chumakov 1983). The uppermost unit of the Abdulino Group, the Shikhan Formation (thickness at least 360 m), is composed of greenish-grey and reddish-grey clayey limestones and stromatolitic limestones, dolostones and marls. All these deposits were formed predominantly under shallow marine and near-shore conditions. Upper Vendian. Correlatives of the Lower Vendian deposits of the Bashkirian Anticlinorium are absent in the Volga-Urals region (Lisovski et al. 1981; Sokolov & Fedonkin 1990). In the southern and central parts of this territory, in the Kamsk-Belsk Aulacogen, the Upper Vendian sedimentary sequence includes the Baikibash, Staropetrovo, Salikhovo and Karlin formations (Sokolov & Fedonkin 1990; Maslov et al 1997) (see Fig. 3). The Baikibash Formation comprises mainly green-grey and pink, quartzfeldspar and polymictic sandstones, siltstones, gritstones and argillites. The thickness of the Baikibash Formation varies up to 70 m. The overlying Staropetrovo Formation (300-350 rn) is represented by green-grey and dark argillites and siltstones with thin

RIPHEAN AND VENDIAN: TIMANIDES AND URALIDES

25

layers of glauconite-bearing quartz sandstones. Glauconites from the Staropetrovo Formation have yielded K-Ar ages ranging from 595 to 545 Ma (Sokolov & Fedonkin 1990). The Salikhovo Formation (thickness varying from 30 to 320 m) includes red polymictic and quartz-feldspar sandstones. The uppermost unit, the Karlin Formation, comprises thin-bedded grey and brownishgrey argillites and siltstones which have a maximum thickness of c. 650 m in the easternmost part of the Volga-Urals region. In the northern part of the Volga-Urals region there is a slightly different Upper Vendian sedimentary sequence (the Las'va Group) in the Verchne-Kamsk depression. This includes the Kykva, Vereschagino, Velva and Krasnokamsk formations. The Kykva Formation (thickness up to 200 m) is represented by polymictic and quartz-feldspar conglomerates, sandstones, gritstones, siltstones and argillites. It overlies both the Riphean sedimentary rocks and the crystalline basement of the East European Craton with angular unconformity. The Vereschagino Formation includes grey, brown-grey and dark argillites with thin layers of sandstones, siltstones and three horizons of vitroclastic tuffs. The KAr whole rock isotopic age of the tuffs is c. 580 Ma (Sokolov & Fedonkin 1990). The maximum thickness of the Vereschagino Formation is about 300-380 m. The Velva Formation (300 m) comprises diverse greenish-grey and grey sandstone-siltstone and siltstone-argillite packets. The Krasnokamsk Formation (200-500 m) includes red and mottled polymictic and quartz feldspar sandstones, siltstones and argillites with desiccation cracks. All these Late Vendian siliciclastic successions of the Volga-Urals region were derived from the interior parts of the East-European craton (Aksenov 1967; Keller & Chumakov 1983; Sokolov & Fedonkin 1990) and deposited in shallow marine and flu vial-deltaic environments.

The Middle Urals Along the western slope of the Middle Urals, Riphean and Vendian sedimentary sequences are known in the core of Kvarkush-Kamennogorsk Anticlinorium. The oldest Riphean deposits (Fig. 4) crop out in the eastern zone (around Visimo-Utkinsk and Sinegorsk, see IIIb on Fig. 1), where they are represented by the Sinegorsk and Klyktan formations of the Upper Riphean Kedrovka Group (Ablizin et al 1982; Ancygin et al 1994; Maslov et al 1996). Upper Riphean. The Sinegorsk Formation has a total thickness of more than 2000 m (lower boundary not exposed); in its lower part it is composed of grey and yellowish-grey quartzites, and in its upper part, of shales and phyllites with thin interbeds of sandstone and gravellite. Accumulation of these siliciclastic deposits of the Sinegorsk Formation occurred in alluvial, near-shore and shallow-marine environments (Ablizin et al. 1982; Kurbatskaya 1985; Maslov et al. 1996). The overlying Klyktan Formation is subdivided into three members. The lower one (500-700 m) comprises shales (partly phyllitic) with interbeds of quartzitic sandstones near the base and limestones and marls higher up. The middle member is a 300-400 m thick sequence of limestones and dolostones containing in some areas (for example in the basin of the Koiva river) stromatolites, comparable to those from the Minyar-Uk level of the Bashkirian Anticlinorium (Ablizin et al. 1982). The upper member of the Klyktan Formation (230-250 m) is represented mainly by shales. The siliciclastic and carbonate deposits of this formation accumulated mainly in nearshore and shallow-marine environments (Kurbatskaya 1985; Maslov et al. 1996). In the western part of the Kvarkush-Kamennogorsk Anticlinorium (see IIIa on Fig. 1), the above-mentioned sedimentary sequences are overlain by volcaniclastic assemblages of the Basegi Group—the Oslyanka, Fedotovo and Us'va formations of Ablizin et al (1982) and Ancygin et al (1994). The Oslyanka

Fig. 4. Stratigraphic correlation chart for the Riphean and Vendian deposits in the Kvarkush-Kamennogorsk Anticlinorium and Poludov Range (from Ablizin et al. 1982; Keller & Chumakov 1982; Ancygin et al 1994; Maslov et al 1996).

Formation (100-300 m) consists of light, thin- and thick-bedded, fine- and medium-grained quartzitic sandstones and quartzites, among which phyllitic interbeds occur locally. The Fedotovo Formation comprises mainly dark shales with rare interbeds of siltstones and sandstones and in its upper part there are thin interbeds of limestones. In its lower part, in several places, there are basic and acidic effusive rocks (Ablizin et al. 1982). These volcanic rocks are attributed by some researchers to a separate unit, the Schegrovit Formation. The total thickness of the Fedotovo Formation varies from 500 m up to at least 2000 m. The overlying Us'va Formation consists of greenish-grey and mottled shales and argillites with interbeds of fine-grained quartz and feldspar-quartz sandstones. The thickness of the Us'va Formation varies from 280-300 m in its western sections, to more than 1200 m further east. In the easternmost parts of the Kvarkush-Kamennogorsk Anticlinorium, a unit known as the Kyrmin Formation, has been regarded by some researchers as a correlative of the Basegi Group. The Kyrmin Formation includes phyllites, siltstones, quartzitic sandstones and volcaniclastic shales and reaches up to 1000 m in thickness. In the Urals unified Stratigraphic scheme (Ancygin et al. 1994), the Kyrmin Formation has been correlated with the Krivaya Luka Formation of the southeastern zone of the Bashkirian Anticlinorium.

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All above-mentioned Upper Riphean deposits of the KvarkushKamennogorsk Anticlinorium are more metamorphosed than the overlying Vendian rocks. Lower Vendian. The Lower Vendian is represented in the Middle Urals by the Serebryanka Group (see Fig. 4), which includes the Tanin, Garevka, Koiva, Buton and Kernos formations (Ablizin et al 1982; Ancygin et al 1994; Raaben 1994; Maslov et al 1996; Maslov 2000&). The Tanin Formation (thickness from 360 to 800 m) contains dark diamictites (tillite-like conglomerates), quartz-feldspar sandstones and siltstones, shales, limestones and volcanic rocks; it conformably overlies the rocks of the Upper Riphean Us'va Formation (Ablizin et al 1982). There are three members in the Tanin Formation. The lower member can be divided into two parts, with a 50-80m thick unit of volcaniclastic shales, dolostones, limestones and shales with occasional beds (thickness 4-5 m) of tillite-like conglomerates, overlain by 250-500 m of dark and greenish-grey, matrix-supported, tillite-like conglomerates. The middle member (80-170m) comprises grey, greenish-grey and greyish-brown, quartz-feldspar and quartz sandstones, gritstones and shales. The upper member of the Tanin Formation (10-80 m) is represented mainly by dark matrix-supported conglomerates. According to Chumakov (1996), the tillite-like deposits of the Tanin Formation can be regarded as epicontinental glacial units associated with tectonically active troughs. They have been correlated with the Kurgashlya and Tolparovo formations of the Bashkirian Anticlinorium (Ancygin et al. 1994). The overlying Garevka Formation (350-700 m) includes banded, greenish-grey and dark phyllites, carbonaceous shales and greenish-grey quartz-feldspar sandstones (Ablizin et al. 1982; Ancygin et al 1994; Maslov 20006). The rocks of the Tanin and Garevka formations are cut by granosyenites of the Troitsk Massif which have yielded a Rb-Sr whole rock isotopic age of 621 ± 12 Ma (Semikhatov 1989). The overlying Koiva Formation (200-250 m) is composed of mottled shales and silty shales, siltstones, cream and claretcoloured limestones, dolostones and dolomitized limestones. In the northern part of the Kvarkush-Kamennogorsk Anticlinorium, the Koiva Formation is represented by tillite-like conglomerates and volcanic rocks (the Dvoretsk volcanic suite). The thickness of the Koiva deposits increases in this zone up to 550-700 m (Ablizin et al 1982). The Dvoretsk suite is represented by highTi trachybasalts, trachyandesites, trachyliparites and limburgites. The isotopic age of these rocks has been determined by the KAr whole-rock method at 577 ± 25 Ma, by Sm-Nd whole-rock and clinopyroxene fraction method to 568 ± 42 Ma, and by Rb-Sr whole-rock and clinopyroxene fraction method to 559 ± 15 Ma (Karpukhina & Pervov 2000). The overlying Buton Formation (300-350 m) is dominated by dark and black shales; in its upper part there are thin beds of fine-grained sandstones, siltstones and limestones. These pass up into the Kernos Formation (thickness up to 900 m), composed of arkosic and quartz-feldspar sandstones, gritstones, carbonaceous shales and thin limestones. In the eastern limb of the Kvarkush-Kamennogorsk Anticlinorium (see IIIb on Fig. 1), there is a unit of tillite-like conglomerates, thin-bedded shales and fine-grained sandstones, the Vilva Formation, that has been correlated with the Tanin and Koiva formations, described above (Ablizin et al 1982; Ancygin et al 1994). Upper Vendian. The Upper Vendian in the Middle Urals overlies the Serebryanka Group unconformably (Ablizin et al 1982; Ancygin et al 1994; Maslov et al 1996) and is represented by the Sylvitsa Group, which includes the Starye Pechi, Perevalok, Chernyi Kamen and Ust-Sylvitsa formations (Ablizin et al 1982; Ancygin et al 1994; Raaben 1994; Maslov et al 1996). The Starye Pechi Formation (thickness up to 500 m) is composed

of greenish-grey, yellow-green and yellow-grey siltstones and argillites with minor fine-grained sandstones. In the lower part of this formation, there are argillites with beds of matrix-supported conglomerates. The overlying Perevalok Formation (200-300 m) comprises dark grey quartz and quartz-feldspar siltstones and silty argillites. In its upper part, fine-grained quartz-feldspar sandstones predominate. The Chernyi Kamen Formation (15001600 m) includes rhythmic interbedded polymictic sandstones, siltstones and argillites. Becker (1980) described the following Metazoa from the formation: Tirasiana disciformis Pal., T. concentralis Beck, and Nemiama simplex Pal etc. K-Ar whole rock isotopic ages of gabbro-diabase dykes, intruding the Chernyi Kamen Formation, range from 630 ±15 to 590 + 20 Ma (Becker 1980; Raaben 1994). The uppermost unit in the Sylvitsa Group, the Ust-Sylvitsa Formation (250-350 m) is represented mainly by mottled fine- and medium-grained sandstones with minor siltstones and argillites. All these Upper Vendian deposits of the Kvarkush-Kamennogorsk Anticlinorium have undergone lower grade metamorphism than the rocks of the Serebryanka Group.

The Poludov Range On the Poludov Range (see IV on Fig. 1), as in the Middle Urals, only Upper Riphean and Vendian sedimentary sequences are known (Fig. 4). These are represented by the Kamen Rassolny, Demino, Nizva, Churochnaya, Il'yavozh and Kocheshor formations (Vladimirskaya 1955; Stratigraphy 1963; Keller & Chumakov 1983; Sokolov & Fedonkin 1990; Ancygin et al 1994). Based on their lithological composition and stromatolites, these formations are correlated with the Zilmerdak and the Katav formations and the Inzer-Uk levels of the Karatau Group of the Bashkirian Anticlinorium and also with certain horizons of the Lower and Upper Vendian (Raaben 1975; Smirnov 1977; Keller & Chumakov 1983; Ancygin et al 1994). Upper Riphean. The lowermost unit described from the Poludov Range is the Kamen Rassolny Formation; it is subdivided into three members. The lower member (thickness 450-550 m) consists of grey, greenish- and brownish-grey or pink arkose and feldspar-quartz sandstones and quartzitic sandstones, with some thin interbeds of gravellites and conglomerates in the lower part. The middle member (c. 50-200 m) is composed of thin-bedded greenish- or dark-grey siltstones, in some sections with desiccation cracks, and shales. The upper member (up to 400 m) consists mainly of red siltstones with thin interbeds of shales. According to Vladimirskaya (1955), accumulation of these siliciclastic deposits took place in near-shore and shallow-marine parts of a shelf basin. The best indicators of such environments are thin subparallel-bedding in all rock types, the presence of glauconite and siderite and, occasionally in the upper part of the formation, carbonate rocks with stromatolites. K-Ar isotopic ages of glauconite from siltstones of this member vary from 910 to 930 Ma (Raaben & Zhuravlev 1962; Becker 1968). The Demino Formation comprises dark cherry-red, greenishand bluish-grey and violet thin-bedded, clayey limestones and marls with relatively thin interbeds of calcareous siltstones, redcoloured argillites and intraclastic carbonate breccias. In the lower part of the formation, there is a unit of limestones (2040 m) with Upper Riphean stromatolites (Raaben & Zhuravlev 1962). The total thickness of the Demino Formation is about 300-350 m. Limestones of this formation pass up transitionally into the Nizva Formation (Stratigraphy 1963; Keller & Chumakov 1983) that consists of grey and dark-grey dolostones and partly dolomitized limestones; among the latter yellowishgrey and pink carbonate rocks occur locally. The total thickness of the Nizva Formation in its type area on Kamen Rassolny is

RIPHEAN AND VENDIAN: TIMANIDES AND URALIDES

about 1700 m; to the north and south, it decreases gradually to 800-1200m. In the most complete sections of the formation four members can be distinguished. The lower one (thickness up to 700 m) consists of grey, thin- and medium-bedded dolostones with thin interbeds of dark dolomites and shales; Upper Riphean stromatolites are characteristic of this member. The second member (c. 240 m) is represented by grey dolostones with lenses and interbeds of intraclastic carbonate breccias and numerous chert concretions, as well as Upper Riphean stromatolites. The third member (c. 350 m) comprises dolostones with interbeds and lenses of quartz sandstones and gravellites. In this member, stromatolites similar to those in the Inzer and Minyar formations of the Bashkirian Anticlinorium are characteristic (Raaben 1975; Ancygin et al 1994). The fourth member (thickness up to 300 m) consists of alternating argillites, siltstones and mottled clay limestones, with sporadic interbeds of fine-grained quartz sandstones. Thus the Nizva Formation includes a considerable thickness of carbonates and an abundance of stromatolites, indicating that deposition took place in shallow-marine environment in a warm climate (Vladimirskaya 1955). Lower Vendian. In the Poludov Range, the Lower Vendian (see Fig. 4) is represented by the Churochnaya and Il'yavozh formations (Sokolov & Fedonkin 1990; Raaben 1994; Ancygin et al. 1994). The Churochnaya Formation (500 m) is represented by grey tillite-like conglomerates, greenish-grey siltstones, argillites and dolostones. It overlies eroded rocks of the Upper Riphean with angular unconformity (Sokolov & Fedonkin 1990; Ancygin et al. 1994). According to Becker in Sokolov & Fedonkin (1990) and Raaben (1994), the Churochnaya Formation corresponds to the Kurgashlya Formation of the Bashkirian Anticlinorium. The overlying Il'yavozh Formation (600 m) includes grey and light-coloured sandstones with thin layers of siltstones and argillites. The K-Ar isotopic dating of authigenic glauconite from this formation yielded an age of c. 615-620 Ma (Sokolov & Fedonkon 1990; Raaben 1994). The Il'yavozh Formation corresponds to the Kernos Formation of the Kvarkush-Kamennogorsk Anticlinorium or to the lowest levels of the Sylvitsa Group (Raaben 1994). Thus, based upon the isotopic data, these two formations would be of Early Vendian age. Upper Vendian. The Upper Vendian sedimentary sequence in this part of the Urals comprises only the Kocheshor Formation (thickness up to 600 m). It consists of greenish-grey and grey conglomerates, sandstones, siltstones and green or claret argillites. K-Ar isotopic dating of authigenic glauconite from the Kocheshor Formation yielded an age of c. 560-590 Ma (Sokolov & Fedonkin 1990). In the regional stratigraphic scheme of the Urals, the Kocheshor Formation is correlated with the lower parts of the Sylvitsa and Asha Groups of the Middle and South Urals (Ancygin et al. 1994; Raaben 1994).

The Subarctic and Polar Urals Riphean. A large number of rock units of Riphean age are represented in the western part of the Subarctic and Polar Urals. They include the Schokur'aya, Oschiz, Puiva, Khobeaya, Ochetyvis, Morozovo and Amderma formations (Ancygin et al. 1994; Dushin 1997). Some are composed of metamorphosed sandstones, siltstones and limestones; others of carbonaceous cherts, phtanites and basic magmatic and jasperoid rocks with lenses and layers of marbles, occurring in various proportions. In general, these formations lack fossils and isotopic age data; thus correlation with the Riphean standard sections of the Bashkirian Anticlinorium has not been possible.

27

Fig. 5. Stratigraphic correlation chart for the Vendian sedimentary sequences in the Subarctic and Polar Urals (from Belyakova et al. 1992; Ancygin et al. 1994; Dushin 1997).

Vendian. In these northern regions of the Urals, both Lower and Upper Vendian deposits have been indentified (Fig. 5). The Lower Vendian includes the S ably a Gora, Enganepe and Kyzgey formations and the Upper Vendian is represented by the Laptopay, Man'ya and Khoidyshor formations (Ancygin et al. 1994; Dushin 1997). Lower Vendian. The S ably a Gora Formation (200 m) of the Subarctic Urals comprises quartz porphyries, dacites, metabasalts, basic tuffs, tuffaceous sandstones, turfites and welded tuffs. U-Pb isotopic dating of zircons from rhyolites of this formation yielded ages of 640 Ma (Chervyakovskii et al. 1992) and Rb-Sr dating of dacites gave an isochron age of 535 + 10 Ma (Andreichev 1999). Lavas and sub volcanic rocks of the Malda volcanic complex in the lower part of the Sablya Gora Formation has given a Rb-Sr isotopic age of 586 + 21 Ma (Chervyakovskii et al. 2000). Further north, in the Polar Urals, a complex of pre-Ordovician sedimentary and igneous rocks occur in the Enganepe Anticline. This Enganepe complex (thickness more than 1000 m) includes siltstones, polymictic sandstones, turfites, basic effusive rocks, basalts, andesites, quartz diorites, gabbro-amphibolites, welded tuffs and tuff-breccias and serpentinites (Dushin 1997). According to Bochkaryov & Yazeva (2000) all the igneous rocks occur in a tectonic melange. The Kyzgey Formation (2000 m) is represented by rhyolites, rhyodacites, basalts, tuff conglomerates, gritstones and polymictic sandstones. Late Vendian. The Laptopay and Khoidyshor formations are represented mainly by conglomerates, gritstones and sandstones, with some thin units of basalt, felsite and trachybasalt. The Laptopay Formation overlies the Sablya Gora Formation with angular unconformity (Ancygin et al. 1994; Raaben 1994; Dushin 1997). K-Ar whole-rock isotopic ages of the effusive rocks of the Laptopay Formation yielded an age of c. 560 Ma (Olovyanishnikov 1998a). The Man'ya Formation includes greenish-grey and mottled shales, tuffs, and basic and acid volcanic rocks. The Timan-Pechora region In the northeastern part of the East European Craton, a number of zones can be distinguished within the Timan-Pechora region

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A. V. MASLOV

(Va-Vf on Fig. 1). From SW to NE these are Mezen-Vychegodsk, Obdyr-Chetlas (including Obdyr-Nivshera and Chetlas-Dzhezhimparma subzones), Tsil'ma-Ropchino, Vymsk-Volsk and Kanin-Pechora zones (Olovyanishnikov 19980). Different Riphean and, in some cases, Vendian sedimentary sequences are known in each of these zones (Getsen 1987, 1991; Olovyanishnikov 1995, 1998Z?), but the Vendian deposits are prevalent only in the Mezen Syneclise and Timan, Vymsk-Volsk and KaninPechora zones (Fig. 7). There are a number of different opinions on the stratigraphical position of the Meso- and Neoproterozoic deposits of the Timan-Pechora region; in this paper, we mainly follow Olovyanishnikov (19980, b). The Mezen-Vychegodsk zone. In this zone (see Va on Fig. 1), the Upper Riphean deposits are represented by the Ust-Nafta Group overlain by the Safonovo Group. The Ust-Nafta Group consists of four formations. The first (thickness up to 400 m) consists mainly of dark shales, argillites and siltstones and, in some sections, fine-grained sandstones. The second formation (180-200 m) comprises predominantly dark-grey and black shales and argillites (this unit was earlier known as the Pez Formation). Thick-bedded medium- and fine-grained sandstones with rare thin argillite interbeds are the main rock types in the third formation of the Ust-Nafta Group. The fourth (400 m) is composed mainly of mottled fine and medium-grained sandstones. The third and fourth formations of the Ust-Nafta Group have been combined into the Dorogor Formation in previous literature (Dedeev & Getsen 1987). The Safonovo Group comprises the Omen and Nafta formations. The Omen Formation (500-700 m) is represented by a rhythmic interbedding of shales, siltstones, fine-grained sandstones, marls, dolostones and limestones. The Nafta Formation (c. 200 m) consists of dark grey and greenish-grey argillites, siltstones, marls, dolostones and limestones. The Upper Vendian deposits (Fig. 7) are represented by the Ust-Pinega, Mezen and Padun formations (Dedeev & Getsen 1987) in the Mezen Syneclise. The Ust-Pinega Formation (300650 m) includes, in its lower part, dark brown argillites with rare siltstone and sandstone layers and tuff horizons. Middle and upper parts of this formation combine grey and dark grey siltstones and argillites, and sandstones and siltstones with grey, green and brown argillites. The Mezen Formation (200-900 m) is represented, in the lower part, by alternating thin dark grey siltstones and argillites and, in the upper part, by greenish-grey sandstones, siltstones and argillites. The Padun Formation (150-350 m) comprises brown and greenish-grey sandstones and siltstones; argillites are known only in its uppermost part (Olovyanishnikov 19980, b). The Obdyr-Chetlas zone (including the Obdyr-Nivshera and ChetlasDzhezhimparma subzones). In the Obdyr-Nivshera subzone (see Vb on Fig. 1), the most complete sections of Upper Precambrian deposits are known from deep boreholes in the Obdyr uplift (Olovyanishnikov 1998&). Here, the so-called Obdyr Group is defined, which correlates with the Chetlas Group of the ChetlasDzhezhimparma subzone, located further east. The Obdyr Group

Fig. 6. Stratigraphic correlation chart for the Riphean deposits in the TimanPechora region (from Olovyanishnikov 19980, b).

Fig. 7. Stratigraphic correlation chart for the Vendian sedimentary sequences in the Timan-Pechora region and Amderma Anticline (from Ancygin et al. 1994; Dushin 1997; Olovyanishnikov 19980, b).

comprises two formations. The lower one (>450 m) consists of interbedded siliciclastic and carbonate rocks, comprising packets of alternating thin argillites, shales, siltstones, fine-grained sandstones and muddy limestones. Thin horizontal wavy bedding and cross-bedding are characteristic in these rocks; less commonly, desiccation cracks and various sole marks are observed. The upper formation (80-90 m) consists of massive, or thin parallel-bedded black shales. The Chetlas-Dzhezhimparma subzone (see Vc on Fig. 1). Riphean deposits are represented by the Chetlas Group, which includes the Svetlino, Novobobrovsk and Vizinga formations. The Svetlino Formation (up to 600 m) consists of grey and greenish-grey quartzitic sandstones, among which thin interbeds of dark-grey and black shales are observed. Lens-like, wavy- and cross- bedding are rather characteristic of sandstones and siltstones for this stratigraphic level of the Chetlas Group (Olovyanishnikov 19980). The Novobobrovsk Formation (100-500 m) comprises mainly shales and argillites with thin interbeds of siltstone and feldspar-quartz quartzitic sandstones. In some sections, there are also lenses and lens-like interbeds of gravellites and conglomerates. Thin, horizontal, rippled and cross-bedded units with desiccation cracks and sole marks of various types are characteristic of these sandstones and siltstones. These structures indicate that accumulation took place in near-shore, non-marine and shallowmarine environments (Plyakin 1972; Dedeev & Getsen 1987; Getsen 1987; Olovyanishnikov 1998&). The Vizinga Formation (650-1000 m) overlies the Novobobrovsk Formation with a local Stratigraphic break (Olovianishnikov 19980, b). It comprises black shales and siltstones, with subordinate quartzitic sandstones and, in some sections, thin interbeds of these rocks and lenses of conglomerates and breccias. Horizontal and wavy bedding, cross-bedding, lenticular bedding, wave and current ripples, sole marks, convolute bedding and desiccation cracks are characteristic of this formation. In the upper part, there are prominent intraclastic shale breccias and desiccation cracks (Dedeev & Getsen 1987). According to Olovyanishnikov (19980), accumulation of the Chetlas Group sediments took place in a shallow-marine basin, which at different times changed into an open lagoon, bay or large lake. The Tsil'ma-Ropchino zone. In the Tsil'ma-Ropchino zone (see Vd on Fig. 1), carbonate sequences of the Upper Riphean Bystrino Group have been divided into the Rochug, Pav'yug and Paun formations. The Rochug Formation (c. 400 m) is represented by dark, greenish-grey and mottled phyllitic shales with rare thin interbeds of limestones, dolostones, siltstones and fine-grained sandstones. The lower part of the formation displays characteristic wavy and horizontal bedding (Dedeev & Getsen 1987) that indicate

RIPHEAN AND VENDIAN: TIMANIDES AND URALIDES

shallow-marine environments. The Pav'yug Formation (10001400 m) comprises limestones and dolostones with numerous stromatolites Conophyton garganicus var ikeni Raab. et Kom., Gymnosolen ramsayi Steinm., G. asymmetricus Raab., Inzeria cf. djejimi Raab., Parmites nubilosis, Baicalia ex. gr. prima, B. cf. lacera, Poludia mutabilis Raab et Kom. (Raaben & Oparenkova 1997). These sediments were formed mainly in shallowmarine and sublittoral environments (Raaben 1975; Getsen 1987, 1991) but back-reef, reef and basinal facies are also present (Bogdanov & Plyakin 1999). The Paun Formation (900-1000 m) is composed mainly of dark phyllites and shales. Limestones, marls and dolostones here play a subordinate role. Horizontal and wavy bedding are characteristic of these rocks (Dedeev & Getsen 1987). In some carbonate interbeds of the Paun Formation, Upper Riphean stromatolites and microfossils have been found (Olovyanishnikov 1998/?). Accumulation of the Paun deposits probably occurred in a shallow-marine environment. The Vymsk-Volsk zone. In the Vymsk-Volsk zone (see Ve on Fig. 1), sedimentary successions of Vymsk Group occur which are similar to the Pav'yug and the Paun formations of the Bystrino Group (see above). The Vymsk Group is subdivided into Pok'yus, Lunvozh and Kykvozh formations (Olovyanishnikov 19980); based on rich microfossil assemblages (Belyakova et al. 1992) only the first formation belongs probably to the Upper Riphean. The Pok'yus Formation is presented in its lower part (350-400 m) by grey and yellowish-grey quartzitic sandstones with rare shale interbeds, middle units of the formation (up to 1400 m) consist of dark shales and phyllites with thin interbeds and lenses of limestones, and the upper part (1000-1100 m) comprises interbeds of siltstones, shales and fine-grained sandstones with lenticular layers of black limestones and marls (Olovyanishnikov 1998Z?). The overlying Lunvozh and Kykvozh formations (Fig. 7) may be of Early Vendian age. The Lunvozh Formation is represented, in its lower part, by dark-coloured argillites and siltstones (more than 900 m thick); in its middle and upper parts (thickness up to 2000-2500m) there are a few thick macrorhythms, ranging from sandstone-dominated packets to argillitic units. The Kykvozh Formation (700-800 m) overlies the Lunvozh deposits with an erosional contact. It comprises light grey quartz and quartz-feldspar sandstones at the base and packets of sandstones, siltstones and dark argillites in the middle and upper parts. The Kanin-Pechora zone. Upper Precambrian sedimentary and metasedimentary associations (Mikulkino, Tarkhanovo and Tabuevo Groups) are known from the territory of the KaninPechora zone (Fig. 6) on the Kanin Peninsula (see Vf on Fig. 1). On the Kanin Kamen ridge, the Mikulkino Group is represented mainly by quartzites, amphibole and chlorite-epidote-amphibole schists and marbles. According to Olovyanishnikov (1998£), the lower part of the Mikulkino Group comprises amphibolite facies psammites and semipelites, whereas the middle part consists of alternating schists, metasiltstones and metasandstones. Higher up in the section, the rocks are gradually enriched in carbonate material and, in the upper part, there are carbonate-bearing rocks (thin-banded, so-called 'scarnoids'). The thickness of the Mikulkino Group is estimated at 1500 m and the base is not seen. The Tarkhanovo Group (4500-5000 m) is composed, in its lower part, of schists and micaceous quartzites; in its middle part monotonous micro- and/or mesorhythmic successions of dark schists (quartz, biotite, plagioclase etc.) occur, among which thin interbeds of quartzitic sandstones are sometimes present. The upper part of the group consists of alternating schists, metasiltstones and quartzitic sandstones. Deposition of the Tarkhanovo Group took place in continental slope or rise environments (Olovyanishnikov 19980, b). The Tabuevo Group consists of three formations. The lower Bolvan Creek Formation is represented in its lower part (c. 900 m) by quartzitic sandstones and metasiltstones, and its

29

middle and upper parts (up to 2000 m) are dominated by greenish, bluish and dark grey quartz-sericite and quartz-sericite-chlorite phyllites with subordinated interbeds of siltstones and marls. The Yanei Formation (c. 400 m) includes mottled quartzites with shale interbeds; the latter increase in number and thickness up the section. The Gnilsk Formation (thickness is c. 800 m) mainly comprises greenish and dark grey chlorite and sericitechlorite-bearing volcanogenic phyllites with thin interbeds of limestones, metasiltstones and quartzitic sandstones. In the upper, c. 100m interval, quartz-bearing dolomites, shales and marls predominate, with stromatolites in the carbonate blocks. The Izhma-Pechora depression. In the Izhma-Pechora depression (VI on Fig. 1) only the Lower Vendian sedimentary successions of Seduaykha Formation (c. 500-600 m) are known (Fig. 7). This formation consists of dark-coloured shales with subordinate layers of fine-grained sandstones and siltstones. The Khoreyver depression. The Khoreyver depression (VII on Fig. 1) is located in the far northeastern part of the TimanPechora region. The Lower Vendian Vozey and Sandivey formations are known from here, based on drillcore data (Dedeev & Getsen 1987). The Vozey Formation (c. 200m) consists of tuffs, quartz porphyries, welded tuffs, liparite-dacites and sericite-quartz-pyrophyllite schists. K-Ar isotopic ages (wholerock method) of the Vozey quartz porphyries range from 520 to 620 Ma. The Sandivey Formation (100-450 m) overlies the Vozey rocks unconformably, and comprises polymictic sandstones, tuff-sandstones, siltstones and acidic ash-fall tuffs. Along the southern coast of the Kara Sea, in the vicinity of Amderma (Fig. 1), there is an Upper Vendian succession known as the Sokol'nino Formation (Dedeev & Getsen 1987; Sokolov & Fedonkin 1990). It includes cherts, sandstones, gritstones and conglomerates with acidic lavas and tuffs. The thickness of the Sokol'nino Formation is 2300-2500 m. According to Olovyanishnikov (19980), the Riphean and Vendian deposits of the eastern part of the Timan-Pechora region were formed in moderately deep-water conditions, probably in outer shelf or upper the part of continental slope environments. Stratigraphic correlation of the sedimentary sequences Correlation of the Riphean and Vendian sedimentary sequences of the western part of the Urals and adjacent regions with typesections in the Bashkirian Anticlinorium is shown in Figures 8, 9 and 10. It is based on: (1) general Stratigraphic similarity of the sequences; (2) similarity of lithological composition of several formations and members; (3) identification of key horizons; (4) isotopic ages; and (5) fossil content (fauna and flora). With regard to the Lower Riphean of the Volga-Urals, the Prikamsk Formation of the Kyrpin Group is correlated with the Ai Formation of the Bashkirian Anticlinorium, and the Kaltasa Formation with the Satka Formation (Keller & Chumakov 1983; Isherskaya & Romanov 1993; Maslov 20000; Romanov & Isherskaya 2001). The Nadezhdino Formation corresponds either to the Bakal or to the Mashak formations (Romanov & Isherskaya 1994, 2001), the latter being the more likely, considering that the Kaltasa Formation is coeval with the entire Satka-Bakal interval of the Burzyan Group (Fig. 8). In the case of the Middle Riphean, the Tukaevo Formation of the Volga-Urals region corresponds to the Zigalga Formation of the Bashkirian Anticlinorium, and the upper part of the Olkhovo Formation has been compared with the Avzyan Formation of the Yurmatau Group (Romanov & Isherskaya 1994, 2001; Maslov & Isherskaya 1998; Kozlov et al. 1999; Maslov 20000). The Akberdino Horizon, in the opinion of most researchers, is

30

A. V. MASLOV

Fig. 9. Vertical and lateral architecture of the Middle Riphean sedimentary successions in the southern segment of the eastern periphery of the East European Craton. Legend as in Fig. 8. R{, Lower Riphean sedimentary sequences; msr^, lower and middle parts of Mashak Fm.; msh2, upper part of the Mashak Fm.; zg, Zigalga Fm.; zk, Zigazino-Komarovo Fm.; av, Avzyan Fm.; tk, Tukaevo Fm.; ol, Olkhovo Fm.

Fig. 8. Vertical and lateral architecture of the Lower Riphean sedimentary successions in the southern segment of the eastern periphery of the East European Craton. ai t , lower part of the Ai Fm.; ai2, upper part of the Ai Fm.; st^, first and second members of the Satka Fm.; st3, Polovinka Member of the Satka Fm.; st4.5, fourth and fifth member of the Satka Fm.; b } , lower member of the Bakal Fm.; b2, upper member of the Bakal Fm.; bin2, upper member of the Bolshoi Inzer Fm.; sr2-4, second, third and fourth members of the Suran Fm.; jsh, Jsha Fm.; prk, Prikamsk Fm.; kit, Kaltaza Fm.; nd, Nadezhdino Fm.; arl, Arlan member of the Kaltasa Fm.

correlated with the Zigazino-Komarovo Formation of the western slope of the Southern Urals (Fig. 9). Upper Riphean correlation indicates that the Usa Formation of the Volga-Urals region is similar to the Biryan-Nugush level of the Zilmerdak Formation of the Bashkirian Anticlinorium (Maslov & Isherskaya 1998). The quartz sandstones of the Leonidovo Formation correspond to the quartz and quartzitic sandstones of the Lemeza Member of the Zilmerdak Formation. The Prijutovo Formation is correlated with the Bederysh Member of the same formation, and the pink and red-coloured limestones of the Shikhan Formation are similar to the red-coloured limestones of the Katav Formation. Further north, in the Kvarkush-Kamennogorsk Anticlinorium, the Kedrovka and Basegi groups have been correlated with the Karatau Group of the Bashkirian Anticlinorium (Maslov et al. 1996). In contrast to the Upper Riphean successions of the Southern Urals, the upper third of the Basegi Group in the Middle Urals area is composed of a substantial volcanogenic assemblage (Fig. 10).

With regard to the Vendian, microfossils from the shales of the Starye Pechi Formation of the Kvarkush-Kamenogorsk Anticlinorium are characteristic for the Upper Vendian Redkino horizon (Ablizin et al. 1982; Raaben 1994). Becker (1980) described casts of Aruberia banksi (Glaes. and Walt.) from the bedding surfaces of the Ust-Sylvitsa siltstones. Together with finds of metazoa from the underlying Chernyi Kamen Formation and the isotopic age of the diabase dyke swarm, it is reasonable to consider the Sylvitsa Group as an Upper Vendian succession. The Upper Riphean sedimentary successions of the Southern and Middle Urals and those in the Poludov Range are similar. There are conglomerate and sandstone sequences in the basal parts of all the Upper Riphean groups, and sandstones, sandstone-si Itstoneshale packages and carbonate or carbonate-shale units predominate in the middle and upper parts of these stratigraphic subdivisions. Stromatolites from the carbonate rocks of the fourth member of the Nizva Formation are comparable with those from the Uk Formation of the Bashkirian Anticlinorium, which make it possible to correlate the upper part of the Nizva Formation with the highest levels of the Upper Riphean in the type locality (Raaben & Zhuravlev 1962; Raaben 1994). On the basis of the compositional similarities of the Ust-Nafta and Safonovo groups of the Mezen Basin (Fig. 10) with the Upper Precambrian deposits of the Central Timan and the finds of the microfossils, the age of these two groups is accepted as Late Riphean to perhaps Early Vendian (Olovyanishnikov 1998Z?). In the fine-grained siliciclastic rocks of the lower unit of the Obdyr Group, there are microfossils (Klidinella hyperboreica and K. sinica) which are comparable with those found in the Vizinga Formation of the Chetlas and Safonovo groups (Olovyanishnikov 1998a, b) and which indicate a Late Riphean, or perhaps Early Vendian age. Raaben and Oparenkova (1997) undertook a special study of stromatolites from carbonate deposits of the Tsilmen Kamen. According to their data, in the Tsilma River basin, the Pav'yug Formation is subdivided into three members. The first one (c. 200-250 m) combines mainly limestones and dolostones with specific associations of the stromatolites Conophyton garhanicus var. ikeni Raab et Kom., Baicalia ex. gr. prima and

RIPHEAN AND VENDIAN: TIMANIDES AND URALIDES

31

compared with the Katav and the Demino formations. The Pav'yug Formation of the Tsilmen Kamen region stratigraphically corresponds to the Nizva Formation of the Poludov Range and is correlated with the Late Riphean Inzer-Minyar interval of the Bashkirian Anticlinorium. There are three or four levels of tillites or tillite-like conglomerates in the Upper Riphean(?)-Vendian deposits of the Urals. The first of them is located in the Tanin Formation in the Middle Urals (Keller & Chumakov 1983; Becker 1988). The second level of tillite-like deposits (the Koiva Formation) is known only on the Middle Urals (Ablizin et al 1982). The third level corresponds to the Kernos Formation of the Middle Urals and the Tolparovo and Suirovo formations on the Southern Urals (Sokolov & Fedonkin 1990). Finally, the uppermost level of tillite-like deposits is located in the Starye Pechi Formation of the Kvarkush-Kamennogorsk Anticlinorium. At present there are no agreed correlations of all these tillite-like horizons in the Urals with the Varangerian horizons. Some of them, for example tillite-like conglomerates of the Tanin Formation, may be synchronous with Late Riphean glacial deposits.

Riphean-Vendian tectonic and sedimentary events of the eastern East European Craton

Fig. 10. Vertical and lateral architecture of the Upper Riphean sedimentary successions along the eastern and northeastern peripheries of the East Europena Craton. Legend as in Fig. 8. zl b Biryan member of the Zilmerdak Fm.; kt, Katav Fm.; in, Inzer Fm.; mn, Minyar Fm.; uk, Uk Fm.; us, Usa Fm.; In, Leonidovo Fm.; prt, Prijutovo Fm.; sn, Sinegorsk Fm.; klj, lower part of the Klyktan Fm.; k!2, upper part of the Klyktan Fm.; os, Oslyanka Fm.; fd, Fedotovo Fm.; usv, Us'va Fm.; rs, Kamen Rassolny Fm.; dm, Demino Fm.; nzv, Nizva Fm.; R3us-nf + saf, Ust-Nafta and Safonovo groups; R3rch + pav + paun, Rochug, Pav'yug and Paun Formations; R3mk + tar + tab, Mikulkino, Tarkhanovo and Tabuevo groups.

B. cf. lacera. The second member (up to 300-350m) does not contain stromatolites, and the third one (450-500 m) is represented by grey and mottled, massive and medium-bedded dolostones with intraclastic carbonate breccias and the stromatolites Gymnosolen giganteus Raab., G. ramsayi Steinm., G. levis Kryl., G. asymmetricus Raab., Minjaria sp., Inzeria cf. djejimi Raab., Parmites nubilosus Raab. et Komar, P. concrescens Raab. Tungussia perforata Raab. and Poludia polymorpha Raab. As these authors note, the Pre-Pav'yug part of the Upper Riphean sequence in the vicinity of the Tsilmen Kamen can be

There are a variety of viewpoints on the tectonic and sedimentary history of the territory under review during the Riphean-Vendian time. In the opinion of the present author, in the Southern Urals there were several short episodes of 'diffuse' and 'linear' rifting at the beginning of both the Early and the Middle Riphean, with development of intracratonic sedimentary basins (1.65-l.OGa) and long intervals in between of quasi-static conditions (Maslov 1994, 20000; Maslov et al 1997). During the Late Riphean, judging from the great similarity of Upper Riphean formations along the eastern edge of the East European Craton, a major shallow-marine basin developed. This stretched from the South Urals up to the Poludov Range over a distance of more than 1500km. Finally, during the Vendian, along the western slopes of the Southern and Middle Urals and the Volga-Urals region, epicratonic sedimentation in relatively wide, shallow-marine basins was dominant, in some periods in terrestrial environments. According to Becker (1968, 1988) and Puchkov (2000), the Late Vendian sedimentary sequences of the Southern and Middle Urals constitute a molasse complex related to Timanian orogeny. Considering the general features of Late Precambrian sedimentation along the northeastern part of the East European Craton, Olovyanishnikov (1998&) pointed out that, in Early and Middle Riphean times, the Timan-Pechora region was characterized by small riftogenic basins. In the Late Riphean, a thick accumulation of carbonate rocks with stromatolites formed along the shelf-edge, within the Tsil'ma-Ropchino zone, which is considered by many researchers to mark the transition zone from the shelf or platform to continental slope and basinal environments. To the SW of this zone in the peri-cratonic domain, in the Late Riphean, mainly near-shore and shallow-marine siliciclastic sediments were deposited. The regions located to the NE from the Tsil'ma-Ropchino zone were characterized by the accumulation, during the Late Riphean, of volcaniclastic sequences, many of which show features of turbidite deposition. Thus, the Late Riphean sedimentary basin on the northeastern edge of the East European Craton can be compared with the present-day Arctic Ocean (Getsen 1991), which is characterized by huge sedimentary prisms along a wide and gentle shelf and continental slope. Dushin (1997) believes that at the end of the Late Riphean and during the Early Vendian, in the Polar Urals, oceanic and volcanic island-arc magmatic complexes were formed (e.g. the Enaganepe ophiolite). The Upper Vendian Laptopay, Man'ya, Khoidyshor and Sokol'nino formations of the Subarctic and Polar Urals

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A. V. MASLOV

represent a volcanogenic coarse-grained molasse. At present, the Timan-Pechora region and the northern part of the Urals form a mosaic of terranes made up of continental blocks and relics of oceanic crust (Belyakova & Stepanenko 1991; Dushin 1997).

Conclusions Based on a variety of basic parameters such as rock types and colour, sandstone composition, sedimentary structures, thickness and distribution of deposits within the basins, character of cyclicity, typical facies associations and their vertical and lateral architecture within the sedimentary basins, depositional environments and the character of the source zones, we can define several types of Riphean sedimentary sequences in the territory under review (Maslov 1998). As noted above, the Lower Riphean deposits are located only in the southern segment of the eastern periphery of the East European Craton. The main features of the Burzyan and Kyrpino sedimentary sequences are: (1) predominance of shallow-marine and peri-littoral siliciclastic and carbonate deposits with only a minor amount of terrestrial material; (2) thickness increases from proximal to distal zones of the Lower Riphean sedimentary basin; and (3) depocentres of all three main, Lower Riphean lithostratigraphic subdivisions (the Ai-Prikamsk, Satka-Kaltasa and Bakal-Nadezhdino levels) connected with the central zone of the sedimentary basin. Breaks and hiatuses are characteristic for the outer zones of the Lower Riphean sedimentary basin, whereas the most complete sedimentary successions are found in its inner parts. A different facies association is characteristic of the Middle Riphean deposits of the eastern periphery of the East European Craton. These deposits, like those of the Lower Riphean, are located only in the southern segment (Bashkirian Anticlinorium and the Volga-Urals region) of the territory under review. The maximum thickness (up to 5000-6000 m) of these sedimentary and volcano-sedimentary deposits is known from the eastern edge of the Bashkirian Anticlinorium, whereas in its western part and in the Volga-Urals region, the total thickness of the Middle Riphean Serafimovo Group is no more than HOODOO m. For the lithostratigraphic equivalents of the Yurmatau and Serafimovo Groups (such as the Zigalga and Tukaevo formations, the Avzyan Formation and the upper part of the Olkhovo Formation) the depocentres varied considerably (Maslov 20000). For the Upper Riphean sedimentary sequences of the TimanPechora region some general features are characteristic (Maslov 1996) including the lateral combination of near-shore and shallow-marine deposits adjacent to moderately deep to deepwater deposits and the thickest sedimentary sequences being typical for the distal zones of the basin. There were two laterally connected zones in the Timan-Pechora region during the Late Riphean. The first, having a predominance of shallow-marine sandstones and carbonate deposits in the southwestern parts of this territory (total thickness not more than 2200 m) and the second, with moderately-deep marine (continental slope and rise) turbidite sediments (up to 9000-10 000 m) in its northeastern parts. The Upper Riphean deposits of the southern segment of the eastern East European Craton are characterized by a different type of vertical and lateral architecture. A thick sequence of terrestrial and near-shore massive arkose, grits and gravel-sandstones is characteristic of the lower part of the Karatau Group (the Biryan Member of the Zilmerdak Formation). Middle and upper parts of the Karatau Group consist of shallow-marine siliciclastic and carbonate deposits, which are very similar to the so-called 'Grand Cycles' of the Windermere Supergroup (Aitken 1989; Narbonne & Aitken 1995).

It is probable that, during the Late Riphean, there were two different zones in the sedimentary basin. The first of them, a proximal shelf zone, was located on the territory of the modern Bashkirian Anticlinorium, and the second, an outer shelf and upper slope zone, located in the Timan-Pechora region. The Vendian sedimentary sequences on the northeastern and eastern parts of the East European Craton and in the western megazone of the Urals probably were formed in a broad basin in front of the Timanide collisional zone (Becker 1968; Keller & Chumako 1983; Gee et al 2000; Puchkov 2000; Roberts & Siedlecka 2002; Grazhdankin 2004). Special thanks are due to the Europrobe Programme and especially to David G. Gee and Victor Puchkov for discussions at many workshops. For constructive reviews and linguistic help I also thank David Roberts and David G. Gee, and for editorial aspects Olga K. Bogolepova. The Russian Foundation for Basic Research (RFBR) is acknowledged for financial support (grants 00-05-64497 and 03-05-64121).

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SOKOLOV, B. S. & FEDONKIN, M. A. (eds) 1990. The Vendian System, Vol. 2, Springer-Verlag, Berlin, 273 pp. [Stratigraphy of the USSR. Upper Precambrian]. 1963. Gosgeoltekhizdat, Moscow, 2, 761 pp [in Russian]. TUGARINOV, A. I., BIBIKOVA, E. V., et al. 1970. [Geochonology of the Urals Precambrian]. Geochemistry, 4, 501-509 [in Russian]. VLADIMIRSKAYA, E. V. 1955. [Pre-Devonian deposits of Kolva-Vishera Region]. In: NALIVKIN, D. V. (ed.) [Stratigraphy of Palaeozoic deposits of Timan and western slope of the Urals]. Gostoptekhizdat, Leningrad, 225-280 [in Russian]. WEISS, A. F., KOZLOVA, E. V. & VOROB'EVA, N. G. 1990. [Organicwalled microfossils of the Uralian type section (South Urals)]. Izvestiia Akademii nauk SSSR, Seriia geologicheskaia, 9, 20—36 [in Russian]. ZAIDES, B. B. 1973. [Use of hydromica minerals for investigation of catagenesis and metamorphism of rocks]. PhD thesis, Kiev [in Russian]. ZAITSEVA, T. S., IVANOVSKAYA, T. A., et al. 2000. [Rb-Sr age and NGR-spectrums of glauconites from Uk Formation, Upper Riphean, South Urals]. In: LAVEROV, N. P. (ed.) [Isotopic dating of geological processes: new methods and results]. GEOS, Moscow, 144-147 [in Russian].

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Late Neoproterozoic sedimentation in the Timan foreland D. GRAZHDANKIN Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK (e-mail: [email protected])

Abstract: The late Neoproterozoic Vendian succession fills a peri- to epicratonic Mezen Basin in front of the Timan Orogen, and is exposed along the northwestern flank of the basin. In the SE White Sea area, the siliciclastic succession demonstrates a wide range of lithofacies that define a transition from a low-energy shallow-marine muddy shelf to a braid-delta plain. Significantly, the prodelta, distributary-mouth bar, and delta plain lithofacies all demonstrate a remarkably tight clustering of palaeocurrent trends suggesting that the sediment was sourced from the NE. This pattern is interpreted as representing deposition in a distal setting in the Timan foreland basin. U-Pb zircon dates of 558 + 1 Ma and 555.3 + 0.3 Ma for volcanic tuffs in the Vendian succession provide lower age constraints for the emergence of the Timanian hinterland.

The Neoproterozoic tuffaceous-siliciclastic Vendian succession, constituting a thick (250-2500 m) sediment prism that fills the Mezen Basin in front of the Timan Orogen, extends into the Ural Basin, and forms a vast epicratonic inland tongue extending into the Moscow Basin (Fig. 1). Recent progress in understanding the deep structure of the Timan-Pechora collisional zone (Olovyanishnikov 1998; Olovyanishnikov et al 1996, 2000), coupled with radiometric dating (Gee et al. 2000; Gorokhov et al. 2001), have revealed that, during late Neoproterozoic time, the northeastern margin of the East European Craton experienced compressive stress from the Timan Orogen. That the Vendian shallow marine, fluviodeltaic and alluvial deposits were related to orogeny has long been appreciated (Shatsky 1952; Keller 1963; Becker 1968); however, the cause and timing of Vendian subsidence is still obscure (Aksenov 1985; Nikishin et al. 1996; Puchkov 1997; Maslov et al. 1997; Ivanov & Rusin 2000). A link between the geodynamic evolution of the Timanides and the Neoproterozoic depositional history of the craton may therefore be expected, but such a relationship has been difficult to prove given the limited exposure and the poor understanding of Vendian sedimentology. In this paper, I present preliminary results of an on-going sedimentological and stratigraphic investigation of the Vendian succession, focusing on exposures in the SE White Sea area on the northwestern flank of the Mezen Basin. The Vendian succession in the White Sea area is unmetamorphosed, undeformed, abundantly fossiliferous, well known from extensive drilling, contains volcanic tuffs with U-Pb zircon dates, and provides the only unambiguous outcrop of Vendian strata within the Mezen Basin (Stankovsky et al. 1981, 1985; Martin et al. 2000; Grazhdankin 2003, 2004). Stratigraphically, the succession is subdivided into four formations (Fig. 2): Lamtsa, Verkhovka, Zimnegory and Yorga (Grazhdankin 2003). A 10 cm thick tuff layer in the Zimnegory Formation in the SE White Sea area yielded a population of up to 275 jjim long, euhedral, doubly terminated prismatic zircons, many of which contain abundant inclusions. Four multigrain and 15 single-grain analyses define a normally discordant array (207pb/206pb > 207pb//235U > 206pb//238u) wkh ^ uppef intercept

date of 555.3 ± 0.3 Ma (Martin et al. 2000). This constrains a major sequence boundary in the Vendian succession. It separates a lower, mostly marine depositional sequence from an upper, mostly alluvial sequence. Lower in the sequence, a tuff at the base of the Verkhovka Formation has a U-Pb zircon date of 558 ± 1 Ma (Grazhdankin 2003). This date is the closest constraint for the timing of the drowning of the craton. The Neoproterozoic-Cambrian boundary has not been documented in the section, although peculiar brecciated rocks in the White Sea area, interpreted as kimberlite-hosted xenoliths, have yielded late Cambrian Ungulate brachiopods and late Cambrian

to early-middle Ordovician acritarchs (Verichev et al. 1990; Popov & Gorjansky 1994). The Vendian succession is intruded by subvolcanic pipes (diatremes) and sills of late Devonian age (Mahotkin et al. 2000), and is unconformably overlain by middle Carboniferous sediments. The present day northwestern limits of the Vendian succession are erosional (Jakobson & Nikulin 1985).

Pre-Vendian setting Pre-Vendian sediments are confined to a vast pericratonic Mezen Basin on the northeastern margin of the East European Craton, but also fill deep depressions in the eastern slope of the Baltic Shield, the Onega and Zimnegory Grabens (Fig. 1). The sediments are poorly dated; nevertheless their early Neoproterozoic age is constrained by associated microfossils and regional correlations (Sivertseva & Stankovsky 1982; Dedeev & Keller 1986; Jakobson et al. 1991; Sivertseva 1993; Nikishin et al. 1996). Early Neoproterozoic sedimentology is also poorly known because it is based on sporadic and incomplete borehole data. In the Mezen Basin, the pre-Vendian succession is divided into the Ust-Nafta and Safonovo Groups and the Uftuga Formation (Aksenov et al. 1978; Dedeev & Keller 1986; Olovyanishnikov 1998). The Ust-Nafta Group reaches 1200 m in thickness and consists of interbedded shales, siltstones and fine-grained sandstones. The overlying Safonovo Group is a flysch-like carbonatesiliciclastic sequence that thickens progressively to the NW, from 490 m to 735 m. It is truncated and unconformably overlain by poorly-sorted sandstones, with pebble-size lithic and volcanic clasts of the Uftuga Formation (100-1200 m thick). The Onega Graben extends NW-SE, from Finland through the head of Kandalaksha Bay on the White Sea, to the Onega Peninsula, parallel to the craton margin (Fig. 1). The oldest drilled sequence is represented by tholeiitic basalts with volcaniclastic breccia and volcanic bombs (Stankovsky et al. 1972). The volcanic rocks are overlain by the siliciclastic Nenoxa Formation (350 m thick) of presumed early Neoproterozoic age. The basal unit (80m thick) is represented by grey, poorlysorted, polymictic, cross-bedded sandstones. It is followed by variegated, reddish-brown to pink, medium- to coarse-grained, quartzose sandstones, interbedded with grey, cross-bedded gravelstones. The sandstones are massive along the southwestern side of the graben, but grade northeastwards into thin-bedded and crossbedded varieties, interstratified with shales; the gravelstones wedge out in the same northeasterly direction (Zoricheva 1963; Stankovsky et al. 1981). The rocks are intruded by basalts of pre-Vendian age (Stankovsky et al. 1977).

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 37-46. 0435-4052/047$ 15 © The Geological Society of London 2004.

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D. GRAZHDANKIN

Pre-vandian sedinent

in the Mezen Basin

Fig. 1. Present-day setting of the Vendian succession on the East European Craton. Palaeocurrent data for the South Urals are from Becker (1968).

Vandian succession And coeval srata

Palaecurren direction Duting vandial

The Zimnegory Graben is located to the NE and parallel to the Onega Graben (Fig. 1). The southwestern part of the graben is filled with interbedded sandstones and shales of the Chidvia Formation (drilled thickness 524 m). The sandstone beds are maroon to pink fine- to coarse-grained arkoses exhibiting erosional bases, shrinkage cracks, shale clasts and wave ripple cross laminations; they are very similar to the Nenoxa Formation of the Onega Graben (Jakobson et al. 1991). In contrast, the coeval northeastern part of the graben consists of maroon to dark grey, laminated shales and marls of the Tuchkino Formation (drilled thickness 162 m), which are very similar lithologically to the Safonovo Group of the adjacent Mezen Basin (Fig. 1).

Sedimentology of the Vendian succession In the course of my fieldwork in the SE White Sea area, over 150 sections and six boreholes (drilled during 1993-1996 on the Onega Peninsula, with 85% mean core recovery) were logged and correlated (Fig. 2). Vendian stratification is characterized by a pronounced and recurring cyclic arrangement of lithotypes, allowing long-distance correlation between boreholes and outcrops (Igolkina 1959; Grazhdankin 2003). The Vendian succession can also be divided into a series of lithofacies, which can be grouped into five process-related facies assemblages: (1) a laminated shale; (2) an alternating shale/siltstone; (3) an interstratified sandstone and shale; (4) a channelized sandstone; and (5) a trough-cross-bedded sandstone (Fig. 2). Laminated shale Description. This assemblage comprises maroon laminated shale, grey to pale brown thin shale-siltstone alternations, grey

wave-rippled siltstone in shale, and very fine sandstone. Units of shale are up to 20 m thick and consist of laminations to very thin (< 1 mm) beds, accentuated by bedding-parallel partings, sapropel-like films, and thin (1-2 mm to 10 cm) graded beds of volcanic ash (Fig. 3a). The shale-siltstone alternations are about 1 mm to 10 mm thick; siltstone makes up less than a quarter of the lower couplets. Siltstone and sandstone beds (15-30 cm) are rare and laterally discontinuous. They have sharply defined lower boundaries with tool and swing marks, crescentic scour casts, flute casts, and frondescent casts (Fig. 3b). These beds appear to have been amalgamated, preserve wave-rippled lamination with several truncation surfaces, and have rippled tops. Interpretation. The laminated shale facies, containing undisturbed ash beds and sapropel-like films, is interpreted as the product of suspension fallout below wave base. The sharp interbedding of rare wave-rippled amalgamated siltstone and sandstone beds with sharp, locally scoured erosional bases suggests deposition in alternating wave-influenced and quiet-water conditions. Considered together, these facies characterize a low-energy shallowmarine muddy shelf setting with storm influence. Alternating shale/siltstone Description. This assemblage is dominated by thick monotonous units of greyish-green graded siltstone-shale couplets, generally 4-6 mm to 15 mm thick; siltstone makes up the lower half to three quarters of the couplets (Fig. 3c). Siltstone laminae are characterized by sharp bases and fine graded bedding, but thicker siltstone laminae tend to have cross-bedding and starved wave ripples. The starved wave ripples, formed of siltstone coarser than the couplets themselves, have heights of 20-30 mm and a wavelength of 10-20 cm. Cross-strata sets have a mean

SEDIMENTATION IN THE TIMAN FORELAND

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Fig. 2. Stratigraphy of the Vendian succession in the White Sea area, showing the sections studied, lithofacies assemblages, and volcanic tuff beds with radiometric dates in Ma. Numbers beside logs indicate stratigraphic levels (thickness in metres).

orientation vector of 210° (n = 9). Overall, bedding is parallel or lenticular, although siltstone-shale couplets occasionally drape across shallow scours. The thicker couplets are continuous on the scale of the outcrop (tens of metres). Laterally discontinuous, parallel- to wavy-laminated, fine-grained sandstone beds (5-20 cm), with erosional bases and undulating tops, are another element of the stratification (Fig. 3d). Interpretation. Graded couplets could be the result of progressive sorting of fine-grained material by storm-generated currents and suggest alternating weak density flows with quiet-water

conditions. The presence of starved ripples is evidence of stronger flow, perhaps amplified by storms, with limited sediment supply. The sharp base of the siltstone and sandstone beds and common wave-formed structures suggest deposition on a storm-influenced shelf. Interstratified sandstone and shale Description. This assemblage comprises packages (1.0-1.5 m) of grey and yellowish-grey, fine-grained, thin-bedded sandstone

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Fig. 3. Common shallow marine to prodelta facies associations in the Vendian succession in the SE White Sea area, (a) Thin beds of volcanic ash of the laminated shale facies association. Verkhovka Fm, Agma Section. (1 m scale bar) (b) Lower bedding plane view of a wave-rippled sandstone bed with swing marks. Lamtsa Fm, Lamtsa Section. (0.1 m scale bar) (c) Alternating shale and siltstone facies association with thicker siltstone laminae showing cross-bedding and starved wave ripples. Zimnegory Fm, Winter Mts Section. (0.1 m scale bar) (d) Wave-rippled sandstone facies of the alternating shale and siltstone facies association. Verkhovka Fm, Suzma Section. (1 m scale bar) (e) Ball-and-pillow structures. Verkhovka Fm, Solza Section. (1 m scale bar) (f) Isolated putter casts of the mterstratified sandstone and shale facies association. Zimnegory Fm, Winter Mts Section. (1 m scale bar) (g) Cyclothem-like sequences of the mterstratified sandstone and shale facies association. Yorga Fm, Winter Mts Section. (5 m scale bar) (h) Overflow channel casts of the channelized sandstone facies association. Zimnegory Fm, Winter Mts Section. (5 m scale bar)

SEDIMENTATION IN THE TIMAN FORELAND

units (0.1-0.5 m) interbedded with intervals (0.3-0.5 m up to 2 m thick) of graded siltstone-shale couplets. The sandstone beds have sharp bases, fine upwards, and have rippled tops. Thinner beds tend to consist of fine horizontal laminations. Some of them contain gently curved lamina-sets that probably represent hummocky stratification. However, thicker units exhibit finingupward textures, convoluted laminations, amalgamation surfaces, ball-and-pillow structure (Fig. 3e), isolated shale clasts, and waveripple laminations. In addition, the intervals of graded siltstone shale couplets host isolated sandstone gutter casts (0.3-0.4m thick, up to 1 m wide) (Fig. 3f). The gutter casts are uniformly aligned (240°-60°, n = 21), although locally their traces are sinuous in plan, and they meander gently with < 1 m amplitudes and 2.0-2.5 m wavelengths. The gutter casts are presumed to extend beyond the limit of the outcrop, but their steep wedgeshaped terminations have been observed on several occasions. Deformed lamination in the host sediment suggests that formation of the gutter casts was associated with scouring of the liquefied substrate. Therefore, the gutters must have originated under sand-saturated flows and become immediately cast by coarser sediment, as is evident from their steep and overhanging sides. Interpretation. The fades assemblage contains diagnostic structures of modern upper-shelf prodelta deposits (Swift et al. 1991). The thin-bedded sandstone packages could be the consequence of the oscillatory and rapid progradation of the sandy shoreface, whereas the amalgamated units with soft-sediment deformation features suggest episodes of a high rate of sediment supply, causing local destabilization of slopes and promoting slumping. Although the gutter casts may have formed under fluvial conditions, they may also have formed under storm wave action.

Channelized sandstone Description. This assemblage comprises greenish-grey, finegrained, thin-bedded, cross-bedded, and planar-laminated sandstones hosted by purple-grey, massive to parallel-laminated siltstone. The sandstones are channel fills that appear in lenses (0.3-1.8 m thick and up to 10m wide) and laterally discontinuous packages (up to 14 m thick) with convex downward bases and nearly flat upper surfaces (Fig. 3h). The cross-bedded channelized sandstones tend to consist of multistoried cross-laminations that have a mean orientation vector of 230° (n = 25), which is in accord with the channel alignment (250°-70°, n = 22). Bases of the lenses are ornamented with gutter and scour casts (less than 0.4 m deep). Some large channel casts exhibit several 'waterways' at their base implying anastomosing of the channels. These lenses, in turn, occur on surfaces that are also characterized by numerous isolated sand-filled scour casts and wave ripples (Fig. 4c). Individual sandstone beds (0.1-0.4m) are another element of stratification. They record strong current influence and exhibit a variety of sole marks, including load casts, chevron, drag, and flute marks, current crescents, swing marks, and gutter casts with a mean alignment vector of 245° (n = 12) (Fig. 4a, b). Also common are small isolated shale clasts found as lag deposits. A final, but not least common siltstone facies comprises intervals of massive to parallel-laminated coarse siltstone, centimetres to a few metres thick, with no apparent grading, although with bedding-parallel partings. Interpretation. Individual sandstone beds are interpreted as the product of single flood events, whereas sandstone lenses and packages could represent sand-filled overflow and distributary channels. The strong correspondence between the orientations of erosional markings, current-formed parting lineation, and multistoried cross-bedding suggests that the channels may be derived

41

from bed erosion by marine hyperpycnal inflows that flushed out distributary systems during river inundation events (Normark & Piper 1991; Swift et al. 1991). Hence, the stratification is regarded as fluviomarine, and the beds are primarily inundites deposited in a distributary-mouth bar setting. Trough-cross-bedded sandstone Description. This assemblage is dominated by thick (0.5-3.5 m) and wide (several metres) lenses of pink, medium- to coarsegrained sandstone with multistoried, medium- and large-scale trough cross-bedding (set thicknesses 0.1-1.0m) (Fig. 4d-f). Beds are characterized by upward-decreasing grain size and scale of cross-bedding, as well as intraformational recumbent folds. Lower bedding contacts are commonly undulatory scours filled with massive sandstone, locally with lenses of maroon shale clasts; clasts as large as 0.2m have been observed (Fig. 4g). Isolated shale clasts are also present along foresets. Upper divisions consist of fine sandstone with planar- and waveripple laminations. The sandstone lenses host siderite concretions. Measurements of the high-angle, trough cross-strata sets reveal the same uniform pattern of palaeocurrent trend (240°, n = 5), which correlates with a current-formed, parting-lineation trend (215°-35°, n = 8). In a vertical succession, these sandstones alternate with grey intervals of interbedded sandstone and shale characterized by abundant oscillation wave ripples and evaporite pseudomorphs. Interpretation. This facies assemblage represents filling of broad and shallow distributary channels by migrating sinuous-crested subaqueous dunes in a braid-delta plain setting. The intervals of interbedded sandstone and shale evidently represent interdistributary areas of delta plains. In this respect, the shale clasts indicate scouring of mud-covered overbank areas cannibalized during fluvial channel migration. The variegated colour of sediments, presence of siderite nodules, and occasional evaporite pseudomorphs in this facies indicate strongly fluctuating or variable salinity. Depositional sequences The Vendian succession in the White Sea area can be divided into three depositional sequences: the Lamtsa-Verkhovka, Zimnegory, and Yorga. The bounding surfaces are recognized in the field by evidence of valley incision and an abrupt superposition of disparate facies (Grazhdankin 2003). The Lamtsa-Verkhovka sequence comprises low-energy shallow-marine shales that interfinger with intervals of interstratified sandstone and shale exhibiting a clastic wedge filling pattern sourced from the craton (Fig. 2). The cycles present are parasequences, since they coarsen upwards and rest on simple flooding surfaces, and can be traced with certainty in the subsurface for many tens of kilometres (Grazhdankin 2003). Each parasequence consists of a transgressive laminated shale, a condensed section with carbonate interbeds representing peak transgressive conditions, and a regressive package of graded siltstone-shale couplets and sandstone storm beds. Drill-core sections on the Onega Peninsula demonstrate that the intervals of alternating shale and siltstone merge westward with the intervals of interbedded sandstone and shale, which in turn thicken and coalesce in the same direction, towards the Baltic Shield. The channelized sandstone facies wedged in-between the wave-dominated shelf facies in the middle part of the Verkhovka Formation probably record the initial fluvial influence (Grazhdankin & Bronnikov 1997) (Fig. 2). The strong correspondence between palaeocurrent directions here and the trough cross-bedding in the overlying braid-delta facies suggests that channels may have formed by erosive currents that

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D. GRAZHDANKIN

Fig. 4. Common distributary-mouth bar to braid-delta plain facies associations in the Vendian succession in the SE White Sea area, (a) Channelized sandstone facies association. Plano-convex and biconvex cross-bedded lenses of sandstone embedded in shale represent load casted ripple marks. Zimnegory Fm, Winter Mts Section. (0.1 m scale bar) (b) Erosional scour casts at the base of a planar-laminated bed of the channelized sandstone facies association. Zimnegory Fm, Winter Mts Section. (0.1 m scale bar) (c) Wave ripples of the channelized sandstone facies association. Yorga Fm, Winter Mts Section. (0.5 m scale bar) (d) Trough-cross-bedded facies with medium scale of multistoried trough cross-bedding. Yorga Fm, Winter Mts Section. (0.1 m scale bar) (e) Distributary channel cast of the trough-cross-bedded facies association. Yorga Fm, Zolotitsa Section. (10 m of visible thickness) (f) Trough-cross-bedded facies with large scale of multistoried trough cross-bedding. Yorga Fm, Zolotitsa Section. (5 m scale bar) (g) Shale clasts of the trough-cross-bedded facies. Yorga Fm, Winter Mts Section. (0.1 m scale bar) (h) Trough-cross-bedded facies association. Numerous distributary channel casts are seen. Padun Fm (terminal Neoproterozoic?), Zolotitsa Section. (20 m of visible thickness)

SEDIMENTATION IN THE TIMAN FORELAND

were strengthened by flood discharge from the prograding delta in the NE. The base of the Zimnegory sequence is an erosional unconformity that is revealed in drill-core sections of the Winter Mountains. The erosional relief is at least 100 m and incises the interstratified sandstone and shale of the Verkhovka Formation. The valley fill begins with thin (0.1-0.2m) lenticular conglomeratic beds followed by quartzitic sandstone of high mineralogical and textural maturity, with lenticular/flaser laminations and channel casts (Fig. 2). This is overlain by laminated shales that prograde into laminated siltstones with packages of channelized cross-bedded and planar-laminated sandstone deposits (Figs 3h & 4a). Measured trough-cross-set azimuths are unimodal and directed SW, indicating strong current influence with a fluvial origin. The cross-stratified sandstone is interbedded with planarlaminated, very shallow-marine, sheet sandstone, which supports a fluviomarine interpretation (Fig. 4b). Within the valley, the initially fluviomarine sandstone and shale deepen upsection into marine, storm-influenced alternating siltstone and shale (Fig. 2). Apparently, the landward migration of a coastline occurred within the valley, resulting in the accumulation of transgressive deposit. The interfingering prodelta sandstones and shales with gutter casts provide evidence for continued fluvial influence (Fig. 3f). The isolated gutter casts are more common in the shallowest facies and exceed the typical size for storm erosional features. Their isolated occurrence, without connection to continuous sandstone beds, may be the result of partial discharge from bypassing sediment-laden flows. Having discharged some material, the flows may have become more buoyant, ascended as plumes, and transported the rest of the suspended material into more distal settings. Strong inertiadriven flows of fluvial nature are likely to have been responsible for origin of the gutter casts. The top of the Zimnegory sequence is truncated and marked by incision with erosional relief up to 25 m. The incised valley fill at the base of the succeeding Yorga sequence is well exposed in coastal cliffs of the Winter Mountains (Fig. 2). Further north, near the mouth of the Zolotitsa River, there is an additional incised valley. Where the incision has not occurred, such as in the section along the Torozhma River (Fig. 2), the base of the Yorga sequence is lined with packages of quartzitic sandstone of high mineralogical and textural maturity. The lithofacies in the valley fills are organized into 1.3-3.2 m thick, fining-upward cyclothem-like sequences (Fig. 3g). Each cyclothem begins with channel casts or thick (0.5-0.6m) packages of laterally discontinuous thin-bedded sandstones often exhibiting soft-sediment deformations. Then follows a package (0.4-0.7 m) of interbedded thinner wave-rippled sandstones, progressively thinning towards the top of the package. The upper part of each cyclothem-like sequence is represented by an interval of alternating siltstone and shale. As the sequences are traced up-section and in the source ward direction (NE, based on palaeocurrent data), sandstone beds thicken, mud interbeds disappear, and thick (3-4 m) amalgamated units with soft-sediment deformation features develop. The cyclothem-like sequences of the valley fill are of parasequence scale and formed by a coalescing series of wave-dominated deltas prograding seaward along a straight prodelta front. Hence, the depositional system of the lower Yorga Formation has all the attributes of a highstand system tract. The valley fill is overlain by a thin (17 m) wedge of distributarymouth bar channelized sandstone. The rest of the Yorga sequence consists of trough-cross-bedded sandstone deposited in a braiddelta plain setting (Fig. 2). Transitions from prodelta facies to distributary-mouth bar facies, and from distributary-mouth bar facies to braid-delta plain facies are equally sharp. The cause of the abrupt shift is uncertain, but is consistent with the interpretation of an erosional unconformity, with incision of valleys, at the base of the Yorga sequence. The vertical persistence of facies (> 100 m) and the tight clustering of palaeocurrent trends indicate

43

that the Yorga sequence represents a large alluvial apron with a uniform SW palaeoslope draining directly out of an orogenic hinterland.

Discussion Ever since the first lithological descriptions of the Vendian succession, it has been interpreted as sub-storm wave-base pelagic deposit, laid down in an epeiric embayment under transgressive and highstand conditions in response to deglaciation and the opening of a new oceanic basin (Sokolov 1952; Keller 1963; Aksenov 1985; Bessonova et al 1980; Nikishin et al 1996). On the other hand, the configuration of the Vendian Basin conforms to the pattern of an early Neoproterozoic rift system in the craton basement. This has been interpreted as direct evidence of tectonic control, and thereby offers a basis for interpretation of the Vendian setting in connection to an early post-rift thermal subsidence phase, when variations in the amount of subsidence would most strongly reflect variations in the magnitude of overlying extension (Stankovsky et al 1985; Kostyuchenko et al 1999). However, Shatsky (1952) suggested synorogenic deposition of the Vendian succession in relation to the Baikalian Orogeny. The Vendian succession of the White Sea area provides a test for these competing hypotheses. The lithofacies succession in the lower part of the LamtsaVerkhovka sequence is here interpreted as an expression of the oscillatory progradation of the lower shoreface across the mudtype, low-energy inner shelf in a shallow epeiric sea. Muddy sediment may have originated in nearshore mud streams, as a product of sediment resuspension, and have escaped across the shelf during storms as the frontal surface became disrupted by storm currents (McCave 1972; Holmes 1982; Sahl et al 1987). Graded siltstone-shale alternations were probably deposited from suspension in a transitional zone between the coastal sand belt and an offshore mud belt. Subordinate, thin-graded storm beds of fine sandstone, suggesting intermittent storm wave deposition, reflect an increase in wave resuspension and sediment bypass in shallower settings as the depositional system prograded. The epeiric sea facies constitutes the lower one third of the White Sea area section. The bulk of the Vendian succession, however, is dominated by prodelta, distributary-mouth bar, and braid-delta plain depositional systems, where facies distribution remains obscure on account of limited lateral observations. Nevertheless, all three demonstrate a remarkably tight clustering of palaeocurrent trends. The wide extent of flood stratification (inundites) in the upper two thirds of the coarsening upward Vendian succession suggests that the main mechanism of clastic sediment input was provided by quasi-steady, inertia-driven hyperpycnal plumes (Grazhdankin 2003). Interfingering between the prodelta and distributary-mouth bar facies is indicated, with excellent resolution, on the coastal cliff face of the Winter Mountains. In the case of the SE White Sea area, it is clearly a river mouth setting; therefore, the cyclicity of the prodelta interstratified sandstone and shale may be nearly autocyclic in nature, forming in response to channel avulsion and shifts in the distributary-mouth bar location during progradation of the delta. The bulk of the Vendian succession in the White Sea area was deposited in delta-related marine environments, with deltaic coastal landforms located NE of the present day line of outcrops. Thus the Vendian sedimentary environments in the White Sea area represent distal settings bordered by land to the NE. The Mezen Basin extends southwestwards as a vast epicratonic inland tongue, to the Moscow Basin (Fig. 1). Palaeocurrent orientations in the White Sea area suggest that the Moscow Basin is located along the distal side of the Vendian sediment dispersal system. The Vendian succession of the Moscow Basin, as defined by Sokolov (1952), comprises 250-800 m of

44

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tuffaceous-siliciclastic sediments known exclusively from boreholes. These Vendian rocks were divided into two formations, the Redkino and Kotlin (Kirsanov 1968c; Solontsov et al 1970), which were subsequently upgraded into the Redkino and Kotlin stratohorizons to accommodate various correlative strata in the sedimentary cover (Fig. 5) (Aksenov et al. 1978). The Redkino stratohorizon in the Moscow Basin comprises laminated shales of the Gavrilovyam and Nepeitsino Formations deposited in low-energy epeiric sea settings, whereas the Kotlin succession consists of variegated shales and trough-cross-bedded sandstones of the Lubim Formation deposited in coastal plain settings with strongly fluctuating or variable salinity (Fig. 5) (Bessonova et al. 1980; Aksenov 1985). The Kotlin stratohorizon also includes the Reshma and Padun Formations (Figs 4h & 5); however their stratigraphic age is uncertain. The Redkino and Kotlin stratohorizons in the central and western parts of the Moscow Basin are separated by a disconformity lined with quartzitic sandstones (10-20m thick) of high mineralogical and textural maturity (Kirsanov 19680). These are correlated with the Makariev Formation (<200m thick) in the northeastern part of the basin (Fig. 5) (Kirsanov 19680; Solontsov et al. 1970; Aksenov 1985; Kuzmenko & Burzin 1996). Early in the research, a relative sea-level fall and major erosion was invoked to explain the apparent loss of 200 m of sediment in the central and western parts of the Moscow Basin (Fig. 5); however, the lack of extensive conglomerate deposits and the absence of subaerial erosion surfaces impose serious limitations on the evidence for a hypothetical pre-Kotlin sea-level drop. On the other hand, the Vendian succession in the northeastern part of the Moscow Basin comprises distributary-mouth bar and prodelta facies implying depositional settings similar to the incised Zimnegory sequence in the SE White Sea area. Lateral discontinuity of the Makariev Formation in drill-cores suggests that it too fills incised valleys. It is proposed here that during late Redkino time most of the terrestrial clastic material supplied from the NE was trapped in incised valleys and mouth bar depositional systems developed along the northeastern periphery of the Vendian setting. The

Fig. 5. Vendian lithostratigraphy and comparative thickness of the formations in the Moscow Basin (Kuzmenko & Burzin 1996) and correlation with the White Sea area. See Fig. 1 for geographical reference.

interior Moscow Basin thereby experienced an inadequate supply of elastics (sediment starvation) and was instead subjected to submarine shelf erosion by waves and currents that resulted in the condensed interval of quartzitic sandstone. In some cases, data on the previous structural history can be used to resolve geological problems, particularly those concerning the nature of the subsidence, and the Vendian succession of the East European Craton is thought to be of no exception (Stankovsky et al. 1981, 1985; Kostyuchenko et al 1999). However, the Neoproterozoic succession gives no evidence of post-rift thermal subsidence as a possible mechanism for creating accommodation space during the Vendian. What is evident, is a significant gap between the end of early Neoproterozoic rifting and the beginning of the late Neoproterozoic Vendian subsidence (Nikishin et al. 1996). Unless the available temporal constraints for the graben deposits are erroneous, other factors must be invoked to explain the Vendian subsidence. The Vendian succession of the Mezen Basin extends along the Timan Orogen into the Ural Basin (Aksenov 1967; Kirsanov 1968&), and correlates with the Asha Group of the western flank of the Urals (Becker 1990). Here too, a marine-alluvial transition is recognized within the c. 1700 m thick siliciclastic succession of the Asha Group (Keller 1963). The Basa Formation (up to 1050 m thick) in the lower part of the Group consists of interbedded shales, siltstones and sandstones that represent shallow-marine shelf facies. The succeeding Kuk-Karauk and Zigan Formations (up to 700 m thick) consist of poorly-sorted sandstones, interstratified with shales and conglomerates with cobble- and boulder-size clasts. This succession represents a gravelly braided plain depositional setting and is interpreted as a molasse-type deposit (Becker 1968, 1988). During Asha time, detritus continued to be eroded from the craton NW of the basin, whereas the second, more important dispersal system derived sediment from the metamorphic complex of Beloretsk, which had already been emplaced and exhumed along the northeastern craton margin (Fig. 1; Becker 1968; Maslov et al. 1997; Willner et al. 2001). The relationship between the Beloretsk Terrane and the Timan Orogen remains to be seen (Maslov et al. 1997; Glasmacher et al. 2001). The Asha Group in the South Urals is unconformably overlain by late Ordovician conglomerates and fossiliferous grits (e.g. Maslov et al. 1997). In conclusion, the Neoproterozoic sedimentation of the northeastern part of the East European Craton is most simply explained as deposition in a distal foreland setting in front of the TimanPechora collisional zone. The Vendian Basin architecture suggests differential subsidence, with higher rates of accommodation space growth along the proximal northeastern side bordered by the Timan-Ural Orogen and lower subsidence rates in the distal Moscow Basin. The lack of apparent structural complexity and the simple sediment geometries suggest that the Vendian succession represents the later erosion products of post-collisional exhumation of the Timanide hinterland. The timing of the Timan Orogenesis is constrained by the late- to post-tectonic granite ages of c. 550560 Ma from the basement of the hinterland (Gee et al. 2000), as well as by the early burial diagenesis ages of c. 560 Ma from the Neoproterozoic shales (Gorokhov et al. 2001; Roberts & Siedlecka 2002). This is compatible with the U-Pb zircon dates of 558 ± 1 Ma and 555.3 + 0.3 Ma from the volcanic tuffs associated with the Vendian succession (Grazhdankin 2003). Considerable work must be done before we are able to map the three-dimensional distribution of Vendian lithofacies and bounding surfaces accurately, and thereby to document the geological history of the Neoproterozoic sedimentation in the Timan foreland more precisely. In addition, analyses of the mineralogical composition of sandstones and secular variations in provenance may resolve the evolution of the source terrains and the style of sedimentation. The history of ancient orogenic belts is inferred mainly from the sedimentary records within the associated foreland basins. The Vendian succession of the Mezen Basin

SEDIMENTATION IN THE TIMAN FORELAND

provides an excellent and sensitive tool that needs to be calibrated, so that we can detect and measure the dynamics of the Timanian Orogenesis. This research is based on observations and collections made in the field with the help of A. A. Bronnikov and the Novodvinsk Expedition (Arkhangelsk). I thank the Natural Environment Research Council (United Kingdom) for funding this research (grants NER/B/S/2000/00316 and NER/A/2001/01049 to N. J. Butterfield). The fieldwork was also funded by the Alfred Toepfer Stiftung F.V.S. (Germany), the 1997 Nauka/Interperiodica International Academic Publishing Company Prize, and the Russian Foundation for Basic Research (grants to M. A. Fedonkin and B. S. Sokolov). P. F. Friend, A. Siedlecka, V. Pease, and an anonymous referee critically read a first edition of the submitted manuscript. This is Cambridge Earth Sciences contribution ES.7667.

References AKSENOV, E. M. 1967. [Vendian Complex in the eastern part of the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 9, 81-91 [in Russian]. AKSENOV, E. M. 1985. [Vendian of the East European Platform]. In: SOKOLOV, B. S. & FEDONKIN, M. A. (eds) [The Vendian System, 2 Regional Geology}. Nauka, Moscow, 3-34 [in Russian; English translation published in 1990 by Springer-Verlag, Berlin, 1-37]. AKSENOV, E. M., KELLER, B. M., SOKOLOV, B. S., SOLONTSOV, L. F. & SHULGA, P. L. 1978. [Generalized Upper Precambrian stratigraphy of the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 12, 17-34 [in Russian]. BECKER, Y. R. 1968. [Late Precambrian molasse in the South Urals]. Nedra, Leningrad [in Russian]. BECKER, Y. R. 1988. [Precambrian molasse]. Nedra, Leningrad [in Russian]. BECKER, Y. R. 1990. Vendian of the Urals. In: Sokolov, B. S. & Fedonkin, M. A. (eds) The Vendian System, 2 Regional Geology. SpringerVerlag, Berlin, 88-101. BESSONOVA, V. Y., VELIKANOV, V. A., KELLER, B. M. & KIRSANOV, V. V. 1980. [Valdai Series]. In: ROZANOV, A. Y. & LYDKA, K. (eds) [Palaeogeography and lithology of the Vendian and Cambrian of the western East European Platform]. Nauka, Moscow, 15-24 [in Russian; English translation published in 1987 by Wydawnictwa Geologiczne, Warsaw, 20-29]. DEDEEV, V. A. & KELLER, B. M. (eds) 1986. [Upper Precambrian in the northern European part of the USSR: explanatory notes]. Academy of Sciences, Syktyvkar [in Russian]. GEE, D. G., BELIAKOVA, L., PEASE, V., LARIONOV, A. & DOVSHIKOVA, L. 2000. New, single zircon (Pb-evaporation) ages from Vendian intrusions in the basement beneath the Pechora Basin, northeastern Baltica. Polarforschung, 68, 161-170. GLASMACHER, U. A., BAUER, W., GIESE, U., REYNOLDS, P., KOBER, B., PUCHKOV, V., STROINK, L., ALEKSEYEV, A. & WILLNER, A. P. 2001. The metamorphic complex of Beloretzk, SW Urals, Russia—a terrane with a polyphase Meso- to Neoproterozoic thermo-dynamic evolution. Precambrian Research, 110, 185-213. GOROKHOV, I. M., SIEDLECKA, A., ROBERTS, D., MELNIKOV, N. N. & TURCHENKO, T. L. 2001. Rb-Sr dating of diagenetic illite in Neoproterozoic shales, Varanger Peninsula, northern Norway. Geological Magazine, 138, 541-562. GRAZHDANKIN, D. V. 2003. [Structure and depositional environment of the Vendian Complex in the Southeast White Sea area]. Stratigrafiia. Geologicheskaia korrelyatsiia, 11(4), 3-24 [in Russian; English translation: Stratigraphy and Geological Correlation, 11, 313-331]. GRAZHDANKIN, D. V. 2004. Patterns of distribution in the Ediacaran biotas: facies versus biogeography and evolution. Paleobiology, 30, 203-221. GRAZHDANKIN, D. V. & BRONNIKOV, A. A. 1997. [A new fossil locality of the late Vendian soft-bodied organisms on the Onega peninsula]. Doklady Akademii nauk, 357(6), 792-796 [in Russian; English translation: Transactions (Doklady) of the Russian Academy of Sciences, Earth Sciences Sections, 357A, 1311-1315]. HOLMES, C. W. 1982. Geochemical indices of fine sediment transport, northwest Gulf of Mexico. Journal of Sedimentary Petrology, 52, 307-321.

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IGOLKINA, N. S. 1959. [Tentative recognition of the lower Cambrian Baltic Complex in the north of the Russian Platform]. Informatsionnyi sbornik VSEGEI, 11, 17-23 [in Russian]. IVANOV, S. N. & RUSIN, A. I. 2000. Late Vendian tectonic evolution of the Urals. Geotectonics, 34, 187-197. JAKOBSON, K. E. & NIKULIN, S. N. 1985. [Vendian deposits on the Vetreny Belt]. Sovetskaia Geologia, 4, 80-83 [in Russian]. JAKOBSON, K. E., KUZNETSOVA, M. Y., STANKOVSKY, A. F., GRIB, V. P., MEDVEDEV, V. A. & TRETIACHENKO, V. V. 1991. [Riphean sequence of the White Sea Winter Coast]. Sovetskaia geologiia, 11, 44-48 [in Russian]. KELLER, B. M. 1963. [General structural aspects of the Upper Precambrian: palaeogeography and geological history]. In: KELLER, B. M. (ed.) [Stratigraphy of the USSR. Upper Precambrian]. Gosnauchtekhizdat, Moscow, 615-631 [in Russian]. KIRSANOV, V. V. 19680. [New data on Precambrian stratigraphy in the central regions of the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 4, 98-113 [in Russian]. KIRSANOV, V. V. 1968£. [Stratigraphy and correlation of the Vendian Complex on the eastern margin of the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 6,86-103 [in Russian]. KIRSANOV, V. V. 1968c. [Precambrian stratigraphy in the axial zone of the Moscow Syneclise]. Doklady Akademii nauk SSSR, 178, 1160-1163 [in Russian]. KOSTYUCHENKO, S. L., EGORKIN, A. V. & SOLODILOV, L. N.

1999.

Structure and genetic mechanisms of the Precambrian rifts of the East-European Platform in Russia by integrated study of seismic, gravity, and magnetic data. Tectonophysics, 313, 9-28. KUZMENKO, J. T. & BURZIN, M. B. 1996. [Vendian stratigraphic chart for the Moscow Syneclise: explanatory notes]. Interdepartmental Stratigraphic Committee of Russia, Moscow [in Russian]. MAHOTKIN, I. L., GIBSON, S. A., THOMPSON, R. N., ZHURAVLEV, D. Z. & ZHERDEV, P. U. 2000. Late Devonian diamondiferous kimberlite and alkaline picrite (proto-kimberlite?) magmatism in the Arkhangelsk Region, NW Russia. Journal of Petrology, 41, 201-227. MARTIN, M. W., GRAZHDANKIN, D. V., BOWRING, S. A., EVANS, D. A. D., FEDONKIN, M. A. & KIRSCHVINK, J. L. 2000. Age of Neoproterozoic bilaterian body and trace fossils, White Sea, Russia: implications for metazoan evolution. Science, 288, 841-845. MASLOV, A. V., ERDTMANN, B.-D., IVANOV, K. S., IVANOV, S. N. & KRUPENIN, M. T. 1997. The main tectonic events, depositional history, and the palaeogeography of the southern Urals during the Riphean-early Palaeozoic. Tectonophysics, 276, 313-335. McCAVE, I. N. 1972. Transport and escape of fine-grained sediment from shelf areas. In: SWIFT, D. J. P., DUANE, D. B. & PILKEY, O. H. (eds) Shelf Sediment Transport: Process and Pattern. Dowden, Hutchinson & Ross, Stroudsburg, Pa, 225-248. NlKISHIN, A. M., ZlEGLER, P. A., STEPHENSON, R. A., CLOETINGH, S. A.

P. L., FURNE, A. V., FOKIN, P. A., ERSHOV, A. V., BOLOTOV, S. N., KOROTAEV, M. V., ALEKSEEV, A. S., GORBACHEV, V. I., SHIPILOV, E. V., LANKREIJER, A., BEMBINOVA, E. Y. & SHALIMOV, I. V. 1996. Late Precambrian to Triassic history of the East European Craton: dynamics of sedimentary basin evolution. Tectonophysics, 286, 23-63. NORMARK, W. R. & PIPER, D. J. W. 1991. Initiation processes and flow evolution of turbidity currents: implications for the depositional record. In: OsBOURNE, R. H. (ed.) From Shoreline to Abyss: Contributions in Marine Geology in Honor of Francis Parker Shepard. SEPM, Tulsa, Oklahoma, 207-230. OLOVYANISHNIKOV, V. G. 1998. [Upper Precambrian of Timan and Kanin Peninsula]. Russian Academy of Sciences, Ekaterinburg [in Russian]. OLOVYANISHNIKOV, V. G., BUSHUEV, A. S. & DOKHSAN'YANTS, E. P. 1996. [Conjugation zone structure between the Russian and Pechora plates from geological and geophysical data]. Doklady Akademii nauk, 351(1), 88-92 [in Russian; English translation: Transactions (Doklady) of the Russian Academy of Sciences, Earth Science Section, 351, 1228-1232]. OLOVYANISHNIKOV, V. G., ROBERTS, D. & SIEDLECKA, A. 2000. Tectonics and sedimentation of the Meso- to Neoproterozoic Timan-Varanger Belt along the northeastern margin of Baltica. Polarforschung, 68, 267-274.

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POPOV, L. & GORJANSKY, V. 1994. First record of Upper Cambrian from the eastern White Sea coast: new evidence from obolids (Brachiopoda). Geologiska Foreningens i Stockholm Forhandlingar, 116,31-35. 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. ROBERTS, D. & SIEDLECKA, A. 2002. Timanian orogenic deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian-Cadomian connections. Tectonophysics,352, 169-184. SAHL, L. E., MERRELL, W. J., McGRAiL, D. W. & WEBB, J. A. 1987. Transport of mud on continental shelves: evidence from the Texas shelf. Marine Geology, 76, 33-43. SIVERTSEVA, I. A. 1993. [Microbiota of the Upper Riphean Tuchkino Formation in the White Sea area]. Doklady Akademii nauk, 332, 621-623 [in Russian]. SIVERTSEVA, I. A. & STANKOVSKY, A. F. 1982. [New data on geology of the Upper Precambrian deposits in the north-western Arkhangelsk Region]. VestnikLeningradskogo universiteta, 12, 30-40 [inRussian]. SHATSKY, N. S. 1952. [On the Proterozoic/Palaeozic boundary and the Riphean deposits on the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 5, 36-49 [in Russian]. SOKOLOV, B. S. 1952. [Age constraints of the oldest sedimentary cover of the Russian Platform], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 5, 21-31 [in Russian]. SOLONTSOV, L. F., AKSENOV, E. M., ANDREEV, S. P. & POLIKARPOVA, N. T. 1970. [Valdai Group of the Moscow Basin: lithology and indexing of sandstone-siltstone beds and packages]. Trudy Geologicheskogo instituta (Kazan), 30, 324-345 [in Russian]. STANKOVSKY, A. F., SINITSYN, A. V. & SHINKAREV, N. F. 1972. [Buried traps of the Onega Peninsula in the White Sea area]. Vestnik Leningradskogo universiteta, 18, 12-20 [in Russian].

STANKOVSKY, A. F., VERICHEV, E. M., KONSTANTINOV, Y. G., SKRIPNICHENKO, V. A. & YUZHAKOV, V. M. 1977. [First record of effusive rocks among Redkino deposits in the north of the Russian Platform]. Doklady Akademii nauk SSSR, 234,661-664 [in Russian]. STANKOVSKY, A. F., VERICHEV, E. M., GRIB, V. P. & DOBEIKO, I. P. 1981. [Vendian sequence of the Southeast White Sea area], hvestiia Akademii nauk SSSR, Seriia geologicheskaia, 2, 78-87 [in Russian; English translation (1983): International Geology Review, 25, 897-905]. STANKOVSKY, A. F., VERICHEV, E. M. & DOBEIKO, I. P. 1985. [Vendian of the south-eastern White Sea area]. In: SOKOLOV, B. S. & FEDONKIN, M. A. (eds) [The Vendian System, 2 Regional Geology]. Nauka, Moscow, 67-76 [in Russian; English translation published in 1990 by Springer-Verlag, Berlin, 76-87]. SWIFT, D. J. P., PHILLIPS, S. & THORNE, J. A. 1991. Sedimentation on continental margins, IV: lithofacies and depositional systems. In: SWIFT, D. J. P., OERTEL, G. F., TILLMAN, R. W. & THORNE, J. A. (eds) Shelf sand and sandstone bodies: geometry, fades and sequence stratigraphy. Special Publications of the International Association of Sedimentologists, 14, 89-152. VERICHEV, E. M., VOLKOVA, N. A., PISKUN, L. V., SIVERTSEVA, I. A. & STANKOVSKY, A. F. 1990. [Ordovician acritarchs from northern parts of the Russian Platform], hvestiia Akademii nauk SSSR, seriia geologicheskaia, 7, 152-155 [in Russian]. WILLNER, A. P., ERMOLAEVA, T., STROINK, L., GLASMACHER, U. A., GIESE, U., PUCHKOV, V. N., KOZLOV, V. L & WALTER, R. 2001. Contrasting provenance signals in Riphean and Vendian sandstones in the SW Urals (Russia): constraints for a change from passive to active continental margin conditions in the Neoproterozoic. Precambrian Research, 110, 215-239. ZORICHEVA, A. L 1963. [Northern parts of the Russian Platform]. In: ZORICHEVA, A. I. (ed.) [Geology of the USSR, 2 Arkhangelsk and Vologda regions and Komi ASSR, Geological description]. Gosnauchtekhizdat, Moscow, 79-99 [in Russian].

Structural and tectonic development of the Timanide orogen DAVID ROBERTS1 & VSEVOLOD OLOVYANISHNIKOV2 Geological Survey of Norway, N-7491 Trondheim, Norway (e-mail: [email protected]) 2 Institute of Geology, Komi Research Centre, Ural Division RAS, Pervomayskaya 54, Syktyvkar, 167000 Russia 1

Abstract: The northeastern margin of the East European Craton developed passively in an extensional regime from late Mesoproterozoic through to the later stages of Neoproterozoic time. Along the exposed parts of the Timan-Varanger Belt, a major fault zone separates pericratonic (platformal) and basinal domains. Successions of the basinal domain can be traced beneath the Pechora Basin, via drillcore and geophysical data, to where intra-oceanic subduction systems with island arcs are inferred to have existed in the later stages of the Late Riphean. In terminal Riphean to Vendian time, inferred subduction polarity reversal resulted in a progressive telescoping, dissection and accretion of these diverse magmatosedimentary assemblages against the northeastern margin of the craton, culminating in Mid to Late Vendian, Timanian orogenesis. The Timan Range exposes SW-verging upright folds with anchizone to lower greenschist-facies cleavages. Higher-grade rocks in the Kanin-North Timan area occur in anticlinal cores and thrust slices. Isotopic dating constraints suggest that peak Timanian metamorphism occurred during the time interval 600-550 Ma.

Extending over a distance of 1800 km from the Polyudov Ridge just west of the northern Urals to the northeasternmost tip of Norway (Fig. 1), the Timanide Orogen of the Timan Range sensu stricto comprises a topographically modest, fold-and-thrust belt deriving from the SW-directed accretion of Neoproterozoic, basinal and oceanic rock assemblages onto the pericratonic, platformal successions of the northeastern margin of the East European Craton during Vendian time. Deformation of this age is also registered in the western parts of the northern and Polar Urals, where it has been termed the pre-Uralide Orogeny. Important evidence bearing on the extent of Timanian deformation and metamorphism has also been retrieved from deep drillholes penetrating the Palaeozoic and younger cover in the Pechora Basin (Getsen 1987; Belyakova & Stepanenko 1991; Bogatsky et al 1996; Gee et al 2000). Historically, recognition of this important event in the Timans goes back just over a century to the pioneering studies of Ramsay (1899) and Tschernyschev (1901) who coined the term Timanian mountain chain', summarized in a sketch map by

Reusch (1900, reproduced in Roberts & Siedlecka 2002, Fig. 3). These early workers firmly believed that similar rocks and structures could be followed northwestwards via coastal areas of northern Kola into northeasternmost parts of Norway. Although these ideas were largely disregarded for nearly a century, they had been discussed from time to time (e.g. Schatsky 1958; Siedlecka 1975) and were revived a few years ago following some of the investigations and results of Norwegian-Russian collaborative research (Roberts 1995, 1996). At the same time, the term Timan-Varanger Belt was introduced for the on-land expression of this ancient mountain chain (Olovyanishnikov et al. 1997). Although Schatsky (1935) generated some terminological confusion by introducing the name 'Baikalian' for this fold belt, he also referred to the Timan Range (Fig. 1) as the 'Timanides'. For many years, in the Russian literature, the name Baikalian held favour, but more recently Timanian' has been accepted as a more appropriate term, and with a century-old precedence (Gee & Ziegler 1996; Puchkov 1997; Gee et al. 2000; Olovyanishnikov et al. 2000). In this contribution we present a summary of the main features of structural development of the Timanide Orogen in the TimanVaranger Belt and Pechora region. We also discuss the links between the juxtaposed domains of Timanian and Caledonian structures on Varanger Peninsula, and show where this zone of juncture is likely to occur beneath the Late Palaeozoic and younger sedimentary successions of the Barents Sea.

Pre-orogenic basinal evolution

Fig. 1. Simplified map centred on the Barents Sea showing the location of the Timan Range and Pechora Basin. The horizontal-lined ornament shows the extent of the Caledonides on land and the inferred continuation northeastwards across the Barents Sea.

The Precambrian basement to the Neoproterozoic rock assemblages along this northeastern margin of the Fennoscandian Shield and East European Craton consists largely of high-grade, polymetamorphic crystalline complexes of Archaean to Palaeoproterozoic age (Getsen 1987; Gaal & Gorbatschev 1987; Dobrzhinetskaya et al. 1995). The Timan-Varanger Belt and its Riphean to Vendian stratigraphical successions trend NW-SE, mirroring the structural orientation of the principal terrane boundaries in the crystalline basement (Marker 1985). Regional studies, complemented by geophysical data, have revealed the presence of important strike-parallel faults (Fig. 2), one of which (see below) separates distinctive sedimentation domains along this passive continental margin. In the Timan-Kanin-Pechora region, the main NW-SE-trending faults (Fig. 2) may thus reflect a reactivation of known Palaeoproterozoic and older lineaments during

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 47-57. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 2. The Timan-Pechora Region showing the principal structural zones in the pre-Palaeozoic basement. The major faults are as follows: WTF, West Timan Fault; CTF, Central Timan Fault; ETF, East Timan Fault; PKF, Pechora-Kozhva Fault. The composite Timan-Izhma zone embraces the basinal slope-and-rise deposits between the CTF and the ETF, and the 'Izhma microplate' or zone of Olovyanishnikov et al (1996) between the ETF and the PKF.

the Mid to Late Riphean extension that characterized this margin of proto-Baltica. Although the stratigraphic record along the Timan-Varanger Belt is discontinuous and fragmentary, the successions that are available for study, seen in relation to the overall fault pattern, have led to a basic subdivision into two distinctive, elongate, sedimentation domains on either side of an important NW-SEtrending fault zone. These have been termed the pericratonic and basinal domains (Siedlecka & Roberts 1995; Olovyanishnikov et al. 2000), broadly corresponding to platform (or shelf) and continental slope-and-rise situations, respectively. In the Timan Range, the principal fault structure separating these domains is the Central Timan Fault (CTF) (Fig. 2). Farther to the NW along the northern coast of Kola Peninsula, the same basin-bordering structure is known as the Sredni-Rybachi Fault Zone (SRFZ) (Roberts 1995). This extends onto the Varanger Peninsula in Norway where it is termed the Trollfjorden-Komagelva Fault Zone (TKFZ) (Siedlecka & Siedlecki 1967; Karpuz et al 1993). In addition to these main strike-parallel faults, in the TimanKanin region there are also several transverse faults. These are inferred to have caused segmentation of the developing basin (Roberts & Siedlecka 2002), and served as loci for later magmatic activity. Biostratigraphic evidence, based on microfossils and stromatolites, has indicated that in the pericratonic domain, fluvial and shallow-marine sedimentation commenced in the Late Riphean (Getsen & Pihova 1977; Lyubtsov et al. 1989; Olovyanishnikov 1998). A distinct basal unconformity is exposed in some areas. Although no indisputable Vendian-age rocks, including tillites, are known from the successions in the Timan Range (Getsen &

Pihova 1977; Samuelsson 1997; Olovyanishnikov 1998), a richly fossiliferous Vendian stratigraphy is recorded in the adjacent Mezen Basin SE of the White Sea (Fig. 1). This succession is considered by Grazhdankin (2004) to have been deposited in a foreland basin that developed in front of the Timanide orogen. On Varanger Peninsula, sedimentation continued into the Vendian, and Varangerian tillites are widely exposed (Edwards 1984). Shortly after the Varangerian, source areas for erosion and sediment dispersal changed abruptly from southerly to northeasterly (Banks et al. 1971). This is a feature that Gorokhov et al. (2001) relate to exhumation and erosion of the rising Timanide orogen, with Mid to Late Vendian sediment accumulation in a small foreland basin ahead of the deformation front. In the Timans, a stromatolitic carbonate formation formed along the outer margin of the platform (Fig. 3), at the fault-controlled (CTF) break of slope (Getsen 1970); these are termed carbostromes in Russian literature. Comparable Late Riphean, stromatolitic carbonates have been reported from the KolaVaranger region. Detailed lithostratigraphies from different areas in the Timan-Varanger Belt with discussion on sedimentary facies and environments of deposition are contained in papers by Getsen (1968, 1975), Lyubtsov et al. (1978, 1989), Siedlecka et al. (19950), Olovyanishnikov (1998), Olovyanishnikov et al. (2000), as well as in another contribution in this volume (Siedlecka et al. 2004). In the basinal domain, sedimentation began in latest Mesoproterozoic (late Mid Riphean) time in the Timan Range, but started later, in the Late Riphean, in the Kola-Varanger region (Olovyanishnikov et al. 2000). A characteristic feature of lithostratigraphic successions in northwestern parts of the Timan-Varanger Belt is the widespread presence of submarine turbidite systems (Siedlecka 1972; Pickering 1983, 1985; Siedlecka et al. I995b). The deep-water depositional systems of the Timan-Kanin region, on the other hand, show contrasting sedimentary facies to those on Rybachi and Varanger (Roberts et al. 2004). Furthermore, a sporadic volcanic component is present, as tuffs or rare basaltic lavas, in some of the basinal successions in the Timan Range (Olovyanishnikov 2001). Such rocks have not been recorded in the Kola-Varanger region. Much of the deeper marine parts of the basinal domain are now concealed beneath the Palaeozoic and younger rocks of the Pechora Basin (Fig. 1), but there is evidence from many deep drillcores pointing to the presence of both primitive and evolved arc volcanites (Fig. 3) during the Late Riphean to Vendian period (Getsen 1987, 1991; Olovyanishnikov et al. 1995; Gee et al. 2000; Dovzhikova et al. 2004), as well as diverse Mid to Late Riphean, bimodal plutonic rocks (Getsen 1991; Bogatsky et al. 1996). Ophiolitic associations have been inferred, but not proven, except in the Polar Urals (Scarrow et al. 2001). An additional feature of this sub-Pechora, deep-basinal regime, according to Getsen (1991), is the presence of small microcontinental blocks and slivers (e.g. Khoreyver, Kolguev, forming part of the Bolshezemelskaya terrane; Fig. 2) that are considered by this same author to have drifted away from mainland protoBaltica in Mid Riphean times during the main stage of crustal extension and rifting. Taking the Timan-Varanger Belt and its northeastward, hinterland extension beneath the Pechora Basin as a whole, the evidence suggests that a Mid to early Late Riphean passive margin was ultimately converted to an active one in Late to terminal Riphean time as a new volcanic arc, or arcs, developed in response to an inferred subduction reversal from seaward- to landward-facing (Scarrow et al. 2001; Roberts & Siedlecka 2002) in an intra-oceanic environment (the Timan-Urals ocean of Bogatsky et al. 1996; also called the pre-Uralian or Palaeo-Asian ocean; Zonenshain et al. 1990). This change of crustal stress regime, from extensional or transtensional to compressional, heralded the initiation of the Timanian orogenic cycle that reached its tectonothermal peak in Mid to Late Vendian time.

STRUCTURE OF THE TIMANIDE OROGEN

49

Fig. 3. Palaeogeographic/palaeotectonic sketch map showing the inferred distribution of sedimentary and volcanogenic belts and basement blocks along the northeastern margin of the East European Craton in the Late Riphean; modified from Getsen (1991). CTF, Central Timan Fault; PKF, PechoraKozhva Fault; Kol. T., Kolguev terrane; Khor. T., Khoreyver terrane; Nov. T., Novozemelskaya terrane.

Timanian orogenesis Extent and character Throughout the exposed Timan-Varanger Belt, structural transformation of the late Mesoproterozoic to Neoproterozoic terrigenous successions during the processes of collision and accretion that are ascribed to Timanian orogenesis was comparatively simple. In almost all of the basin-domainal areas along this margin of Baltica, there are NW-SE-trending, upright to SW-verging, mesoscopic and large-scale folds. These tight to open folds invariably carry a steep, NE-dipping, penetrative axialplanar cleavage (Figs 4 & 5) denoting that metamorphic grade was generally restricted to anchizone to lowest greenschist facies. Locally, folding was more intense (Fig. 4b). Only in parts of the Kanin Peninsula, especially in the SE near Cape Mikulkin, do amphibolite-facies mineral parageneses occur (Getsen 1987; Lorenz et al 2004). Within the pericratonic domain, Timanian structures are even simpler, in the form of gentle folds and warps (also called 'arches' in Russian literature), generally lacking cleavage (except close to major faults); however, a compactional fabric of diagenesis grade is ubiquitous. Structural details from different areas along the Timan-Varanger Belt are provided in Malkov & Puchkov (1964), Getsen (1987), Roberts (1995, 1996), Roberts & Karpuz (1995), Mitrofanov et al (1999) and Roberts & Siedlecka (2002). The basin-bordering faults and other belt-parallel structures that had functioned as normal faults during the pre-orogenic

extensional regime were reactivated as oblique- to re verse-slip faults during the Timanian orogeny. SW-directed thrusting has been reported, for example, along the Central and East Timan Faults. Seismic data from the central Timan Range indicate that higher-grade basinal facies rocks have been thrust by some tens of kilometres along the CTF (Olovyanishnikov et al. 1995). Our recent collaborative work on Kanin, in 2000, in the hangingwall of the ETF also supports the notion that the Timanides are a foreland fold-and-thrust belt. The geothermobarometry studies of Lorenz et al. (2004) led to the inference that thrust-stacking is the most likely explanation for the observed relationships in the amphibolite-facies, Mikulkinskaya Group rocks of the Mikulkin Antiform in southeastern Kanin. Likewise, in coastal tracts of northwestern Kanin there are slivers of epidoteamphibolite facies, garnetiferous schists and amphibolites in tectonic contact with anchizone to lower greenschist-facies turbidites in the adjacent footwall of a strike-parallel steep fault, pointing to SW-directed thrusting (unpublished data). These amphibolite-facies rocks may be upthrust correlatives of the Mikulkinskaya Group, or higher-grade variants of overlying formations in the Kanin Kamen Supergroup (Roberts et al 2004). Possible conjunctive (i.e. layer-parallel) tectonic contacts have also been reported, but not confirmed, from within other parts of the generally low-grade, lithostratigraphical successions on the Kanin Peninsula (Olovyanishnikov 2001). In the Rybachi-Sredni-Varanger region, SW-directed thrust translations have been reported from the basin-bordering TKFZ (Karpuz et al 1993) and SRFZ (Roberts 1995). On the northern

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Fig. 4. (a) Graded and ripple-drift cross-laminated silty sandstones and pelites with a penetrative spaced cleavage axial planar to an open fold. Rumyanichnaya Formation, coastal section of North Timan, looking NNW. (b) Strongly deformed and schistose Rumyanichnaya Formation, cut by aplite-syenite dykes; near Cape Bolshoy Rumyanichny, North Timan; looking NW.

side of the TKFZ, mesoscopic high- and low-angle reverse faults are common, indicating SW-directed thrusting (Herrevold 1993; Gjelsvik 1998). Elsewhere in this particular region, a steeply NE-dipping, slaty to spaced cleavage is ubiquitous, also in northeastern Varanger Peninsula (Fig. 5). As Timanian deformation is

inferred to have waned gradually at this northwestern end of the Timan-Varanger Belt, as described in Roberts (1996), the signs of thrusting should also be less evident in these consistently anchizone-grade (Rice & Roberts 1995) basinal successions, and none has been proven. This does not deny the possibility that

Fig. 5. (a) Graded-bedded sandstones, siltstones and mudstones with penetrative, steeply dipping spaced (to locally slaty) cleavage, Tsypnavolokskaya Formation, Rybachi Peninsula; looking approximately NW. This prominent anchizone-grade cleavage is axial planar to common, open to close, NW-SE-trending folds in pelitic formations throughout most of Rybachi Peninsula, (b) Interbedded greywacke-sandstone and strongly cleaved mudstone of the Basnaering Formation, near Svartnes, east coast of Varanger Peninsula, north of the TKFZ; looking approximately NNW. The locality is situated on the 'eastern' limb of a NNW-SSE-trending open fold.

STRUCTURE OF THE TIMANIDE OROGEN

conjunctive thrust repetitions may be present, and particularly in the 4-8 km thick wedge of Neoproterozoic rocks that occur in a zone up to 80 km wide just offshore from the Rybachi Peninsula (Simonov et al 1998). In the Timan- Pechora region, geophysical data and about 70 deep drillholes have provided insight into the hinterland extension of the basinal domain, and its Timanian structures, beneath the Pechora Basin. Many important faults or fault zones have been defined, including the NW-SE-trending, East Timan (ETF) and Pechora-Kozhva Faults (PKF) (Figs 2 & 3). The latter has several different names in the Russian literature (e.g. PrePechora), and is recognized as a fundamental deep-seated fault zone against which the thick package of Riphean-Vendian metasedimentary rocks is terminated, in its footwall. The hanging wall of the PKF is characterized by a variety of magmatic rocks, including arc-type products. These were all strongly deformed during the Timanian Orogeny in what is termed the Pechorskaya collision zone (Olovyanishnikov et al. 1995) (Fig. 2), in the immediate hanging wall of the PKF; and drillcore data have revealed the presence of late- to post-tectonic, Late Vendian granites (Gee et al. 2000). Drillholes penetrating different parts of the Pechora Basin, mostly within the Pechorskaya collision zone, have also recovered evidence of greenschist-facies to, in one case, amphibolite-facies metamorphism (Ermolenko & Sobolev 1978; Belyakova & Stepanenko 1991) akin to that exposed on Kanin and in northern Timan.

Age of deformation Although there is evidence of local unconformities and even gentle folding of latest Riphean age in some sections of the Timan Range (Getsen 1987, 1991), there are sufficient biostratigraphic constraints to suggest that the principal Timanian deformation and metamorphism occurred during the Vendian period (Getsen 1987, 1991; Malkov 1992; Olovyanishnikov et al. 2000). In the inner foreland of the pericratonic zone, in the Mezen Basin or syneclise, a deltaic depositional system developed in approximately Mid Vendian time, subsequent to a marine flooding episode (Grazhdankin 2004), a progradation that relates to erosion of the Timanian mountains. Farther to the SE, to the west of the Southern Urals, Late Vendian conglomerates and sandstones in the foreland basin are interpreted as molasse (Becker 1968) deriving from the Timanide orogen. Over much of what is now the Pechora Basin, Ordovician sedimentation was widespread, locally with pockets of Late Cambrian deposits (MoczydlowskaVidal et al. 2004) providing an upper age constraint on Timanian deformation. Farther to the NW along the Timan-Varanger Belt, biostratigraphic control is less definitive. In SW Varanger Peninsula, although sedimentation continued throughout Vendian time, the significant changes in Mid Vendian palaeocurrent dispersal patterns noted earlier (Banks et al. 1971) are inferred to relate to encroaching Timanian orogenesis (Gorokhov et al. 2001). In the extreme SE of the Timan-Varanger Belt, in the Polyudov Ridge area, sedimentation continued into early Late Vendian time. There, the Kocheshor Formation has yielded a minimum K-Ar age for authigenic glauconite of c. 560 Ma (Sokolov & Fedonkin 1990; Maslov 2004). The current geochronological database pertaining to Timanian deformation and metamorphism is, in general, hampered by insufficient precision. There is an abundance of K-Ar dates from the Timan-Pechora region, but the analytical data and decay constants are not always available. Compilations of isotopic data include those of Akimova (1980), Getsen (1987) and Malkov (1992); and for the drillcores of the Pechora Basin, Gee et al. (2000). An assessment of pertinent dates reported over the last three decades is given in Olovyanishnikov et al. (2000).

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Summarizing the K-Ar data from the Timan-Kanin tract, Akimova (1980) considered that the Timanian deformation and metamorphism occurred during the broad time interval 680-570 Ma. Malkov (1992), employing recalculated and some more recent dates, narrowed down the main tectono-thermal event to around 600-570 Ma, a time range also favoured by Getsen (1991); and a second metamorphic stage was put at around 505-470 Ma. In northern Timan, alkaline plutonic and hypabyssal rocks intruded mainly at around 620-600 Ma (Andeichev 1998; Andreichev & Larionov 2000). Many of the syenitic dyke rocks are clearly involved in the main folding and are foliated, and also affected by boudinage (unpublished data). However, younger pegmatitic and aplitic syenite variants transect the folds and foliation. These relationships suggest that the alkaline magmatism was largely pre- to early syntectonic, with pegmatitic differentiates coming in slightly later. This also appears to date the Timanian tectono-thermal event to around 610-600 Ma in this particular part of the Timan Range. Some implications of this are discussed below. In the nearby Mezen foreland basin, U-Pb dating of zircons from tuff layers, in conjunction with sequence stratigraphy, constrain the main period of emergence of the Timanide mountains to c. 555-558 Ma (Martin et al. 2000; Grazhdankin 2004). Constraints on the age of the main tectono-thermal event in the successions of the Rybachi and Sredni Peninsulas are poor. On Sredni, the compactional fabric in the Late Riphean succession has an age of 620-610 Ma, based on a Rb-Sr study on illites (Gorokhov et al. 1995, 2002), and this is cut, in one area, by an undeformed mafic dyke with a 40Ar-39Ar mineral isochron, minimum age of 546 Ma (Roberts & Onstott 1995). An attempt to date the anchizone metamorphic fabric on Rybachi was unsuccessful (Roberts et al. 1998), due to the presence of a consistent detrital phyllosilicate component. A rather loose age range of c. 610-546 Ma has thus been accorded to this Timanian deformation. However, Gorokhov et al. (1995) reported a maximum age of 570 Ma for a fine, authigenic illite subtraction from a site adjacent to the SRFZ. This generation of illite was regarded as having formed during fluid transport associated with fault movement; and the maximum 570 Ma age has been suggested (Roberts 1996), to date the basinal inversion and SW-directed thrust deformation. Geochronological data from Varanger Peninsula also provide relatively poor constraints on the timing of Timanian deformation. Based on 40Ar-39Ar and Rb-Sr studies on illite (Dallmeyer & Reuter 1989; Gorokhov et al. 2001), a diagenetic event has been defined in the Vestertana Group in southern Varanger Peninsula at around 635-630 Ma, an age which is not dissimilar to that recorded on Sredni. In addition, intra- and post-tillite formations acquired their burial, compactional fabric at about 560-570 Ma. Interestingly, this is broadly coeval with a major reversal in palaeocurrent vectors (Banks et al. 1971) and with sediment infill from a northeasterly source accumulating in a small foreland basin (Gorokhov et al. 2001). In the northeastern part of Varanger Peninsula, a mafic dyke with a U-Pb zircon, upper intercept age of c. 567 Ma transects a cleavage in Upper Riphean, low-grade sedimentary rocks (Roberts & Walker 1997). Although the constraints are poor, this dyke may also provide an upper age limit for the Vendian, Timanian deformation. Isotopic investigations on drillcore samples taken from beneath the Pechora Basin had earlier been largely reliant on the K-Ar method, and had provided evidence of greenschist-facies metamorphism spanning the wide time interval Vendian to Cambrian. Gee et al. (2000), employing the single-zircon Pb-evaporation technique, showed that granites and diorites, cutting cleavage, yielded ages of c. 560-550 Ma. These results lend support to the notion of Timanian orogenesis and late- to post-tectonic, Vendian magmatism some 200-400 km NE of the Timan Range. Taking the isotopic dating evidence for the Timan-Varanger Belt as a whole, there are indications that the main Timanian

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event may have been diachronous along the belt, perhaps slightly younger in the far northwestern and southeastern areas. It would also be expected to have occurred earlier, in sea-floor situations, in the oceanic realm, farther to the NE. Today's geography shows that the North Timan-Kanin area forms a minor promontory along the Timan-Varanger Belt, marking a fairly abrupt strike deviation of some 15-20°. Assuming that this area had existed as such a recess (cf. Rankin 1976) in Neoproterozoic time, which is quite possible as all continental margins are rarely perfectly linear, then the SW-directed Timanian deformation would have affected this part of the cratonic margin first, and the Rybachi-Varanger and Polyudov Ridge areas slightly later. However, in view of the scanty nature of our current geochronological database, this idea can rank only as an interesting speculation at the present time.

Timanian-Caledonian interrelationships on Varanger Peninsula The prominent NW-SE fold axial and cleavage trend that characterizes the structural geology of Rybachi Peninsula can be followed into the eastern part of Varanger Peninsula (Fig. 6). This NW-SE structural grain is also clearly visible on LandsatTM satellite imagery (Roberts & Karpuz 1995) and, for the most part, represents the surface expression of the penetrative cleavage. It is convenient to consider the structural trends on Varanger Peninsula in three separate areas: (1) Vard0 district, (2) Vads0 district and (3) western Varanger Peninsula. In this way, evident links with Rybachi-Sredni are immediately apparent. Only the principal features are presented here; structural details concerning each sub-area have been given in Roberts (1996). In the Vard0 district (Fig. 6), north of the TKFZ, the low-grade, Late Riphean metasedimentary rocks of the Barents Sea Group show a fold axial trend and associated steep cleavage oriented between NW-SE and NNW-SSE in coastal areas (Fig. 5b), both swinging gradually towards N-S farther to the west. Folds are either upright or verge WSW, and there are many smallscale reverse faults with the same WSW to SW thrust sense (Karpuz et al 1993; Gjelsvik 1998). This structural grain is interrupted by cross-folds of ENE-WSW to NE-SW trend (Roberts

1972; Gjelsvik 1998). Comparative field studies, both here and in western Rybachi, aided by satellite imagery, provided convincing evidence that the main folds and associated cleavage belong to one and the same, Timanian event. The cross-folds, on the other hand, have been interpreted to relate to an episode of Caledonian buckling (Roberts 1996). In earlier explanations of the anomalous, c. NNW-SSE structural grain in this part of Varanger, a Scandian age was assumed, the folds arising either through buffering in front of a concealed basement 'ridge' (Roberts 1972) or as tip-folds that developed by late-stage backthrusting (Rice et al. 19890). Only one attempt at dating the cleavage in this subarea has been made, providing an imprecise Rb-Sr whole-rock isochron date of 520 ± 4 7 Ma (Taylor & Pickering 1981) from pelites near Hamningsberg. Recalculations of these analytical data have tended to favour an older, Vendian age (B. Sundvoll, pers. comm. 2003). In the Vads0 district, the diagenesis-grade, Late Riphean to Vendian sedimentary rocks dip at low angles to the NNE and show a diffuse structural grain (Fig. 6), reminiscent of that on Sredni. Large-scale, gentle, upright folds of east-west trend, devoid of cleavage, occur in the north of this subarea. Closer to the TKFZ, there are some mesoscopic, NW-SE-trending, SWfacing folds and related, small-scale thrusts (Herrevold 1993; Karpuz et al 1993). In the south, bedding intersection with topography shows an approximate WNW-ESE trend, and a bedding-parallel compactional fabric is widespread. In the western part of this subarea, there are sporadic WSW-ENE-trending, small-scale folds, verging SSE. In western Varanger Peninsula as a whole, the characteristic structural features are those of NE-SW to ENE-WSW-trending, SE-verging folds and thrusts (Fig. 6), the folds carrying an axialplanar spaced or slaty cleavage. These structures can be followed over long distances along strike to the SW, where they clearly form an integral part of the Caledonian allochthon, notably the Gaissa Nappe Complex (Rice et al 1989Z?; Siedlecka & Roberts 1996). In the southwestern part of this subarea, the succession ranges from Lower Vendian to Lower Cambrian, but to the west, on Digermul Peninsula (Fig. 6), the youngest fossiliferous rocks to be affected by the NE-SW-trending folds and cleavage are Early Tremadoc in age. A precise age for the anchizone to epizone cleavage has yet to be acquired. A K-Ar maximum age

Fig. 6. Principal structural trends in the Neoproterozoic rocks of Rybachi and Sredni Peninsulas and Kildin Island, and the Neoproterozoic to Cambrian rocks of Varanger Peninsula. The long broken lines represent either fold axial traces or penetrative cleavage traces, the latter detectable even on satellite imagery. The shorter broken lines represent the strike of bedding surfaces in the diagenesis-grade, pericratonic domains. DP, Digermul Peninsula.

STRUCTURE OF THE TIMANIDE OROGEN

of c. 440 Ma has been reported from a part of the Gaissa Nappe (Dallmeyer et al 1989), suggesting Scandian deformation, but there are also indications of possible Finnmarkian (latest Cambrian-Early Ordovician) low-grade metamorphism (Sturt et al 1975; Sundvoll & Roberts 2002). In the western two-thirds of the region north of the TKFZ, the trend of the common folds and thrusts, and associated cleavage, in the Late Riphean to Vendian rocks is generally closer to ENE-WSW (Fig. 6). The central parts of this particular region mark a zone of transition where the Caledonian structures die out eastwards and, conversely, the Timanian structures disappear, or are difficult to detect, westwards. Field evidence of fold or cleavage interferences has denoted that the Timanian structures are definitely the older of the two (Herrevold 1993; Roberts 1995; Gjelsvik 1998). In western areas, the precise age of the folding and related cleavage is not known, but there is a preliminary report of an Early Ordovician cleavage age in the extreme NW, near Berlevag (Frank & Rice 1999). Kinematic analysis of a variety of mesoscopic and microscopic structures and vein arrays along parts of the TKFZ (Karpuz et al. 1993; Herrevold 1993; Gjelsvik 1998) has indicated that the oldest component of movement along this major lineament relates to a SW-directed compressive stress regime. Although no isotopic dating is yet available, these structures are in harmony with the folds and cleavage described above and relate readily to the Timanian story. Subsequently, the principal compressive stress, (T!, shifted anticlockwise, reflected in a variety of kinemaic indicators pointing to dextral strike-slip. This component of deformation is inferred to be Caledonian, probably Scandian; and based on the age of a mafic dyke cutting the TKFZ in one area, all significant strike-slip displacive movement had ceased by Late Devonian time (Guise & Roberts 2002). In summary, the prominent NW-SE fold axial and cleavage trend seen on Rybachi Peninsula is clearly Timanian and can be followed without difficulty into the northeastern part of Varanger Peninsula (Fig. 6). The more diffuse but broadly parallel, structural grain seen on Sredni is discernible in the Vads0 district, but farther west this is overprinted by the pervasive NE-SW Caledonian trend. Likewise, north of the TKFZ, a Caledonian structural trend is ubiquitous in western areas. Varanger Peninsula is unique in exposing mutually interfering structures from two disparate fold belts, Timanian and Caledonian, the latter most likely, but not definitively, of Scandian age. North of the TKFZ, the Late Riphean, lithostratigraphical succession (Barents Sea Group) that is affected by Timanian structures can itself be followed as far west as the Kongsfjord area; but there, Caledonian structures prevail. Farther west, these rocks are concealed beneath Caledonian nappes composed of different rock assemblages of unknown, but probable Late Riphean age. Thus, we do not know the original northwestward extent of the basinal-domain, turbiditic rocks of Rybachi and Varanger. The TKFZ, however, can be followed some 150-200 km farther to the WNW (Lippard & Roberts 1987; Gabrielsen & Faerseth 1989; Karpuz etal 1993), pointing to multiple reactivation. Since the fault zone transects and offsets the Caledonian nappes and thrust sheets onshore, its prolongation probably represents a fundamental, deep, linear structure in this area, and may thus preserve basinal Riphean rocks in its concealed hanging wall beneath the nappe cover.

Timanian and Caledonian trends in the Barents Sea Several deep seismic profiles immediately offshore from the Barents Sea coast of the Kola Peninsula, together with drillcore material, have suggested that the Neoproterozoic rock successions can be traced as a 4-8 km thick, NW-SE-trending package extending up to c. 80 km into the southern Barents Sea, beneath Mesozoic cover (Fig. 7) (e.g. Simonov et al 1998; Sharov

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Fig. 7. Regional palaeotectonic setting in Late Permian time, modified and simplified from Gudlaugsson et al (1998), with small additions, showing the NE and NW branches of the Caledonides, and the distribution of the Timanide and Uralide orogens. The question marks are those of Fichler et al. (1997) who regard these particular areas of the Barents Sea as being dominated by a Timanide trend. The dashed lines with tags, marked A and B, are as follows: A, Approximate trace of a zone of SE-directed thrusting, interpreted as a possible Caledonian suture (Gudlaugsson et al. 1987); B, Approximate trace of an interpreted, SE-dipping, Caledonian suture zone (Breivik et al. 2002). The dotted line passing between the Barentsian microplate and Franz Josef Land is part of an inferred Caledonian suture, according to Gee et al. (2000); this is believed to link up, across the Barents Sea, with the North Norwegian Caledonides.

2000). Exactly how far northwestwards along strike this wedge of Timanian-deformed metasedimentary rocks extends is not known as it is difficult to recognize beneath the inferred northeastward extension of the Caledonian nappes. The Timanian trend is also reflected in several faults in the southwestern Barents Sea (Gabrielsen 1984; Lippard & Roberts 1987) which are considered to have inherited the structural grain of the sub-Palaeozoic basement. The same trend can be recognized from gravity and magnetic fields over the greater part of the eastern Barents Sea and Pechora Basin, and this has also been supported by satellite radar-altimetry gravity data (Fichler et al 1997). Tracing the extension of the Caledonian nappes of Finnmark into the offshore areas of the Barents Sea is controversial. Harland & Gayer (1972) believed that the Caledonides swung sharply eastward, paralleling the north coast of Kola, effectively transposed upon the Timanian trend, an interpretation that lacks supporting evidence. Siedlecka (1975) preferred a solution whereby the Finnmark Caledonides continued in a northeasterly direction across the Barents shelf to the sea area between Franz Josef Land and northern Novaya Zemlya, separating a Barents craton to the NW (later called Barentsia—Gudlaugsson et al 1998) from a Pechora craton (part of Baltica) to the SE. In her model, Siedlecka also recognized a branch of the Caledonides extending towards western Svalbard, reflecting an inferred triple

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junction of original pre-Iapetus Ocean rifting. This model was also adopted by Roberts & Gale (1978) in marking nodal areas of carbonatite and related alkaline magmatism during the rifting stage of lapetus. In other interpretations, both Ziegler (1988) and Nikishin et al. (1996) favoured a northwestward swing in the Caledonian grain, linking up with the Innuitian fold belt via the Caledoniandeformed terranes on Svalbard. Over the years, abundant seismic and refined, potential field data have tended to support Siedlecka's (1975) view of a bifurcation of the Caledonides, the main arm extending northeastwards to west of the northern tip of Novaya Zemlya, and a separate arm reaching to Svalbard (Gudlaugsson et al 1987, 1998; Dore 1991; Johansen et al 1994; Fichler et al 1997; Breivik et al 2002) (Fig. 7). Although favouring this bifurcate model, Fichler et al (1997) restricted the northeasterly trending branch of the Caledonides to the SW Barents Sea, noting that NW-SE-trending gravity highs cause the NE-SW trend to terminate in central areas of the Barents Sea. From this same area of the SW Barents Sea, Gudlaugsson et al (1987) reported deep-seismic reflection profiling that revealed a pattern of reflections consistent with east- to SE-directed thrusting at middle and lower crustal levels (Fig. 7). It was speculated that this west- to NW-dipping feature might represent the main Caledonian suture, an interpretation supported by Dore (1991). More recently, however, Breivik et al (2002) presented oceanbottom seismometric data from roughly the same part of the SW Barents Sea which, in their interpretation, denote the existence of a SE-dipping Caledonian suture between Barentsia and Baltica (Fig. 7). Thus, there are indications of the presence of a suture, or sutures, dipping in diametrically opposed directions. In another interpretation, Gee (2000) and Gee et al (2000) show a possible suture extending from the Scandinavian Caledonides northwards through the Barents Sea to separate eastern Svalbard (Kvit0ya) from Franz Josef Land.

Palaeotectonic setting: a synopsis Models of the Timanian fold-and-thrust belt have been presented by, e.g. Scarrow et al (2001) and Roberts & Siedlecka (2002). Here, we provide a brief description of the geodynamical development of the orogen, based on the outlines presented above and on many years' work on drillcores (by V.O.) taken from the Pechora Basin. Throughout Mid to Late Riphean time, this particular margin of the East European Craton, or proto-Baltica, was characterized by crustal extension, leading to incipient rifting and the generation of an array of NW-SE-trending deep faults that are considered to have developed by reactivation of older megalineaments. One long fault in particular (the composite CTF, SRFZ and TKFZ) functioned as a border structure and by early Mid Riphean time separated pericratonic or platformal deposition on the craton side from deeper-water, slope-and-rise, mainly terrigenous sedimentation in an evolving, deeper marine, basinal domain. This bipartite division marked a fundamental, passive, northeastern margin of the East European Craton. Indications that this elongate 'basin' gradually deepened northeastwards into an oceanic domain are seen, in the Timan Range, in Mid to Late Riphean time, in the form of sporadic tuffs and rare basaltic lavas in thick turbiditic successions. Drillcores from beneath the Pechora Basin, in different zones, confirmed extensive, Late Riphean, magmatic activity, which is also bimodal and largely of island arc character (Getsen 1987, 1991; Belyakova & Stepanenko 1991; Dovzhikova et al 2004). Thus, a subduction system is inferred to have been operational at about this time in a fairly mature oceanic domain. Since the cratonic margin represented by the Timan-Varanger Belt remained passive, initial subduction polarity is likely to have faced away from Baltica

(e.g. Scarrow et al 2001; Roberts & Siedlecka 2002). By Vendian time, possibly even terminal Riphean, the presence in drillcores of evolved, calc-alkaline, granitic and dioritic plutonic rocks with Late Vendian crystallization ages show that the situation had changed (Gee et al 2000). This may have followed subduction slab break-off, and a subduction reversal, that converted the passive margin to an active, compressional one (Roberts & Siedlecka 2002), and heralded the Timanian orogenesis. Defining the location, or locations, of true oceanic crust in this region, in Riphean time, is difficult, since the only definite exposed remnant is that of the arc-related ophiolite of the Engane-Pe range in the Polar Urals, U-Pb (zircon) dated to 670 ± 5 Ma (Dushin 1997; Khain et al 1999; Scarrow et al 2001). However, there seems to be general agreement that oceanic spreading conditions occurred in what is now the Varandey-Adzhva zone (Fig. 2), where there is also geophysical evidence suggesting the presence of mafic and ultramafic bodies at depth (Dedeyev & Zaporozhtseva 1985; Kostyuchenko 1994). This zone may have extended towards the southern part of Novaya Zemlya where Timanian deformation is also recorded (Lopatin et al 2001). Oceanic crust may also have existed in narrow zones in the Bolshezemelskaya zone between the microcontinental blocks that are inferred to characterize this particular zone (Getsen 1991). At some stage during the Vendian, and perhaps in terminal Riphean time in some arc-related situations, the Timanian compressional to transpressional, orogenic cycle commenced. During this lengthy process, the Riphean to Early Vendian collage of island arc and ocean-floor magmatic rocks, slope-andrise volcanosedimentary assemblages, bimodal plutonic rocks and blocks of crystalline continental character were telescoped and accreted against this northeastern margin of Baltica. Deformation, with peak amphibolite-facies metamorphism, is considered to have been most intense within the Pechorskaya collisional zone (Olovyanishnikov et al 1995) and probably started in sea-floor environments. Lower-grade conditions characterize the Timan zone, with higher-grade paragenesis appearing in the cores of anticlinoria and in upthrust fault slivers as, for example, on Kanin and in northernmost Timan close to the ETF, or in the hanging walls of adjacent subparallel faults. Constraints on the timing of Timanian deformation indicate that the peak tectono-thermal event along the Timan-Varanger Belt occurred around 600-570 Ma. This may be slightly earlier in the vicinity of Kanin/North Timan, where a minor continentalmargin promontory may have existed and thus acted as a buttress to the SW-directed compression, with deformation progressively younger both to the NW and to the SE. Elsewhere in the Timan-Pechora-Urals region, Puchkov (1997) concluded that, taking the Urals as a whole, this pre-Uralide or Timanian orogeny ranged from c. 600 to 550 Ma, though it may have started as early as 630 Ma in some thrust complexes. The age of the Timanian orogenesis is similar to that recorded in the dispersed, Avalonian-Cadomian, arc-related terranes of eastern North America and western and central Europe. Many of these terranes show cycles of Neoproterozoic slope-and-rise sedimentation, intra-oceanic subduction and phases of magmatic arc development comparable to those that are thought to have existed in the Timan-Pechora-Urals region. The distribution of these fragmented terranes and their inferred palaeotectonic relationship in a peri-Gondwanan context have been discussed by, e.g. Scarrow et al (2001) and Roberts & Siedlecka (2002). The first author acknowledges support from the Geological Survey of Norway (NGU) and the Norwegian Science Research Council for fieldwork on Kola Peninsula during the years 1990-1993; and from NGU for field studies in central Timan in 1995. Fieldwork in northern Timan and Kanin Peninsula, for both authors, was supported largely by the Swedish Polar Research Secretariat (SPRS). The Timanide Orogen has been the subject of wide-ranging discussion at Europrobe (Timpebar project) workshops. We are grateful to everyone behind the scenes at SPRS, and to the expedition leader Professor David Gee.

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Constructive comments by David Gee on the two original manuscripts which were eventually welded into the present single contribution, reviewed by David Gee and Roy Gabrielsen, are much appreciated. Irene Lundquist is thanked for her assistance with figure preparation.

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Northeast Norway, and Avalonian-Cadomian connections. Tectonophysics,352, 169-184. ROBERTS, D. & WALKER, N. 1997. U-Pb zircon age of a dolerite dyke from near Hamningberg, Varanger Peninsula, North Norway, and its regional significance. Norges geologiske unders0kelse Bulletin, 432,95-102. ROBERTS, D., COSCA, M. A. & RICE, A. H. N. 1998. A 40Ar/39Ar dating study of very low-grade metamorphism from the Rybachi and Stedni Peninsulas, NW Russia - a preliminary report. Norges geologiske undersfikelse Report 98.168, 17 pp. ROBERTS, D., SIEDLECKA, A. & OLOVYANISHNIKOV, V. G. 2004. Neoproterozoic, passive-margin, sedimentary systems of the Kanin Peninsula, and northern and central Timan, NW Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30,5-17. SAMUELSSON, J. 1997. Biostratigraphy and palaeobiology of Early Neoproterozoic strata of the Kola Peninsula, Northwest Russia. Norsk Geologisk Tidsskrift, 77, 165-192. ScARROW, J. H., PEASE, V., FLEUTELOT, C. & DUSHIN, V. 2001. The Late Neoproterozoic Enganepe ophiolite, Polar Urals, Russia: an extension of the Cadomian arc? Precambrian Research, 110, 255-275. SCHATSKY, N. S. 1935 [On tectonics of the Arctic. Geology and economic deposits in northern USSR]. Trudy, First geology exploration conference, Glavsevmorputy, Lenigrad, 1, 476-509 [in Russian]. SCHATSKY, N. S. 1958. Les relations du Cambrian avec le Proterozoigne et les plissements Balkaliens. Les relations entre Precambian et Cambrian. CNRS Colloques Int. Paris, 91-101. SHAROV, N. V. 2000. [Deep structure of the junction zone of the Baltic Shield and the Barents shelf plate from seismic data]. Regional Geology & Metallogeny 10, Vsegei, St Petersburg [in Russian, with English abstract]. SIEDLECKA, A. 1972. Kongsfjord Formation - a Late Precambrian flysch sequence from the Varanger Peninsula, Finnmark. Norges geologtiske undersfikelse, 278, 41-80. SIEDLECKA, A. 1975. Late Precambrian stratigraphy and structure of the northeastern margin of the Fennoscandian Shield (East Finnmark Timan region). Norges geologiske undersfikelse, 316, 313-348. SIEDLECKA, A. & SIEDLECKI, S. 1967. Some new aspects of the geology of the Varanger Peninsula (northern Norway). Norges geologiske undersfikelse, 247, 288-306. SIEDLECKA, A. & ROBERTS, D. 1995. Neoproterozoic sedimentation and subsequent tectonic deformation in the northern coastal areas of Norway and Russia, (extended abstract) Norges geologiske undersfikelse, Special Publication, 7, 331-332. SIEDLECKA, A. & ROBERTS, D. 1996. Finnmark Fylke. Berggrunnsgeologi (Bedrock geology map) M 1:500,000. Norges geologiske undersfikelse.

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SIEDLECKA, A., LYUBTSOV, V. V. & NEGRUTSA, V. Z. 1995a. Correlation between Upper Proterozoic successions in the TanafjordenVarangerfjorden Region of Varanger Peninsula, northern Norway, and on Sredni Peninsula and Kildin Island in the northern coastal area of Kola Peninsula in Russia. Norges geologiske undersfikelse, Special Publication, 7, 217-232. SIEDLECKA, A., NEGRUTSA, V. Z. & PICKERING, K. T. \995b. Upper Proterozoic turbidite system of the Rybachi Peninsula, northern Russia—a possible stratigraphic counterpart of the Kongsfjord Submarine Fan of the Varanger peninsula, northern Norway. Norges geologiske undersfikelse, Special Publication, 7, 201-216. SIEDLECKA, A., ROBERTS, D., NYSTUEN, J. P. & OLOVYANISHNIKOV, V. G. 2004. Northeastern and northwestern margins of Baltica in Neoproterozoic time: evidence from the Timanian and Caledonian orogens. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London Memoirs, 30, 169-190. SIMONOV, A. P., GUBERMAN, D. M. & YAKOVLEV, Y. N. 1998. Tectonics of the Kola-Kanin monocline: a multidisciplinary interpretation of geological and geophysical data in the ocean-continent transition zone. In: MITROFANOV, F. P. & SHAROV, N. V. (eds) Seismogeological model of the lithosphere of northern Europe: the Barents Region. Kola Science Centre, RAS, Apatity, 159-197. SOKOLEV, B. S. & FEDONKIN, M. A. (eds) 1990. The Vendian System, 2. Regional geology, Chapter 3. Vendian of the south-eastern White Sea area. Springer Verlag, Berlin, 76-87. SOKOLEV, B. S. & FEDONKIN, M. A. 1990. The Vendian System. Springer Verlag, Berlin, 273 pp. STURT, B. A., PRINGLE, I. R. & ROBERTS, D. 1975. Caledonian nappe sequence of Finnmark, northern Norway, and the timing of orogenic deformation and metamorphism. Bulletin of the Geological Society, America, 86, 710-718. SUNDVOLL, B. & ROBERTS, D. 2002. Rb-Sr dating of cleaved mudstones from eastern and western parts of the Gaissa Nappe Complex, Finnmark. Norges geologiske undersfikelse Report 2002.110, 20 pp. TAYLOR, P. N. & PICKERING, K. T. 1981. Rb-Sr isotope age determination on the Late Precambrian Kongsfjord Formation, and the timing of compressional deformation in the Barents Sea Group, East Finnmark. Norges geologiske undersfikelse, 367, 105-110. TSCHERNYSCHEV, F. N. 1901. [On the geological structure of the Timans and on relation of the Timan Fault to other regions of northern Europe]. Zapiski Russiskogo Mineralogicheskogo Obshchestva, 34, 29-33 [in Russian]. ZIEGLER, P. A. 1988. Evolution of the Arctic-North Atlantic and the Western Tethys. American Association of Petroleum Geologists Memoir, 43, 198 pp. ZONENSHAIN, L. P., KUZMIN, M. I. & NATAPOV, L. M. 1990. Geology of the USSR: a plate tectonics synthesis. American Geophysical Union, Geodynamics Series 21, 242.

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Neoproterozoic high-grade metamorphism of the Kanin Peninsula, Timanide Orogen, northern Russia H. LORENZ1, A. M. PYSTIN2, V. G. OLOVYANISHNIKOV2 & D. G. GEE1 Uppsala University, Department of Geosciences, Villavdgen 16, 752 36 Uppsala, Sweden (e-mail: [email protected]) 2 Russian Academy of Sciences, Ural Division, Komi Science Centre, Institute of Geology, 54, Pervomayaskaya str., 167982, Syktyvkar, Komi, Russia 1

Abstract: Throughout most of the exposed Timanide Orogen, the grade of regional metamorphism does not exceed greenschist facies. Only on the Kanin Peninsula and in the northernmost Timan and in a few of the drillcores that sample basement beneath the Pechora Basin, are amphibolite facies rocks present. This study presents new P- T estimates by garnet-biotite geothermometry and GASP geobarometry from the high-grade rocks of the Kanin Peninsula, together with a new structural interpretation of the SE-Kanin area. Thick metaturbiditic successions show an increase of metamorphic grade downwards from low greenschist to high amphibolite facies. On the SE Kanin Peninsula, peak metamorphic conditions are estimated to have reached c. 0.72 GPa and c. 610 °C. These inferred Late Proterozoic successions were thickened by thrust-stacking and emplaced southwestwards onto low-grade metamorphic pericratonic sedimentary rocks of the East European Craton during the Timanian Orogeny. Timing of metamorphism is inferred to be late Neoproterozoic, but radiometric ages are not well constrained and more work is needed. The new field observations and analytical results suggest that high-grade regional metamorphism is not confined to the Timan-Kanin area. This is of importance when considering the geophysical data over the Pechora Basin and Barents Sea, particularly the interpretation of crustal velocity structure. Higher velocities, often inferred to represent pre-Riphean basement, may well be related to Neoproterozoic and Mesoproterozoic complexes.

The NW-trending Timanide Orogen extends from northernmost Norway to the Urals, a distance of nearly 3000 km, marking the northeastern margin of the East European Craton. Neoproterozoic platform successions, often with stromatolite-bearing carbonate rocks, pass northeastwards into basinal siliciclastic formations, dominated by turbidites. Further NE towards the hinterland, the Timanides are covered by thick Phanerozoic successions of the Pechora Basin, reappearing only in the foreland fold belt of the Urals, where Neoproterozoic fragmented ophiolites and arc-volcanites are preserved. The character of the Timanide Orogen beneath the Pechora Basin is known from about seventy deep (3-5 km) drillholes (Belyakova & Stepanenko 1991) and interpreted with the help of potential field and seismic data (Kostyuchenko 1994). The turbidite dominated successions exposed in the Timan Range (Fig. 1)

extend nearly 100km northeastwards beneath the Early Palaeozoic unconformity towards the hinterland, being intruded by c. 560 Ma granites (Gee et al. 2000). These metasedimentary successions are apparently cut off abruptly by faulting along the western boundary of the Pechora Zone, to the NE of which the pre-Palaeozoic bedrock is dominated by igneous rocks of oceanic affinity, also intruded by Vendian granites (Dovzhikova et al. 2004). Between the Pechora Zone and the Urals, other terranes have been recognized in deep drillholes, also intruded by late Neoproterozoic granitoids (Gee et al 2000). The age of the sedimentary successions of the Timanian foreland fold-and-thrust belt is known to be mainly Neoproterozic, based on reports of acritarchs (Olovyanishnikov 1998); however, the succession may, in part, be older. The metaturbidites are thrust southwestwards onto the platform of the Mezen Basin, where Late Vendian northeasterly-derived deltaic successions rest with angular unconformity on the underlying Neoproterozoic rocks (Grazhdankin 2004). Throughout the Timanide Orogen, as it is known at the surface and beneath the Early Palaeozoic unconformity of the Pechora Basin, the metamorphic grade is no higher than greenschist facies (Getsen et al. 1987). Where the turbidite successions are exposed, for example in the Timan Range and along the Barents Sea coast and on Kanin Peninsula, sedimentary structures are often well preserved, but penetrated by a strong, generally high angle NE-dipping cleavage. Sericite and chlorite grew in the cleavage of these slates, which to the NE are refolded and phyllitic, developing biotite and some Mn-rich garnet. Only in the northernmost Timan Range and on Kanin Peninsula are higher grade metasedimentary rocks exposed at the surface in a major fold culmination, providing insight into the character of Timanian regional metamorphism at depth.

Kanin Peninsula and northern Timan Regional geology Fig. 1. Tectonic elements of the western Eurasian Arctic in the Early Tertiary. Timan-Kanin region marked with a box (from Gee 2004).

A major gentle anticline (the Kanin-North Timan Megaanticlinorium of Olovyanishnikov 2001) arches the Palaeozoic

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 59-68. 0435-4052/047$ 15 © The Geological Society of London 2004.

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H. LORENZETAL.

Fig. 2. Simplified geological map of the Timan-Kanin region, based on Olovyanishnikov (2001, fig. 2). CTF, Central Timan Fault (after Olovyanishnikov et al. 2001); RTF, East Timan Fault; CB, Cape Barmin; CBR, Cape Bolshoy Rumyanichny; CLe/Clw, Cape Ludovatye east/west; CM, Cape Mikulkin.

formations of the Kanin Kamen (ridge) and northern Timan (Figs 2 & 3). Beneath the Palaeozoic unconformity, with Devonian 'red beds' and basalts and locally Silurian limestones, metasedimentary successions are recognized together with a range of intrusions, including dolerites and an alkaline igneous complex (Ivensen 1964). The metasedimentary rocks are dominated by turbidites, generally in greenschist facies, but both in parts of Kanin Kamen, and, locally, in northernmost Timan, metamorphic grade is higher. In the Timan area, however, relationships between amphibolite and greenschist facies rocks are not well defined; their presence on Cape Barmin and also in a drillhole (no. 112 in Fig. 2) beneath Devonian strata near the east coast of Cheshskaya Bay have proved difficult to relate to the neighbouring lower grade rocks. The stratigraphy of Kanin Peninsula is subdivided into three major units, from base upwards the Mikulkinskaya, Tarkhanovskaya and Tabuyevskaya Groups ('Series' in the Russian literature, e.g. Getsen 1975; Novitsky 1976), all dominated by metaturbidites (Figs 3 & 4). The Mikulkinskaya Group, exposed on the north side of Cheshskaya Bay along the coast of southeasternmost Kanin Peninsula (Fig. 2), has an estimated thickness of 1500 m and consists mainly of amphibolite facies metaturbidites, with a prominent formation of quartzites and calcsilicate rocks (skarnoids) in the upper part. This group has a conformable contact with the overlying Tarkhanovskaya

Fig. 3. Geological map of the Kanin Peninsula, based on Olovyanishnikov (2001, fig. 18).

METAMORPHISM OF KANIN PENINSULA

Fig. 4. (a) Parallel-bedded metaturbidites of the Mikulkinskaya Group on the east coast of Kanin Peninsula, near the locality of sample G76 (Fig. 6). (b) Folded metaturbidites in the core of the Mikulkin Antiform.

Group, a c. 5000 m thick sequence metamorphosed to garnet mica schists. Although the Tarkhanovskaya and overlying Tabuyevskaya Groups are separated by a tectonic contact in many places, stratigraphic continuity has been inferred (Olovyanishnikov 2001). The c. 4000m thick Tabuyevskaya Group exposed along the Barents Sea coast consists of various mica and calcareous schists and volcano-sedimentary successions. Neoproterozoic acritarchs have been reported from this group, the upper part of which may be a separate tectonic unit, as discussed below. All the metaturbiditic units of the Kanin Peninsula and northern Timan are thrust southwestwards onto the margin of the East European Craton along the Central Timan Fault (CTF). The latter is inferred to pass through Cheshskaya Bay (Fig. 2). The northwestern continuation of the CTF, the Sredni-Rybachi and Trollfjorden-Komagelva Fault Zones, separate the basinal siliciclastic successions from the platformal successions of the continental margin (e.g. Olovyanishnikov et al. 2000). Neoproterozoic carbonate rocks are exposed c. 130 km SE of Cape Mikulkin between the Ludovatye Capes (Fig. 2). Olovyanishnikov (1998) inferred that another fault zone (the East Timan Fault) separates the high-grade rocks on the Kanin Peninsula from lower grade metaturbidites to the SW.

Range and epidote-amphibolite facies rocks from Kanin Peninsula. A detailed petrographic description of the metamorphic rocks from the northern Timan Range and Kanin Peninsula was given by Novitsky (1976), who estimated peak-metamorphic conditions at Cape Mikulkin to 725 °C and 0.8 GPa. Getsen et al (1987) published the first metamorphic map of the Timan Range and Kanin Peninsula. The metamorphic evolution was divided into three stages, interpreted to reflect the same orogenic cycle. A post-diagenetic low-grade metamorphism, caused by loading during basin subsidence, was followed by a second stage of regional dynamo-thermal metamorphism during Timanian (in their terminology 'Baikalian') orogeny. A third stage of local late-orogenic contact metamorphism concluded the metamorphic evolution. A comprehensive study of the metamorphic zoning in the Kanin region was carried out by Kasak et al. (1989) based on samples collected by other workers (A. P. Kasak pers. comm. 2000) and compilations by previous authors (e.g. Getsen 1975; Getsen et al. 1987; Novitsky 1976). Metamorphism during the Timanian orogeny was responsible for the 'zonal metamorphism in the area' (Kasak et al. 1989, p. 74). They described the metamorphic facies zones to be subparallel to the stratigraphic boundaries and only discordant in the Cape Mikulkin area (Fig. 3) because of the influence of a 'thermal dome'. Metamorphic grade decreases generally up-section; only locally in the ME, in the upper part of the Tabuevskaya Group, is an increase from lower to upper greenschist facies documented. On Kanin Peninsula, Kasak et al. (1989) distinguished five metamorphic zones. Located at the top of the Tabuyevskaya Group on the northeastern Kanin Peninsula, the chlorite zone is represented by sericite-chlorite-bearing schists; subordinate minor biotite has a low Fe/(Fe + Mg) ratio. Glaucophane is reported from these greenschists, growing in the schistosity; its presence was explained by locally increased pressure near a thrust zone (perhaps the tectonic contact between the middle and upper formations of the Tabuyevskaya Group, Fig. 5). In central and northwestern parts of the Kanin Peninsula, biotite in the biotite zone is characterized by Fe/ (Fe + Mg) ratios between 0.63 and 0.73; it grows in the foliation together with fine-grained spessartine-almandine garnet. Garnet contributes up to 30% to the total rock volume and has a Fe/(Fe + Mg) ratio of 0.86-0.89. Metamorphic grade increases downwards; most of the underlying Tarkhanovskaya and Mikulkinskaya Groups are within the staurolite zone, predominantly consisting of schists with subordinate gneisses and calcareous horizons. Kasak et al. (1989) report rocks characterized by assemblages containing kyanite-staurolite-sillimanite and kyanite-almandine-sillimanite with high titanium (3.4%) biotite with a Fe/(Fe + Mg) ratio of 0.52-0.60. The sillimanite zone, present only in the south westernmost part of Cape Mikulkin, is

Previous work on metamorphism Ivensen (1964) was the first to describe the metamorphic facies distribution in the Timan Range and on the Kanin Peninsula. He reported greenschist facies rocks from the northern Timan

61

Fig. 5. Interpretation of the Mikulkin Antiform as a thrust-related structure (sketch is not to scale).

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H. LORENZCTAL.

characterized by parageneses containing sillimanite and orthoclase (Kasak etal. 1989). Cordierite is common in some rocks; staurolite is relict and replaced by muscovite. Garnet is of pyrope-almandine composition with c. 65% almandine. Kasak et al. (1989) refer to three post-metamorphic stages of alteration on Kanin Peninsula: high temperature alkaline metasomatism, acid leaching and, finally, ferromagnesian metasomatism. Pegmatite dykes, associated with muscovitization and intense alteration of carbonaceous rocks and development of calcitediopside-amphibole rocks (the so-called 'skarnoids') are, among others, thought to be the consequences of this late-stage alteration. In northernmost Timan, the biotite-quartz, muscovite-quartz and garnet-staurolite-biotite-quartz schists of the Barminskaya Group near Cape Barmin (Fig. 2) have been correlated with the Tarkhanovskaya Group of Kanin Peninsula (Getsen 1975) and are said to belong to the muscovite, biotite and garnet zones. Near the eastern coast of Cheshskaya Bay, from drillcore 112 (Fig. 2), garnet-biotite and garnet staurolite-biotite-muscovite 'plagioschists' have been reported, underlying the Devonian deposits. These rocks may be part of the staurolite zone, as observed at Cape Mikulkin on the Kanin Peninsula.

Previous work on geochronology Much geochronological data, mostly obtained by the K-Ar and Rb-Sr methods and published during the last four decades, provide a somewhat ambiguous basis for interpreting the ages of the rocks and the timing of Timanian orogeny. A comprehensive summary of these data is beyond the scope of this paper, but studies which are important for this work (cf. Olovyanishnikov et al 2000) are summarized below. Metamorphosed gabbro and diabases of the Timan-Kanin region have been studied using the K-Ar method (Malkov 1966). He concluded that regional metamorphism took place between 640 and 620 Ma. Recalculation of his results by Andreichev (1998) with modern decay constants (source not mentioned) gives an average of c. 665 Ma. Andreichev's K-Ar dating of Mikulkinskaya Group rocks resulted in 970 ± 25 Ma and 521 ± 13 Ma for biotite and muscovite, respectively. Andreichev (1998) also provided Rb-Sr and K-Ar ages on the alkaline igneous complex of northernmost Timan of c. 600 Ma, but with a wide spread of ages. Larionov et al. (2004) obtained new data by the Pb/Pb single zircon evaporation method on the intrusive massifs of the northern Timan Range of 610-620 Ma.

Fig. 6. Geological map of the Cape Mikulkin area, southeastern Kanin Peninsula. Modified from Olovyanishnikov (2001).

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METAMORPHISM OF KANIN PENINSULA

From parts of the Timanide orogen, buried beneath the Pechora Basin, new data from drillcores recovered from granitoids of lateto post-tectonic character yielded Pb/Pb single zircon evaporation ages of c. 560 Ma (Gee et al 2000).

Fieldwork and sampling on southeastern Kanin Peninsula Fieldwork was carried out in the northern Timan Range and on the Kanin Peninsula. The Cape Mikulkin area, southeastern Kanin Peninsula, is well exposed along the coast; inland outcrops are scarce (Fig. 6). The coastal zone has been mapped and a wide range of samples has been collected for laboratory investigations of petrography and metamorphic evolution.

Structure The structure of the Neoproterozoic rocks of the Cape Mikulkin area is dominated by the doubly-plunging Mikulkin Antiform (Fig. 6). Referred to as a gneiss dome (e.g. Olovyanishnikov 2001), it exposes rocks of the highest metamorphic grade within the Timanides in its culmination. The structure suggests a thrustrelated origin (Fig. 5) as part of an antiform-synform pair, overlying a SW-vergent thrust-zone (Figs 3 & 5). At least two generations of isoclines, the later with amplitudes of more than 100m, can be identified. These isoclinal folds show axial planar foliations and are refolded by upright to west-vergent tight folds apparently related to the development of the Mikulkin Antiform. Individual metaturbiditic units, consisting of a quartz-rich base grading up into a pelitic top, usually have a thickness of 1-3 m in the lower stratigraphic levels; they become thinner (0.1-0.5 m) in the upper part of the Mikulkinskaya Group and in the Tarkhanovskaya Group (Fig. 4). Metamorphic grade is

highest in the hinge of the Mikulkin Antiform, i.e. in the lowest exposed structural level. Despite the high metamorphic grade, the development of a gneissic texture is subordinate and most of the exposed rocks are kyanite-garnet schists. Within some of the gneissic rocks at Cape Mikulkin, small pods of neosome have been observed. Kasak et al (1989) reported moderate-P, high-r assemblages (cordierite, sillimanite and orthoclase bearing) from the same area. Metamorphic grade decreases up-section, staurolite-garnet schists and garnet-mica schists successively predominate. Little deformed pegmatites of decimetres to several tens of metres thickness cross-cut foliations in the Mikulkinskaya Group.

Lithologies Samples were taken throughout the section in both limbs of the Mikulkin Antiform. Thin-section examination allowed the selection often representative samples from different stratigraphic levels and metamorphic grades for electron microprobe investigations. Short descriptions of the samples (eight, Uppsala University plus two, Komi Science Centre, Syktyvkar) are given in Table 1.

Thermobarometry Methods Garnet, biotite and plagioclase have been analysed for thermobarometry. Most of the mineral analyses have been made at the National Microprobe Laboratory, Uppsala University, using a Cameca SX50 WDS electron micropobe at 20.0 kV acceleration voltage and 15nA beam current. A few analyses were also performed at the Komi Science Centre, Syktyvkar on a Joel

Table 1. Sample descriptions Sample

Lithology

Paragenesis

Accessories

Equillibrium assemblage

Texture

G48

grt-mica schist

apa, op. min.

grt + bt + mu + qtz + chl

G55

grt amphibolite

grt + bt + mu + pi + qtz + chl hbl + grt + bt + pl + qtz + cc + chl

G56

grt-mica schist

apa, op. min.

grt + bt + mu + qtz + chl

G72

kya-stt-grt schist

grt + bt + mu + pi + qtz + chl stt + grt + bt + mu + pi + kfs + qtz; relict kya

tor, apa, tit

stt + grt + bt + mu + qtz

G76

kya-stt-grt schist

stt + grt + bt + mu + pi + qtz; relict kya

tor, apa

stt + grt + bt + mu + qtz

G77

kya-stt-grt schist

kya + grt + bt + mu + pi + qtz; relict stt

tor, apa

kya + grt + bt + mu + qtz

G78

kya-grt schist

apa

kya + grt + bt + mu + qtz

G87

kya-grt schist

kya + grt + bt + mu + pi + kfs + qtz kya + grt + bt + mu + pi + qtz

rough anastomosing foliaton, fine-grained matrix with small (<0.5 mm) euhedral grt-porphyroblasts rough anastomosing foliation, fine-grained matrix with 2-3 mm long bt in the foliation; large subhedral grt-porphyroblasts with cc and chl inclusions spaced anastomosing foliation, subhedral to anhedral grt-porphyroblasts in fine-grained matrix clear, euhedral to subhedral stt- and large (> 1 cm) anhedral grt-porphyroblasts in coarse-grained matrix; relict kya in form of grain aggregates rough, spaced foliation, clear, euhedral to anhedral stt-(up to 2 mm) and smaller anhedral grtporphyroblasts in coarse-grained matrix preferred fabric orientation marked by bt (c. 20% of the rock), white mica and kya; anhedral grt (< 1 mm) and large mu crystals small anhedral grt-porphyroblasts in fine-grained matrix

tor, apa

kya + grt + bt + mu + qtz

TK26

grt schist

grt + mu + pi + qtz

apa

grt + pi + qtz

TK34

stt-grt schist

stt-f grt + bt + pi + kfs + qtz + chl

ilm

stt + grt + bt + qtz

hbl + pi + grt

fine-grained equigranular matrix; crenulated mica-rich bands with elongated grt in preferred orientation; anhedral to skeletal grt elsewhere spaced foliation with concentration of mu; coarse-grained matrix with up to 2 mm large corroded grt large prismatic stt- and small corroded grt-porphyroblast in fine-grained equigranular matrix

apa, apatite; bt, biotite; cc, calcite; chl, chlorite; grt, garnet; hbl, hornblende; ilm, ilmenite; kfs, K-feldspar; kya, kyanite; mu, muscovite; op. min., opaque minerals; pi, plagioclase; qtz, quartz; stt, staurolite; tit, titanite; tor, tourmaline

64

H. LORENZCTAL.

Table 2. Garnet compositions from microprobe analyses (cations per 12 oxygens) Sample number

Si

Al

Ti

Mg

Fe*

Mn

Ca

Na

K

^Oxide

Xpyr

aim

G48A G55A G56 Al G56A2 G72A G76B G77B G77D1 G77D2 G77F G78B1 G78B2 G78C1 G78C2 G78C3 G78C4 G87B G87C G87D1 G87D2 TK26 TK34

2.970 2.988 3.016 3.013 3.016 2.954 3.009 3.001 2.974 2.990 3.016 3.029 2.995 3.005 3.035 2.993 2.994 2.948 2.966 2.972 2.924 2.939

1.981 1.960 1.951 1.949 1.946 2.020 1.971 1.971 1.969 2.011 1.970 1.988 1.956 1.966 1.952 1.962 1.994 2.015 2.030 2.034 1.940 1.951

0.003 0.004 0.006 0.004 0.001 0.003 0.002 0.005 0.003 0.000 0.000 0.001 0.000 0.000 0.000 0.002 0.000 0.000 0.002 0.001 0.003 0.000

0.231 0.230 0.182 0.158 0.377 0.364 0.365 0.322 0.330 0.343 0.252 0.273 0.270 0.291 0.266 0.305 0.358 0.386 0.372 0.359 0.325 0.205

2.216 2.024 2.001 1.999 2.332 2.371 2.268 2.269 2.331 2.305 2.122 2.064 2.103 2.083 2.073 2.094 2.293 2.361 2.356 2.365 2.544 2.467

0.119 0.042 0.046 0.120 0.136 0.138 0.175 0.204 0.200 0.208 0.434 0.379 0.409 0.375 0.412 0.344 0.103 0.094 0.102 0.104 0.115 0.153

0.517 0.778 0.801 0.764 0.200 0.181 0.207 0.234 0.231 0.148 0.204 0.240 0.293 0.284 0.251 0.325 0.260 0.240 0.190 0.175 0.208 0.371

0.000 0.004 0.000 0.004 0.004 0.005 0.013 0.005 0.000 0.000 0.000 0.000 0.000 0.016 0.000 0.000 0.010 0.002 0.000 0.000 0.090 0.000

0.001 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.003 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.001 0.000 0.000 0.000 0.000

99.179 100.240 99.418 98.364 99.550 98.973 99.265 99.600 100.972 99.097 97.357 99.350 99.869 97.831 100.352 98.680 101.135 98.876 99.924 99.985 102.120 99.590

0.075 0.075 0.060 0.052 0.124 0.119 0.121 0.106 0.107 0.114 0.084 0.092 0.088 0.096 0.089 0.099 0.119 0.125 0.123 0.120 0.102 0.064

0.719 0.658 0.660 0.657 0.766 0.776 0.752 0.749 0.754 0.767 0.705 0.698 0.684 0.687 0.691 0.683 0.761 0.766 0.780 0.788 0.797 0.772

^sps

V A

Fe/ (Fe + Mg)

0.039 0.014 0.015 0.039 0.045 0.045 0.058 0.067 0.065 0.069 0.144 0.128 0.133 0.124 0.137 0.112 0.034 0.031 0.034 0.035 0.036 0.048

0.168 0.253 0.264 0.251 0.066 0.059 0.069 0.077 0.075 0.049 0.068 0.081 0.095 0.094 0.084 0.106 0.086 0.078 0.063 0.058 0.065 0.116

0.906 0.898 0.917 0.927 0.861 0.867 0.861 0.876 0.876 0.870 0.894 0.883 0.886 0.877 0.886 0.873 0.865 0.859 0.864 0.868 0.887 0.923

gro

* ferrous iron pyr, pyrope; aim, almandine; sps, spessartine; gro, grossular

JSM-6400 electron microscope with energy dispersive system (EDS) detector at 15.0kV acceleration voltage and 1 nA beam current. In most samples, garnet homogeneity has been evaluated either by traverses or several spot analyses. One to three mineral associations (garnet + biotite; garnet + biotite + plagioclase with stable kyanite) have been selected in each sample for analysis. The results are presented in Tables 2-4. For thermobarometric estimates, the garnet-biotite geothermometer (model 5AV, Holdaway 2000) and the GASP geobarometer (Holdaway 2001) have been applied. Both tools are

calibrated on the basis of recent thermodynamic models. Correction for ferric iron has been carried out according to average values for pelitic schists (Holdaway et al. 1997), and for Ti + and Mn2+ in biotite according to the measured contents. Due to difficulties with error propagation within complex thermodynamic models, errors are estimated by the ability to reproduce the input values for the thermodynamic model (Holdaway 2000, 2001). Derived 2 a errors of c. 25 °C for geothermometry and c. 0.08 GPa for geobarometry (Holdaway 2000, 2001) have been used in this study.

Table 3. Biotite compositions from microprobe analysis (cations per 12 oxygens) Sample number

Si

Al

Ti

Mg

1Fe*

Mn

Ca

Na

K

^Oxide

Xphi

^ann

X]MgBT

Fe/ (Fe + Mg)

G48A G55A G56A1 G56 A2 G72A G76B G77B G77D1 G77D2 G78B1 G78B2 G78C1 G87B1 G87B2 G87C G87D1 G87D2 TK26 TK34

3.006 2.995 2.984 2.991 2.986 2.944 2.986 2.942 2.926 2.965 2.931 2.968 2.978 2.930 2.921 2.927 2.913 2.930 2.960

1.767 1.669 1.770 1.754 1.876 1.860 1.817 1.868 1.842 1.802 1.818 1.804 1.873 1.913 1.927 1.900 1.903 1.829 1.834

0.106 0.155 0.160 0.163 0.115 0.124 0.130 0.136 0.159 0.159 0.165 0.150 0.159 0.146 0.159 0.153 0.157 0.150 0.143

1.198 1.244 1.039 1.040 1.233 1.247 1.245 1.227 1.243 1.132 1.115 1.103 1.244 1.263 1.204 1.256 1.259 1.043 0.804

1L462 1.464 .499 .499 .269 .337 .303 .329 .326 .414 .453 .444 .212 .235 .274 .252 .273 .482 .663

0.003 0.003 0.008 0.009 0.000 0.001 0.001 0.003 0.004 0.011 0.011 0.014 0.002 0.000 0.001 0.001 0.000 0.002 0.003

0.000 0.001 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.002 0.000 0.001 0.000 0.001 0.000 0.000 0.000 0.001 0.008

0.026 0.011 0.024 0.020 0.057 0.049 0.075 0.044 0.051 0.024 0.018 0.019 0.019 0.068 0.024 0.035 0.054 0.043 0.016

0.898 0.957 0.998 1.008 0.909 0.933 0.909 0.923 0.933 0.959 0.981 0.974 0.899 0.895 0.920 0.922 0.893 1.095 1.115

95.057 94.091 94.878 95.030 94.332 94.162 96.055 93.850 93.849 95.054 95.197 95.399 94.081 95.283 93.200 94.406 94.056 95.860 96.540

0.450 0.459 0.408 0.408 0.493 0.482 0.488 0.480 0.483 0.443 0.432 0.431 0.506 0.506 0.486 0.501 0.497 0.413 0.326

0.549 0.540 0.589 0.588 0.507 0.517 0.511 0.519 0.515 0.553 0.563 0.564 0.493 0.494 0.514 0.499 0.503 0.586 0.673

0.001 0.001 0.003 0.003 0.000 0.000 0.000 0.001 0.002 0.004 0.004 0.005 0.001 0.000 0.000 0.000 0.000 0.001 0.001

0.550 0.541 0.591 0.590 0.507 0.517 0.511 0.520 0.516 0.555 0.566 0.567 0.494 0.494 0.514 0.499 0.503 0.587 0.674

* ferrous iron phi, phlogopite; ann, annite; MgBT, Mg-biotite

65

METAMORPHISM OF KANIN PENINSULA Table 4. Plagioclase composition from microprobe analyses (cations per 8 oxygens) Sample number

Si

iy

Ti

Mg

Fe*

Mn

Ca

Na

K

^Oxide

Xan

Xab

G72 A G76B G77B G77D1 G77D2 G77F G78B1 G78B2 G78C1 G78C2 G78C3 G78C4 G87B G87C G87D1 G87D2

2.707 2.741 2.780 2.777 2.786 2.776 2.749 2.766 2.736 2.708 2.682 3.016 2.674 2.713 2.719 2.718

]L290 ][.258 .229 .236 .212 .221 .234 .234 .254 .270 .310 ().976 L.322 L.291 L.273 L.284

0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.001 0.000 0.001 0.000 0.000 0.000 0.000 0.000

0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.000

0.000 0.003 0.000 0.001 0.003 0.017 0.010 0.007 0.011 0.018 0.014 0.015 0.011 0.000 0.001 0.001

0.000 0.000 0.000 0.000 0.000 0.000 0.002 0.000 0.001 0.002 0.003 0.002 0.001 0.000 0.000 0.000

0.283 0.280 0.226 0.221 0.216 0.222 0.262 0.238 0.266 0.308 0.331 0.005 0.302 0.278 0.289 0.272

0.734 0.693 0.738 0.737 0.780 0.757 0.754 0.732 0.730 0.692 0.636 0.942 0.703 0.712 0.716 0.724

0.003 0.002 0.004 0.004 0.003 0.001 0.002 0.007 0.004 0.011 0.009 0.018 0.006 0.007 0.006 0.005

99.468 99.365 98.894 100.461 102.200 99.520 99.561 98.952 97.590 99.545 100.277 98.976 99.949 98.891 98.828 100.664

0.277 0.287 0.233 0.230 0.216 0.227 0.257 0.244 0.266 0.305 0.339 0.005 0.299 0.279 0.286 0.272

0.720 0.711 0.763 0.766 0.781 0.772 0.741 0.750 0.730 0.684 0.652 0.976 0.695 0.714 0.708 0.723

V

^or

0.003 0.002 0.004 0.004 0.003 0.001 0.002 0.007 0.004 0.011 0.009 0.019 0.006 0.007 0.006 0.005

* ferrous iron an, anorthite; ab, albite; or, orthoclase

Analytical data and interpretation The results of the electron microprobe studies are summarized in Table 5 and Fig. 7. The composition of the analysed minerals in the various samples are very similar (Tables 2-4). Garnets are almandine rich (65-75%, occasionally up to 80%) with a Fe/ (Fe + Mg) ratio of c. 85%. Linear traverses over one of the larger (c. 2 cm) garnets of G72 and spot analyses on all other samples analysed at Uppsala University show that individual garnets are compositionally homogeneous, with slight changes only in the outermost rim. Changes are within a few percent; generally, a Xsps and Fe/(Fe + Mg) increase and a Xpyr decrease is recognized in the samples near the culmination of the Mikulkin Antiform; a Xpyr increase and a Fe/(Fe + Mg) decrease occurs in the samples farther up-section. Biotite contains c. 50%, occasionally up to 60%, annite and 50%, occasionally down to 40%, phlogopite. Plagioclase is generally oligoclase, close to the compositional range of andesine. Consistent results have been obtained for the two locations of highest metamorphic grade (kyanite-bearing) in the core of the Mikulkin Antiform. Three analysed mineral associations of sample G77 (Fig. 7a) give pressure estimates at the P-Tintersection points between 0.68 and 0.71 GPa, with temperatures between 588 °C and 603 °C. Two similar analyses of sample G78 result in somewhat higher estimates: 0.74-0.76 GPa at 615 °C (Fig. 7b). In one of these two analyses, minerals from a garnet-biotite plagioclase triple-junction have been used; the other was performed on minerals that were not in contact with each other. In a third analysis of G78, a substantially lower pressure of 0.60 GPa was obtained, but a comparable temperature of 590 °C. Comparison of all the pressure estimates results in a narrow overlap within 2a error (0.08 GPa, Holdaway 2001). Another kyanite-bearing sample (G87) located a few hundred metres higher in the section (Fig. 6), yields somewhat lower P-T estimates of 0.63 GPa at 587 °C and 0.65 GPa at 604 °C (for matrix biotite and biotite in contact to garnet, respectively; Fig. 7c). In contrast to these results, the pressures obtained from the crenulated mica-rich bands within G87 are as low as 0.51-0.53 GPa (Fig. 7d), but with temperatures of 581-586 °C comparable to the other results obtained from this sample. Kyanite is unstable or totally lacking in other analysed samples. Therefore, the GASP geobarometer was not used and temperatures were calculated for 0.50 GPa and 0.60 GPa (Table 4). The 0.60 GPa values are reported here, with temperature slopes of c. l-2°C/GPa(Fig. 7e).

Temperature estimates of G77 and G78 are supported by calculations for sample TK26 (Fig. 6) which yields a temperature of 607 °C (assumed 0.60 GPa; slope 2.5 °C/0.1 GPa). Garnet-biotite geothermometry results in 595 °C for G76 and 599 °C for G72 (Fig. 7e). The three samples from the two locations highest in stratigraphy yield lower temperatures (Fig. 7e). To the east, sample G56, a garnet-mica schist from the Tarkhanovskaya Group, yields temperatures of 556 °C and 575 °C. A similar result of 579 °C was obtained from the associated garnet amphibolite (G55). Farther NE, from high in the Mikulkinskaya Group, the garnet schist G48 yields a temperature of 568 °C. Finally, a garnet-mica schist from the top of Peschanaya Hill yields a temperature estimate from the staurolite-garnet schist TK34 of 604 °C. Plagioclase inclusions have been found in garnet in the samples G77 and G78. In both samples, pressure estimates for the inclusions are close to the estimates for matrix plagioclase.

Discussion Conditions of metamorphism As expected from field evidence, the kyanite-garnet schist G77 and the garnet schist G78, both from the same location in the core of the Mikulkin Antiform, yield the highest P-T estimates, 0.68 to 0.76 GPa at 588 to 615 °C. These are higher than the estimations for G87 (0.62 to 0.65 GPa at 587 to 604 °C) which is located c. 500 m higher in the structure. The significance of this estimate is difficult to assess; because of intensive folding, there is no basis to estimate how much the sequence was thickened by isoclinal folding at the time of peak metamorphism or subsequently thinned. Analyses from the crenulated mica-rich band in G87 (0.51 and 0.53 GPa at 581 and 586 °C) are entirely within the sillimanite stability field. Kyanite is in close proximity to the analysed spot. These low pressure values are local within the sample and not observed in any other analysis; they might be related to the moderate-P high-7 cordierite-sillimanite assemblages reported by Kasak et al. (1989) for the same area. The temperature estimates of c. 610 °C presented above are supported by the results from sample TK26 (Table 4). The results of the two kyanite-staurolite-garnet schists, G72 and G76, are comparable with the calculated temperatures of G77, G78 and G87 (Table 4). These samples apparently reached similar peak-metamorphic conditions. The location of Kasak's

66

H. LORENZ£TAL.

Fig. 7. Results of thermobarometric work, (a) to (d) Samples with stable kyanite from the centre of Mikulkin Antiform. (e) Samples without aluminosilicate minerals, (f) Pressure estimations from plagioclase inclusions in garnet. For sample location, see Fig. 6.

(1989) thermal dome (number 10 of the map on p. 66 in his article) coincides with the location of the above-mentioned five samples with kyanite-bearing assemblages. We interpret it as the deepest exposed level of the doubly-plunging Mikulkin Antiform (Fig. 3). Samples G55 and G56, from stratigraphically higher levels within the Mikulkin Antiform, yield similar results. Lower temperature estimates for the garnet-mica schist and the garnet amphibolite (Table 4), the only samples from the Tarkahanov Group, meet our expectations from field observations of decreasing metamorphic grade at higher levels in the structure (Figs 3 & 5). Kasak et al.'s (1989) metamorphic map shows a metamorphic break across the southeastern Kanin Peninsula, interrupting the general trend of the stratigraphically conformable facies boundaries. We have found no evidence to support this interpretation.

The isograds are folded by the Mikulkin Antiform and related folds. According to our P-Tdata, the Mikulkinskaya Group reached a depth of at least 25 km. This estimate assumes an average crustal density of 2.8 g cm~ 3 (Philpotts 1990) and implies a low geothermal gradient (here c. 24 °C km"1) typical for crust thickened by thrust-stacking during orogeny (Philpotts 1990). Deformation under these conditions resulted in the isoclinal folding. A local heat source, as suggested by some workers (thermal dome, Kasak et al. 1989; gneiss dome, Olovyanishnikov 2001), is not necessary to explain these typical regional metamorphic parageneses and hinterland-style structures. Evidence for tectonic thickening on the Kanin Peninsula are thrusts within the Tarkhanovskaya and Tabuyevskaya Groups and the apparantly thrust-related nature of the Mikulkin Antiform (Fig. 5). Thrust-stacking is also favoured

67

METAMORPHISM OF KANIN PENINSULA

Table 5. Results from thermobarometric calculations

Sample number

P (GPa)

T(°C)

Sample number

T! (°Q at 0.50 GPa*

T2 (°C) at 0.60 GPa*

Sample number

Pi (GPa) at 600 °Ct

p2 (GPa) at 650 °Cf

G77B G77D1 G77D2 G78B1 G78B2 G78C G87B1 G87B2 G87C G87D1 G87D2

0.6795 0.6926 0.7087 0.6023 0.7423 0.7624 0.6278 0.6290 0.6514 0.5309 0.5139

603.08 591.85 588.38 589.83 615.54 614.90 587.23 587.87 603.82 585.92 581.25

G48 A G55A G56A1 G56 A2 G72A G76B TK26 TK34

565.79 576.94 572.84 553.65 595.75 592.73 604.92 601.56

568.02 579.18 574.95 555.72 598.90 594.99 607.31 603.89

G77F G78C2 G78C3 G78C4

0.5681 0.6717 0.5934 2.4084

0.6647 0.7730 0.6890 2.6178

* Temperature estimate at an assumed pressure, no pressure estimate for these samples f Pressure estimate at an assumed temperature, no temperature estimate for theses samples

by the report of glaucophane (Kasak et al 1989) in greenschist facies units overlain by garnet-biotite bearing units (in the uppermost part of the Tabuyevskaya Group). Late Neoproterozoic (or earliest Cambrian) glaucophane also occurs c. 700 km to the SE in the Uralian foreland fold-belt (Beckholmen et al 2004) and has also been identified as a detrital mineral in Devonian sandstones at Tsilma river, NE of the central Timan Range (O. S. Kochetkov pers. comm. 2000); these occurrences suggest that the high-P, low-T parageneses were developed regionally within the Timanide Orogen at approximately the same structural level. The turbidite-dominated successions of the Timan Range comprise the footwall to the thrust-emplaced igneous complexes with oceanic affinities, located beneath the Pechora Basin to the NE of the Pechora Fault. It is probable that the high-grade regional metamorphic assemblages of the Kanin Peninsula, northern Timan and the southern Pechora basement were developed during thrusting beneath these allochthons. Our investigations lack evidence for the cordierite-sillimanitebearing assemblages at Cape Mikulkin, reported by Kasak et al. (1989). A local heat-source ('thermal dome') remains a plausible explanation for a superimposed moderate-P, high-7 metamorphism and the development of pegmatites in this area, perhaps related to the intrusive rocks reported from Cape Bolshoy Rumyanichny (Fig. 2), northern Timan (Andreichev 1998, Larionov et al. 2004). However, rapid, more or less isothermal uplift after peak metamorphism is another possible explanation; this finds support in the lower pressures associated with the crenulated mica-rich bands of sample G87. High-grade regional metamorphism of similar grade as on the southeastern Kanin Peninsula has not been recorded elsewhere in the Timan Range, but is known from drillcores in the southern and central parts of the Pechora Basin (Ermolenko & Sobolev 1978). The new observations and interpretations of the structure and metamorphism reported here suggest that Timanian highgrade metamorphism is of regional importance, extending far beyond the borders of the Timan-Kanin area. This is of particular importance for interpretations of the seismic velocity structure of the Pechora Basin and Barents Sea regions. Assumptions are often made (e.g. Verba & Sakoulina 2001) that Timanian lithologies are only low-grade and have relatively low /7-wave velocities, concluding that underlying crust with higher velocities must be composed of pre-Riphean Palaeoproterozoic or Archaean crystalline basement. This is probably correct for the higher velocity crust below the Timanides of the Timan Range and along the northern coast of the Kola Peninsula, in the footwall of the main Timanian thrusts. It cannot be assumed for hinterland regions further to the north and east.

Timing of metamorphism Many attempts have been made to date the orogenic phases within the Timanides. Intrusion of the 610-620 Ma (Larionov et al. 2004) alkaline igneous complexes of northernmost Timan and the c. 560 Ma (Gee et al. 2000) granites intruding Pechora basement apparently preceded the final uplift and emplacement of the Timan Range allochthons onto the platformal successions of the craton margin (Grazhdankin 2004) and imply a late Vendian age for orogenesis in the frontal allochthons of the TimanKanin area. The large number of K-Ar muscovite ages of c. 665 Ma (Andreichev 1998) favour a late Neoproterozoic age for the high-grade metamorphism. K-Ar ages (Andreichev 1998) of biotite (e.g. 970 Ma) greatly exceeding those of muscovite (e.g. 520 Ma) in the same rock, suggest a problem with excess argon in the former. In the Timan-Pechora region, the oldest Palaeozoic strata above the Timanian unconformity are Late Cambrian in age (Moczydlowska-Vidal et al. 2004; Bogolepova & Gee 2004). Little metamorphosed red beds and acid volcanites have been discovered overlying other volcano-sedimentary complexes in drillcores penetrating beneath the Palaeozoic unconformity (Bogdanov et al. 1996); these have also been considered to be of Vendian age (Bogdanov et al. 1996).

Conclusions The structural development on southeastern Kanin Peninsula can be divided into two stages. Isoclinal folding closely related to the high-grade metamorphism and thrust-related structures (e.g. Mikulkin Antiform) developed during the emplacement of the allochthons onto the continental margin. New geothermobarometric studies, reported here, indicate that peak metamorphic conditions in the core of the Mikulkin Antiform were as high as c. 0.72 GPa and c. 610 °C. This implies a depth of at least 25 km reached by thrust-stacking during orogeny. Lower P and high-7 assemblages (Kasak et al 1989) superimposed on kyanite-staurolite grade schists and gneisses suggest isothermal decompression during uplift. The evidence that these rocks are emplaced on only low-grade metamorphosed pericratonic footwall successions of the East European Craton to the SW (Getsen et al 1987) implies rapid exhumation during thrust emplacement of the metamorphosed turbidites and igneous complexes. The new observations on metamorphism and structure point to the regional importance of Timanian high-grade metamorphism far beyond the borders of the Timan-Kanin region. This has

68

H. LORENZCTAL.

importance for interpretation of geophysical data and particularly the seismic velocity structure of the northeastern margin of the East European Craton. This publication is a contribution to the Europrobe Timpebar project. The fieldwork in northern Timan and on Kanin Peninsula (TIMAN International Expedition 2000) has been supported by the Swedish Polar Research Secretariat and the INT AS programme. We acknowledge Hans Harry son's help with the electron microprobe at Uppsala University and helpful comments on the manuscript by Per Gunnar Andreasson, Victor Melezhek, Alasdair Skelton and Victoria Pease. We also thank our field assistants Timur Babushkin, Robert Eriksson and Denis Kalinin.

References ANDREICHEV, V. L. 1998. [Isotope geochronology of intrusive magmatism in the northern Timans]. Russian Academy of Sciences—Ural Division, Ekaterinburg, 89 p [in Russian]. BECKHOLMEN, M. & GLODNY, J. 2004. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125-134. BELYAKOVA, L. T. & STEPANENKO, V. I. 1991. [Magmatism and Geodynamics of the Pechora Syneclise's Baikalid Basement]. News (Izvestiya) of the Academy of Sciences of the USSR, Moscow, 12, 106-117 [in Russian]. BOGDANOV, N. A., KHAIN, V. E., BOGATSKY, V. L, KOSTYUCHENKO, S. L., SENIN, B. V., SHIPILOV, E. V. & SOBOLEV, S. F. 1996. Tectonic Map of the Barents Sea region and the northern parts of the European Russia. Russian Academy of Sciences, Moscow. BOGOLEPOVA, O. K. & GEE, D. G. 2004. Early Palaeozoic unconformity across the Timanides, NW Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 145-157. DOVZHIKOVA, E., PEASE, V. & REMIZOV, R. 2004. Neoproterozoic island arc magmatism beneath the Pechora Basin, NW Russia. GFF, 126, 353-362. ERMOLENKO, Y. P. & SOBOLEV, V. K. 1978. [On the garnet and staurolite source of the Givetian deposits of the Timans]. Proceedings of the Institutes of Higher Education, Geology and Exploration, 9, 58-61 [in Russian]. GEE, D. G. 2004. The Barentsian Caledonides: Death of the High Arctic Barents Craton. In: SMELROR, M. & BRUGGE, T. (eds) Arctic Geology, Hydrocarbon Resources and Environmental Challenges. Abstracts and Proceedings of the Geological Society of Norway, 2, 48-49. GEE, D. G., BELIAKOVA, L., PEASE, V., LARIONOV, A., & DOVSHIKOVA, L. 2000. New, single zircon (Pb-evporation) ages from Venian Intrusions in the basement beneath the Pechora Basin, Northeastern Baltica. Polarforschung, 68, 161-170. GETSEN, V. G. 1978. [Structural evolution of the metamorphic complex of the Kanin Peninsula]. Geology and minerals of the northeast European part of the USSR. Syktyvkar, 60-64 [in Russian]. GETSEN, V. G., ANDREICHEV, V. L., & STEPANENKO, V. I. 1987. [Metamorphic evolution of the upper proterozoic complex of Timan, as inferred from geologic and geochronologic data.] Transactions (Doklady) of the U.S.S.R. Academy of Earth Sciences/Earth Sciences Section, 285, 103-106 [published in Russian 1985].

GRAZHDANKIN, D. 2004. Late Neoproterozoic sedimentation in the Timan foreland. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 37-46. HOLD AW AY, M. J. 2000. Application of new experimental garnet Margules data to the garnet-biotite geothermometer. American Mineralogist, 85, 881-892. HOLD AWAY, M. J. 2001. Recalibration of the GASP geobarometer in light of recent garnet and plagioclase activity models and versions of the garnet-biotite geothermometer. American Mineralogist, 86, 1117-1129. HOLDAWAY, M. J., BlSWAJIT, M., DYAR, M. D., GUIDOTTI, C. V. &

DUTROW, B. L. 1997. Garnet-biotite geothermometry revised: New Margules parameters and a natural specimen data, set from Maine. American Mineralogist, 82, 582-595. IVENSEN, Y. P. 1964. [Magmatism of the Timan and Kanin Peninsula]. Nauka, Moscow-Leningrad, 123 pp [in Russian]. KASAK, A. P., DYMNIKOVA, N. G., GORONSTAY, B. A. & YAKOBSON, K. E. 1989. Metamorphic Zoning of Riphean Formations in the TimanKanin Region. Sovetskaya Geologia, 7, 65-74. KOSTYUCHENKO, S. L. 1994. Crustal structure and tectonics model of the Tyman-Pechora basin by geological-geophysical data. In: LEONOV, Yu G., ANTIPOV, M. P., MOROZOV, A. F. & SOLODILOV, L. N. (eds) Tectonics and Magmatism of East-European Platform. Russian Academy of Sciences, Moscow, 121-133. LARIONOV, A. N., ANDREICHEV, V. L. & GEE, D. G. 2004. The Vendian alkaline igneous suite of northern Timan: ion microprobe U-Pb zircon ages of gabbros and syenite. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 69-74. MALKOV, B. A. 1966. [New data about the age of the pre-Silurian intrusive complexes of Timan Range and Kanin Peninsula.] Lectures of the Academy of Sciences of the USSR, 170(3) [in Russian]. MOCZYDLOWSKA, M., STOCKFORS, M. & POPOV, L. 2004. Late Cambrian relative age constraints by acritarchs on the post-Timanian deposition of Kolguev Island, Arctic Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 159-168. NOVITSKY, I. P. 1976. [Petrology of the metamorphic complex of the Kanin Peninsula and North Timan]. Abstract of candidate dissertations. Moscow State University, 24 pp [in Russian]. OLOVYANISHNIKOV, V. G. 1998. [Upper Precambrian of Timan and Kanin Peninsula]. Russian Academy of Sciences, Syktyvkar, 163 p [in Russian]. OLOVYANISHNIKOV, V. G. 2001. [Neoproterozoic of the North Timan and Kanin Peninsula]. Russian Academy of Sciences, Syktyvkar, 46pp. OLOVYANISHNIKOV, V. G., ROBERTS, D. & SIEDLECKA, A. 2000. Tectonics and sedimentation of the Meso- to Neoproterozoic Timan-Varanger Belt along the northeastern margin of Baltica. Polarforschung, 68, 267-274. PHILPOTTS, A. R. 1990. Principles of Igneous and Metamorphic Petrology. Prentice Hall, Upper Saddle River, 498 pp. VERBA, M. L. & SAKOULINA, T. S. 2001 The reconstruction of the early paleozoic structure of the Barents Sea sedimentary basin inferred from geophysical surveys along Profile I-AR. Polarforschung, 69, 85-94.

The Venetian alkaline igneous suite of northern Timan: ion microprobe U-Pb zircon ages of gabbros and syenite ALEXANDER N. LARIONOV1, V. A. ANDREICHEV2 & D. G. GEE3 Isotope Research Centre, VSEGEI, Sredny Prospect 74, 199106 St Petersburg, Russia (e-mail: [email protected]) Geological Institute of the Komi Science Centre RAS, 54 Pervomaiskaya str., 167610 Syktyvkar, Komi Republic, Russia ^Institution for Earth Science, Geocentrum, Uppsala University, Villav. 16, 752 36, Uppsala, Sweden (e-mail: [email protected]) 1

Abstract: Near the Barents Sea coast in northern Timan, turbidites of probable Neoproterozoic age are intruded by pre-tectonic dolerites and a major suite of gabbros, granites and syenites (some nepheline bearing). Zircon ion microprobe dating of three plutons has yielded well-defined ages of 613-617 Ma. This alkaline igneous activity apparently represents a final phase of Vendian extensional magmatism prior to Timanian Orogeny. Previous work on late to post-orogenic calc-alkaline granites in the basement beneath the Pechora Basin, three hundred kilometres towards the SE, has yielded c. 550-560 Ma single zircon Pb-evaporation ages. These compositionally different intrusive suites are inferred to constrain the main phase of Timanian Orogeny to c. 610-560 Ma.

Throughout the southwestern margin of the Timanide Orogen from the Ural Mountains to the Varanger Peninsula, Neoproterozoic (possibly also Mesoproterozoic) successions are intruded by preorogenic dolerites and so-called gabbro-diabases. Their distribution varies greatly and their age has been difficult to constrain. In addition, there are a variety of deformed plutons of alkaline affinity and the one in northern Timan is the largest and the most variable in composition (Ivensen 1964; Kostyukhin & Stepanenko 1987). During the SWEDARCTIC 2000 expedition to the Timan and Kanin Peninsula (Gee 2000), samples of the Northern Timan Igneous Suite (NTIS) were collected for new isotope age studies from a wide range of compositions. The results from three intrusions are reported here; they throw light on the age of igneous activity and provide a better constraint on the timing of the Timanian Orogeny.

Geological setting Beneath a Devonian (locally Early Silurian) unconformity in the northernmost part of the Timan Range, metasedimentary successions, largely turbidites, are complexly folded and metamorphosed to greenschist facies. Higher grade, garnet and staurolite-bearing schists are known from further north on the Kanin Peninsula and also locally from drillholes that pass through the Palaeozoic succession and penetrate the underlying basement. The greenschist facies metasedimentary rocks are thought to be of Neoproterozoic, or perhaps late Mesoproterozoic (mid-late Riphean), age (Getzen 1975; Olovyanishnikov 2001 and references therein) and represent a continental slope-rise association deposited marginal to the peri-cratonic platform of the East European Craton (Roberts et al 2004). The NTIS (Fig. 1) ranges in composition from essexitic and olivine-kaersutite gabbros to granites, syenites and nepheline syenites with associated pegmatites and aplites. They are variably deformed and metamorphosed. Differences in the degree of alteration have led previous authors to believe that the NTIS might represent a period of intrusion that started prior to orogeny, continued through the latter and even post-dated the main tectono-thermal activity. An alternative hypothesis relates the variability in metamorphism and deformation to lithological heterogeneity and competence. Semi-penetrative deformation and high strain gradients leaves some intrusions little deformed and metamorphosed,

whereas others are highly sheared and recrystallized. Locally, deformation fabrics have been reported to be cut by felsic intrusions, but the relationships of these early fabrics to regional deformation and metamorphism have not been established. However, each of the wide range of rock types experienced some Timanian strain and recrystallization, ranging from the development of intense marginal foliations, boudinage and concordance with the dominating schistosity of the host metasediments, to preservation of primary intrusive relationships disturbed only by shear zones. The igneous petrology of the NTIS has been documented previously (Cherny et al. 1972) based on detailed mapping and extensive laboratory studies. No attempt has been made here to amplify these studies; the alkaline character of the suite is well defined and unambiguous. Primary mineral assemblages in the mafic rocks are locally well preserved, but widely altered to hornblende, epidote, biotite, microcline, muscovite and chlorite; in the felsic intrusions, secondary albite, microcline, muscovite, epidote and chlorite are typical. Many previous attempts to date the igneous rocks in the northern Timan Range were made by K-Ar (e.g. Malkov 1966; Ivensen 1964) and more recently, by Rb-Sr methods (Andreichev 1998). Gabbro-diabase dykes have yielded Rb-Sr whole-rock isochrons of 1100 ± 3 9 Ma and 693 ± 1 7 Ma. An olivine-kaersutite gabbro gave a K-Ar whole-rock age of 595-615 Ma, and a Rb-Sr whole-rock age of 702 ± 45 Ma. Syenites and granites yielded 610 ± 11 Ma and 587 ± 4 Ma, respectively and some alkaline gabbro dykes yielded c. 534 Ma all by the Rb-Sr whole-rock method. Various younger Rb-Sr and K-Ar ages on the intrusions, ranging down to c. 455 Ma, were interpreted as later Palaeozoic fault-related alteration. Older ages, in particular the c. 1100 Ma age for a metadiabase, raise the possibility that at least some of the metasedimentary host rocks might be of Mesoproterozoic age. Both the Rb-Sr and K-Ar methods, when applied to these rocks, are clearly susceptible to regional and local metamorphic and metasomatic processes and it was therefore decided to prioritize zircon dating.

Samples Samples from two gabbro bodies, a granite and two syenites were collected for zircon dating. Their locations are shown on Fig. 1. The samples are referred to as:

¥rom\ GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 69-74. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Geological map of northern Timan showing the location of the main instrusions of the Northern Timan Igneous Suite (NTIS) (after Olovyanishnikov 2001). Circled figures: 1, Krayny Kameshek massif; 2, Maly Kameshek massif; 3, Bolshoy Kameshek massif; 4, Sopka Bolvanskaya massif; 5, Kamennye Sopki massif.

a)

b)

The Rumyanishnaya River olivine-kaersutite gabbro (195 and G00:92). These samples were collected from the largest (c. 100 m across) outcrop of this intrusion, located in the mouth of Rumyanichnaya River. The Krayny Kameshek (137) and the Rumyanichny Cape (G00:91) syenites. The former was collected from the northern part of the Krayny Kameshek massif, whereas the latter was sampled on the shoreline, c. 200 m north of Rumyanichny Cape.

c)

The Bolshoy Kameshek granite (108) was collected from the southern part of this composite massif. d) The Bolshoy Kameshek gabbro (G00:40 & G00:42). These two compositionally equivalent specimens were sampled in the eastern part of the massif.

Zircons were separated following standard procedures using heavy liquids and an isodynamic magnetic separator. Optical microscopy revealed the following details. Most of the grains

VENDIAN NORTHERN TIMAN IGNEOUS SUITE

71

Fig. 2. Typical zircons from the dated rocks in transmitted light: (a) & (b) from the Bolshoy Kameshek granite (b possibly contains an inherited core); (c) & (d) from the Krayny Kameshek syenite; (e) & (f) from the Bolshoy Kameshek basite; (g) & (h) from the Rumyanichnaya River olivine-kaersutite gabbro.

are euhedral to subhedral and show no traces of resorbtion or overgrowth. In transmitted light, they exhibit oscillatory growth zoning (Fig. 2). The only inhomogeneity observed was in the zircons from the Bolshoy Kameshek granite, where a few grains with possible inherited cores have been noted (Fig. 2b). The rest of the samples show no inherited domains. During handpicking of the zircons for analysis, xenocrystic grains and those with complex structure were avoided. Cathodoluminescence (CL) imaging of sectioned zircons (Fig. 3) revealed that most of the zircons are CL-dark to very dark. They have pronounced to faint oscillatory zoning, characteristic for zircons of magmatic origin (e.g. Hanchar & Miller 1993). Isotope dating: methods and results As the first stage of this study, single-zircons from four plutons were analysed by the Pb-evaporation technique following the procedure described by Kober (1986) and many others (e.g. Hellman et al. 1997 and references therein). Zircons, free of inclusions and fractures, were handpicked and analysed at the Laboratory for Isotope Geology, Swedish Museum of Natural History, using a Finnigan MAT 261 mass spectrometer. These analyses yielded the following range of plateau ages of individual zircons (Larionov, unpublished data): 1) The Bolshoy Kameshek gabbro (G00:40 & G00:42): 595-630 Ma. 2) The Bolshoy Kameshek granite (108): 612-640 Ma. 3) The Rumyanishnaya River olivine-kaersutite gabbro (195): 600-624 Ma. 4) The Krayny Kameshek syenite (137): 594-619 Ma. In an attempt to obtain more precise control on the timing of intrusion of the NTIS, zircons from three of the intrusions were analysed using an ion microprobe facility (SIMS). In-situ U-Pb analyses were performed on a SHRIMP-II in the Centre of Isotopic Research (CIR) at VSEGEI, applying a secondary electron multiplier in peak-jumping mode. A primary beam of molecular oxygen was employed to bombard zircon in order to extract secondary ions. A 70 |xm Kohler aperture allowed focusing of the primary beam so that the ellipse-shaped analytical spot had a size c. 15 x 10 juim, and the corresponding ion current varied from

Fig. 3. Typical Cathodoluminescence images of the dated zircons from: (A) the Rumyanichny Cape syenite; (B) the Bolshoy Kameshek gabbro; (C) the Rumyanichnaya River olivine-kaersutite gabbro.

1.5 to 2.2 nA. The sputtered secondary ions were extracted at 10 kV. The 80 |jim wide slit of the secondary ion source, in combination with a 100 |xm multiplier slit, allowed mass-resolution M/AM > 5000 (1% valley); thus, all the possible isobaric interferences were resolved. Two minute rastering over a rectangular area of c. 30 x 20 |xm was employed before each analysis in order to remove the gold coating and possible surface, common Pb contamination. The following ion species were measured in sequence: 196 (Zr20)-204Pb-background (c. 204 AMU)-206Pb-207Pb208pb_238u_248Tho_254uo j^ cydes for each

analysed spot

were acquired. Apart from 'unknown' zircons, each fourth measurement was carried out on the zircon standard 91500. It has accepted values of U content of 81.2 ppm and a 207Pb/206Pb age of 1065.4 + 0.3 Ma (Wiedenbeck et al 1995). CL-dark areas of the 'unknown' zircons were chosen for analysis. The results collected were then processed with the SQUID vl.02 (Ludwig 2001) and ISOPLOT/Ex3.00 (Ludwig 2003) software, with decay constants of Steiger and Jager (1977). The common lead correction was applied on the basis of measured 204 Pb/206Pb and modern (i.e. 0 Ma) Pb isotope composition, according to the model of Stacey and Kramers (1975). The results of the zircon analyses are shown in Table 1 and Fig. 4. Eight analyses from both samples (G00:40 & 42) of the Bolshoy Kameshek gabbro were pooled and yielded a concordia

Table 1. Results of SIMS U-Pb analyses of the zircons from the NTIS

Spot

G0092-4-1 G0092-11-1 G0092-7-1 G0092-9-2 G0092-3-2 G0092-5 G0092-1 G0092-4 G0092-7 G0092-3 G0092-10 G0092-11 195-3 195-7 195-9 G0091-1 G0091-2 G0091-3 G0091-4 G0091-6 G0040-1 G0040-2 G0040-3 G0040-4 G0042-1 G0042-2 G0042-3 G0042-4

206

Pbc % 0.09 1.50 0.00 0.03 0.06 0.08 0.02 0.07 0.09 0.04 0.07 0.11 0.07 0.36 0.04 0.19 0.23 0.14 0.22 0.12 0.04 0.03 0.01 0.02 0.04 0.03 0.71 0.07

Concentration (ppm) 206pb*

107 97 149 171 265 131 159 256 148 257 229 153 221 88 226 51.2 56 72 97 79.4 162 134 229 140 228 397 72.4 208

U

Th

1255 1123 1740 1983 3062 1541 1867 2985 1718 2965 2658 1776 2601 1021 2600 598 639 837 1120 925 1861 1586 2632 1632 2657 4472 842 2423

90 113 120 251 405 295 318 144 107 183 801 422 1095 213 3017 190 148 294 333 410 3596 2032 4528 2501 1747 2662 658 1602

232Th

206pb

207pb

238U

238JJ

206pb

Age± (Ma)

Age ±: (Ma)

0.07 0.10 0.07 0.13 0.14 0.20 0.18 0.05 0.06 0.06 0.31 0.25 0.44 0.22 1.20 0.33 0.24 0.36 0.31 0.46 2.00 1.32 1.78 1.58 0.68 0.62 0.81 0.68

609 610 612 615 619 610 611 612 617 619 615 614 608 616 622 612 625 614

618 613 623 604 621 615 613 634 611 614

4.1 3.1 3.6 2.5 2.2 2.2 1.9 1.7 4.6 1.8 7.7 7.8 2.8 4.6 8.1 8.1 9.6 8 7.9 8.1 8.2 8.1 8.1 7.9 7.7 7.9 8.2 7.8

652 664 683 601 632 572 603 611 610 629 604 588 604 661 612 601 595 587 575 591 610 601 648 621 613 624 564 610

D % 40 84 29 20 28 20 17 18 43 15 14 23 28 74 23 34 33 26 25 26 15 16 22 15 13 9.6 59 14

7 9 12 -2 2 -6 -1 0 -1 2 -2 -4 -1 7 -2 -2 -5 -4 -7 -4 -2 -1 4 1 0 -2 -8 -1

238

U

±%

207pb*

±%

207pb*

±%

206pb*

206pb*

206pb*

235U

238U

t

t

t

t

10.095 10.073 10.050 9.992 9.925 10.080 10.067 10.034 9.964 9.930 9.990 10.010 10.115 9.972 9.870 10.040 9.820 10.000 9.940 10.020 9.860 10.180 9.880 9.990 10.020 9.680 10.070 10.010

Note: Errors are 1-sigma; Pbc and Pb* indicate common and radiogenic lead, respectively. Error in the standard calibration was 0.46. t Common Pb corrected using measured 204Pb. D%, Discordance, percent EC, Error correlation

0.7 0.53 0.62 0.43 0.37 0.37 0.32 0.29 0.78 0.31 1.3 1.3 0.48 0.78 1.4 1.4 1.6 1.4 1.3 1.4 1.4 1.4 1.4 1.3 1.3 1.3 1.4 1.3

0.0614 0.0617 0.0623 0.0599 0.0608 0.0591 0.0600 0.0602 0.0602 0.0607 0.0600 0.0596 0.0600 0.0616 0.0602 0.0599 0.0597 0.0595 0.0592 0.0596 0.0602 0.0599 0.0612 0.0605 0.0603 0.0606 0.0589 0.0602

1.9 3.9 1.4 0.9 1.3 0.9 0.8 0.8 2.0 0.7 0.6 1.1 1.3 3.5 1.1 1.6 1.5 1.2 1.2 1.2 0.7 0.7 1.0 0.7 0.6 0.4 2.7 0.6

0.838 0.845 0.854 0.827 0.845 0.809 0.822 0.827 0.833 0.843 0.828 0.821 0.818 0.852 0.841 0.823 0.839 0.821 0.821 0.821 0.842 0.812 0.854 0.835 0.829 0.863 0.807 0.829

1.99 3.96 1.50 1.01 1.37 0.99 0.84 0.89 2.14 0.75 1.5 1.7 1.40 3.54 1.7 2.1 2.2 1.8 1.8 1.8 1.5 1.6 1.7 1.5 1.4 1.4 3.1 1.5

0.099 0.099 0.100 0.100 0.101 0.099 0.099 0.100 0.100 0.101 0.100 0.099 0.099 0.100 0.101 0.099 0.102 0.100 0.101 0.099 0.101 0.098 0.101 0.100 0.099 0.103 0.099 0.099

±%

EC

0.70 0.53 0.62 0.43 0.37 0.37 0.32 0.29 0.78 0.31 1.3 1.3 0.48 0.78 1.4 1.4 1.6 1.4 1.3 1.4 1.4 1.4 1.4 1.3 1.3 1.3 1.4 1.3

0.35 0.13 0.41 0.42 0.27 0.38 0.38 0.32 0.37 0.41 0.90 0.78 0.35 0.22 0.79 0.67 0.73 0.75 0.76 0.75 0.89 0.89 0.81 0.89 0.91 0.95 0.46 0.90

VENDIAN NORTHERN TIMAN IGNEOUS SUITE

73

Fig. 4. U-Pb concordia plots showing zircon analytical results for the Bolshoy Kameshek gabbro, the Rumyanichnaya River olivine-kaersutite gabbro, and the Rumyanichny Cape syenite.

age of 617 ± 6 Ma (MSWD = 3.6, probability of concordance (POC) = 0.058). Fifteen analyses on zircons from the Rumyanishnaya River gabbro (195 & G00:92) gave a concordia age of 614 ± 2 Ma (MSWD = 0.075, POC = 0.78). Finally, from five analyses of the Rumyanichny Cape syenite, a concordia age of 613 ± 7 Ma was obtained (MSWD = 3.4, POC = 0.067). The previously obtained Pb-evaporation age estimations are in good agreement with these SIMS data.

Conclusions The ages obtained on these NTIS intrusions define a narrow time interval from c. 613 to 617 Ma. This suggests that the variability in deformation and metamorphic alteration of the intrusive suite is unlikely to be related to a changing tectonic environment, but is rather a heterogeneous response of the intrusions to subsequent Timanian deformation and metamorphism. Analysis

74

A. N. LARIONOV ETAL.

of a wider range of intrusions will be necessary to test this conclusion. These new ages on the NTIS are closely comparable to those of a similar alkaline igneous suite in the southern Urals (Alekseev & Alekseeva 1980). Zircons probably derived from this Early Vendian suite (Willner et al 2001, 2003) are found in the Late Vendian molasse of the Timanian foreland basin in the southern Urals (Puchkov 1997). The alkaline character of the NTIS favours a pre-orogenic origin, perhaps marking a final phase of crustal extension prior to the onset of Timanian deformation. This conclusion is consistent with that of other authors (e.g. Roberts & Siedlecka 2002; Olovyanishnikov et al. 2000) who have based their interpretations on the variety of published K-Ar and Rb-Sr mineral and wholerock ages. However, the existence of a few older ages by these methods on metamorphic rocks allows a more complex hypothesis involving earlier Neoproterzoic (probably early Vendian) metamorphism, separated from post-610 Ma orogeny by the NTIS extensional episode. Further east in the Timanide Orogen, in the basement beneath the Pechora Basin, single-zircon, Pb-evaporation dates of granites defined a concentration of intrusions at c. 560 Ma (Gee et al. 2000). These granites occur in drillcores from depths of about four kilometres. From observations of contact relationships to the metasedimentary host rocks in the cores, it has been inferred that the granites were late to post-tectonic in origin. Thus, it seems likely that the timing of Timanian Orogeny can be bracketed between c. 610 Ma and 560 Ma, perhaps continuing to the end of the Vendian (c. 545 Ma) further to the SW in the front of the orogen (Beckholmen & Glodny 2004). We thank the Swedish Polar Research Secretariat for support of the fieldwork and the Laboratory for Isotope Geology at the Swedish Museum of Natural History and Centre of Isotopic Research at VSEGEI for the use of laboratory facilities. Comments by MJ. Whitehouse, V. Pease, K. Ludwig and A.P. Willner resulted in major improvements to the manuscript, which is a contribution to the Europrobe program and the INT AS-HALE project no. 96:1941.

References ALEKSEEV, A. & ALEKSEEVA, G. 1980. [Vendian alkaligabbros and their potassic suite in the SW Urals.] Doklady Akademii Nauk SSSR, 255, 954-957 [in Russian]. ANDREICHEV, V. A. 1998. [Isotope geochronology of intrusive magmatism in Northern Timan.} Ekaterinbourg, 1-90 [in Russian]. BECKHOLMEN, M. & GLODNY, J. 2004. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia. In: GEE, D. G. & PEASE, V. (eds). The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125-134. CHERNY, V. G., SMIRNOV, V. G. & CHERNAYA I. P. 1972. [Formational analysis of the Timan igneous rocks.] Materials on geology and resources of North-East of the European part of USSR. Syktyvkar, 167-179 [in Russian]. GEE, D. G. 2000. Timan-Kanin international expedition. Cruise report. Polarforskningssekretariatets arsbok 2000, 69-73. GEE, D. G., BELIAKOVA, L., PEASE, V., LARIONOV, A. & DOVZHIKOVA, L. 2000. New, single-zircon (Pb-evaporation) ages from Vendian intrusions in the basement beneath the Pechora Basin Northeastern Baltica. Polarforschung, 68, 161-170. GETZEN V. G. 1975. [Structure of the basement of Northern Timan and Kanin peninsula.} Leningrad, Nauka, 1-144 [in Russian]. HANCHAR, J. M. & MILLER C. F. 1993. Zircon zonation patterns as revealed by cathodoluminescence and backscattered electron

images—implication for interpretation of complex crustal histories. Chemical Geology, 110, 1-13. HELLMAN, F. J., GEE, D. G., JOHANSSON, A. & WITT-NILSSON, P. 1997. Single-zircon Pb evaporation geochronology constrains basement-cover relationships in the Lower Hecla Hoek Complex of northern Ny Friesland, Svalbard. Chemical Geology, 137, 117-134. IVENSEN, Yu. P. 1964. [Magmatism of Timan and Kanin Peninsula.] Moscow, Leningrad, Nauka, 1-126 [in Russian]. KOBER, B. 1986. Whole-grain evaporation for 207Pb/206Pb-ageinvestigations on single zircons using a double-filament thermal ion source. Contributions to Mineralogy and Petrolology, 93/4, 482-490. KOSTYUKHIN, M. N. & STEPANENKO, V. I. 1987. [Baikalian magmatism of the Kanin-Timan Region.] Leningrad, Nauka, 1-232 [in Russian]. LUDWIG, K. R. 1999. Using Isoplot/Ex, Version 2.01: a geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication, la, 47. LUDWIG, K. R. 2001. SQUID 1.02, A User's Manual. Berkeley Geochronology Center, Special Publication, 4, 19 pp. LUDWIG, K. R. 2003. User's Manual for Isoplot 3.00. A Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center, Special Publication, 4, 74 pp. MALKOV, B. A. 1966. [New data on the age of the pre-Silurian intrusions of Timan and Kanin.] Transactions, Academy of Sciences, USSR, 170/3, 669-672 [in Russian]. OLOVYANISHNIKOV, V. G. 2001. [Neoproterozoic of the Northern Timan and Kanin Peninsula.} Syktyvkar, 1-46 [in Russian]. OLOVYANISHNIKOV, V. G., ROBERTS, D. & SIEDLECKA, A. 2000. Tectonics and sedimentation of the Neoproterozoic TimanVaranger Belt along the northern margin of Baltica. Polarforschung, 68, 267-274. 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. ROBERTS, D. & SIEDLECKA, A. 2002. Timanian orogenic deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian-Cadomian connections. Tectonophysics,352, 169-184. ROBERTS, D., SIEDLECKA, A., & OLOVYANISHNIKOV, V. G. 2004. Neoproterozoic passive-margin sedimentary systems of the Kanin Peninsula and northern and central Timan, NW Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30,5-18. STAGEY, J. S. & KRAMERS, J. D. 1975. Approximation of terrestrial lead isotope evolution by a two-stage model. Earth and Planetary Science Letters, 26, 207-221. STEIGER, R. H. & JAGER, E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. WlEDENBECK, M., ALLE, P., CORFU, F., GRIFFIN, W. L., MEIER, M.,

OBERLI, F., VON QUADT, A., RODDICK, J. C. & SPIEGEL, W. 1995. Three natural zircon standards for U-Th-Pb, Lu-Hf, trace element and REE analyses. Geostandards Newsletter, 19, 1-23. WILLNER, A. P., ERMOLAEVA, T., STROINK, L., GLASMACHER, U. A., GIESE, U., PUCHKOV, V. N., KOZLOV, V. I. & WALTER, R. 2001. Cotrasting provenance signals in Riphean and Vendian sandstones in the SW Urals (Russia); constraints for a change from passive to active continental margin conditions in the Neoproterozoic. Precambrian Research, 110, 215-239. WILLNER, A. P., SINDERN, S., METZGER, R., ERMOLAEVA, T., KRAMM, U., PUCHKOV, V. & KRONZ, A. 2003. Typology and single grain U/Pb ages of detrital zircons from Proterozoic sandstones in the SW Urals (Russia): early time marks at the eastern margin of Baltica. Precambrian Research, 124, 1-20.

Late Neoproterozoic granitoid magmatism in the basement to the Pechora Basin, NW Russia: geochemical constraints indicate westward subduction beneath NE Baltica 1

V. PEASE1, E. DOVZHIKOVA2, L. BELIAKOVA2 & D. G. GEE3 Department of Geology and Geochemistry, Stockholm University, SE-106 91 Sweden, and Laboratory for Isotope Geology, Swedish Natural History Museum, Stockholm, Sweden (e-mail: [email protected]) 2 Timan-Pechora Scientific Research Centre, Pushkina 2, 169400 Ukhta, Russia ^Uppsala University, Villavdgen 16, SE-752 36 Uppsala, Sweden

Abstract: A unique collection of deep (up to 4.5 km) drillcores from across the Pechora Basin of NW Russia indicate that the Precambrian basement beneath Palaeozoic sediments mainly comprise Neoproterozoic metasedimentary and magmatic complexes. These Neoproterozoic rocks are variously deformed and metamorphosed. They are intruded by granitoids of latest Neoproterozoic age (557 + 6 Ma, 95% confidence). The granitoids are variably deformed, metamorphose their country rock, and represent syn- to post-tectonic magmatism. Major element, trace element, and Sm-Nd and Rb-Sr isotopic data indicate that these intrusions represent continental arc magmatism. Their geochemistry indicates assimilation of older crust by a magma derived from a depleted mantle source. We infer that the granitoids were generated by westward subduction of oceanic crust beneath Baltica (present day coordinates) at c. 560 Ma.

Neoproterozoic complexes of diverse lithology occur throughout the 2000 km long eastern margin of the East European Craton (Zonenshain et al 1990). Within the foreland fold belt of the Uralide orogen, in the cores of late Palaeozoic anticlines, these Neoproterozoic successions generally occur beneath a major unconformity that separates early Palaeozoic platform (margin) deposits of Baltica from underlying Neoproterozoic (and locally older) complexes. The Proterozoic rocks within these anticlines provide evidence of folding and thrusting prior to Palaeozoic deposition. This late Neoproterozoic deformation along the eastern margin of Baltica is known as Timanian; it is generally accompanied by sub-greenschist to greenschist facies metamorphism, though locally it reaches amphibolite facies. Throughout the Southern and Middle Urals, the trend of Timanian-age structures is longitudinal. In the Northern Urals, the Timanian strike swings to the NW, across the Timan Range and along the northern edge of Baltica (Fig. 1), via the Kanin Peninsula and the southern Barents Sea to the eastern part of northern Norway (Roberts & Siedlecka 2002 and references therein). The Timanide Orogen (Tschernyschev 1901; Schatsky 1935; Getsen 1987) is a SW-verging fold and thrust belt in which Neoproterozoic basin successions are thrust over platform deposits of Baltica. In the Timan Range, folded and cleaved turbidites are unconformably overlain by Devonian sandstones (locally, Silurian limestones) and associated mafic volcanic rocks. To the NE in the Pechora Basin, hydrocarbon exploration drilling sampled basement on structural highs, in some cases at depths of 4.5 km (Beliakova & Stepanenko 1991). These drillcores, together with potential field data and seismic profiling (Kostiuchenko 1994), constrain interpretations of the basement beneath the Pechora Basin (Fig. 2). The Phanerozoic succession thickens to the NE (Beliakov 1994) and, beneath the late Palaeozoic and Mesozoic sequences of the Pechora Basin, both Silurian and Ordovician platform successions occur—sandstones (quartzites), carbonates, and shales, deposited along the passive margin of Baltica. Devonian rifting was associated with predominantly mafic magmatism. Rifting was followed by partial inversion in the Permo-Carboniferous, prior to deposition of thick foreland basin successions in the front of the Polar Urals (Lobkovsky et al 1996; Ismail-Zadeh et al. 1997). Neoproterozoic or older oceanic crust lies beneath c. 4 km of Palaeozoic sediments of the Pechora Basin (Dovzhikova et al.

2004). Further east in the Urals foredeep, Palaeozoic, Mesozoic and younger strata reach 10-12 km in thickness and the basement is of unknown character. In the Polar Urals, 400 km NE of the Timanian thrust front, pre-Ordovician complexes occur both in Palaeozoic allochthons and in the cores of Uralian foreland folds. In the Engenape anticline (Dushin 1997; Scarrow et al. 2002), island arc volcano-sedimentary successions and fragmented ophiolites are present. The latter are dated to c. 670 Ma (conventional U-Pb zircon; Khain et al. 1999). In the Polar Urals, both the Dzela and the Marun Keu complexes document accretion during Timanian orogenesis of Neoproterozoic oceanic crust (Remizov & Pease 2004; Glodny et al. 2004). Thus, it appears that late Neoproterozoic (Timanian) accretion and associated magmatism resulted in significant crustal growth of the NE East European Craton. To understand the Neoproterozic accretionary margin of the East European Craton better, new geochemical data are presented from the c. 560 Ma, late- to post-tectonic granitoids intruding the Precambrian basement beneath the Palaeozoic and younger deposits of the Pechora Basin. The tectonic significance of these geochemical results is then discussed. Precambrian geology of the Timan-Pechora region A great diversity of names referring to different structures in the basement of the Timan-Pechora region occur in the literature (e.g. Bogatsky et al. 1996). We refer to five major NW-trending belts from SW to NE, the Pericratonic region, the Timan Range, the Izhma zone, the Pechora zone, and the Bolshezemel zone (Figs 1 & 2). These zones are distinguished on the basis of a unique collection of drill core material from across the Pechora Basin and geophysical characteristics.

Pericratonic region SW of the Timan Range, the East European Craton is covered by Neoproterozoic platform successions. These are known from drill cores and outcrops, e.g. SW of the Kanin Peninsula, where socalled pericratonic carbonate, shale, and sandstone are in fault

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 75-85. 0435-4052/047$ 15 © The Geological Society of London 2004.

75

76

V. PEASE £7 AL.

Fig. 1. Tertiary regional setting and simplified Precambrian basement geology of the Pechora Basin. Geochemical sample distribution shown by intrusion number. Triangle indicates the location of the Enganepe anticline.

contact with basinal shale and turbidite. Devonian strata overlie, with near-parallel unconformity, late Neoproterozoic red sandstone and shale ('molasse'), which in turn rest unconformably on the underlying earlier Neoproterozoic successions (Olovyanishnikov et al 1997).

Timan Range Exposures in the Timan Range are dominated by basinal facies low-grade metasediment with steep NE-dipping cleavage (Olovyanishnikov et al. 1997; Roberts et al. 2004 and references

therein). The base of the succession is not exposed and the deepest structural levels are seen in the core of a major anticline in the northernmost part of the Timan Range and on the adjacent Kanin Peninsula. At these locations, a thick succession increases in metamorphic grade downwards to amphibolite facies (Olovyanishnikov et al 1997; Lorenz et al. 2004). Gabbro, granite, and syenite intrusions occur in the core of the structure (Olovyanishnikov et al. 1997); basement has not been recognized. The existing seismic data indicate that these metasediments have been thrust to the SW at least some tens of kilometres and that the East European Craton extends northeastwards beneath the Timan Range towards the Pechora Basin (Kostiuchenko 1994; Olovyanishnikov et al. 1995).

Fig. 2. Schematic SW-NE cross-section across the Pechora Basin in the Devonian. Ages shown on granitoid rocks from Gee et al. (2000). Note that late Neoproterozoic intrusion of post-tectonic granitoids occurred during passive margin sedimentation further west along the Baltoscandian margin of Baltica.

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77

Izhma zone

Sampling, isotopic ages, and petrography

Basinal facies of the Timan Range are known from drillholes to extend NE beneath the Palaeozoic cover of the Izhma zone (Figs 1 & 2). This folded biotite shale is variously intruded by both diabase-gabbro suites and granitoids. The former are generally pre-tectonic, i.e. deformed; the latter have been demonstrated in drillcores to be late- to post-tectonic. In one drillcore (Malaya Pera-11), shale unconformably overlying granite yielded Cambrian microfossils (Beliakova 1988).

The samples for this investigation are exceptional, coming from the few deep drillcores that penetrate basement c. 4 km beneath overlying Palaeozoic and younger sediments in the Pechora Basin. The seven granitoid samples have a relatively small volume (drillcore diameter 7 cm), are fine- to medium-grained, and distributed over a wide region. They provide the only detailed geochemical information on granitoid magmatism in the Pechora basement. Their age is well-known (discussed below). Most of the samples are of similar age and therefore probably document a single tectonic event. Consequently, information on the genesis of this magmatism and the basement hidden beneath the Pechora Basin, even from this small geochemical data set, is invaluable. Granitoid intrusions in this rare drillcore material are spread across the Izhma, Pechora, and Bolshezemel zones and have been sampled from below a regional Lower Ordovician (locally Cambrian) unconformity (Bogolepova & Gee 2004). Gee et al (2000) documented consistent single zircon Pb-evaporation ages for six of these granitoids with a mean of 557 ± 6 Ma (95% confidence), noted additional younger and older ages of c. 380 and 620 Ma, and identified significant zircon inheritance (summarized in Table 1). The granitoids superimpose contact metamorphism on cleavage of their host rocks and consequently represent a pulse of late- to post-tectonic Neoproterozoic III magmatism across the basement of the Pechora Basin. For our geochemical analyses, we used an aliquot of the same samples that Gee et al. (2000) used for zircon geochronology. The only exception is intrusion number eight in Tables 1-4, in which all of the initially collected material was used for zircon geochronology. Additional material of this sample was obtained from the drillcore. For intrusion number 6 (Table 1), there was no remaining material in the drillcore and geochemical analysis was not possible. We have used the same intrusion/sample nomenclature as Gee et al (2000; cf. Table 1).

Pechora zone Potential field data and aeromagnetic anomaly maps show a marked change in the character of the pre-Palaeozoic basement across the Izhma zone (Kostiuchenko 1994). Strong positive magnetic anomalies in the Pechora zone are related to pre-Ordovician basement magmatism, which increases in volume and changes in character (from solely intrusive to both extrusive and intrusive) and chemistry (intermediate to mafic) towards the NE. Detailed analyses of seismic profiling (mainly wide angle), in combination with the potential field data, define significant changes in the deeper basement and a major fault zone (perhaps a suture) has been inferred to separate the Izhma and Pechora zones (Beliakova & Stepanenko 1991). The Pechora zone intrusive complexes are intermediate to mafic in composition (Beliakova & Stepanenko 1991). For example, the Novaya-1 drillhole located near Pechora penetrated c. 270m of gabbro-diorites with associated plagiogranites (after the Russian usage). The host of volcano-sedimentary associations are tightly folded and metamorphosed to greenschist facies; mafic intrusions are altered to amphibolites. Bolshezemel zone The depth to basement generally increases towards the Polar Urals, NE of the Pechora zone. A broad basement uplift, however, located within c. 3 km of the surface defines the Bolshezemel arch. This structure gives its name to a zone that differs greatly in character from those to the SW. Several drillholes in the Bolshezemel arch penetrated pre-Ordovician basement, providing evidence of a volcano-sedimentary sequence of red sandstone and shale, volcaniclastic conglomerate, tuff and tuffaceous interlayers, rhyolite, and subvolcanic porphyritic and granophyric intrusions. Two-mica granite and gabbro intrude this volcanosedimentary association and the former contain xenoliths of the surrounding country rock.

Izhma zone From the Izhma zone drillcores, three intrusions were analysed for geochemistry (No. 1, South Charkaya-10; No. 3, Malaya Pera-11; and No. 5, East Charkaya-1). These granitoids may not be restricted to the Izhma zone given their late- to post-tectonic character; indeed they may have intruded after fault juxtaposition of the Izhma and Pechora zones. Izhma zone intrusions yielded a weighted mean age of 553 ± 4 Ma (95% confidence). The timing of fault juxtaposition is poorly constrained to be pre-Cambrian in

Table 1. Summary of sample numbers and age data from Gee et al. (2000) Intrusion no.

Izhma zone 1 3 5 6 Pechora zone 2 8 Bolshezemel zone 4

7

Drillhole ID

Sample depth (m)

Sample no.

Rock type

Age (2a) Magmatic

Inherited

>1013 + 9

South Charkaya-10 Malaya Pera-11 East Charkaya-1 Palyu-21

2952 3318 3219 3360

18 22 27 30

granitic granitic granitic dioritic

553 + 6 551 ± 8 557 + 15 560 ± 15

Mytni Materik-2 Novaya-1

3097 4320

19 62

granodioritic dioritic

378 ± 15 565 ± 8

>964 + 19

East Kharyaga-26

4450

26

granitic

567 ± 36

>1269 + 11 >1447 + 66

Veyak-2

4395

31

granodioritic

618 + 6

2708 + 26

78

V. PEASE ETAL

age on the basis of overlying sediments (Fig. 2) and the age of the cross-cutting granitoids. Petrographically, these samples are generally hypidiomorphic, with plagioclase (oligoclase, 40%) sometimes idiomorphic and dominating over K-feldspar (microcline, 15-20%), quartz (30%) and green (hydrated) biotite (5-10%). Apatite and zircon occur as accessory minerals. Primary hornblende may be present near the contact of the Pechora zone. Retrogression is generally minor (e.g. chloritization of biotite and saussuritization of plagioclase), though secondary muscovite and epidote (1-2%) occur in association with biotite in the Malaya Pera-11 sample (No. 3). Pechora zone Two intrusions from the Pechora zone were sampled. One (No. 8, Novaya-1) is a quartz diorite, a characteristic component of the mafic to intermediate igneous suites from this zone. The crystallization age of this intrusion was 565 ± 8 Ma (2a). The second (No. 2, Mytni Materik-2) is a highly altered (hybrid?) granodiorite, apparently of Devonian age (c. 380 Ma). The Novaya-1 sample (No. 8) is dominated by idiomorphic plagioclase (70%) with compositional zoning from andesine to oligoclase, subordinate quartz (15-20%), biotite (5-15%), and hornblende (2%). Accessory minerals include titanite, zircon, garnet, and magnetite. Sericite, clinozoisite, and chlorite are present as alteration products. Bolshezemel zone Two intrusions from the Bolshezemel zone were investigated (No. 4, East Kharyaga-26; No. 7, Veyak-2). Granitoids underlie Ordovician strata in the East Kharyaga-26 drillcore and are

567 ± 36 Ma (2a) in age. They are similar to those in the Izhma zone being hypidiomorphic, plagioclase-dominated (45%) intrusions, with quartz (30%), K-feldspar (20%), plus biotite and subordinate muscovite (5%). Accessory minerals include apatite, zircon, and magnetite. Alteration is extensive with pervasive saussuritization of plagioclase, chloritization of biotite, and some calcite. The sample contains abundant xenoliths of andesite porphyry. The Veyak-2 sample (No. 7) is more complex than the Kharyaga-26 granite. The drill core displays a variety of intrusive relationships into amphibole-bearing, diabase-gabbro with local hybridization. This granodioritic intrusion is c. 620 Ma, somewhat older than the late Neoproterozoic intrusions from elsewhere in the basement of the Pechora Zone. It is dominated by plagioclase (50%) and quartz (40%), with secondary epidote (10%) and accessory green hornblende, apatite, magnetite, and zircon.

Geochemistry Major element, trace (including rare earth) element, and isotopic (Rb-Sr and Sm-Nd) data from the late- to post-tectonic granitoids which intrude across the Pechora Basin are presented. Six of these samples are of similar age (c. 560 Ma) and therefore likely to represent a single tectonic event (cf Section 3). In addition, because of the rarity of the samples, we include the geochemical analyses from the Devonian dyke and the older (c. 620 Ma) intrusion (Tables 2, 3 & 4). These younger and older samples, however, are not directly relevant to Timanian orogeny and are omitted from subsequent diagrams and the following discussion. Sample locations are shown (Fig. 1) and analytical results are summarized (Tables 2, 3 & 4). Major element abundances in the text and figures that follow have been recalculated to 100% (volatile free).

Table 2. Major element oxides and normative mineralogy ofsyn- and post-tectonic granitoids Intrusion No. Sample No.

SiO2 TiO2 A1203 Fe2O3T MnO MgO CaO Na2O K2O P205 LOI Total

Pechora

Izhma

Bolshezemel

1 10SC18

3 11MP22

5 1EC27

2 2MM19

8 1N63

4 26EK26

7 2V31

65.97 0.86 14.52 5.61 0.08 1.46 2.74 3.24 3.72 0.28 1.31 99.79

68.84 0.40 15.42 3.67 0.04 1.18 0.74 3.36 3.67 0.14 2.43 99.89

66.86 0.68 14.30 4.83 0.07 1.67 2.86 3.16 3.62 0.23 1.52 99.80

68.03 0.61 11.67 7.59 0.06 1.79 1.55 4.60 0.67 0.15 3.12 99.84

55.76 1.54 15.15 8.74 0.14 4.24 5.29 3.88 2.53 0.41 2.03 99.71

72.50 0.19 12.94 1.74 0.06 0.64 0.57 0.31 8.87 0.11 1.89 99.82

63.73 0.91 12.47 6.90 0.08 1.93 3.26 3.53 4.14 0.25 2.61 99.81

5.4 0.0

34.2

0.5 21.4 26.7 12.7

15.0 32.8 16.4

52.5

CIPW normative mineralogy 23.7 q c 0.8 or 22.0 27.4 ab 11.8 an di 0.0 hy 8.0 2.2 mt il 1.6 ap 0.6

24.9

0.0 8.0 1.9 1.3 0.5

5.9 14.7

3.0 2.9 0.9

2.0 2.6 2.1 0.0 3.1 0.7 0.4 0.2

Note: All analyses by inductively coupled plasma atomic emission spectrometry (ICP-AES). Oxides as weight percent. Theoretical normative mineralogy (CIPW) calculated after Cox et al. (1979) for samples with LOI <2 wt%, using recommended values of Middlemost (1989) for oxidation state of iron.

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79

Table 3. Trace and rare earth element analyses of syn- and post-tectonic granitoids Izhma

Bolshezemel

Pechora

Intrusion no. Sample no.

1 10SC18

3 11MP22

5 1EC27

2 2MM19

8 1N63

4 26EK26

7 2V31

As Ba Be Bi Cd Co Cr Cs Cu Ga Ge Hf In Mo Nb Ni Pb (total) Rb

0.66 473 2.95 0.33 0.19 10.6 36.5 7.69 19.3 22.3 1.5 6.48 0.15 1.04 13.2 19.6 20.9 161 0.22 4.41 206 1.18 11.2 3.77 60 0.63 25.1 67.1 260 23.9 51.1 6.08 24.0 5.26 1.39 4.48 0.79 4.34 0.85 2.35 0.37 2.50 0.39

0.83 264 4.27 1.9 0.04 8.2 27.8 9.92 6.8 22.1 1.9 4.24 0.12 0.73 12.7 15.7 20.0 168 0.18 9.27 67 1.86 13.9 6.74 33 0.82 28.0 45.2 149 29.2 62.2 7.33 27.6 6.27 0.75 5.40 0.89 4.75 0.96 2.59 0.40 2.84 0.38

0.64 599 2.80 0.15 0.22 8.6 49.2 4.05 10.7 18.9 1.4 6.53 0.15 1.28 13.1 22.2 20.7 139 0.16 3.94 169 1.34 14.1 3.00 52 0.55 27.2 62.5 234 36.7 74.3 8.40 33.4 6.39 1.46 5.27 0.87 4.83 0.99 2.43 0.38 2.56 0.42

0.48 108 1.57 0.02 0.21 3.7 34.1 0.50 7.63 27.4 1.7 12.10 0.22 1.65 12.0 18.2 2.28 24 0.13 4.98 103 1.54 6.72 1.71 1.1 0.80 88.0 30.9 438 30.2 72.7 9.75 46.6 12.8 3.26 12.81 2.27 14.67 3.29 8.56 1.32 8.76 1.34

0.39 722 1.38 0.13 0.17 25.5 150 3.13 35.0 20.0 1.4 6.36 0.12 1.71 17.5 23.9 6.67 120 0.14 2.67 327 1.26 7.62 1.40 177 1.2 32.1 103.3 280 36.3 74.2 8.58 33.9 6.93 1.89 5.92 0.88 5.19 1.12 2.93 0.45 2.91 0.45

1.58 691 1.77 0.14 0.00 2.2 34.4 9.11 8.20 14.4 1.9 2.95 0.07 1.35 11.3 18.1 8.28 135 0.18 1.86 39 1.20 8.91 1.31 11 0.58 32.8 12.6 90.0 45.9 81.1 8.49 31.0 6.40 1.36 5.52 0.84 4.78 1.02 2.71 0.41 2.89 0.44

0.83 571 1.45 0.30 0.00 9.3 53.9 2.58 18.3 21.0 1.5 10.6 0.19 1.69 14.7 20.9 14.9 133 0.12 5.12 75 1.27 17.0 2.94 61 2.25 64.0 69.0 407 50.7 118.2 14.6 56.4 12.7 1.99 10.74 1.84 11.21 2.30 6.02 0.95 6.43 0.94

Sb Sn Sr Ta Th U V

w

Y Zn Zr La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Note: Element abundances in parts per million (ppm). All analyses by inductively coupled plasma mass spectrometry (ICP-MS).

Analytical methods Major elements were measured by inductively coupled plasma atomic emission spectroscopy (ICP-AES) and trace and rare earth elements by inductively coupled plasma mass spectrometry (ICP-MS) at the CNRS commercial facility in Nancy, France. Whole-rock powders were digested with a concentrated acid mixture (HF, HNO3) and H3BO3 was added to dissolve the precipitate fluoride and neutralize the excess HF. Calibrations were made using reference samples and international standards. Relative standard deviations are about 1% for SiO2 and about 2% for the other major elements, except MnO and P2O5 (<5%). Relative standard deviations for trace elements are generally about 5%. The limits of determination were calculated on the basis of six times the average background after 60 measurements. Major element abundances are reported as weight percent oxide (wt%) and trace elements as parts per million (ppm). Total Fe as FeOT was calculated from total Fe measured as Fe2O3 (normalized to

100%) by molecular weight. Theoretical normative mineralogy (CIPW) was calculated after Cox et al (1979) for unaltered samples, i.e. loss on ignition less than about 2 wt%, with the modification that FeO and Fe2O3 follow the recommended values (by silica content) of Middlemost (1989). Chemical separation and mass spectrometric analyses of Rb, Sr, Sm, and Nd were carried out at the Laboratory for Isotope Geology, Naturhistoriska Riksmuseet, Stockholm. Typically 0.3-0.5 g of whole-rock powder was mixed with a 147Sm- 150Nd spike solution and dissolved with HF and HNO3 in Teflon bombs. Sr, Sm, and Nd were separated using standard chromatographic ion-exchange procedures. Isotopic ratios were determined on a Finnegan MAT261 mass spectrometer equipped with multiple Faraday cups. Rb and Sr concentrations (ppm) were determined by ICP-AES (see above). Precision for Sr and Rb concentrations is ±3% (2cr) and ± 8% (2cr), for 50 ppm and 5 ppm respectively. Samples were not spiked for Sr and the 87Rb-86Sr ratios were calculated using the measured 87Sr/86Sr ratios and determined

80

V. PEASE ETAL Table 4. Isotope geochemistry ofsyn- and post-tectonic granitoids Izhma

Pechora

Bolshezemel

Intrusion no. Sample no.

1 10SC18

3 11MP22

5 1EC27

2 2MM19

8

Rb (ppm) Sr (ppm) 87 Rb/86Sr 87 Sr/86Sr 87 Sr/86Sri Sm (ppm) Nd (ppm) 147 Sm/144Nd 143 Nd/144Nd eNdi

160.6 206.0 2.257 0.72554 0.70775 5.0 23.0 0.1326 0.51225 -3.0

167.9 66.6 7.330 0.75624 0.69867

139.1 169.0 2.386 0.72849 0.70954

24.2 130.0 0.538 0.70956 0.70666 12.5 44.8 0.1689 0.51281

119.9 326.6 1.063 0.71425 0.70569 6.8 33.9 0.1216 0.51244 1.6

5.9 29.9 0.1365 0.51212 -6.0

6.6 33.5 0.1193 0.51207 -6.2

4.7

1N63

4 26EK26

7

135.2 39.5 9.977 0.77709 0.69644

132.9 75.4 5.125 0.75149 0.70632 12.6 56.9 0.1343 0.51225 -2.6

6.6 31.5 0.1261 0.51248

2.0

2V31

Note: Chemical separation of Rb, Sr, Sm, and Nd used standard chromatographic ion-exchange procedures. Refer to text for details of instrument operating conditions, precision, and data reduction. Decay constants follow the convention of Steiger & Jager (1977) and Lugmair & Marti (1978). eNd parameters were calculated relative to CHUR (143Nd/144Nd = 0.512638; 147Sm/144Nd - 0.1966; Jacobsen & Wasserburg 1984).

concentrations of Rb and Sr. Within-run mass fractionation was corrected by normalization to an 88Sr/86Sr ratio of 8.37521. Error associated with 87Rb-86Sr ratios is about ±6% (2a). Replicate analyses of unspiked USGS BCR1 standard yielded 87 Sr/86Sr ratios indistinguishable from the accepted value of Gladney et al (1990) and no further correction was applied. External precision of 87Sr/86Sr ratios is ± 0.01 %. Initial 87Sr/86Sr ratios (87Sr/86Srj) were calculated by correcting for the amount of 87Sr produced by 87Rb decay since the formation of the rock. Decay constants follow the convention of Steiger & Jager (1977) and Lugmair & Marti (1978). Sm and Nd concentrations (ppm) were determined by isotope dilution. Within-run mass fractionation was corrected by normalization of Sm to a 149Sm/152Sm ratio of 0.51686 and Nd interference was monitored at mass 146; Nd was normalized to a 146 Nd/144Nd ratio of 0.7219, with Sm interference monitored on

mass 149. Error on 147Sm-144Nd ratios is less than 0.1%. Error on 143Nd/144Nd ratios is less than 0.002%. Replicate analyses of La Jolla standard yielded 143Nd/144Nd ratios somewhat lower (0.511826 + 0.000023, 2or) than the accepted value of 0.511854. Consequently, a correction factor of +0.000028 has been applied to all analyses. eNd parameters were calculated relative to chondritic unfractionated reservoir (CHUR) (143Nd/144Nd - 0.512638; 147Sm-144Nd = 0.1966; Jacobsen & Wasserburg, 1984); depleted mantle model ages (tDM) are after DePaoloetal. (1991). In the text and figures which follow, trace element data are normalized to normal mid-ocean ridge basalt (NMORB) compositions of Saunders and Tarney (1984) and Sun (1980). Rare earth element (REE) data are normalized to the chondrite values ofNakamura(1974). Alteration The granitoid rocks can be altered, as seen petrographically by the saussuritization of plagioclase and the chloritization of biotite. This alteration may be reflected in high values for loss on ignition (LOI) during the determination of structurally bound water (Table 2). Samples are regarded as altered, at least with regard to 'mobile' elements, if LOI is greater than about 2 wt%. Normative mineralogy has not been calculated for altered samples (cf. Table 2). Major and trace elements

Fig. 3. Classification of Pechora Basin granitoid rocks (after de la Roche et al 1980). Using millication proportions, Rl = 4Si -1 l(Na + K) - 2(Fe + Ti) and R2 = 6Ca + 2Mg + Al. Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

The late Neoproterozoic (c. 560 Ma) granitoid rocks intruding the Precambrian basement of the Pechora Basin are heterogeneous. They include syenogranite, monzogranite, granodiorite, and monzodiorite (Fig. 3; after de la Roche et al. 1980). Potassium varies from 2.5 to 9 wt%, at 57 and 74 wt% SiO2, respectively (Fig. 4). The major oxides, with the exception of K^O and possibly A^Qs, show coherent negative correlations with silica (Fig. 4) and suggest that the 'less mobile' elements have not been significantly affected by alteration. The inverse correlation of Rb and Sr (Fig. 5) is consistent with plagiocase fractionation. These granitoids show Fe-depletion characteristic of the magnesian differention trend of Frost et al. (2001), i.e. the 'traditional' calc-alkaline magmatic trend of Irvine & Baragar (1971) (Table 2, Fig. 6). They are calc-alkaline to alkali-calcic according to the more comprehensive modified alkali-lime index (MALI) of

WESTWARD SUBDUCTION BENEATH BALTICA

81

Fig. 4. Major element oxides of Pechora Basin granitoid rocks, as weight percent (wt%). Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

Frost et al (2001). They are transitional (0.8-1.4 molecular % Al/Ca + Na + K) from metaluminous to peraluminous and are similarly metaluminous, peraluminous, and peralkaline according to the modified alumina saturation index (ASI) of Frost et al. (2001). They have relatively low (<2%) normative corundum abundances (Table 2), consistent with their low degree of alumina saturation and reflecting a low alumina source. The c. 560 Ma granitoids show large and relatively smooth NMORB-normalized trace element variations (Fig. 7). They document a large increase in large-ion lithophile (LIL) concentrations, about two orders of magnitude from Sr to Rb, relative to NMORB (Fig. 7). This is characteristic of continental crust (Taylor & McLennan 1981; Weaver & Tarney 1984), as well as subduction-related magmatism (e.g. Holm 1985). They also have high Ce concentrations (50-100ppm), consistent with the involvement of continental crust (Wilson & Davidson 1984). Negative Ba and Ti anomalies (relative to adjacent elements)

Fig. 6. Calc-alkaline affinity (after Irvine & Baragar 1971) of Pechora Basin granitoids. Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

Fig. 5. Large ion lithophile abundances in Pechora Basin granitoids. Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

Fig. 7. Trace elements normalized to NMORB (Saunders & Tarney 1984; Sun 1980). Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

V. PEASE ETAL.

Fig. 8. Rare earth elements normalized to chondrite (Nakamura 1974). Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

and pronounced Ta-Nb trough, however, indicate similarities with modern subduction-related arc magmatism. Other high field strength (HFS) element enrichments (e.g. Th) are also associated with modern subduction-related arc magmatism.

Fig. 9. Initial eNd versus initial 87Sr/86Sr isotope ratios (calculated using Pb/Pb ages of Gee et al. 2000). Arrow indicates intrusions number 3 and 4 plot off the graph. Depleted mantle array (DePaolo et al. 1991), solid line. Bulk earth (DePaolo & Wasserburg 1976), dashed lines. Reference fields after Morris and Hart (1983): IOA (intra-oceanic arcs), OIB (ocean island basalts). Continental arc of the Sierra Nevadas, western US (DePaolo 1981) shown for reference. Diamonds, samples from Izhma zone; squares, from Pechora zone. Open symbols have high loss on ignition. Only c. 560 Ma samples plotted.

Rare earth elements The c. 560 Ma granitoids generally show a moderate to high degree of fractionation in their chondrite normalized rare earth element patterns (Fig. 8). Absolute values of light rare earth elements (LREE) are high (La, 24-45), c. 100 times chondrite. Heavy rare earth element (HREE) concentrations are flat and similar to NMORB (Lu, 0.4-1.4). These samples show moderate enrichment in chondrite normalized LREE/HREE (LaN/LuN 6-11; Table 3). Steep patterns from La to Eu, with flatter HREE patterns are a characteristic feature of mature arc magmas (Green 1980).

Rb-Sr and Sm-Nd isotopes Initial ratios were calculated using the Pb evaporation ages from Gee et al (2000) and summarized in Table 1. In an initial 87 Sr/86Sr versus initial eNd plot (Fig. 9), the samples generally plot to the right of the depleted mantle array (DePaolo et al 1991) and towards the enriched 87Sr/86Sr and t43Nd/144Nd field. Intrusions numbered three and four have unrealistically low 87 Sr/86Sri ratios which plot off the diagram to the left of the mantle array. This may reflect LILE mobilization due to alteration. Consequently, no significance is attached to the Sr isotope ratios of these two samples. The other samples have 87Sr/86Srj ratios of 0.7057 to 0.7095 and eNdi of +1.6 to -6.2; they are consistent inasmuch as they reflect a notable 'crustal' component. Initial 143Nd/ Nd ratios (expressed in eNd notation in Table 4) for samples from the Izhma zone have negative eNdi values (-3 to —6.2). The sample from the Pechora zone has a positive value (+1.6), as does the sample from the Bolshezemel zone (+2). The negative eNdi values associated with the Izhma zone may suggest a longer crustal residence time (relative to the Pechora zone samples) and are within the range observed for isotopic variations of granites from active continental margins (e.g. +6.5 to —7.6 in the Cretaceous arc of western North America; cf. DePaolo 1981). The positive eNdi values from the Pechora and Bolshezemel zones suggest that these samples, at least, were derived from a depleted mantle source with a high time-integrated Sm-Nd ratio. Depleted mantle model age calculations (tDM; assuming the model of DePaolo et al. 1991) are similarly

correlative, with Izhma zone samples giving ages >1.6Ga, and the Pechora and Bolshezemel zones ages of 1.1 Ga (Fig. 10). The intrusion from the Bolshezemel zone has an isotopic signature (eNd of +2 and tDM of 1.1 Ga) similar to the Pechora zone samples and is close to the Pechora zone boundary (Fig. 1). This perhaps suggests that the boundary between the Bolshezemel/ Pechora zones is either broader than currently envisaged or is not precisely located. Discussion Sources, mantle and crust Though the data are few, there is good uniformity between the chondrite normalized REE patterns and NMORB-normalized trace element patterns of the granitoids. This, combined with

Fig. 10. Nd evolution diagram (after DePaolo et al. 1991). Depleted mantle (DePaolo et al. 1991) indicated by heavy line. Compositional fields after Morris and Hart (1983). Intrusion numbers indicated along right margin. Diamonds, samples from Izhma zone; squares, from Pechora zone; triangles, from Bolshezemel zone. Open symbols have high loss on ignition.

WESTWARD SUBDUCTION BENEATH BALTICA

their similar ages, suggests that they may be genetically related. The unfractionated and NMORB-like concentrations of HREE suggests that the c. 560 Ma granitoids were derived from a shallow (spinel stability field) depleted mantle source, most consistent with the positive initial sNd +1.6 from the Pechora zone sample. The marked enrichment in LILE (relative to NMORB) and enriched LREE relative to HREE, in conjunction with the isotopic data (see below), suggests assimilation of continental crust. In addition, the noted presence of xenocrystic zircons (Gee et al. 2000), as well as high Ce concentrations, are also consistent with assimilation of continental crust. Variations in the petrology and mineralogy of these granitoids probably reflect differences in either the amount or composition of crust assimilated. Strongly peraluminous melts, for example, are generally thought to have formed from a sedimentary source (Chappel & White 1974), but they may also form by melting of biotite-bearing metaluminous felsic rocks (Miller 1985) or even by water-excess melting of mafic rocks (Ellis & Thompson 1986). The range in initial e Nd for the c. 560 Ma granitoids, their tDM ages, their generally parallel Nd evolution curves, and their high 87 Sr/86Srj suggests that these intrusives could be related by mixing between a depleted mantle component (best represented by the Pechora zone sample with the most positive initial sNd value) and assimilation of either heterogeneous crust, or variable assimilation of older than 1.6 Ga crust. Gee et al (2000) noted the presence of xenocrystic Mesoproterozoic zircons in these granitoids (Table 1) and suggested the presence of Mesoproterozoic crust at depth. This is consistent with the negative eNd values and tDM ages presented here and is further evidence that material of probable Mesoproterozoic age was assimilated during the genesis of these granitoids. Isotopic data from the basement rocks hosting the granitoids is sparse. The Neoproterozoic (possibly older) passive margin sediments exposed in the Timan Range, and inferred to continue beneath the Izhma zone, are best characterized but lack similar isotopic data (see Precambrian geological description above). Mafic igneous rocks of the Pechora zone have OIB-affinities, with initial sNd values (assuming a minimum 560 Ma age) of +7 to -2 and initial 87Sr/86Sr of 0.703 to 0.711. Little is known of the deeper crust or the basement zones to the east.

Palaeotectonic setting Perhaps the most distinctive geochemical signature of arc magmas is their consistent enrichment of LILE, relative to Nb-Ta and other HFS elements. This feature is generally ascribed to enrichment of LILE in the asthenospheric mantle wedge above subduction zones, either from aqueous fluids that have been driven off by dehydration of the subducting oceanic crust and possibly accompanying sediments, or from partial melt derived from the same sources (Perfit et al. 1980; Pearce 1982, 1983; Arculus & Powell 1986). Active continental margin magmas share some of the petrological and geochemical characteristics of oceanic island arcs. Diagnostic features of active continental margin magmas seem to be higher absolute concentrations of the most incompatible HFS elements, such as Th, Ta, Nb, Zr, P and LREE. The c. 560 Ma granitoid intrusions from the basement beneath the Pechora Basin have major, trace, and rare earth element geochemistry, as well as Rb-Sr and Sm-Nd isotopes, more typical of a continental arc associated with subduction than with intraplate, rift-related volcanism. Negative Nb-Ta (and Ti) anomalies, typical of subduction-related magmas, constitute a clear feature of the NMORB-normalized trace element pattern of the c. 560 Ma granitoid intrusions (Fig. 7) which have NbN-CeN generally < 1.0 (0.5 to 1.1). Such a signature may also be generated by partial melting of an amphibolitized lower continental crust in a rift-related

83

tectonic environment. However, alkaline rocks and a ferroan magmatic differentiation trend typical of rift environments, but are not associated with the Pechora basin granitoids. Furthermore, the range in isotopic signatures suggests involvement (assimilation) of heterogeneous crust, or variable assimilation of >1.6Ga crust, yet negative Nb-Ta (and Ti) anomalies are common to all granitoids. These magnesian granitoids plot in the volcanic arc granitic fields of most tectonic discrimination diagrams (e.g. Fig. 11). In general, with increasing arc maturity, arc magmas are increasingly enriched in LILE and Th. Absolute concentrations of LILE + Th for the c. 560 Ma Pechora intrusions are moderate. Overall Fe and Mg abundances are greater than those associated with an immature arc, but less than those associated with a mature arc (Brown 1982). Similarly, their Rb/Zr ratios (0.41.5) and low Nb, but high Y concentrations, also suggest an arc of 'intermediate' maturity. Consequently, these late- to posttectonic granitoids are interpreted to represent in situ, subductionrelated magmatism which has assimilated continental crust in a continental arc setting. Their chemical composition and lithological association most likely represent the central arc or inboard of the central arc (Frost et al. 2001). In order for arc magmatism to intrude basement beneath the Pechora Basin, subduction at c. 560 Ma must have been directed westward beneath the Baltica margin. The across-arc compositional variation is non-systematic, e.g. negitive initial sNd and high initial 87Sr/86Sr occur both east and west of the Pechora zone, and reflects compositional differences of the host basement beneath the Pechora Basin. The well-defined and restricted age of arc magmatism suggests that it was short-lived. Arc magmatism was followed by uplift, erosion and deposition of Early Palaeozoic passive margin successions until the onset of Uralian orogenesis. It has been proposed (e.g. Olovyanishnikov et al. 1995; Kostiuchenko 1994) that the Pechora zone, with its concentration of mafic igneous rocks, marks the site of a late Neoproterozoic suture. The igneous rocks of the Pechora zone represent transitional oceanic to ocean island arc crust (Dovzhikova et al. 2004). Their crystallization age is unknown, but must be older than the 560 Ma granotoids intruding them. Pechora zone island arc rocks were probably thrust onto already accreted Mesoproterozoic basement (documented by the xenocrystic zircons and Nd model ages of the granitoids). Consequently, they must represent a pre-560 Ma, post-Mesoproterozoic suture between oceanic crust and the NE East European Craton. Arc-related granitoids intrude basement underlying most of the Pechora Basin. At c. 560 Ma, after the accretion of Pechora zone oceanic crust (Dovzhikova et al. 2004) and the lesserknown acid volcanic rocks of the Bolshezemel zone, shallow west-dipping subduction beneath the NE East European Craton occurred. Clearly, the trench associated with c. 560 Ma arc magmatism must have been located outboard of the granitoids. Waning continental arc magmatism at c. 560 Ma appears to coincide with the end of Timanian orogenesis.

Fig. 11. Tectonic discrimination diagram (after Pearce et al 1984). Note that all unaltered, c. 560 Ma granitoids plot in the volcanic arc field.

84

V. PEASE ETAL.

Conclusions Samples of c. 560 Ma, late- to post-tectonic granitoids intrude across about 200 km in basement to the Pechora Basin. They define short-lived magnesian, subduction-related magmatism typical of a continental arc, or inboard of it, and imply that west-dipping subduction was occurring beneath Baltica's NE accretionary margin in the late Neoproterozoic. The trench associated with this subduction would have been outboard/ oceanward. Subduction occurred towards the end of Timanian orogeny. It was followed by Cambrian uplift, erosion, and a return to passive margin sedimentation, recorded by the overlying Cambrian to Ordovician angular unconformity. The rare earth element patterns of the granitoids imply that they were derived from a depleted NMORB-type shallow mantle source, probably enriched by a subduction-related component. Isotopic differences between the granitoids likely reflect variable degrees of crustal assimilation of heterogeneous crust. This is consistent with depleted mantle model ages, ages of xenocrystic zircons, and high stronium isotope ratios indicating significant involvement of Mesoproterozoic crust in the genesis of these arc rocks. This paper is the result of an EuROPROBE-Timpebar project collaboration and has been supported by the Natural Science Research Council, the Royal Society, Uppsala University, the Swedish Museum of Natural History, an EU INTAS grant (No. 96-1941), and an EU Marie Curie postdoctoral research grant. The assistance of Marina Fisherstrom is gratefully acknowledged. Constructive reviews by G. Fershtater, J. Scarrow, and an anonymous reviewer improved the quality of the manuscript.

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Basin formation mechanisms and subsidence history. Tectonophysics, 266, 251-285. LORENZ, H., PYSTIN, A., OLOVYANISHNIKOV, O. & GEE, D. 2004. Neoproterozoic high-grade metamorphism of the Kanin Peninsula, Timanide Orogen, northern Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogeny of Eastern Baltica. Geological Society, London, 30 Memoirs, 59-68. LUGMAIR, G. & MARTI, K. 1978. Lunar initial 143Nd/144Nd: Differential evolution of the lunar crust and mantle. Earth and Planetary Science Letters, 35, 273-284. MIDDLEMOST, E. 1989. Iron oxidation ratios, norms and the classification of volcanic rocks. Chemical Geology, 77, 19-26. MILLER, C. 1985. Are strongly peraluminous magmas derived from pelitic sedimentary sources? Journal of Geology, 93, 673689. MORRIS, J. D. & HART, S. R. 1983. Isotopic and incompatible element constraints on the genesis of island arc volcanics from Cold Bay and Amak Island, Aleutians, and implications for mantle structure. Geochimica et Cosmochimica Acta, 47, 2015-2030. NAKAMURA, N. 1974. Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and ordinary chondrites. Geochimica et Cosmochimica Acta, 38, 757-775. OLOVYANISHNIKOV, V. G., BUSHUEV, A. S. & DOKHSANYANTS, E. P. 1995. [The structure of the conjugation zone of the Russian and Pechora plates from geological and geophysical data.] Transactions (Doklady) of the Russian Academy of Sciences, Earth Science Section, 351(8), 1228-1232 [in Russian]. OLOVYANISHNIKOV, V. G., SIEDLECKA, A. & ROBERTS, D. 1997. Aspects of the geology of the Timans, Russia, and linkages with Varanger Peninsula, NE Norway. Norges Geologiske Undersokelse Bulletin, 433, 28-29. PEARCE, J. 1982. Trace element characteristics of lavas from destructive plate boundaries. In: THORPE, R. S. (ed.) Andesites. Wiley, Chichester, 525-548. PEARCE, J. 1983. The role of sub-continental lithosphere in magma genesis at destructive plate margins. In: HAWKESWORTH, C. & NORRY, M. (eds) Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, 230-249. PEARCE, J., HARRIS, N. & TINDLE, A. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983. PERFIT, M., GUST, D., BENCE, A., ARCULUS, R. & TAYLOR, S. 1980. Chemical characteristics of island arc basalts: implications for mantle sources. Chemical Geology, 30, 227-256. REMIZOV, D. & PEASE, V. 2004. The Dzela complex, Polar Urals, Russia: a Neoproterozoic island arc. In: GEE, D. G. & PEASE, V. (eds) The

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Neoproterozoic Timanide Orogeny of Eastern Baltica. Geological Society, London, 30, Memoirs, 107-123. ROBERTS, D. & SIEDLECKA, A. 2002. Timanian orogenic deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian-Cadomian connections. Tectonophysics, 352, 169-184. ROBERTS, D., SIEDLECKA, A. & OLOVYANISHNIKOV, V. G. 2004. Neoproterozoic passive-margin sedimentary systems of the Kanin Peninsula, and Northern and Central Timan, NW Russia. In: GEE, D. G. & PEASE, V. (eds) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30,5-17. SAUNDERS, A. & TARNEY, J. 1984. Geochemical characteristics of basaltic volcanism within back-arc basins. In: KOKELAAR, B. & Ho WELLS, M. (eds) Marginal Basin Geology, Geological Society, London, Special Publications, 16, 59-76. SCARROW, J. H., HETZEL, R., GOROZHANIN, V., DINN, M., GLODNY, J., GERDES, A., AY ALA, C. & MONTERO, P. 2002. Four decades of geochronological work in the southern and middle Urals: a review. In: BROWN, D., JUHLIN, C. & PUCHKOV, V. (eds) Mountain Building in the Uralides: Pangea to present. AGU Geophysical Monograph, 132, 233-256. SCHATSKY, N. 1935. On Cambrian-Proterozoic relations and Baikalian Orogeny. Nauk SSSR, 1, 581-587 [in Russian]. STEIGER, R. & JAGER, E. 1977. Subcommission on geochronology: Conventions on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. SUN, S. 1980. Lead isotopic study of young volcanic rocks from mid-ocean ridges, ocean islands and island arcs. Philosophical Transactions of the Royal Society, A297, 409-445. TAYLOR, S. & MCLENNAN, S. 1981. The composition and evolution of the continental crust: Rare earth element evidence from sedimentary rocks. Philosophical Transactions of the Royal Society, A301, 381-399. TSCHERNYSCHEV, T. 1901. [On geological structure of the Timans and on relation of the Timan Fault to other regions of northern Europe.] Zapiski Mineralogicheskogo Obstchestva, 34, 29-33 [in Russian]. WEAVER, B. & TARNEY, J. 1984. Empirical approach to estimating the composition of the continental crust. Nature, 310, 575-577. WILSON, M. & DAVIDSON, J. 1984. The relative roles of crust and upper mantle in the generation of oceanic island arc magmas. Philosophical Transactions of the Royal Society, A310, 661-674. ZONENSHAIN, L., KUZMIN, M. & NATAPOV, L. 1990. Uralian Foldbelt. In: PAGE, L. (ed.) Geology of the USSR: A Plate-Tectonic Synthesis. Geodynamic Series 21, American Geophysical Union, Washington, District of Columbia, 27-54.

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Protolith ages of eclogites, Marun-Keu Complex, Polar Urals, Russia: implications for the pre- and early Uralian evolution of the northeastern European continental margin J. GLODNY1'4, V. PEASE2, P. MONTERO3, H. AUSTRHEIM4 & A. I. RUSIN5 GeoForschungsZentrum Potsdam, Telegrafenberg C2, 14473 Potsdam, Germany (e-mail: [email protected]) 2 Institute of Geology and Geochemistry, Stockholm University & Swedish Museum of Natural History, Box 50007, S-104 05 Stockholm, Sweden ^Department of Mineralogy and Petrology, University of Granada, Campus Fuentenueva, E-18002 Granada, Spain 4 lnstitutt for Geologi, Universitetet i Oslo, Postboks 1047, Blindern, 0316 Oslo, Norway 5 Institute of Geology and Geochemistry, Ubr RAS, Pochtovy Per. 7, Yekaterinburg 62015, Russia

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Abstract: Single zircon U-Pb dating by ion microprobe and Pb-evaporation techniques, combined with Sr-Nd isotopic and geochemical data, constrain the pre- and early Uralian evolution of the Marun-Keu eclogite-facies metamorphic complex, Polar Urals, Russia. The complex is interpreted as a part of the Early Palaeozoic East European passive margin, subducted, metamorphosed and exhumed during Uralian arc-continent collision. In the Marun-Keu complex, a variety of meta-igneous and metasedimentary rocks documents different evolutionary stages of the northeastern margin of Baltica. Both the oldest, polystage detrital zircon population, which is 570 to 700 Ma in age, and the dominant zircon population in an island arc-type magmatic suite (c. 550 Ma) are related to the Neoproterozoic Timanian orogeny, which resulted in continental growth of northeastern Europe by accretion of juvenile oceanic and island arc material. A distinct magmatic event at c. 490 Ma is documented in zircons from metagranites, and is interpreted as evidence for incipient postTimanian rifting which finally led to formation of both the Proto-Uralian ocean and the Early Palaeozoic East European passive margin. A convergent setting was re-established in Silurian times, culminating in Uralian arc-continent collision by east-dipping subduction of the East European passive margin. Metamorphic zircon ages reflect the collision event at 360 to 355 Ma.

The Urals form an orogenic belt stretching from the Arctic coast (68°N) towards the latitude of the Caspian Sea (48°N), a distance of c. 2500 km. It marks Devonian to Triassic collisional events amalgamating the East European plate (Baltica) with island arc domains and Siberian-Kazakhian terranes further east (Hamilton 1970; Ivanov et al 1975; Zonenshain et al 1984). Although the Uralian orogenic evolution has been well studied, much less is known regarding the pre-Uralian history of the East European continental margin. However, understanding the pre-Uralian history of the East European margin is of crucial importance for models of the Uralian Orogeny: Uralian structures are superimposed on pre-existing lithologies and structures, and the rheology of the pre-Uralian crust influenced the Palaeozoic Uralian orogenic evolution. Due to the ubiquitous cover of Phanerozoic platform and basin sediments, exposure of the pre-Uralian East European crust is largely restricted to outcrops immediately to the west of, or within the Urals. Here, different units and structures of the continental margin were reworked, metamorphosed and partly obliterated by Uralian processes. This paper investigates Late Devonian—Early Carboniferous eclogite facies rocks of the Marun-Keu metamorphic complex, Polar Urals (Fig. 1), with particular focus on the protolith history. The Marun-Keu complex forms part of a discontinuous, but persistent, high pressure metamorphic belt along the western flank of the Urals, in the footwall of the Main Uralian Fault. There are striking similarities between the different metamorphic complexes of this belt with respect to metamorphic grade, structure and tectonic position (Dobretsov & Sobolev 1984; Sobolev et al 1986; Puchkov 1997; Lennykh et al 1997). The Maksyutov complex (Southern Urals), Middle to Late Devonian in age (375 ± 2 Ma; Glodny et al 2002), and structurally equivalent to the Marun-Keu complex, has been interpreted as part of the leading edge of the East European continental margin during the east-directed Uralian subduction (Hetzel et al 1998; Hetzel 1999; Brown & Spadea 1999; Brown et al 2000). From this structural equivalence between the metamorphic complexes we expect the Marun-Keu complex to be metamorphosed crust of the pre-Uralian East European margin.

We present new ion microprobe U-Pb and Pb-evaporation zircon ages, Sm-Nd and Rb-Sr isotopic data, and geochemical data from the Marun-Keu complex. The objective of the study is to characterize the eclogite protoliths, and to contribute to a more profound understanding of the pre-Uralian and early Uralian geodynamic evolution of the East European continental margin in near-Polar regions. The data also provide a test of the response of the U-Pb system of zircons to eclogite facies metamorphic conditions. Geological setting A WSW-ENE profile across the Polar Urals shows the following structural entities (Fig. 1): the Timan-Pechora basin, the western foreland and slope of the Polar Urals, a high pressure (HP)-low temperature (LT) metamorphic belt, and the Main Uralian Fault (MUF). An ophiolite belt along the eastern flank of the Urals is overlain by island arc lithologies, representing the easternmost exposed part of the Uralides. Further east, Uralian structures are buried beneath sediments of the West Siberian Basin. We outline the current state of knowledge regarding these structures and their evolution below. In the following, age data are cited with 2a errors. The IUGS International Stratigraphic Chart (2000) is used throughout this paper. The Timan-Pechora region, west of the Polar Urals, is dominated by Late Cambrian and younger sedimentary rocks of the Pechora basin. For the underlying basement, seismic imaging reveals thinned crust, in contrast to the adjacent East European platform crust and to the marked crustal root of the Urals (Schueller et al 1997). The Timan-Pechora basement was assembled during the late Neoproterozoic Timanian orogeny, and peneplained by the Cambrian (Raznitsyn 1973). It consists of several, mostly geophysically defined domains (cf. Bogdanov et al 1996) and is characterized by a general NW-SE direction of strike. From drillcores and sparse outcrops it is known that island arc and oceanic rocks are widespread. Calc-alkaline effusives, together with a great variety of felsic to mafic plutonic

From. GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 87-105. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Geological map of the Polar Urals and the Timan-Pechora region (after Bogdanov et al 1996). MUF, Main Uralian Fault; E, Engane-Pe Anticline.

rocks, like granites with high proportions of mafic inclusions, gabbros, gabbro-diorites and K-poor granites predominate (Belyakova & Stepanenko 1991). Available isotopic age determinations from the Pechora basement (cf Gee et al. 2000) indicate a Neoproterozoic III time frame for Timanian orogeny, between 620 Ma (the age of the oldest granitic intrusions in the Pechora basin) and 550 Ma. Zircon ages for Pechora basement granitoids cluster at around 560 Ma (Gee et al 2000) and suggest that the Timanian orogeny migrated from the Timan Range towards the NE. Convergence, with SW-directed subduction, led to successive accretion of material to the northeastern margin of the East

European Craton (Scarrow et al. 2001; Pease et al. 2004). Timanian processes were followed by locally intense acidic volcanism, the products of which in places occupy the highest part of the pre-Ordovician sections, in particular in the Bolshezemelsky arch basement domain (Fig. 1; Belyakova & Stepanenko 1991; Bogdanov et al. 1996). Precise age data for this volcanism are not available. In the western foreland of the Polar Urals, Timanian basement crops out in the cores of Uralian frontal anticlines. In the Engane-Pe Anticline, some 60 km west of the Marun-Keu complex (Fig. 1), Timanian structures trend NW-SE, highly

PROTOLITH AGES OF ECLOGITES, MARUN-KEU, POLAR URALS

obliquely to the direction of strike in the Polar Urals foldbelt. The Engane-Pe basement rocks, in general, have oceanic and island arc affinity. Fragments of ophiolitic sequences are present (Dushin 1997; Scarrow et al 2001). Plagiogranite from a serpentinite melange yielded an U-Pb zircon age of 670 + 5 Ma (Khain et al. 1999), interpreted as the age of formation of oceanic crust. Scarrow et al. (2001) showed that obduction of the Engane-Pe ophiolite, and thereby termination of westward subduction of oceanic crust under the East European continent, must have occurred between 560 and 530 Ma. On a regional scale, the Timanian basement, as exposed on the western flank of the Polar Urals, is made up of flysch-like volcaniclastic sediments, volcanic rocks (andesites, porphyritic basalts, dacites and rhyolites), ophiolite-related rocks, greenschists, amphibolites and granitoids (Miklukho-Maklai & L'vov 1960; Yudovich et al. 1993; Dushin 1997, and own observations). The basement is largely covered by Early Ordovician and younger Pechora basin sediments. Early Ordovician tectonic activity is recognized from the sedimentological and palaeontological record (e.g. Yeliseyev 1978). In the limbs of the Engane-Pe Anticline (Fig. 1), Early Ordovician sedimentation started with basal conglomerates on folded and variably metamorphosed basement, and was accompanied by local formation of basaltic dykes. Later on, subsidence and sedimentation continued quietly until Middle Devonian times (Raznitsyn 1973). This geological evidence is interpreted in terms of Early Ordovician rifting, followed by passive margin formation, and deposition of continental margin and slope sediments (Savelieva & Nesbitt 1996; Nikishin et al. 1996; Saveliev 1997; Antoshkina 1999; Saveliev et al. 1999). With respect to the Uralian Orogeny, the western foreland of the Urals structurally represents an Uralian arc-continent collision accretionary complex, overprinted by later formation of a westfacing foreland fold and thrust belt (Brown & Spadea 1999;

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Saveliev et al. 1999). The Marun-Keu metamorphic complex relates to the Uralian arc-continent collision and most likely was, in analogy to the Maksyutov complex, exhumed in the rear of the accretion system. The Marun-Keu complex (Fig. 2), the target of this study, was described in detail by Udovkina (1971) and Molina et al. (2002). The complex is lens-shaped, with the long axis parallel to the regional strike of the Urals (NNE-SSW). It covers an area of approximately 300 km2, and is delimited by tectonic contacts. In the east, the east-dipping (cf. Savelieva & Nesbitt 1996; Bogdanov et al. 1996) Main Uralian Fault juxtaposes it against ultramafic rocks of the Syum-Keu complex. In the west, it is thrust over a suite of low-grade metasediments containing abundant quartz-graphite schists and some granitic and dioritic intrusions. The Marun-Keu complex itself comprises different lithologies showing blueschist- to eclogite-facies metamorphism (Fig. 2). In its southern and central part, it is dominated by metamorphosed plutonic rocks with minor metasediments. Eclogitefacies metamorphic conditions are estimated at 14-17kbar and 520-690 °C (Udovkina 1971; Molina et al 2002). The northernmost, blueschist-facies part of the complex ('Pike river complex') is made up of metasediments, metarhyolites and metabasic rocks (Udovkina 1971; Sobolev et al. 1986), and in places resembles a meta-olistostrome (Kazak 1981). Metamorphic conditions here are estimated at c. 10-11 kbar and 500 °C (Dobretsov & Sobolev 1984). The age of HP metamorphism, as constrained by Rb-Sr internal mineral isochrons on eclogites, is 355.5 ± 1.4 Ma (Glodny et al. 2003). Internal structures and lithologies in the southern segment of the Marun-Keu complex are presented in more detail below. The Syum-Keu ophiolite complex, in the hanging wall of the Main Uralian Fault, consists of ultramafic rocks in the west, and gabbroic-basaltic rocks further east. The Syum-Keu ophiolite and

Fig. 2. Geological and tectonic sketch map of the Marun-Keu complex (modified after Udovkina 1971). MUF, Main Uralian Fault. Box identifies the study area in the southern part of the complex.

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the Voykar/Rai-Iz ophiolite (about 100 km further south) both are erosional relics of one huge ophiolite nappe (Bogdanov et al 1996). A Sm-Nd isotope study by Sharma et al (1995) showed that these mantle rocks experienced extraction of basaltic melts at 387 ± 34 Ma. Tonalites, crosscutting the eastern margin of the Voykar ophiolite, were dated at 400 ±10 Ma (Buyakayte & Vinogradov 1983). 'Uralian' obduction of the ophiolite suite was already in progress at 373 ± 5 Ma (Glodny et al. 2003). A similar minimum age for the beginning of obduction is also defined by the presence of ophiolitic detritus in the Middle Devonian sediments of the West Uralian foreland (Savelieva & Nesbitt 1996). Island arc type rocks occurring east of the Syum-Keu complex have Late Silurian to Mid-Devonian ages (Saveliev 1997 and references therein) and overlie older oceanic basement (Saveliev 1997). They represent an analogue to the Tagil-Magnitogorsk zone in the Middle and Southern Urals. More eastern parts of the Polar Urals are largely unknown, due to burial beneath the Mesozoic to Cenozoic sediments of the West Siberian basin. There is no clear evidence for the presence of pre-Palaeozoic continental crust in the northern West Siberian basin (Aplonov 1995). The Marun-Keu complex: lithologies and sampling Samples were collected in the southern part of the Marun-Keu complex, south of the Great Khadyta river (Fig. 2). In this area, felsic to mafic eclogites predominate, interpreted as a metamorphosed island-arc type association of plutonic rocks with minor metasediments (Molina et al. 2002). All lithologies were affected by metamorphic mineral reactions to variable degrees, whereas deformation is partitioned into distinct shear zones. Eclogitization was fluid controlled. Proximal to eclogite facies fluid pathways, eclogitization was complete. Distal to these pathways, which are marked by vein-type mineralizations, primary igneous assemblages are partly preserved (Molina et al. 2002; Glodny et al. 2003). Post-eclogite facies retrogression is mainly confined to haloes around post-eclogite facies fluid pathways in which amphibolitization of the eclogites occurred (Glodny et al. 2003). Large volumes of rock were eclogitized under static conditions such that pre-metamorphic structures, e.g. migmatitic schlieren, features of magma mingling, or sedimentary layering, are well preserved. Partial inheritance of primary grain size distribution facilitates discrimination between lithologically different protoliths of similar chemical compositions, e.g. between metadolerites and metagabbros. Locally, the rocks even display pre-metamorphic microstructures, e.g. porphyric textures in some hybrid intrusives. From field observations, the rocks can be broadly grouped into four different categories. 1) There is a variegated suite of plutonic rocks, which is bimodal in composition. A mafic end-member is represented by olivine gabbros. The gabbros are surrounded by more felsic, dioritic-quartzdioritic to granodioritic lithologies. The felsic end-member is tonalitic to granitic, in places with migmatitic nebulitic schlieren. Textures are variable: both coarsegrained plutonic rocks, and porphyritic rocks suggesting subvolcanic intrusion levels, are present. All kinds of hybrid compositions occur in this suite, as well as widespread indications for magma mingling and mixing. 2) Leucocratic metagranites occur which are lithologically homogeneous at the outcrop scale. Mafic intercalations or migmatitic schlieren were not observed in these rocks. In places, the metagranites form felsic dykes and/or present intrusive contacts with their surroundings (V. Lennykh, pers. comm.). 3) Metasediments are rare and can be recognized from local preservation of compositional layering. In places, metasediments also occur as raft-like xenoliths within granitoids. 4) Eclogite facies veins, i.e. fractures filled with mainly quartz-dominated eclogite facies mineral assemblages are

widespread, and interpreted as indications for fluid activity at eclogite facies conditions (Glodny et al. 2003). The variegated intrusive suite is, by volume, the most abundant group of rocks. During fieldwork it was not possible to map in detail the regional distribution and the local geological context of the different rock types. In particular, the size and shape of the leucocratic metagranite bodies remains unclear, since the variegated suite comprises some leucocratic rocks as well. Therefore, our sampling focused on a nearly representative collection of all distinguishable lithologies. Sample details are presented in the Appendix. Analytical procedures Analytical work was performed on material from all four of the above rock categories. For Rb-Sr and Sm-Nd isotope analyses and elemental chemistry, rock powders were prepared using standard procedures. For zircon geochronology, we selected seven samples of the variegated intrusive suite, three metaleucogranites, one metasediment and one eclogite facies vein sample. In order to ensure full recovery of zircon, samples (0.1-1 kg) were crushed manually in a steel mortar. Zircon was then concentrated by magnetic separation and heavy liquids only, followed by hand-picking of grains suitable for single zircon analysis. Cathodoluminescence imaging was used to characterize zircon crystals, to check for evidence of magmatic zonation, inherited zircon cores or new overgrowth, and to choose analytical locations for the ion microprobe. Single zircon Pb evaporation Single zircon Pb-evaporation analyses were performed on a Finnigan MAT 262 mass spectrometer at the University of Granada, Spain. For analysis, a double filament ion source arrangement was used (Kober 1987). Selected zircon grains were mounted on canoe-shaped Re evaporation filaments. Pb evaporation was triggered by heating until the Pb+ ion beam was intense enough (200400 counts 206Pb per second). Pb was collected on the ionization filament for 20-30 minutes, and then analysed in 5 blocks with 7 scans each. Analysis proceeds by repeated evaporation, deposition and isotope ratio determination, until the zircon is exhausted. Data acquisition was performed using a secondary electron multiplier (SEM) detector. Common Pb correction was applied using Pb as a monitor and assuming an isotopic composition as given by Stacey and Kramers' (1975) model. Mass fractionation was corrected by multiplying measured 207Pb/206Pb ratios by V207/206. The standard error (SE) for each step was calculated as SE — 2v/^/n. A more detailed outline of the analytical procedures and the error treatment is given in Montero et al. (2000). Lead obtained from the first, low-temperature evaporation steps is often characterized by low 206Pb/204Pb ratios and may originate from contamination or from disturbed, metamict sectors of the zircons (Kober 1986). Therefore, this lead was not considered for geochronological evaluation. Lead released at high temperatures is assumed to have originated from undamaged zircon that shows no post-crystallization Pb loss. It is thus expected to yield undisturbed crystallization ages, if inherited zircon cores are absent. Ion microprobe zircon analyses For analysis, zircons were mounted in epoxy, sectioned and polished. U-Th-Pb ion microprobe analyses were carried out on a Cameca IMS 1270 instrument located at the Swedish Museum of Natural History, Stockholm (NORDSIM facility). Methods for sample preparation, analytical procedures and data reduction

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PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS Table 1. Pb evaporation analytical data for zircons from the Marun-Keu complex # *

Sample

grain #

(I) Variegated intrusive suite 1 PU12 1 2 PU12 1 3 PU12 1 4 PU12 1 5 PU12 2 6 J12a 1 7 J12a 1 8 J12a 1 9 J12a 2 (II) Meta-Leucogranites 10 J30 1 11 J30 1 12 J30 1 13 J30 2 14 J30 2 15 J30 2 16 J30 2 17 J30 2 18 J30 2 19 PU22 1 20 PU22 1 21 PU22 2 22 PU22 2 23 PU22 3 24 PU22 3 25 PU22 3 26 PU22 3 (III) Metasediments 27 J25 1 28 J25 1 29 J25 1 30 J25 2 31 J25 2 32 J25 2

Eva. step

Zir size [jmi]

204pb//206pb

207pb/206pb

207pb/206pb

208pb/206pb

meas.

meas.

corr.

meas.

Age [Ma] (t)

1 2 3 4 1 1 2 3 1

250 250 250 250 190 250 250 250 150

x x x x x x x x x

80 80 80 80 75 150 150 150 110

0.000034 0.000019 0.000038 0.000010 0.000045 0.000097 0.000063 0.000007 0.000002

0.056245 0.057051 0.059215 0.059019 0.056301 0.059728 0.058927 0.058087 0.057951

0.055881 0.056916 0.058799 0.058579 0.055774 0.058455 0.058159 0.058132 0.058067

0.0175 0.0415 0.0896 0.1119 0.0075 0.0373 0.0612 0.0815 0.0479

448 ± 11 488 ± 6 560 ± 14 551 ± 10 443 ± 40 547 ± 16 536 ± 6 535 ± 6 532 ± 4

1 2 3 1 2 2b 3 4 5 1 2 1 2 1 2 3 4

250 250 250 300 300 300 300 300 300 250 250 250 250 300 300 300 300

x x x x x x x x x x x x x x x x x

85 85 85 100 100 100 100 100 100 100 100 130 130 90 90 90 90

0.000041 0.000040 0.000052 0.000185 0.000212 0.000196 0.000164 0.000252 0.000188 0.000194 0.000045 0.000088 0.000069 0.000789 0.000432 0.000157 0.000104

0.057683 0.057300 0.057969 0.059784 0.059824 0.059908 0.059460 0.060390 0.059440 0.059536 0.057340 0.057995 0.057537 0.067767 0.063092 0.058821 0.057886

0.057219 0.056856 0.057354 0.057234 0.056880 0.057194 0.057216 0.056856 0.056840 0.056848 0.056820 0.056859 0.056594 0.056419 0.056942 0.056673 0.056516

0.2876 0.2714 0.2783 0.2636 0.2534 0.2613 0.2341 0.2176 0.2425 0.2924 0.3039 0.2083 0.2654 0.2284 0.2392 0.2331 0.2411

500 + 8 486 ±6 505 ± 12 501 ± 10 487 ± 6 499 ± 7 500 + 6 486 + 4 485 + 6 486 + 15 485 + 20 486 + 6 476 + 9 469 ±5 489 ± 10 479 + 7 472 + 4

1 2 3 1 2 3

300 300 300 325 325 325

x x x x x x

80 80 80 100 100 100

0.000057 0.000079 0.000205 0.000023 0.000017 0.000005

0.061967 0.063046 0.065298 0.060251 0.060539 0.060030

0.061297 0.062049 0.062489 0.060007 0.060445 0.060097

0.1175 0.1293 0.1348 0.1074 0.1031 0.1041

650 + 3 676 + 6 691 + 10 606 + 7 619 + 9 607 + 5

Errors are reported at 2am level. *: zircon No. in Fig. 6; (t): 207Pb/206Pb ages for individual evaporation steps; corr.: 207Pb/206Pb ratios are corrected for mass fractionation, and for common lead following the model of Stacey & Kramers (1975). Age error is calculated from error on corrected 207Pb/206Pb ratios. Data acquisition at Granada University.

follow those described in Whitehouse et al (1997, 1999). Common Pb correction was based on the measured 204Pb, and Stacey and Kramers' (1975) model isotopic compositions. Analysis spot selection and spot documentation were carried out using cathodoluminescence (CL) images. For geochronological evaluation, we only used concordant data (i.e. concordancy between 206Pb/238U and 207Pb/206Pb ages within 2a ion-microprobe analytical error). Since 206Pb-238U ages are, for Neoproterozoic and younger zircons, analytically more 'robust' than 207pb/206Pb ages, the apparent 2°6pb_238u ages are considered for discussion. Sm-Nd and Rb-Sr analyses Samples were analysed for Rb, Sr, Sm, and Nd contents by isotope dilution. They were spiked with suitable mixed 87Rb, 84Sr and 149 Sm, 150Nd spike solutions and dissolved. Solutions were then processed by standard cation-exchange techniques. Determinations of Rb and Sr isotope ratios were carried out on a VG Sector 54 multicollector TIMS instrument (GFZ Potsdam). The value obtained for 87Sr/86Sr of the NBS standard SRM987 during the period of analytical work was 0.710266 ± 0.000012 (2a, n = 25, dynamic mode). Determinations of Sm-Nd isotopic ratios were carried out on a Finnigan MAT 262 TIMS instrument (University of Oslo, analysis numbers starting with OS in Table 3)

and on a similar instrument at the GFZ Potsdam (analysis numbers starting with PS in Table 3). Nd was analysed in dynamic mode. The value obtained for 143Nd/144Nd of the La Jolla Nd standard was 0.511852 + 0.000006 (2a, n = 6). No systematic bias between the two instruments was observed.

Elemental chemistry Major and trace elements, except REE, Y, and Sc, were analysed by XRF on fused glass discs and powder pellets at Oslo University and at the GFZ Potsdam. REE, Y, and Sc abundances were determined by ICP-AES at GFZ Potsdam, following methods described in Zuleger & Erzinger (1988). Loss on ignition was determined gravimetrically after igniting the sample powder to 1050 °C for 1 h.

Results and discussion Zircon geochronology Comparative studies of concordant zircons have shown that no significant method-inherent systematic bias exists between single

Table 2. Ion microprobe U-Th-Pb analytical data for zircons from the Marun-Keu complex Sample

*

grain #, spot#

(I) Variegated intrusive suite 1 PU12 1.1 2 PU12 1.2 2.1 PU12 3 3.1 4 PU12 4.1 5 PU12 J12a 6 1.1 2.1 7 J12a 3.1 J12a 8 3.2 J12a 9 4.1 10 J12a 4.2 11 J12a 1.1 12 J3 1.2 J3 13 2.1 14 J3 3.1 J3 15 4.1 16 J3 1.1 17 J23 2.1 J23 18 3.1 J23 19 4.1 20 J23 J23 5.1 21 5.2 J23 22 6.1 J23 23 7.1 24 J23 1.1 25 J12c 2.1 26 J12c 3.1 27 J12c 4.1 28 J12c 5.1 J12c 29 5.2 J12c 30 1.1 31 019 1.2 32 019 2.1 33 019 3.1 34 019 4.1 35 019 5.1 36 019 5.2 37 019 6.1 38 019 7.1 39 019 8.1 40 019

Structural class, (from CL)

U [ppm]

Pb [ppm]

Th/U (meas.)

f206

inner chaotic tip inner dark inner zoned $ inner rim, zoned $ inner zoned outer zoned $ in. + outer, zoned $ outer zoned $ tip a inner; reworked? bright rim $ inner zoned $ inner inner inner zoned $ inner $ inner + incl inner outer bright inner dark $ outer bright $ core, dark inner bright zoned fragment zoned fragment zoned fragment diffuse-dark diffuse-dark diffuse rim inner zoned inner zoned inner zoned inner zoned rim inner zoned inner zoned inner zoned inner zoned t

540 556 2016 411 328 453 483 452 398 384 288 306 282 196 98 143 158 2991 134 100 180 89 3039 57 214 82 218 1009 167 221 265 271 178 513 182 191 84 389 223 246

86 53 302 48 36 45 47 46 38 38 26 35 30 25 11 16 9 206 8 6 11 5 480 4 34 10 32 147 21 33 29 30 17 56 18 18 9 47 26 28

0.46 0.03 0.02 0.48 0.39 0.30 0.30 0.41 0.29 0.24 0.34 0.62 0.51 1.15 0.75 0.93 0.00 0.02 0.01 0.03 0.00 0.00 0.39 0.01 2.47 1.13 1.90 1.89 1.81 2.04 0.71 0.51 0.31 0.47 0.44 0.17 0.47 0.94 0.81 0.76

16.3 0.21 17.0 1.08 0.05 0.05 0.02 0.00 0.02 0.03 0.04 0.06 0.13 0.06 0.15 0.61 0.07 8.59 1.37 8.91 0.04 0.00 22.0 0.17 0.03 0.03 0.15 0.02 0.22 0.12 0.14 0.09 0.82 0.42 0.09 0.96 0.08 0.33 0.11 0.13

206pb/204pb

207pb/206pb

[%] (a) 115 8929 110 1729 40950 39780 100800 >10E5 110600 59920 44660 30960 14890 33140 12270 3090 26800 218 1361 210 48760 >10E5 85 10880 59950 61840 12390 90740 8418 15640 13440 20790 2272 4458 20850 1947 23390 5599 16620 13890

0.17410 0.06319 0.17800 0.06749 0.05759 0.05758 0.05708 0.05623 0.05918 0.05630 0.06029 0.05902 0.05848 0.05156 0.05258 0.05691 0.05456 0.12160 0.06403 0.11460 0.05398 0.05315 0.22380 0.05587 0.05844 0.06074 0.05968 0.05871 0.06030 0.05893 0.05927 0.05882 0.05692 0.05793 0.05945 0.05478 0.06017 0.05717 0.05667 0.05881

±^

207pb/206pb

[%]

Age [Ma]

0.87 1.25 1.40 1.21 1.36 1.28 0.84 1.23 1.09 1.24 1.35 1.96 1.19 1.53 2.49 4.78 1.15 1.45 1.35 1.48 1.56 2.06 2.31 2.14 0.92 1.50 0.88 0.53 1.25 0.90 0.84 0.81 2.11 0.84 1.13 2.45 1.57 1.37 1.01 0.94

2597.5 ± 14.4 714.6 ± 26.4 2634.3 ± 23.1 852.9 ± 25.0 514.2 ± 29.6 513.8 ± 27.9 494.6 ± 18.4 461.5 ± 27.1 573.7 ± 23.5 464.2 ± 27.3 614.0 ± 28.8 567.8 ± 40.0 547.8 ± 25.8 265.9 ± 34.6 310.7 ± 55.8 487.9 ± 102.1 394.2 ± 25.7 1979.8 ± 25.7 742.6 ± 28.2 1873.6 ± 26.5 370.2 ± 34.8 335.2 ± 46.0 3008.2 ± 36.6 447.2 ± 46.9 546.3 ± 19.9 630.0 ±31.9 592.0 ± 18.9 556.4+ 11.4 614.4 ± 26.7 564.5 ± 19.5 577.0 ± 18.1 560.4 ± 17.5 488.4 ± 45.9 527.0 ± 18.3 583.6 ± 24.4 403.2 ± 54.0 609.7 ± 33.5 498.1 ± 29.8 478.7 ± 22.1 560.1 ± 20.3

Disc.% 2(j lim.

206

-76.7 -14.0 -75.5 -25.2

0.0956 0.0881 0.0911 0.0925 0.0926 0.0877 0.0861 0.0872 0.0834 0.0875 0.0775 0.0915 0.0858 0.0901 0.0891 0.0852 0.0564 0.0613 0.0579 0.0495 0.0579 0.0569 0.0908 0.0578 0.0867 0.0823 0.088 0.0895 0.0785 0.0886 0.0865 0.0918 0.0847 0.0913 0.0844 0.0838 0.0877 0.0905 0.0884 0.0886

2.2 1.8 -9.4

68.9 26.9 -75.3 -40.5 -77.2 -74.7

-5.8 -9.4

Pb/238U

±a [%] 1.73 1.55 2.49 1.57 1.61 1.54 1.54 1.54 1.54 1.54 1.55 1.54 1.54 1.81 1.67 1.67 1.69 1.68 1.67 1.67 1.67 1.67 .72 .69 .68 .67 .67 .67 .67 .70 .39 .48 .41 .42 .57 .41 .54 .36 .45 .43

206 Pb/238U Age [Ma]

588.4 ± 9.7 544.2 ± 8.1 562.2 ± 13.4 570.2 ± 8.6 570.8 ± 8.8 542.0 ± 8.0 532.5 ± 7.9 538.9 ± 8.0 516.7 ± 7.6 540.5 ± 8.0 481.1 ± 7.2 564.1 ± 8.3 530.8 ± 7.9 556.3 ± 9.6 550.3 ± 8.8 527.6 ± 8.5 353.4 ± 5.8 383.4 ± 6.2 362.7 ± 5.9 311.3 ±5.1 362.9 ± 5.9 356.9 ± 5.8 560.7 ± 9.2 362.3 ± 5.9 535.8 ± 8.6 509.6 ± 8.2 544.0 ± 8.7 552.3 ± 8.8 487.3 ± 7.8 547.5 ± 8.9 534.7 ± 7.1 566.1 ± 8.0 524.0 ± 7.1 563.4 ± 7.7 522.3 ± 7.9 519.0 ± 7.1 542.1 ± 8.0 558.2 ± 7.3 545.9 ± 7.6 547.4 ± 7.5

9.1 019 41 9.2 019 42 10. 019 43 11. 019 44 12. 019 45 1. 018 46 2. 018 47 3. 018 48 3.2 018 49 4.1 018 50 4.2 018 51 018 52 5.1 5.2 018 53 6.1 018 54 6.2 018 55 7.1 018 56 7.2 018 57 (II) Meta-leucogranites 1.1 PUlOa 58 2.1 PUlOa 59 3.1 PUlOa 60 4.1 PUlOa 61 5.1 PUlOa 62 6.1 PUlOa 63 (III) Metasediments J25 64 1.1 J25 65 1.2 2.1 J25 66 J25 67 2.2 J25 68 3.1 J25 69 3.2 4.1 J25 70 J25 71 4.2 (IV) Eclogite facies vein 4G 72 1. 4G 73 2. 4G 74 3. 4G 75 4. 4G 76 5. 4G 77 6.1 4G 78 7.1 4G 79 7.2

rim inner zoned inner zoned inner zoned inner zoned rim core rim core rim $ core $ rim core rim core core rim

188 463 165 553 221 1711 2924 351 368 3280 347 1383 502 542 240 2445 780

15 49 17 60 23 147 324 35 44 332 36 131 52 51 26 243 75

0.22 0.50 0.29 0.54 0.46 0.17 0.06 0.17 0.32 0.09 0.19 0.09 0.13 0.09 0.26 0.02 0.12

1.89 0.59 0.08 0.09 0.41 1.01 0.10 7.62 0.08 0.29 0.17 0.85 0.03 1.82 0.03 0.07 0.03

992 3190 24510 20760 4558 1858 19250 245 22150 6410 11300 2211 55990 1029 73050 26760 72890

0.04779 0.05406 0.05870 0.05783 0.05688 0.06623 0.06138 0.05984 0.06136 0.05937 0.06113 0.05913 0.05957 0.05927 0.06082 0.05848 0.05871

5.23 1.09 1.05 0.66 1.35 0.36 0.22 0.80 0.63 0.24 0.61 0.38 0.51 0.63 0.69 0.24 0.56

88.7+119.4 373.4 ± 24.4 556.0 ± 22.8 523.3 ± 14.4 487.1 ± 29.6 813.7 + 7.6 652.6 + 4.8 597.8 + 17.1 651.9 + 13.5 580.7 + 5.2 643.8 + 13.1 571.9 + 8.4 588.0+ 11.1 577.0 + 13.6 632.9 + 14.9 547.8 ± 5.2 556.4 + 12.1

inner zoned $ dark inner, reworked? dark in. + zoned outer dark inner dark in., reworked? inner zoned

717 3510 386 2779 709 578

72 267 34 277 59 58

0.78 0.21 0.49 0.53 0.64 0.76

0.86 1.73 5.01 0.03 0.87 0.29

2172 1079 373 69490 2157 6481

0.05712 0.05506 0.05717 0.05646 0.05772 0.04437

1.76 4.86 4.16 0.40 2.34 3.22

496.1 414.5 498.1 470.5 519.3 -90.3

+ ± ± + + +

38.2 105.1 89.2 8.8 50.5 77.1

core, zoned $ inner rim $ core rim, zoned inner, zoned tip tip inner, zoned

236 770 247 468 200 179 271 194

28 91 28 50 23 19 30 24

0.35 0.21 0.15 0.40 0.36 0.17 0.28 0.46

0.10 0.03 0.08 0.28 0.06 0.06 0.01 0.04

18970 58140 24060 6618 29720 30010 129000 46230

0.06323 0.06034 0.06084 0.05761 0.05628 0.06034 0.06080 0.05913

2.05 0.68 1.26 1.34 1.63 1.29 1.39 1.34

716.0 615.8 633.6 515.0 463.4 615.8 632.2 571.9

+ + + + + ± ± ±

43.0 14.5 26.8 29.3 35.8 27.7 29.6 28.8

tip* core tip inner inner inner tip $ core $

148 20 182 272 109 152 153 181

9 2 11 17 7 10 10 28

0.02 1.37 0.02 0.03 0.03 0.02 0.02 2.78

0.17 0.89 0.09 0.17 0.21 0.00 0.03 0.14

11200 2107 20910 10890 8913 >10E5 62230 13270

0.05487 0.06034 0.05508 0.05502 0.05391 0.05465 0.05509 0.05780

1.53 3.43 1.47 1.35 1.74 1.48 1.51 1.21

406.9 615.8 415.5 413.0 367.3 397.9 415.9 522.2

+ + + + + + ± +

34.0 72.3 32.5 29.8 38.7 32.8 33.3 26.4

d 30.5

-37.6

-1.5

2.0 d

14.4

0.0743 0.0884 0.0885 0.0892 0.0888 0.0757 0.1024 0.0881 0.1021 0.0928 0.0928 0.0871 0.0946 0.0860 0.0949 0.0934 0.0873

1.42 1.35 1.57 1.52 1.43 4.84 4.85 4.86 4.85 4.84 4.85 4.84 4.85 4.87 4.86 4.85 4.90

462.0 ± 6.3 546.3 + 7.1 546.7 + 8.2 550.9 + 8.0 548.4 + 7.5 470.1 + 22.0 628.4 + 29.1 544.4 + 25.4 626.6 + 29.0 571.8 + 26.6 571.8 + 26.6 538.1 + 25.0 582.6 ± 27.1 531.5 + 24.9 584.3 + 27.2 575.3 + 26.7 539.3 + 25.4

0.0788 0.0677 0.0728 0.0818 0.0680 0.0799

1.67 1.67 1.67 1.69 1.88 1.67

488.7 422.3 452.8 506.6 423.9 495.8

+ 7.9 + 6.8 + 7.3 + 8.3 + 7.7 + 8.0

0.1030 0.1049 0.1024 0.0892 0.1009 0.0973 0.0960 0.1037

1.57 1.54 1.55 1.56 1.61 1.70 1.54 1.67

632.2 643.0 628.4 550.7 619.7 598.7 591.0 636.0

+ 9.5 + 9.4 + 9.3 + 8.2 + 9.5 + 9.7 + 8.7 + 10.1

0.0583 0.0731 0.0575 0.0599 0.0597 0.0600 0.0595 0.0795

1.68 1.96 1.68 1.68 1.67 1.67 1.69 1.75

365.1 454.5 360.7 374.8 373.5 375.6 372.4 493.3

+ 6.0 + 8.6 + 5.9 + 6.1 + 6.1 + 6.1 + 6.1 + 8.3

Errors are given at the la level, are based on the counting statistics and include a component of the standard error. *, zircon No. in Fig. 7; $, CL image in Fig. 3; Disc., Degree of discordance (%); not reported for analyses which are concordant within 2a error limits; d, highly discordant, data regarded as aberrant; f(a), percentage of 206Pb contributed by common Pb, estimated from 204Pb assuming Stacey & Kramers (1975) model compositions. Data acquisition at NORDSIM facility, Stockholm.

94

J. GLODNYCTAL.

Table 3. Sm-Nd data and Nd model ages

Sample

Material

Sm [ppm]

Nd [ppm]

147

Sm/144Nd

143Nd/144Nd

143Nd/144Nd

2am [%]

eNd [at x Ma]

TDM (*)

TCHUR (f)

(I) Variegated intrusive suite: Neoproterozoic III to Early Cambrian zircon ages (^-550 Ma) WR OS49 17.0 93.9 0.1103 0.512298 PU51 WR OS21 22.2 4.60 0.1260 0.512518 J12a WROS18 2.97 21.0 0.0861 0.512230 PU50 WR OS22 7.32 48.8 0.0914 0.512221 J12c WROS16 0.471 0.1683 0.512603 1.71 PU12 WRPS149 3.33 11.0 0.1830 0.512229 PU62 WR PS385 0.508 2.15 0.1427 0.512587 PU63c WR PS388 2.95 0.1328 0.512485 0.649

0.0023 0.0038 0.0026 0.0026 0.0050 0.0017 0.0016 0.0017

-0.5 [550] 2.6 [550] -0.1 [550] -0.7 [550] 1.4 [550] -7.0 [550] 2.8 [550] 1.6 [550]

1173

601 259 564 604

(II) Meta-leucogranites: Late Cambrian/Early Ordovician zircon ages (~490 Ma) PUlOa 26.6 0.1233 WR OS20 5.38 J30 WROS51 9.16 0.1119 49.9 PU22 36.1 0.1160 WROS31 6.89

0.512386 0.512398 0.512419

0.0032 0.0027 0.0022

(III) Metasediments WR PS495 WR OS30

0.512123 0.512276

0.0012 0.0034

J3

J25 J24

5.46 6.99

26.2 43.7

0.1261 0.0975

998 1030 1087

* *1086

* t

1147

145 365

0.3 [550] 1.3 [550] 1.5 [550]

1194 1040 1051

526 433 415

-4.2 [650] 0.0 [550]

1698 1071

1111

558

(*) TDM: Depleted mantle model, 147Sm/144Nd = 0.2117, 143Nd/144Nd - 0.513079 (De Paolo et al 1991). (t) TCHUR: Assumption of chondritic mantle, 143Nd/144Nd = 0.512638, 147Sm/144Nd = 0.1967 (Jacobsen & Wasserburg 1980). ($) No sensible results due to insufficient spread between sample and model reservoir 147Sm/144Nd ratios. An error interval of + 0.5% (2a) is assigned to the 147Sm/144Nd ratios.

Table 4. Rb-Sr whole rock data 87

Rb/86Sr

87

87

0.249 0.334 0.506 0.629

0.708376 0.707360 0.709112 0.710896

0.0027 0.0027 0.0026 0.0032

0.7064 0.7047 0.7050 0.7059

(II) Meta-leucogranites: Late Cambrian/Early Ordovician zircon ages (~490 Ma) 21.4 PUlOa granite 165 22.6 436 0.318 J30 granite 48.0 380 0.475 PU22 granite 62.3

0.849738 0.707774 0.709676

0.0060 0.0014 0.0027

(nd) 0.7055 0.7064

Sample

Material

Rb [ppm]

Sr [ppm]

(I) Variegated intrusive suite: Neoproterozoic III zircon ages (~550 Ma) 740 J3 granitoid 65.6 364 PU51 metabasite 42.0 174 J12c metabasite 30.4 339 PU50 granitoid 73.6

Sr/86Sr

Sr/86Sr2am[%]

87

Sr/86Sri

(nd), not determinable due to high Rb-Sr ratio. An error interval of ± 1.5% (2a) is assigned to the 87Rb/86Sr ratios.

zircon 207Pb/206Pb-evaporation ages, ion microprobe U-Pb ages and ID-TIMS U-Pb ages (e.g. Kroner & Todt 1988; Kroner et al. 1991; Claoue-Long et al. 1995). This justifies the discussion of all zircon age values in this study. The analytical data are presented in Table 1 and Fig. 6 (Pb evaporation data) and Table 2, Fig. 7 (ion microprobe data). CL images are displayed in Fig. 3. Variegated suite of plutonic rocks. Zircons from seven samples of the variegated intrusive suite have been investigated (samples PU12, J12a, J3, J23, J12c, 018, 019, Tables 1, 2; Fig. 3). With the exception of sample J23, the CL images of the zircons generally show strong, concentric, oscillatory zoning, regarded as igneous in origin (Hanchar & Miller 1993). They frequently display growth features truncated by rounded internal surfaces, indicative of resorption of zircon by a melt. These surfaces are overgrown by bright luminescent bands, merging into new shells of oscillatory zoned zircon. Such multiple growth stages are thought to be due to fluctuations in the Zr saturation state of the magma, and characteristic of the thermally and chemically variable magma mingling environment of calc-alkaline intrusions (Vavra 1994). Optically, it is sometimes difficult to distinguish such resorption events from the presence of inherited xenocrystic cores. Inherited (magmatic) zircon cores with magmatic overgrowth are likely to be present in the granite sample 018, which exhibits migmatitic schlieren.

With the exception of sample J23, zircons from the variegated suite generally yield 206pb-238U ion microprobe and Pb-evaporation ages of around 550 Ma (Figs 6 & 7). We interpret this age as the crystallization age of this hybrid intrusive suite, which formed in an island arc setting (see also Molina et al. 2002). The cluster of ages around 550 Ma nearly coincides with zircon crystallization ages around 560 Ma from calc-alkaline granitoids of the Timan-Pechora basement (Gee et al. 2000). This leads to the interpretation that the Marun-Keu complex contains intrusives related to the Timanian orogeny. A few dates apparently younger than 550 Ma have been obtained by the Pb-evaporation technique for the granitoid sample PU12 (Fig. 6). An inherent problem of this technique is that zircons with a poly stage history may produce mixed ages. This is then reflected by a significant shift of apparent ages between individual evaporation steps (e.g. Dougherty-Page & Bartlett 1999). Such a shift is observed for grain #1 (Fig. 6). Since the age of the magmatic zircons of the sample (550570 Ma) is constrained by the concordant ion microprobe age data, which are consistent with the last two Pb evaporation step ages for grain #1, we regard the younger dates as spurious due to resetting for unknown reasons. In some zircons of the migmatitic granite sample 018, old inherited cores are present. For two crystals we obtained concordant ion microprobe U-Pb ages for the cores of about 630 Ma, while the

95

PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS

Table 5. XRF major and trace element data of representative rocks, Marun-Keu complex

1 J30

2 PUlOa

3 PU22

4 J3

5 J12c

6 J12a

7 PU12

8 PU50

9 PU51

10 PU58

11 PU59

12 J25

SiO2 TiO2 A1203 Fe203 MnO MgO CaO Na2O K2O P205 BaO H2O C02 LOI

71.8 0.46 13.1 3.76 0.06 1.39 5.02 0.54 2.01 0.17 0.05 nd nd 1.06

77.5 0.13 12.2 1.56 0.02 0.10 0.48 3.42 4.96 0.02 0.02 nd nd 0.35

77.2 0.32 12.7 2.41 0.03 0.43 0.95 0.92 2.98 0.05 nd 1.61 0.11 nd

63.0 0.78 17.0 6.28 0.11 1.40 4.85 2.02 2.29 0.19 0.09 nd nd 1.14

47.1 0.42 16.7 9.06 0.15 13.2 7.67 3.45 1.00 0.06 0.02 nd nd 1.15

53.0 0.62 17.9 3.51 0.02 6.67 8.90 4.56 3.01 0.14 0.08 nd nd 1.59

70.5 0.44 16.2 2.54 0.03 0.77 0.70 5.55 2.56 0.06 0.04 nd nd 1.03

72.4 0.25 14.9 3.08 0.02 0.44 1.88 1.56 3.18 0.17 nd 1.37 0.13 nd

47.8 1.14 18.8 10.8 0.18 5.81 9.66 3.08 1.61 0.18 nd 0.96 0.11 nd

61.3 0.91 18.4 4.85 0.06 1.80 5.44 4.44 1.42 0.26 nd 0.75 0.09 nd

50.8 1.62 18.5 9.06 0.11 4.15 6.78 4.14 2.50 0.47 nd 1.42 0.12 nd

55.9 0.83 17.8 9.51 0.13 3.91 7.24 2.19 1.59 0.13 0.04 nd nd 0.54

Sum

99.42

100.8

99.71

99.15

99.98

100.0

100.4

99.38

100.1

99.72

99.67

99.81

6.9 30 62 148 253 17 104 24 53 19 4 8 267 6 9 143 407

5.2 7.1 20 39 462 <5 14 8 44 12 4 27 114 13 13 105 143

Sc La Ce V Cr Co Ni Cu Zn Th U Y Zr Nb Pb Rb Sr

6.6 60 118 25 84 8 17 11 15 13 4 41 388 8 18 56 437

2.1 27 68 7 363 <5 7 <5 29 21 5 50 143 21 21 174 24

5.2 42 86 9 160 <5 6 <5 31 17 3 42 255 19 19 69 371

11 101 208 65 224 14 12 12 58 26 >2 41 462 19 17 69 726

13 1 2.7 87 793 65 379 9 138 <2 >2 13 47 3 7 36 161

3.2 73 156 24 237 5 6 11 42 39 4 26 269 4 14 77 326

33 16 40 299 344 33 18 20 105 <2 3 31 105 4 7 50 371

11 20 51 74 295 14 14 21 64 7 6 19 196 11 14 73 607

13 36 75 199 276 27 36 38 114 3 <2 19 178 9 11 78 908

26 22 48 196 311 21 21 8 91 <2 3 31' 136 5 7 67 321

Analysis of major elements by XRF, either at the University of Oslo (analyses including BaO) or at GFZ Potsdam. Trace element abundances measured at the University of Oslo by XRF. Exception: La, Ce determined by ICP-AES, GeoForschungsZentrum Potsdam, nd, not determined. Samples 1-3: Meta-leucogranites (Late Cambrian-Early Ordovician zircon ages) Samples 4-11: Variegated intrusive suite (Neoproterozoic III calc-alkaline magmatism) Sample 12: Metasediment, eclogite facies

rims of these grains, as well as other grains, show oscillatory zoning in CL and yield ages around 550 Ma. Most likely this granitoid crystallized at around 550 Ma, and inherited zircon material from its (at least partly metasedimentary) protolith. The metagranitic sample J23 is, from field relations, interpreted to belong to the variegated intrusive suite. However, zircons from this sample are different with respect to both CL appearance and age pattern. Zircons from this sample generally have a round, isometric-polyfaceted shape, show pronounced sector zoning, and only rarely display xenocrystic cores (Fig. 3). We interpret these grains to be metamorphic in origin, as similar morphology and CL characteristics are commonly observed in zircons related to high-grade metamorphic events (e.g. Rubatto et al 1999; Hoskin & Black 2000; Pidgeon et al 2000). These zircons are further characterized by their remarkably low Th-U ratios (between 0.00 and 0.03, Table 2; with the exception of one xenocrystic core). The concordant U-Pb age data for the metamorphic zircon domains are between 353 and 362 Ma, coincident with the age of metamorphism as inferred from Rb-Sr internal mineral isochrones (355.5 ±1.4 Ma; Glodny et al. 2003).

widespread. Many zircon crystals, particularly in sample J30 (felsic dyke) have irregular shapes, which may either indicate skeletal growth, irregular growth because of space restrictions, or strong corrosion (Figs 3 & 4). Inclusions of quartz, apatite and a (U, Th)- rich phase are common. Inherited xenocrystic zircon was not observed in these granites. The Pb evaporation data for samples J30 and PU22 consistently yield ages around 490 Ma (Fig. 6). A comparable concordant ion microprobe 206Pb-238U age of 488.7 ± 7.9 Ma was obtained for a zircon crystal of sample PUlOa (highly differentiated metaleucogranite, with brownish zircon). However, other grains of this sample yield ion microprobe U-Pb age values down to 422 Ma. Zircon from this sample is characterized by high U contents and metamictization, which makes it prone to open system behaviour and apparent rejuvenation during post-crystallizational processes. Therefore, we interpret the age values younger than 480 Ma as geologically meaningless. From the data it is evident that the Late Cambrian to Early Ordovician zircon ages (c. 490 Ma) reflect a distinct pulse of granitic magmatism, roughly 60 Ma after the peak of island arc-type magmatism.

Metaleucogranites. Zircons from the metaleucogranites (samples PU22, J30 in Table 1; PUlOa in Table 2) show oscillatory zoning in their outer parts. The inner domains of many crystals are either CL-dark, unzoned, or have poor CL contrast. Irregularly fluctuating growth, with embayments in growth surfaces is

Metasediments (or metatuffitic-volcano-sedimentary rocks). These are represented by the compositionally layered sample J25 and, most likely, partly by the granitoid of sample 018, which contains metasedimentary xenoliths and migmatitic schlieren. CL images of the zircons (Fig. 3) show features similar to those from the

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Fig. 3. Cathodoluminescence images of sections through zircon crystals. Numbers within crystals denote ion microprobe 206Pb-238U spot ages. Grain numbers refer to Table 2. Variegated intrusive suite (Neoproterozoic III): samples PU12, J12a, J3, 018, 019. Metaleucogranites (Late Cambrian to Early Ordovician): samples PUlOa, PU22, J30. Samples PU22 and J30 have been analysed by Pb evaporation only (no spot ages). Metamorphic zircons: samples 4G, J23. Metasediments: sample J25. See text for explanation.

PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS

97

implying growth of the zircon crystal eclogite-facies conditions (Fig. 5). However, these metamorphic zircons were too small and too rare to be analysed. Zircons from sample J25 have been examined by both ion microprobe and Pb evaporation. The Pb-evaporation data reveal ages around 610 Ma for grain #2, whereas grain #1 presents evaporation step ages shifting from 650 to 691 Ma (Table 1, Fig. 6). This shift most probably relates to presence of an 'old' core within the zircon; the apparent step ages reflect a maximum age for the youngest and a minimum age for the oldest zircon domain. The minimum age of 691 ± 10 Ma for a zircon core is the oldest age observed in the entire dataset from the Marun-Keu complex. The ion microprobe U-Pb ages for this sample fall between 643 and 591 Ma, with the exception of one zircon rim that yields an age of 550.7 ± 8.2 Ma. For sample 018, the ion microprobe U-Pb ages for zircon rims scatter around 550 Ma (Table 2). However, some zircons have cores with significantly higher apparent ages of 628.4 and 626.6 Ma, i.e. in the range of ages obtained from sample J25. The metasediments document the presence of pre-550 Ma magmatic zircon. We suggest that the pre-550 Ma zircons crystallized in the early stages of 'Timanian' island arc magmatism. They may have been deposited in volcano-sedimentary sequences which were later affected by the magmatic activity at c. 550 Ma. Fig. 4. Backscatter electron image of irregularly shaped zircon, sample J30. Inclusions of quartz and apatite. Matrix phases: rutile, epidote, garnet, quartz. Scale bar: 100 jxm.

variegated intrusive suite, with characteristic polystage magmatic crystallization histories, and rounded internal surfaces. All zircons from both samples show oscillatory zoning, indicative of a magmatic origin. Distinctive evidence for zircon abrasion by exogenic transport is absent. In sample J25, some tiny zircons form a late, xenomorphic, interstitial phase mantled by garnet and omphacite,

Eclogite fades vein. Zircons from the eclogite facies vein (sample 4G) are round to prismatic in shape. Some grains exhibit distinct xenocrystic cores, with linear banding or oscillatory zoning (Fig. 3). The rims show 'metamorphic' sector zoning, similar to the grains of sample J23 (see above), or weakly developed zoning of an irregular, nebulous, patchy appearance. The CL pattern suggests metamorphic overgrowth on inherited magmatic cores. The sector-zoned zircon domains are characterized by very low Th-U ratios of <0.05 (Table 2) in the same way as the sector-zoned zircons of sample J23. The ion microprobe age of one xenocrystic core with weakly developed linear CL banding (206Pb-238U age of 493.3 ± 8.3 Ma) suggests that this core was formed in a Late Cambrian-Early Ordovician granitoid. The rim ages are only partly consistent with the inferred age of eclogite facies metamorphism of approximately 355 Ma: four grains display ages between 365 and 375 Ma. It remains unclear whether these values reflect mixed ages, or growth of zircon during prograde-metamorphic veining as suggested by Rubatto et al. (1999) for zircon of a comparable vein setting. Zircon stability in eclogite-facies metamorphism

Fig. 5. Backscatter electron image of zircon (bright), sample J25. Zircon is interstitial with respect to surrounding sodium-rich amphibole and garnet, implying eclogite-facies metamorphic growth. Scale bar: 10 jjim.

In most lithologies, zircon proves to be remarkably inert to the eclogite-facies metamorphic overprint. Zircon related to the metamorphic event is, in most samples, either absent, or forms only tiny and rare crystals (sample J25). Apart from the eclogite-facies vein (sample 4G), eclogite-facies metamorphism is unambiguously documented in the zircon ages only in one metagranitic sample (J23). It is an open question as to why zircons from apparently similar metagranitic lithologies do not reflect similar zircon-forming effects of metamorphism. However, there is some evidence for partial reorganization of the U-Pb system of specific zircon grains, like anomalously young ages in particularly U-rich zircons (e.g. sample PUlOa), or a few 'disturbed' zircons with cores apparently younger than the respective rims (e.g. sample J3, grain 1). But it remains unclear whether these effects are due to igneous processes, metamorphism, low-temperature fluid-rock interaction (as described by Krogh & Davis 1975; Gebauer & Griinenfelder 1976), or other means. In any case, eclogite-facies metamorphism may not, in many lithologies, leave a significant imprint in the zircon U-Pb age patterns.

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Fig. 6. Zircon age data (Pb-evaporation data), with individual 2arm errors. Data source: Table 1. Numbers in brackets are zircon grainnumber of the respective samples. Asterisked data points are interpreted as aberrant. See text for discussion.

Sm-Nd data and Nd model ages In order to constrain the protolith history of the Marun-Keu rocks further, we analysed samples of the lithologically defined premetamorphic Marun-Keu rock suites for their Sm-Nd isotope systematics. Results are presented in Table 3. Sm-Nd whole-rock data can be used to estimate model ages, reflecting the time of differentiation and isolation of crustal material from the Earth's mantle (e.g. Miller & O'Nions 1984; De Paolo et al 1991). Model age calculations and interpretations are based on several premises: (1) The model parameters reasonably match the properties of the local mantle. (2) A one-stage differentiation history from the mantle reservoir towards crustal material is assumed. (3) Crustal processes, like erosion, resedimentation, or metamorphism, do not change the Sm-Nd ratios significantly. If these premises are met, the model ages can be interpreted as 'crustal residence ages' (Miller & O'Nions 1984). In the case of mixture of material with different evolutions (e.g. sediments), Sm-Nd model ages reflect an 'average crustal residence age' (Arndt & Goldstein 1987; Goldstein el al 1997). Since the properties of the mantle source(s) for the Marun-Keu rocks are not exactly known, we calculated both TDM (assuming a depleted mantle source; De Paolo et al. 1991) and TCHUR (assuming a local mantle of chondritic composition, as determined by Jacobsen & Wasserburg 1980) model ages. Mantle compositions with Nd isotopic signatures between depleted mantle (DM) and chondritic unfractionated reservoir (CHUR) are widespread. They often form the source of intra-oceanic island arc material (Zindler & Hart 1986). As the meta-igneous Marun-Keu rocks are of island arc type as well (Molina et al 2002), TDM and TCHUR span a range of possible 'accurate' model age values. For Marun-Keu rocks, most TCHUR model ages are <600 Ma. Most TDM model ages range from 1000 to 1200 Ma (Table 3). These model ages correspond to sNd values, calculated for the crystallization ages provided by the U-Pb zircon data, between -0.7 and +2.8. A crystallization age of 550 Ma was similarly used for the Late Cambrian-Early Ordovician leucogranites, since we suspect them to represent reworked Timanian age (c. 550 Ma) material (see below). From the Sm-Nd data, it is

impossible to discriminate between the Neoproterozoic III and Late Cambrian-Early Ordovician intrusive suites. Two samples form significant exceptions in the dataset by exhibiting clearly negative eNd values, i.e. the metasedimentary sample J25 (sNd = - 4.2) and granite PU12 (eNd = - 7). However, all valu have to be considered with caution as eclogite-facies metamorphic processes may mobilize REE and thereby change Sm-Nd ratios (Griffin & Brueckner 1985; Bernard-Griffiths et al 1991). We speculate that sample PU12 has been affected by such processes. Nevertheless, eNd values close to zero predominate in the sample set (Table 3). Assuming that most of the samples have been closed systems for REE since crystallization, we regard these near-zero sNd values as the key information from the metamorphosed island-arc type crust of the Marun-Keu complex. During the generation of new continental crust in a magmatic arc setting, older crustal material is commonly incorporated into and mixed with new mantle-derived magmas (e.g. De Paolo 1980; Hickey et al 1986). From modelled examples of mantle magma v. crust-mixing in active margin settings (De Paolo et al 1991), it is known that in proximity to old continental basement, the proportions of old material in the hybrid arc-type crust may be high, up to more than 50%. For the Marun-Keu, representing Timanian-age basement, major magmatic input occurred at around 550 Ma, as indicated by the spectrum of zircon ages. The East European craton, to the SW of the Timanides, is characterized by Archean to Early Proterozoic basement with negative sNd values and older than Mesoproterozoic Nd model ages. The Marun-Keu rocks yield Nd model ages generally less than 1200 Ma, indicating absence of large fractions of this 'old' material here. This implies either only minor contributions of old crustal material to magma genesis, or that the crustal residence ages of the reworked crustal materials are similar to the crystallization ages of the igneous rocks. Field observations show that in some of the igneous rocks from Marun-Keu, as in the migmatitic granitoids, a high proportion of recycled crustal material must be present. However, Nd model ages of these lithologies are not significantly higher than those of other granitoids. Therefore, it is likely that almost the entire basement in the region, with the possible exception of some metasediment components (sample J25), formed by Timanian (c. 550 Ma), or only slightly older (Early Timanian?) crustal growth, most likely in an intra-oceanic arc setting, distal to any early-to mid-Precambrian craton.

PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS

The distinct, juvenile signature of the Timanian age basement along the northeastern European margin is corroborated by other datasets: the presence of reworked pre-Timanian crust should be evidenced by 'old', pre-Neoproterozoic zircon components, which are not observed. For the Marun-Keu itself, Shatskii et al (2000) report Nd isotopic data near the CHUR value (eNd c. 0 at c. 360 Ma), similar to our dataset. For Timanian rocks from the Engane-Pe Anticline, which were not affected by Uralian metamorphism, Scarrow et al (2001) obtained a Sm-Nd whole rock age of 687 + 210 Ma, together with Sm-Nd isotope systematics which yield model ages in the same range as the data from Marun-Keu. Lower Triassic sandstones from the eastern Barents Sea, interpreted as being derived from the Polar Urals and the Timan-Pechora region, yield Sm-Nd provenance ages, relative to depleted mantle compositions, between 595 and 1280 Ma with an average of 776 Ma (M0rk 1999). These model ages are significantly younger than comparable values from offshore Northern Europe, but similar to our results for the Timanian age rocks of Marun-Keu. Rb-Sr data In metamorphic rocks, Rb-Sr whole rock isotopic data can only be used to constrain protolith characteristics if metamorphism was isochemical, i.e. if no fractionation of Rb from Sr occurred during metamorphism. Near-isochemical conditions can be inferred for some partly eclogitized rocks of the Marun-Keu: eclogitization is fluid-dependent (Austrheim 1987; Engvik et al 2000). Incomplete eclogitization, therefore, indicates fluid deficiency. Since fluids mediate fractionation of Rb from Sr, only partially eclogitized rocks may have experienced only minor changes in Rb-Sr ratios. We therefore preferred such rocks for Rb-Sr analysis. From the data (Table 4) initial 87Sr/86Sr ratios were calculated for the time of crystallization provided by the zircon ages. The rocks show quite primitive apparent initial 87Sr/86Sr ratios, in the range of 0.7047-0.7065. This finding corroborates our interpretation that juvenile Timanian age crust predominates in the Marun-Keu. Obvious differences between the Neoproterozoic III and Late Cambrian-Early Ordovician intrusives are not observed.

99

For the granite sample PUlOa, a very high Rb-Sr ratio permits calculation of a Sr model age for crystallization. Using 87 Sr/86Sr = 0.705 ± 0.002 as an estimate of the initial Sr isotopic composition of the rock, a model age of 474 + 9 Ma (Early Ordovician) is obtained, which is only slightly lower than the oldest zircon ages for this sample (Table 2), implying that little or no alteration of the Rb-Sr system occurred during metamorphism. Comparative geochemistry The geochemistry of metamorphic rocks has to be considered with caution, as chemical signatures may be altered by metamorphic processes. This is particularly a problem with 'mobile' elements, whereas 'immobile' elements, like HFSE (high field strength elements) and REE, and their elemental ratios may normally persist. Data are presented in Table 5 (major and trace elements), Fig. 8 (multi-element plot) and Fig. 9 (REE plot). The chondrite-normalized multi-element plot (Fig. 8) displays pronounced negative HFSE (Ti and Nb) anomalies in most of the samples, characteristic for subduction-related magmas (Floyd & Winchester 1975). Most Marun-Keu granitoids are characterized by moderately LREE-enriched, fractionated REE patterns (Lan/Ybn ratio 4 to 27, with an average of 14.8) and variably pronounced negative Eu anomalies (Fig. 9). Granite PU12 is atypical due to its flatter REE pattern (Lan/Ybn = 2.3), which may be due to an eclogite facies metasomatic overprint. The gabbroic sample PU62, interpreted as part of the Neoproterozoic III intrusive suite, is only slightly enriched in LREE, with low overall REE abundances. It may represent a plagioclase-rich residue of magma differentiation, as suggested by the positive Eu anomaly. The Late Cambrian-Early Ordovician granites are chemically quite similar to the granitic rocks of the Neoproterozoic III intrusive suite. They only show slight tendencies towards lower Al abundances, higher K-Na ratios, and less fractionated HREE patterns (Dyn/Ybn = 0.94 to 1.14 with an average of 1.05 in metaleucogranites, v. Dyn/Ybn = 1.16 to 1.92 with an average of 1.46 for the Neoproterozoic III intrusives). The pronounced negative Sr and P anomalies of the metaleucogranites PUlOa and PU22 correspond to highly negative Eu anomalies and probably reflect advanced fractionation of feldspar and apatite from these magmas.

Fig. 7. Zircon age data (U-Pb ion microprobe data), with individual la errors. Data source: Table 2. Only concordant (at the 2a-level) data points are considered. See text for discussion.

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Fig. 8. Chondrite-normaLlized multi-element plot for selected representative rocks from the Marun-Keu complex. Normalizing values from Thompson (1982); for P, K, Rb: Sun (1982). Samples J3, PU12, PU50, PU51: variegated intrusive suite (Upper Neoproterozoic); samples J30, PUlOa, PU22: metaleucogranites (Late Cambrian-Early Ordovician); sample J25: metasediment.

Despite the chemical similarities between different granitoids, the distinct petrography and age of the Late Cambrian-Early Ordovician granites testify to genetic differences. Chemical and isotopic similarities do not preclude a different tectonic setting for magma generation. Post-orogenic granites may inherit the geochemical signatures of their volcanic arc-related sources (Ajaji et al 1998; Waight et al 1998). The geochemistry of granitoids then reflects the tectonic setting in which the source material was generated instead of the setting in which the melting of the sources occurred. We conclude that the studied meta-igneous Marun-Keu rocks represent former Timanian-age island arc-type crust.

Discussion: regional correlations and geodynamic interpretation The new data presented here clearly reveal a multistage geological history for the Marun-Keu complex, and are important for understanding the pre-Uralian and Uralian geodynamic evolution of the Polar Urals region. We discuss the geodynamic implications of the

data below. Whenever possible comparisons and correlations with other segments of the Timanides or Uralides are given. Early Timanian evolution (7007/650 to c. 570 Ma) The zircon ages from metasedimentary material reflect zirconforming, most likely magmatic events prior to the peak of the Timanian orogeny. The data indicate that at least a large proportion of the crust in the Polar Urals region represents juvenile Timanianage crust that was generated distally from old continental masses, in an intra-oceanic island arc setting. The presence of zircon with ages between 7007-650 Ma and c. 570 Ma suggests a prolonged process of island arc magmatic activity and crust production. Assemblage of the Timanian basement of the Polar Urals region from intra-oceanic lithologies was similarly inferred by Scarrow et al. (2001). Island-arc material and Timanian structures dominate throughout the Timan-Pechora basement. So far, no unquestionably pre-670 Ma lithologies or structures have been identified. However, to the west of the Urals region, in the basement beneath the Pechora basin, Nd isotopes and zircon xenocrysts document assimilation of older, mainly Mesoproterozoic crust (Gee et al. 2000; Pease et al. 2004). Timanian orogeny (c. 550 Ma)

Fig. 9. Chondrite-normalized REE plot illustrating REE spectra of selected rocks of the Neoproterozoic III variegated intrusive suite (solid lines) and of Late Cambrian to Early Ordovician metaleucogranites (dotted lines). Sample PU62 is a metagabbro. REE plot normalized to chondrite values of Evensen et al (1978)

Island-arc type magmatic activity in the Marun-Keu region, i.e. in the easternmost preserved parts of the Timan-Pechora basement, culminated and ceased at roughly 550 Ma. Textural evidence in the rocks, such as magma mingling zones involving mantlederived magmas, points to a synorogenic rather than to a lateorogenic origin of the 550 Ma intrusives. This time marks a first-order event, probably over the entire Timan-Pechora region as it is documented in Northern Timan (e.g. Malkov 1969) and across the Timan-Pechora basement (Gee et al. 2000; Pease et al. 2004). Accretion of the island-arc rocks to Baltica must have occurred at or slightly after 550 Ma, as is evident from the cessation of magmatic activity at that time. The NW-SEi direction of strike of the Timanian structures implies a NE-SW direction of convergence (present day coordinates). Not only in the Polar Urals, but along almost the entire length of the Urals, there is evidence for a Neoproterozoic III collisional

PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS

orogeny along the eastern margin of Baltica (Getsen 1991; Puchkov 1993, 1997). Compressional deformation and metamorphism, including Neoproterozoic III to Early Cambrian blueschist and eclogite-facies metamorphism, took place along the western foreland of the Middle and Southern Urals, and further to the west foredeeps were filled with molasse (e.g. Puchkov 1993, 1996, 1997; Brown et al 1996; Beckholmen & Glodny 2004). Available age data for this orogeny are similar to our new data from the Polar Urals. In the southwestern Urals, U-Pb ages from detrital zircons, mineral composition data, and light and heavy mineral spectra of sandstones indicate onset of convergence at c. 620-610 Ma (Willner et al 2001, 2003). Glasmacher et al. (1999, 2001) presented Ar/Ar muscovite ages of c. 550 Ma from the southwestern Urals, interpreted as being related to a Neoproterozoic orogeny along the eastern margin of the EEC. Metamorphic clasts and jasper in sediments of the southwestern Uralian foreland (Brown et al. 1996) were interpreted as evidence for Neoproterozoic III uplift and erosion of both metamorphic rocks and deep oceanic sediments in the Urals region. Evidence that Neoproterozoic III orogeny may have extended into the Early Cambrian is given by c. 536 Ma minimum ages, derived from Rb-Sr mineral isochrons, for blueschist-facies metamorphism in the Kvarkush anticline, Northern Urals at 60 °30' N (Beckholmen & Glodny 2004). The Timanide orogeny correlates with other contemporaneous orogenic belts, namely the Cadomian-Avalonian orogen (Puchkov 1997; Scarrow etal 2001).

101

1.4 Ma, Rb-Sr mineral data, Glodny et al. 2003) from the Marun-Keu complex date the arc-continent collision and thereby the termination of the subduction process. The MarunKeu rocks lithologically represent an almost complete crustal section. If the leading edge of continental crust enters a subduction zone, obduction of ophiolite nappes, contemporaneous with eclogite facies metamorphism within the subducted continental crust is likely (cf. Searle & Cox 1999). The positive buoyancy of the subducted crust causes blocking up of the subduction process, with immediate subsequent exhumation of the metamorphosed crustal material (e.g. Hynes 2002). In fact, no subduction-related magmatism younger than c. 360 Ma has been recognized in the Polar Urals region. Furthermore, the timing of the arc-continent collision, with subsequent uplift and erosion, is independently constrained by Lower Carboniferous (i.e. post-355 Ma) flysch deposits in the western foreland of the Polar Urals (Puchkov 1997). East-directed subduction is inferred from the here documented eclogite-facies metamorphism of Timanian-age crust of Baltica, and from the striking structural similarity between the Polar Urals (cf Bogdanov et al. 1996; Puchkov 1997; Saveliev et al. 1999) and the Southern Urals (cf Brown et al 1996; Puchkov 1997), with an east-dipping Main Uralian Fault, and west-facing accretionary wedges and HP metamorphic complexes west of the fault.

Conclusions Baltica passive margin rifting (c. 490 Ma) Late Cambrian to Early Ordovician granitic magmatism has not been recognized previously in the Timanian basement of the Polar Urals. Along the Urals, uppermost Cambrian to Early Ordovician time has long been identified as a period of incipient extension, subsidence and marine transgression, finally leading to the opening of the pre-Uralian ocean, parallel to the later Main Uralian fault (Ivanov et al. 1975; Perfiliev 1979; Puchkov 1997). The formation of the Late Cambrian to Early Ordovician granites is contemporaneous with alkaline basalt dyke swarms in Late Cambrian to Lower Ordovician sediments of the West Uralian foreland (Saveliev 1997 and references therein), and most likely relates to incipient rifting of the eastern margin of Baltica, which marks the onset of post-Timanian plate reorganization. Nearly contemporaneous, extension-related magmatism is also known from other places in the Urals. Examples are the Ordovician nepheline-syenites of the Ilmenogorsk-Vishnevogorsk complex, Middle Urals (Kramm et al. 1983) and Lower Ordovician basalts in the Sakmara zone, Southern Urals (Svyazhina et al. 1992). In places, Late Cambrian subalkaline and alkaline basalts and rhyolites, picritic porphyries and tuffs are associated with the extensional regime (Puchkov 1997 and references therein). Indications for Early Ordovician rift formation and development of a passive margin setting along the western slopes of the Urals are widespread in the sedimentary record of the Southern Urals (e.g. Svyazhina et al. 1992; Puchkov 1997). In summary, the extension and rifting which formed the Early Palaeozoic East European passive margin was nearly coeval all along the Urals, including the Polar Urals sector. Uralian convergence and arc-continent collision (c. 360 Ma) A convergent setting at the northeastern European margin was re-established roughly in Silurian to Early Devonian times, when island-arc rocks were formed in an intra-oceanic setting east of todays Main Uralian Fault (cf Ivanov et al. 1975; Saveliev et al. 1999). Our new age data from metamorphic zircons, together with age data on metamorphism (Late Devonian to Early Carboniferous, Sm-Nd mineral data, Shatskii et al. 2000; 355.5 +

In the western flank of the Polar Urals, the basement consists of Timanian age crust with a Timanian structural fabric. The Marun-Keu complex represents continental crust formed in a Neoproterozoic island arc domain, which was attached to Baltica during the Timanian orogeny. Late Cambrian to Early Ordovician rifting left the Marun-Keu crustal segment in a passive margin setting. Eclogite facies metamorphism of the Marun-Keu complex is due to subduction of the leading edge of Baltica during Uralian arc-continent collision in Late Devonian-Early Carboniferous times. Using single-grain zircon age data, in combination with Sr-Nd isotope and elemental chemistry data, the following evolutionary model for the Marun-Keu complex has been established: c. 670-550 Ma: Early Timanian production of juvenile crust, probably in an intra-oceanic island arc setting. The absence of zircon older than c. 700 Ma, in combination with Sm-Nd model ages, indicates juvenile crust formation, uninfluenced by old continental material. c. 550 Ma: Timanian Orogeny formed a 'modern' collisional belt. The magmatic activity of island arc signature culminated at c. 550 Ma with an intense pulse. Most likely this event is related to collisional processes and final accretion of Timanian terranes to the East European craton. Remnants of the Timanian (or pre-Uralide) Orogen are not restricted to the Polar Urals but can be found and correlated throughout the western flank of the Urals. c. 490 Ma: In Late Cambrian to Early Ordovician time, a previously unrecognized pulse of granitic magmatism affected the Timanian basement of the Polar Urals region. The granites have volcanic arc-type elemental chemistry and Sr-Nd isotopic signatures, interpreted to be inherited from their Timanian sources. This magmatism marks the onset of rifting, which finally led to the opening of the Proto-Uralian oceanic basin, parallel to today's Main Uralian fault and immediately east (modern coordinates) of today's Marun-Keu complex. The Silurian (Saveliev et al 1999; Ivanov et al 1975) transition from extension to convergence, marked by incipient island-arc magmatism, is not seen in the Marun-Keu complex. This indicates a passive margin setting of the Marun-Keu rocks from Ordovician times until the start of the Uralian Orogeny in the Devonian.

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J. GLODNYCTAZ,

c. 355 -360 Ma: Uralian eclogite-facies metamorphism occurre during east-directed subduction of Baltica's passive margin, in the course of Uralian arc-continent collision. The metamorphic event had a very limited impact on zircon age systematics. Mos of the investigated zircon crystals were found to preserve their U-Pb signatures through the eclogite-f acies event. Only those crystals for which a synmetamorphic crystallization can be inferred, either from their textural position in eclogite facies fluid precipitates or from their clear 'metamorphic' appearance in CL imaging, document the age of the metamorphic event. The impact of metamorphism on Rb-Sr and Sm-Nd wholerock systematics cannot be quantified exactly, but appears to be unimportant. The Uralian process of arc-continent collision, with subsequent thrust emplacement of allochthons onto the former Baltica margin, is largely responsible for todays structural features of the Polar Urals region. The extinct and exhumed Uralian subduction-accretion system in the Polar Urals resembles similar geological features in the Southern Urals, which have been compared to currently active systems in the Circum-Pacific (Brown & Spadea 1999), e.g. the ongoing arc-continent collision in Papua-New Guinea.

In the Marun-Keu complex, Uralian processes reworked and restructured the Timanian lithosphere but original information on protolith history and crustal evolution was preserved. Understanding this history provides important constraints for the reconstruction of both the Timanian and Uralian orogenies in the Polar Urals region. Main parts of this work were funded by a grant of the European Commission (TMR-URO Programme, research network contract No. FMRX-CT96-0009 (DG 12-ORGS)). Financial support by the Research Council of Norway (NFR) and by the Nansen Foundation is gratefully acknowledged. We are indebted to V. Dushin, V. Koroteev, E. Khain, G. Savelieva and particularly to the late V. Lennykh for excellent expert guidance in the field area. J. G. wishes to thank for the support provided by the GeoForschungsZentrum Potsdam for completion of this work. M. Whitehouse and F. Bea are thanked for great hospitality and invaluable support in analytical work. A. Krasnobaev kindly supplied sample 4G. The efforts of G. Bye Fjeld, J. Herwig, D. Woroschuk, J. L0kken, M. Langanke and T. Enger with sample collection and processing are gratefully acknowledged. Special thanks to E. Kramer for REE analyses, R. Naumann for XRF analyses, H. Kemnitz for some of the CL images, and M. Dziggel for figure design. R. Hetzel and K. Grafe provided helpful suggestions on an earlier version of the manuscript. Careful and constructive reviews by F. Corfu and D. Brown, as well as detailed comments and editorial handling by D. G. Gee are gratefully acknowledged. The NORDSIM facility is funded by the

Appendix: Sample characterization Table A. Samples analysed in this study

Sample

rock type*

GPS coordinates

Variegated intrusive suite (Neoproterozoic III) J3 MGr 67°29.72'N, 66°31.73/E

Mineralogy1^

Notes

qz, fs, gar, pheng, ru, zir, ap, ep/czo

strongly deformed metagranitoid from eclogite-facies shear zone eclogite with strong eel. -facies foliation, rich in phengite, intergrown with J12c amphibolite, without foliation, dominated by amphibole, intergrown with J12a slightly deformed granitoid granitoid, migmatitic-nebulitic appearance granitoid, partly deformed, magmatically intermingled with PU5 1 mafic hybrid, intermingled with PU50. fine-grained, from magma mingling zone dioritic hybrid, exhibits pseudomorphs after pig, with interstitial quartz mafic dyke in magma mingling zone, fine-grained, crosscuts PU58 gabbro, coronitic textures, incomplete reaction to eclogite fine-grained mafic eclogite; probably resembles equivalent to PU62 migmatitic-nebulitic granitoid; locally metasedimentary and igneous enclaves coarse-grained K-spar-dominated metagranite

J12a

ME

67°29.70/N, 66°32.01'E

omp, pheng, amp, pg, q, ru, zir

J12c

GA

67°29.70/N, 66°32.01'E

gar, amp, pg, pheng, q, ru, zir

PU12 J23 PU50

MGr MGr MGr

67°34.85'N, 66°39.92/E 67°27.56N, 66°37.63/E 67°30.00/N, 66°33.55'E

q, pig, K-fs, bi, gar, pheng, ilm, all, ep, amp, ap, zir q, K-fs, ap, zir, pig, ru, ep/czo, chl, sph, gar, pheng q, K-fs, pig, ep/czo, gar, pheng, ilm, zir

PU51

MDo

67°30.00/N, 66°33.55'E

nd

PU58

MQd

67°29.72N, 66°31.98/E

q, ep/czo, amp, omp, di, sph, wm

PU59

MDo

67°29.72N,66°31.98/E

nd

PU62

MG

67°29.04/N, 66°29.61/E

PU63c

ME

67°29.04/N, 66°29.61/E

ol, opx, di, omp, gar, amp, zo/czo, pheng, bi, ky, q, ru, Cr-sp, ilm, s, Cl-ap omp, gar, ep/czo, amp, q, ru

018

MGr

67°30.00/N, 66°33.55/E

pheng, q, pig, gar, zir, bi, ilm

019

MGr

67°30.00/N, 66°33.55/E

pig, q, ilm, gar, zir, ru, K-fs, pheng, ep/czo, bi

Meta-leucogranites (Late Cambrian -Early Ordovician) PUlOa MGr 67°34.11/N, 66°38.80'E PU22 MGr 67°33.75/N, 66°35.90/E

q, pig, K-fs, bi (green), gar, pheng, ilm, ep, all, zir gar, q, K-fs, zir, ap, sph, ep/czo, all

magmatic appearance, no deformation magmatic appearance, slightly deformed, garnet megablasts felsic dyke, crosscuts mafic eclogite, with metasomatic overprint and foliation

MGr

67°33.68/N, 66°36.40/E

q, pheng, zir, ap, gar, amp, ru, ep/czo

Metasediments J24 QFG J25 QFG

67°28.65/N, 66°36.13/E 67°28.65/N, 66°36.13/E

q, ep/czo, pig, gar, pheng, ap, ru q, ep/czo, amp, omp, pig, gar, pheng, ap, ru, zir

meta-migmatite; similar to and intercalated with J25 meta-migmatite; similar to and intercalated with J24

Southern part of Marun-Keu

q, phi, gar, amp, pheng, zir

Sample supplied by A. Krasnobaev (Yekaterinburg)

J30

Eclogite-facies vein 4G MS

* QFG, quartzofeldspatic gneiss; MS, metasomatic rock; GA, garnet-amphibolite; ME, mafic eclogite; MG, metagabbro; MQd, meta-quartzdiorite; MGr, metagranite. t all, allanite; amp, amphibole; ap, apatite; bi, biotite; chl, chlorite; Cl-ap, Cl-apatite; Cr-sp, Cr-spinel; czo, clinozoisite; di, diopside; ep, epidote; gar, garnet; ilm, ilmenite; K-fs, K-feldspar; ky, kyanite; ol, olivine; omp, omphacite; opx, orthopyroxene; pg, paragonite; pheng, phengite; phi, phlogopite; pig, plagioclase; q, quartz; ru, rutile; s, sulphides; sph, titanite; wm, white mica; zir, zircon; zo, zoisite; nd, no description.

PROTOLITH AGES OF ECLOGUES, MARUN-KEU, POLAR URALS

research councils of Denmark, Finland, Norway and Sweden, together with the Swedish Museum of Natural History. This is a EUROPROBE publication, and NORDSIM publication No. 75.

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PECHERSKY, D. M., KHAIN, V. V. & MATVEENKOV, V. V. 1984. Plate tectonic model of the South Urals development. Tectonophysics,W9,95-l35. ZULEGER, E. & ERZINGER, J. 1988. Determination of the REE and Y in silicate materials with ICP-AES. Fresenius' Zeitschrift fur Analytische Chemie, 332, 140-143.

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The Dzela complex, Polar Urals, Russia: a Neoproterozoic island arc D. REMIZOV1 & V. PEASE2 Institute of Geology, Komi Science Centre, Russian Academy of Sciences, 54 Pervomayskaya Street, 167610 Syktyvkar, Russia ^Department of Geology and Geochemistry, Stockholm University, SE-106 91 Stockholm, and Laboratory for Isotope Geology, Swedish Natural History Museum, Stockholm, Sweden (e-mail: [email protected]) 1

Abstract: The Neoproterozoic Dzela complex in the southern part of the Polar Ural Mountains occurs within the Uralian fault zone. It represents residual oceanic lithospheric mantle, partial melt derived from it, and seafloor basalts. In low strain domains, a depleted mantle residue of ultramafic rocks (pyroxenite, Iherzolite and dunite) and mafic rocks (gabbro, gabbro-norite and quartz gabbronorite) are preserved. Distinctive negative europium anomalies in the ultramafic rocks (and their metamorphosed equivalents) indicate their residual nature after basaltic melt extraction. Gabbroic rocks (and their metamorphosed equivalents) have elevated rare earth element (REE) concentrations, but similar overall REE patterns as the ultramafic rocks that host them, suggesting that the ultramafic rocks are the source of the gabbroic melts. Blueschist- and greenschist-facies metabasalts overlie the ultramafic and mafic rocks. Two groups of metabasalts are defined. One has NMORB-like REE patterns, but is uniformly more depleted than NMORB, and the other has slightly enriched (relative to NMORB) LREE/HREE patterns. The latter has high field-strength elements (HFSE) concentrations consistent with formation in an intraoceanic subduction zone setting. Dzela mafic rocks have been dated for the first time to 578 + 8 Ma (2a) (U-Pb zircon from quartz-bearing gabbro-norite), defining a Neoproterozoic crystallization age. The complex was probably accreted to the NE margin of Baltica (sensu lato) during the Timanian Orogeny. Locally reset U-Pb zircon ages of about 350 Ma probably record Uralian orogenesis and recrystallization associated with high-pressure/low-temperature metamorphism. The Dzela complex provides further evidence that ocean-continent collision was the driving mechanism for Timanian Orogeny.

In the southern part of the Polar Urals, metamorphosed allochthons are exposed along tectonic boundaries between Ordovician continental slope deposits of the East European Craton and ophiolite massifs of the Palaeozoic volcanic belt (Fig. 1). The Uralian fault zone has been interpreted to be a thrust zone separating continental slope deposits of the East European Craton to the west, from Palaeozoic ultramafic rocks to the east which includes the Voikar massif (Fig. 1). The eastern limit of this fault zone is defined by serpentinite and the western limit by unmetamorphosed Ordovician sediments. The Dzela complex is located SW of the Voikar ultramafic massif in the southern part of the Polar Ural Mountains (Fig. 1). This allochthon has been variously interpreted as an extension of the Palaeozoic Voikar ophiolite (PernTev 19790; Savelieva 1987), as the northern continuation of the Platinum Belt of the Urals (Efimov & Potapova 1990, 2000) and as part of the early Proterozoic basement of the Russian platform (Pystin 1994). We present new geological, geochemical, and geochronological data from the Dzela ultramafic/mafic complex indicating that the Dzela complex is a partially metamorphosed fragment of a Neoproterozoic island arc. It occurs in the footwall to the Main Uralian Fault (MUF), indicating that it was part of the preUralian East European Craton margin.

Geological setting of the Dzela complex The geology of the Dzela allochthon presented here is based on our new mapping. The internal structure of the Dzela allochthon, unrecognized by previous workers (e.g. Pystin 1994), is an asymmetrical recumbent antiform with a truncated lower limb (Fig. 1). The Dzela complex occupies a similar structural position as the Marun-Keu Complex further north (see Glodny et al. 2004) and the Ufaley Complex in the Middle Urals (Echtler et al. 1997). The fold axis is generally parallel to the strike of the MUF, dying out gently to the SW and NE. The Dzela complex is

predominantly amphibolite facies gabbro, but also contains ultramafic rocks (Fig. 1). To the west the Dzela complex is surrounded by belts of blueschist- and greenschist-facies metamorphism (Fig. 1). In the eastern limb of the fold, the margins of the Dzela ultramafic massif contain melanocratic garnet amphibolite. To the south, in the hinge of the antiform, coarse-grained (5-15 cm), granular hornblende pegmatite is bordered by garnet amphibolite. In its more central part, the core of the antiform is composed of garnet pyroxenite, Iherzolite and dunite. To the SE, serpentinite melange separates low-grade Silurian-Devonian rocks of island arc association from higher-grade rocks of the Dzela massif. Fragments of lilac coloured Lower Ordovician slates, thought to be derived from the western slope of the East European Craton, are present in this melange and suggest a Uralian tectonic association. To the north in the eastern limb of the antiform, uniform granular, meso-leucocratic foliated norite and gabbro-norite grade into quartz gabbro-norite. In the core of the antiform these grade into amphibolite, which preserves relict gabbroic textures (Fig. 2). In the overturned, lower limb of the fold, amphibolite and gabbronorite are bordered by interlayered chlorite schist, mica schist and amphibolite. To the west, glaucophane schist and greenschist are also present in the overturned limb of the antiform. The greenschist comprises metabasalt with relict ophitic textures (Fig. 3), metasedimentary rocks and jasper. It also contains relicts of glaucophane and garnet, in addition to sediments with abundant carbonate and quartz. Relict glaucophane and garnet indicate that greenschist metamorphism retrograded blueschist facies conditions. Greenschist facies metabasalt is in tectonic contact with low grade Early Ordovician metasediments along a thrust fault. In the southern core of the antiform, the foliation associated with the ultramafic and mafic rocks of the Dzela complex is rotated into parallelism with the metabasalt, which is truncated by serpentinite and overlying Silurian-Devonian island arc rocks, indicating that some of the deformational fabric in the region is pre-Uralian in age.

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 107-123. 0435-4052/047$ 15 © The Geological Society of London 2004.

107

Fig. 1. Regional setting, simplified geological map, and schematic cross-section of the Dzela complex. Cross-section has no vertical exaggeration (same scale as the geological map). The Main Uralian Fault zone (MUFZ) occurs between low-grade Early Ordovician shelf sediments of the East European Craton and serpentinized island arc volcanic rocks of Silurian-Devonian age and is delimited by the dashed lines.

NEOPROTEROZOIC DZELA ISLAND ARC

109

Fig. 2. Photomicrograph of gabbro-norite showing magmatic anhedral textures. (A) parallel nichols, (B) crossed nichols. Opx, orthopyroxene; Cpx, clinopyroxene; PI, plagioclase.

Magmatic rocks of the Dzela complex Ultramafic rocks

Fig. 3. Photomicrograph (crossed nichols) of relict igneous textures in glaucophane schist. Ophitic texture (lower, central grain outlined in white) with dark pyroxene enclosing light plagiocase laths. Subophitic texture (grain in upper left outlined in white) with dark pyroxene crystals partially enclosing light plagioclase laths.

The western limb of the antiform is structurally complicated due to an irregularly developed foliation and localized isoclinal folding. Small bodies of ultramafic rocks are preserved in the hinges of these isoclinal folds and are composed of little-altered peridotite and pyroxenite, identical in composition to the rocks of the main Dzela massif.

Magmatic rocks of the Dzela Complex include variably tectonized ultramafic, gabbroic-noritic, quartz-bearing gabbroic-noritic, and hornblende pegmatite compositions. The ultramafic rocks preserved in the core of the Dzela massif are little altered and diverse. They include peridotite and pyroxenite, and dunite to websterite to banded olivine clinopyroxenite (Streckeisen 1976). The transition between these rock types is gradational and expressed by the variation in abundance of olivine, orthopyroxene, clinopyroxene and ore minerals. Interstitial chrome spinel is present up to 12%. Serpentinization of olivine is variable, but generally less than 10-15%. Olivine websterite comprises serpentinized olivine (up to 50%), with orthopyroxene and clinopyroxene in roughly equal amounts. Spinel makes up less than 10% of the rock volume and occurs as inclusions in pyroxene or as disseminated fine-grained granules in olivine altered to serpentinite. The abundance of spinel is proportional to the modal olivine content. These rocks are anhedral granular, massive, spotted, or layered. The layered structure is due to the segregation of olivine in a pyroxene matrix. Selvages of ultramafic rock in the massif are metamorphosed to garnet hornblendite. Relicts of pyroxene are preserved. In websterite, chains of magnetite border large crystals of pre-existing (skeletal) pyroxene, now replaced by green hornblende. Thin magnetite plates follow the orthogonal crystal structure of pyroxene and define a rim of disintegration around it which is preserved within the new overgrowth of hornblende. All amphibolitized

Fig. 4. Gabbroic symplectite. (A) parallel nichols, (B) crossed nichols. Ol, olivine; PI, anorthite; Sp, spinel; Opx, orthopyroxene; Cpx-Sp, clinopyroxene-spinel symplectite.

110

D. REMIZOV & V. PEASE

Table 1. Dzela whole rock geochemical analyses Sample Rock type Si02 Ti02 A1203 Fe203

FeO MnO MgO CaO

Na20

K2O P205

LOI

Total

La Ce Pr* Nd Sm Eu Gel*

Tb Dy* Ho* Er* Tm* Yb Lu Rb Cs Sr Ba Cr Co As Sb Sc Hf Th U Ta Zr Ni Se Zn

DR-525/1 D 33.97 0.03 1.53 5.54 6.82 0.15 43.97 0.94 0.07 0.02 0.05 6.62 99.71 0.20 0.57 0.10 0.58 0.26 0.066 0.32 0.055 0.35 0.075 0.21 0.03 0.18 0.03 24.9

0 0 145 4,330 138.4 1.92 0.18 4.79 0.36 0.18 3.59 0.095

DR-522/2 Hz 37.34 0.24 2.33 6.56 9.98 0.39 35.97 2.97 0.15 0.02 0.05 4.06 100.07 0.51

1.5 0.24

1.3 0.54 0.056 0.51 0.074 0.43 0.081 0.19 0.022 0.099 0.014

14 0.48

54 530 15,785 172.4 10.2 0.34

13 0.46 0.51 1.56

10

0 32

1430

1190

1.9 90

1.5 110

DR-502/2 PHz 38.52 0.05 7.99 7.06 6.21 0.17 28.17 4.99 0.24 0.02 0.05 6.61 100.07 0.35 0.84 0.12 0.54 0.16 0.019 0.23 0.043

DR-685/1 Grpx

DR-521/2 Px

DR-501/1 Gn

— — — — — — — — — — — — —

40.89 1.25 13.56 5.63 8.28 0.15 13.36 12.33 2.09 0.45 0.05 2.03 100.05

40.66 0.57 19.02 7.03 6.57 0.12 8.98 15.08 0.52 0.02 0.05 1.45 100.07

51.25 0.79 21.87 3.54 5.10 0.18 3.18 9.63 3.60 0.07 0.24 0.58 100.02

41.73 1.03 6.59 8.39 10.43 0.32 19.03 10.75 0.59 0.02 0.05 1.15 100.07

2 4.5

— — — —

-

2.9 7.4 1.1 5.5 1.8

1.5 4.0

0.58

0.71

0.38

0.82

3.8 1.6

2.2 1.2

4.3 1.7

0.67

0.24

0.13

0.76

2

1.9

1.4

2.2

0.33

0.33

0.23

0.39

2 0.4 1

2.1 0.43

1.4 0.3

0.55

0.64

2.8 0.73 0.66

1 0.19

0.3

1.3

0.069 0.21 0.033 0.22 0.041

0.29 0.84 0.13 0.84 0.15

14 0 245 250

42.6

0.7 550 510

62.3 109.4 0.83 0.25 5.82 0.26

38.4 27.7 3.95 0.11 33.8

0

0.95 1.47

0.62

0 45 100 3.8 40

0.2 0 35 90 1.9 0

— — — — — — — — — — — — — — — — — — — — — — — — — —

—

— — — — — — — — — — — —

— — — — — — — — — — — — — — — —

DR-513 Qgn

0.13

0.7 0.11

20 0 1,080

285 33.2 26.1

0 1.44 22.3 0.48 0.22 1.10 0.32

120 120 5.1 80

DR-521/3 Amphpx

1.2 0.17 0.97 0.16 20.1 0.38

0 1,890 145.8 80.9 6.51 0.92 69.2 0.92

0.6 1.66

0 45 160 3.2 200

DR-523/2 Gramph 40.43 0.91 13.35

5.1 10.33 0.19 13.45 13.44 0.99 0.12 0.05 1.69 100.05

1.8

DR-526/1 Gramph 48.14 0.99 13.6 5.18 7.55 0.26

9.3 10.56 1.76

0.1 0.08 2.48

100 1.8 5

2.6 1.5

0.79 0.11 0.63

0.23

0.1

0.24

22 1.21

195 165 56.7 53.9 4.24 0.14 62.5 0.45 0.21 2.07

0 47 110 2 10

1.4 29 0.47

735 220 49.6 35.6 10.1

0.6 54.4 0.33 0.13 1.47

0 90 40 0.1 310

Notes: Major element analyses, wt%; trace and REE analyses, ppm. INAA detection limit <2 ppm to ppb, depending on the element. D, dunite; Hz, harzburgite, Pllz, amphibolite; Miamph, mica-bearing amphibolite; Gl, glaucophane schist; Aec, albite-epidote-chlorite meta-basalt. *, REE values interpolated from measured

111

NEOPROTEROZOIC DZELA ISLAND ARC

DR-527/1 Miamph

DR-1056/1 Gl

DR-1056/2 Gl

48.92 0.34 18.01 4.66 3.72 0.17 7.76 11.21 2.05 0.17 0.05 2.99

40.68 0.28 17.74 4.10 2.62 0.10 16.79 10.66 1.90 0.02 0.05 5.01

38.25 0.2 19.28 3.81 2.19 0.08 12.28 16.84 1.59 0.02 0.05 5.46

100.05 8.9 19 2.6 11 2.7 0.55 3.1 0.53 3.3 0.69 1.8 0.26 1.5 0.25

99.95 0.23 0.75 0.12 0.69 0.38 0.18 0.86 0.19 1.3 0.27 0.78 0.12 0.73 0.12

100.04 0.33 2.10 0.16 0.85 0.32 0.14 0.45 0.086 0.58 0.13 0.38 0.058 0.37 0.057

20.7 0 210 155 77.8 18.9 4.43 0.16 33.6 5.18 3.97 0.93 0 110 0 3.1 10

0 — 681 — 45.5 46 — — 45.7 0 0 — 0

— — —

—

0 — 755 — 38.2 24.9 — — 36 0 0 — 0 — — — —

DR-1056/46 Gl — — — — — — — — — — — __

0.28 0.93 0.15 0.87 0.46 0.24 0.84 0.18 1.2 0.25 0.71 0.11 0.66 0.10 0 — 663 — 16.6 34.5 — — 32.2 0 0 — 0 — — — —

DR-113/6 Gl

DR-124/3 Gl

DR-124/10 Gl

DR-1058/1 Aec

DR-124/8 Aec

DR-124/9 Aec

DR-121/3 Aec

DR-113/2 Aec

49.66 0.92 15.06 3.52 5.52 0.16 9.02 6.59 4.83 0.17 0.26 4.40

48.78 1.02 13.66 4.46 5.39 0.15 10.07 7.8 3.78 0.65 0.24 4.00

47.62 1.05 12.86 3.23 5.84 0.31 8.73 10.76 4.21 0.06 0.32 5.00

52.14 1.10 13.89 4.97 8.97 0.36 6.6 4.85 4.49 0.14 0.2 2.3

48.32 1.48 16.34 4.62 6.89 0.24 6.95 5.04 5.44 0.22 0.41 4.09

51.38 1.07 15.69 3.3 5.26 0.2 6.91 6.31 4.86 1.38 0.32 3.34

51.12 1.02 15.22 2.67 5.68 0.13 9.98 4.07 4.71 1.28 0.26 3.87

— — — — — — — — — — — —

100.11 7.5 16 2.4 11 3.8 0.98 3.9 0.61 4 0.86 2.4 0.36 2.2 0.28

100 5.1 13 1.9 9.1 2.5 0.9 3.5 0.65 4.2 0.87 2.4 0.34 2.0 0.28

99.99 8.3 21 2.8 13 3.5 1.2 3.8 0.63 4.1 0.88 2.5 0.36 2.2 0.27

100 1.7 6.3 1.1 6.4 2.6 1.2 4.2 0.82 5.5 1.2 3.6 0.55 3.5 0.52

100.04 16 42 5.4 24 6.4 2.1 7.4 1.2 8 1.7 4.5 0.64 3.8 0.54

100 11 28 4 19 4.2 1.5 4.3 0.68 4.4 0.94 2.6 0.38 2.3 0.32

100 10 26 3.6 17 4.6 1.5 5.8 1 6.5 1.4 3.7 0.52 3.1 0.39

— 18 35 6.3 29 5.6 1.7 5.6 0.86 5.5 1.1 3.1 0.43 2.5 0.29

0 — 0 — 96.7 31.4 — — 38.2 2.5 1.5 — 0 — — — —

0 — 0 — 102 36.2 — — 45.1 1.2 0.95 — 0 — — — —

0 — 0 — 56 33.5 — — 39.8 1.75 0 — 0

0 — 0 — 46.7 27.7 — — 32.8 3.4 1.6 — 0 — — — —

0 — 0 — 70 26 — — 26 2.6 1.4 — 0 — — —

29.2 — 0 — 30 24.9 — — 35.5 1.8 1.5 — 0

45.7 — 0 — 70.5 29.3 — — 34.2 2.9 2.3 — 2.5 — — — —

0 — 0 — 117 30 — — 36 1.8 1.1 — 0 — — — —

— —

— —

—

— — — —

plagioclase Iherzolite; Px, pyroxenite; Grpx, garnet pyroxenite; Gn, gabbro-norite; Qgn, quartz gabbro-norite; Amphpx, amphibolitized pyroxenites; Gramph, garnet elements: Pr from La and Nd, Gd from Sm and Tb, Dy, Ho, Er, and Tm from Tb and Yb. —, not measured; 0, below analytical detection limits.

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D. REMIZOV & V. PEASE

ultramafic rocks contain green chrome spinel. This is distinct from the black spinel present in amphibolite-facies gabbroic rocks of the study area.

Mafic rocks Near the core of the Dzela antiform, a concordant body of leucocratic gabbro-norite with magmatic texture is preserved. Away from the core of the antiform, these igneous textures gradually give way to a well-developed foliation defined by the presence of secondary hornblende. The periphery of the gabbro-norite is composed of amphibolite, visually indistinguishable from other amphibolites in the Dzela complex. The contacts between ultramafic and mafic rocks are generally conformable. The gabbro-norite is light grey, massive (or with a faint layering), and is anhedral to subhedral (Fig. 2). It is composed of plagioclase (An55_70), hypersthene, and clinopyroxene. Plagioclase varies from 45-75% in this rock. Clinopyroxene is more abundant than euhedral hypersthene, which forms prolate, subcollateral idioblasts and defines the compositional layering of the rock. Green hornblende forms aggregates replacing pyroxene crystals. A small amount of fine-grained serpentinized olivine may be present. A symplex of spinel and pyroxene is variably developed (Fig. 4). Magnetite is accessory, though may comprise up to 10% of these rocks. Other accessory minerals can include apatite, zircon, titanite, and leucoxene. Quartz-bearing gabbro (and its amphibolitized equivalents) is composed of subhedral aggregates of saussuritized plagioclase (An25_2v), epidote, and quartz, with relicts of hypersthene. Quartz occurs interstitially or forms lenticular aggregates accompanied by the growth of secondary muscovite. Hornblende pegmatite in the core of the Dzela antiform consists of large crystals of primary brown hornblende, magnetite, and zoisite. The pegmatite probably reflects a melt with an enriched volatile component/high partial pressure of water during the late stages of magmatic differientation. We suspect that the development of zoisite within the hornblende pegmatite is secondary and may be due to the later alteration of plagioclase (see following section).

Metamorphic rocks of the Dzela complex Amphibolite fades Amphibolite-facies metamorphism is widespread in the Dzela complex and the transition between relatively fresh pyroxenite and gabbro to garnet amphibolite and garnet plagioclase amphibolite, respectively, is observed in the field. A large region of garnet amphibolite in the eastern part of the study area contains 10-30% garnet as a major rock-forming mineral. Garnet also occurs as thin veinlets and less frequently as clots of large crystals. Garnet substitutes for plagioclase on the periphery of veinlets and in this context is probably associated with fluid migration during metamorphism. Throughout the Dzela complex, garnet amphibolite shows localized replacement of plagioclase by zoisite. This saussuritization results from the replacement of plagioclase and/or hornblende during retrograde metamorphism (An-rich plagioclase —» zoisite + albite). This low-temperature retrograde metamorphism is variably developed from minor saussuritization of plagioclase in veins (which due to their original monomineralic composition have been preserved), to the extreme development of nearly bimineralic zoisite plus kyanite. This reaction is easily seen in the field, with melanocratic rocks grading into their metamorphosed, more leucocratic equivalents.

Blueschist fades A zone of glaucophane and garnet-glaucophane schist is traced around the overturned limb of the antiform west of the gabbro-norite and amphibolite (Fig. 1). Blueschist is composed of glaucophane and albite (<20%), + garnet and small amounts of muscovite. Banded glaucophane schist is composed of interlay ers of glaucophane-epidote-albite and quartzglaucophane-albite. Greenschist fades In the westernmost exposure of the overturned limb, epidotechlorite-albite-quartz slate occurs with variable amounts of calcite. This fissile, green to grey-green rock is folded and in thin-section is nemato-granoblastic with relict ophitic texture confirming a basaltic protolith. Accessory minerals include magnetite, apatite, and zircon. Pyrite and thin veinlets of calcite are secondary. Within the zone of greenschist metamorphism, relicts of glaucophane and garnet are present, in addition to metamorphosed sedimentary strata with abundant calcite and quartz. Consequently, greenschist metamorphism overprints blueschist metamorphism. Jasper is observed on the northern tributary of the Hojmad River (Fig. 1) as 0.25-0.3 m thick hardpan among greenschist, blueschist and metamorphosed basalt 500-700 m east of the 'basal' thrust fault. It is cherry-red, non-fossiliferous, and has a thin, banded layering. Blueschist and greenschist metamorphism overprints gabbro of the Dzela complex, indicating that this lowergrade metamorphism post-dates the Dzela igneous rocks. Geochemistry of the Dzela complex Samples of the ultramafic, mafic, and metamorphic rocks of the Dzela complex were collected for whole-rock geochemical analysis in order to determine their petrogenetic relationships and to constrain the genesis of the Dzela complex. Geochemical sampling of the ultramafic and mafic rocks of the Dzela complex targeted the freshest available material, as well as some metamorphosed samples when field relationships permitted their correlation to unaltered protolith. Sampling of metamorphic material could not be avoided in the case of the blueschist- and greenschist-facies metabasalts. The samples were analysed for major and trace (including rare earth) elements. High-field strength elements (HFSE) and rare earth elements (REE) are least mobile during metamorphic processes (e.g. amphibolite facies—Polat et al 2003; blueschist facies—Mocek 2001; eclogite facies—Becker et al 2000). Consequently, HFSE and REE are most likely to preserve original protolith signatures of the Dzela complex and are the focus of the following discussion. Analyses were made by neutron activation (Table 1). The values of Boynton (1984) are used for chondrite normalization. Ultramafic rocks Major and trace element abundances of Dzela harzburgite (DR-522/2), Iherzolite (DR-502/2), and dunite (DR-525/1) are generally refractory (Table 1). They have high MgO (28-44wt%, on a water-free basis), coupled with low A12O3 (1.53-7.99 wt%), CaO (1-4.99 wt%), and TiO2 (0.03-0.24 wt%). They also exhibit moderate depletion of Sc (<4 ppm), with correspondingly high Cr, Ni, and Co abundances (up to about 16000, 1400, 170 ppm, respectively). The transition from relatively fresh pyroxenite to amphibolitized pyroxenite or garnet amphibolite is observed in the field.

NEOPROTEROZOIC DZELA ISLAND ARC

113

Consequently, REE compositions of ultramafic rocks and their inferred metamorphic equivalents are plotted together (Fig. 5A). These rocks have absolute REE concentrations similar to chondrite, but with variable chondrite normalized REE patterns (Fig. 5A). Chondrite normalized light REEs are either unfractionated relative to heavy REEs ([La/Lu]N = 0.69-1.38) or enriched relative to heavy REEs ([La/Lu]N = 3.78). Some samples show a weak depletion of chondrite normalized light relative to middle REEs ([La/Sm]N = 0.48-0.59), whereas others display a weak enrichment of chondrite normalized light relative to middle REEs ([La/Sm]N = 1.38-1.72). These rocks define a notable negative Eu anomaly (Eu/Eu* = 0.30-0.70). Garnet pyroxentite (DR-685/1) and amphibolitized pyroxenite (DR-521/3) have similar REE patterns, but with greater absolute chondrite normalized REE abundances between chondrite and NMORB. Only the garnet pyroxenite displays a positive Eu anomaly (Eu/Eu* = 2.36).

Mafic rocks Fresh gabbro-norites (e.g. DR-501/1, DR-513) have c. 4051 wt% SiO2 and high A12O3 (>18 wt%), CaO (>15 wt%), low to moderate MgO (3-9 wt%), and low TiO2 (<0.8 wt%). The metamorphic transformation of gabbro to garnet amphibolite has been observed in the field. Consequently, chondrite normalized REE compositions of gabbro and its inferred metamorphic equivalents are plotted together (Fig. 5B). Their REE patterns, however, are not very similar, suggesting that not all garnet amphibolites represent metamorphosed gabbro. Caution should be used making such correlations. Dzela quartz gabbro-norite (DR-513) has enriched chondrite normalized light REE relative to heavy REEs ([La/Lu]N = 2.74), with absolute concentrations between chondrite and NMORB. Garnet amphibolite has variable chondrite normalized REE patterns, but generally define a weak (DR-526/1) to moderate (DR-523/2) middle REE enrichment ([La/ Sm]N — 0.30). The latter is accompanied by a pronounced negative Eu anomaly (Eu/Eu* = 0.31). In other samples, an Eu anomaly is variably developed, either absent or negative to slightly positive (Eu/Eu* = 0.31-1.20). Mica-bearing amphibolite (DR-527/1) shows chondrite normalized enrichment of light relative to middle REEs ([La/Sm]N = 2.1) and also has a negative Eu anomaly (Eu/Eu* = 0.58).

Blueschist- and greenschist-facies rocks The Dzela blueschist (glaucophane schist) and greenschist (albite + epidote + chlorite metabasalt) samples define two compositional groups on the basis of major element oxide variations. One group has low SiO2 (38-41 wt%), high A12O3 (18-19wt%), low to moderate CaO (ll-17wt%), high MgO (12-17wt%), and extremely low TiO2 (<0.3 wt%). The other group has higher SiO2 (48-51 wt%), with moderate A12O3 (13-16wt%), low to moderate CaO (4-17wt%), moderate MgO (7-10 wt%), and low TiO2 (c. 1 wt%).

Fig. 5. Chondrite normalized rare earth element distributions of the Dzela complex. (A) Ultramafic rocks (and metamorphosed equivalents). Lherzolite, dunite, and harzburgite, filled symbols; pyroxenites, open symbols. Field of spinel Iherzolite after the Basaltic Volcanism Study Project (1981). (B) Gabbro-norite (and metamorphosed equivalents). Fields of the Uralian platinum belt and the Palaeozoic Voikar ophiolite after Savalevia et al (2002). (C) Blueschist and greenschist. Field of NMORB after Regelous et al. (1999) and ultra-depleted pyroxenites associated with ophiloites after Rampone et al (1996). Sample DR-513 is equivalent toVP98-030.

Table 2. U-Th-Pb ion-microprobe analytical data

Sample (grain.spot)

U ppm

Th ppm

Pb ppm

Th/U

/206 %

207pb/206pb

±lCT

206

Pb/238U

%

Age estimates (Ma)

± lCT

%

207

Pb/206Pb

VP97-029 (metamorphosed dyke?) 206 664 Lie 100 324 2.1c 415 3190 3.1m 353 638 4.1c 22 100 5.1r 29 570 6.1c 0 7.1m 18 662 429 8.1c 0 9.1c 69 VP97-030 (quartz gabbro-norite) 1667 Lie 1609 2.1c 234 226 3.1c 117 288 4.1c 322 161 308 141 5.1c 14 6.1c 520 7.1c 471 289 8.1c 203 235 9.1c 126 279 lO.lc 643 328 11. Ic 203 97 12.1c 1136 587 135 13.1c 338 14.1c 95 59 15.1c 133 81 16.1c 268 117 17.1c 211 151 IS.lc 473 409 19.1c 723 420

Disc %

206

Pb/238U

65 31 166 63 5 32) 1 49 (4)

0.31 0.31 0.00 0.58 0.09 0.00 0.037 1.375 0.00

(0.07) (0.07) 0.36 (0.07) (2.22) 0.43 (1.17) 0.2 0.25

0.0578 0.0571 0.0543 0.0570 0.0654 0.0536 0.0564 0.0565 0.0532

0.52 0.65 0.45 0.52 2.67 1.24 3.79 1.00 2.38

0.0852 0.0826 0.0492 0.0801 0.0470 0.0544 0.0572 0.0770 0.0559

3.16 2.63 2.59 2.88 2.63 2.81 3.08 2.61 2.69

521 494 385 491 787 355 467 473 339

11 14 10 11 55 28 82 22 53

527 511 310 497 296 341 358 478 351

16 13 8 14 8 9 11 12 9

0 0 -12.4 0 -40.2 0 0 0 0

215 27 254 37 131 112 52 24 28 70 23 119 39 12 61 64 21 49 82

1.18 1.17 0.45 0.52 0.43 0.03 0.61 1.17 0.47 0.48 0.46 0.55 0.44 0.68 0.63 0.46 0.74 0.93 0.59

(2.45) (2.30) 0.02 0.19 0.04 0.01 0.50 0.07 0.08 0.15 0.23 0.54 0.09 0.00 0.13 0.02 0.19 0.69 0.05

0.0744 0.0722 0.2433 0.0590 0.1138 0.0759 0.0600 0.0571 0.0569 0.0594 0.0589 0.0599 0.0589 0.0590 0.1158 0.0770 0.0587 0.0612 0.0594

0.41 1.52 0.45 0.59 0.25 0.50 0.51 0.94 0.73 0.83 0.83 0.55 0.87 1.65 0.75 0.58 1.13 0.77 0.60

0.0902 0.0805 0.6484 0.0946 0.3462 0.1995 0.0893 0.0821 0.0824 0.0900 0.0931 0.0850 0.0985 0.0983 0.3525 0.1986 0.0780 0.0777 0.0915

2.56 2.66 3.23 2.67 3.09 2.60 3.45 2.59 2.85 2.16 2.14 2.15 2.14 2.15 2.14 1.79 1.79 1.79 1.80

1052

8 31 7 13 5 10 11 21 16 18 18 12 19 36 14 12 25 16 13

556 499 3222 583 1916 1172 551 509 510 555 574 526 605 605 1946 1168 484 482 564

14 13 83 15 51 28 18 13 14 12 12 11 12 12 36 19 8 8 10

-45.3 -40.4 0 3.8 0 1.7 -1.3 0 0 0 3.5 -6.5 0 0 0 0 -1.4 -18.6 0

992 3142

568

1861 1091

604 497 487 581 562 601 562 568

1892 1120

556 645 583

Notes: Analyses were performed on a high-mass resolution, high-sensitivity Cameca IMS 1270 ion-microprobe at the NORDSIM facility in Stockholm, Sweden. Analyses from cores denoted by 'c', from mixed zones by 'm', and from rims by 'r'. Errors in age estimates are quoted at la. All ages are calculated using the decay constants of Steiger & Jager (1977). Pb concentrations in parentheses determined after correction for common Pb. Th/U ratios calculated from 207Pb/206Pb and 208Pb/206Pb ages. /206(%X fraction of common Pb calculated from 204Pb; parentheses indicate no common Pb correction made. Disc.% refers to the degree of discordance at the 2a error limit between 207Pb/206Pb and 206Pb/238U ages. Reverse discordance is indicated by positive numbers. Analyses more than 5% discordant were not used in final age determinations.

NEOPROTEROZOIC DZELA ISLAND ARC

The same groups are defined by chondrite normalized REE patterns (Fig. 5C). The 'LREE-depleted' group has depleted chondrite normalized light relative to heavy REEs ([La/Lu]N = 0.20-0.60) and includes blueschist (DR-1056/1, DR-1056/2, DR-1056/46) and a greenschist with relict glaucophane (DR-1058/1). This light REE depletion relative to heavy REE is NMORB-like, but absolute concentrations of REE are lower than NMORB (Regelous et al 1999). Concentrations of HFSE (Ce, Ta, Th, Hf) are also less than NMORB and generally below analytical detection limits. Heavy REE are unfractionated ([Ho/Lu]N = 1-1.1), with absolute concentrations similar to NMORB. The 'LREE-enriched' group includes both blueschist- and greenschist-facies samples. It is defined by slightly enriched chondrite normalized light relative to heavy REEs patterns ([La/Lu]N = 2-6), with absolute LREE concentrations similar to, or greater than, NMORB (Fig. 5C). Heavy REE concentrations are similar to NMORB, but are slightly fractionated ([Ho/Lu]N = 1.3-1.8). Concentrations of HFSE are variable: Ce (average 26 ppm) is higher than in NMORB (but less than in continental crust; Wilson & Davidson 1989); Th and Hf (a proxy for Zr) are NMORB-like; Ta (a proxy for Nb) much less than NMORB (sometimes below detection limits).

lateral extent and the presence of xenoliths, it is most likely a dyke. It has a vertical layer-parallel foliation striking N55E, concordant with the foliation of the amphibolite. This sample, therefore, either predates deformation or is syntectonic and can provide at least a minimum age for the protolith. The sample comprises a medium to coarse grained equilibrium assemblage of predominantly zoisite and plagioclase. It represents static recrystallization under low-temperature conditions, with mutually inclusive zoisite and plagioclase (Fig. 6). It is composed of medium grained (4 mm), seriate, optically continuous anhedral, twinned plagioclase (49%), with fine- to medium-grained euhedral zoisite (49%) and acicular muscovite (2%) oriented along cleavage planes within the plagioclase (Fig. 6A,B). Zoisite displays pronounced anomalous blue interference colours, parallel

U-Pb dating of the Dzela complex Zircon can be a highly robust geochronometer, even after the effects of polyphase, high-grade metamorphism, e.g. Whitehouse et al. (1999). Therefore, it is the most likely mineral to record protolith ages in the Dzela complex. Sampling of the Dzela complex, a predominantly ultramafic/mafic complex in which zircon is not expected to be abundant, was restricted to the less common but more 'evolved' lithologies present. Their petrography and geological settings are discussed below. The separation of zircons from whole-rock samples was performed using conventional magnetic and heavy liquid mineral separation techniques. Th-U-Pb analyses of zircons from these samples were performed using a Cameca IMS 1270 ion-microprobe at the NORDSIM Facility, Swedish Museum of Natural History, Stockholm. The results are presented in Table 2. Analytical procedures are similar to those described by Whitehouse et al (1997, 1999). Geostandards 91500 with an age of 1065 Ma (Wiedenbeck et al 1995) was used for internal U-Pb ratio calibration. Common Pb corrections were made using the measured 204Pb signal assuming a Stacey and Kramers (1975) composition of appropriate age. The ion beam was typically oval and c. 25 jjim in its long dimension. Ion-beam spot placement and size was verified using a scanning electron microprobe after ion-probe analyses and is indicated by the oval in the cathodoluminesence (CL) images. Due to the small size of the ion-beam and low lead counts from relatively young/unradiogenic grains, 206 Pb/238U ages are reported for grains younger than 1.0 Ga and 207 Pb/206Pb ages are reported for grains over 1.0 Ga. The data are plotted on inverse concordia (Tera-Wasserburg) diagrams (207Pb/206Pb v. 238U/206Pb). Concordia ages (Ludwig 1998) are reported at the 2cr confidence level and are indicated by black ellipses. Analyses that had more than 5% discordance were not included in the final analysis.

VP97-029, metamorphosed dyke (?) Zircons were separated from a foliated leucocratic layer occurring in amphibolite, which is regionally associated with metamorphosed pyroxenite. The leucocratic layer is 3 m thick and laterally extensive. It is difficult to determine the protolith of this sample due to overprinting metamorphism (see below), but given its

115

Fig. 6. Photomicrographs of equilibrium epidote-amphibolite facies assemblage, VP09-029. (A) High relief, euhedral zoisite (Zo) crystals both including and included in medium grained, optically continuous plagioclase (PI); small, thin needles of muscovite (Ms) occur along cleavage planes within the plagioclase (parallel nichols). (B) (crossed nichols). (C) Large zoisite corona in the centre of the field of view surrounds aligned quartz + plagioclase subgrains (parallel nichols).

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D. REMIZOV & V. PEASE

Fig. 7. Cathodoluminesence images of zircons from foliated meta-dyke (?), VP97-029. Ion-beam location and size indicated by the oval.

extinction, both concentric and sector compositional zoning, and simple twins (Fig. 6B). Zoisite often forms a vermicular symplectite with plagioclase (Fig. 6C) indicating some subsolidus re-equilibration. Zoisite also includes accessory (<1%) anhedral rutile. Several medium- to coarse-grained disequilibrium coronas can be seen in thin-section (Fig. 6C). These corona features comprise c. 20% of the thin section and consist of euhedral zoisite crystals mantling cores of aligned subgrains of recrystallized quartz (18%), sericitized (chlorite + muscovite) plagioclase (77%), and zoisite (5%). Petrographically and mineralogically the sample represents an originally plagioclase-rich rock which later statically recrystallized under epidote-amphibolite-facies conditions. The corona features are interpreted to be sedimentary xenoliths in chemical and thermal disequilibrium with the host melt. Nine zircon grains were separated from c. 3 kg of sample. The zircons are stubby to elongate (aspect ratio 1:1 to 4:1, respectively), subhedral, transparent, and light pink with rare inclusions. CL images (Fig. 7) show that some grains preserve a relict internal morphology of oscillatory growth zoning or sector zoning (grains 1, 2, 3, 4, 5, 6, 8, and 9), though this structure can be either partially (grains 5 and 9) or entirely bleached (grain 7). Most grains are mantled by thin bright, amorphous rims (grains 1, 2, 3, 4, 6, 8 and 9). All nine grains were analysed (Table 2) with generally concordant results (7/9 grains). The two grains with more than 10%

Fig. 8. Inverse concordia diagram for foliated metamorphosed dyke (?), VP97-029. All analytical data are plotted with la error bars, while concordia ages (Ludwig 1998) are reported at the 2a confidence level and are indicated by grey ellispses (dashed lines represent analyses > 5% discordant which were not included in the final analysis). The older age population may represent an igneous age, whereas the younger represents the time of metamorphism.

NEOPROTEROZOIC DZELA ISLAND ARC

117

Fig. 9. Photomicrograph of quartz gabbro-norite, VP09-030. Note anhedral magmatic textures. (A) parallel nichols, (B) crossed nichols. Orthopyroxene (Opx), clinopyroxene (Cpx), plagioclase (PI), and quartz (Qtz). Chemical analysis of this sample is represented by DR-513 in Table 1 and Figure 5.

discordance were excluded from the final analysis. Two populations are defined (Fig. 8). An older concordia age of 501 + 9.7 Ma is defined by zircons that preserve internal structure (grains 1, 2, 4, 8). A younger concordia age of 350 ± 1 1 Ma is defined by cracked and bleached grains (grains 6, 7, 9). Note that the analysis of grain 6 (6.1c in Table 2) was located on a crack. The thin bright rims were too small to analyse.

VP97-030, quartz, gabbro-norite Zircons were separated from a fresh quartz gabbro-norite preserved in a low-strain domain within the Dzela complex. This sample is equivalent to DR-513 in Table 1. A locally developed foliation is defined by pyroxene and plagioclase mineral segregation and a preferred plagioclase shape orientation. This foliation

Fig. 10. Cathodoluminesence images of zircons from VP97-030. Ion-beam location and size indicated by the oval.

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D. REMIZOV & V. PEASE

strikes east-west, in contrast to the regional NE-trending foliation, and dips variably. Younger cross-cutting fractures, faults, and hornblende-bearing andesitic dykes trend NE, parallel to Uralian structures and the MUF. The quartz gabbro-norite represents the most evolved compositional variant of gabbroic rocks in the Dzela complex and may provide a protolith age for the Dzela complex. The sample comprises a medium-grained, weakly deformed, inquigranular, subhedral quartz gabbro-norite with 70% plagioclase, 13% orthopyroxene, 11% clinopyroxene, 2% quartz, and 3% oxides (Fig. 9). Oxides are always associated with pyroxene, either along grain boundaries or as inclusions within pyroxene. Plagioclase has serrate to well-developed polygonal grain boundaries indicative of recrystallization. In addition, minor late, brittle deformation of plagioclase (grain bending, twin distortion, and kink banding) is associated with crystals at higher angles to the shape aligned fabric defining the foliation. A good zircon yield was obtained from c. 3 kg of sample. Zircon external morphology is sub- to euhedral, equant to elongate (aspect ratios 1:1 to 3:1), and transparent, light pink to dark pink-purple in colour with occasional inclusions. In CL (Fig. 10), some grains show concentric oscillatory growth zoning mantled by thin bright, amorphous rims (grains 3, 7, and 11). Other grains show sector zoning (grain 5) or appear partially (grain 8) or totally bleached (grains 9). Nineteen grains were analysed (Table 2) with complex results (Fig. 11). A concordia age of 578 ±11 Ma is defined by a significant proportion of grains which preserve concentric oscillatory growth zoning (grains 4, 7, 10, 11, 13, 14, and 19). In addition, a large percentage of concordant older dates (1.09 to 3.14Ga) were obtained from grains with both oscillatory and sector zoning (grains 3, 5, 6, 15, and 16). Three grains (grains 8, 9, and 17) yield apparently concordant dates at c. 510 Ma, but appear to be bleached (Fig. 10). Many significantly discordant (>5%) analyses are associated with either high-U and/or high-Th zircons (grains 1 and 12) or bleached zircons (grains 2, 12, and 18). Significantly discordant (>5%) results are excluded from the final analysis. The results associated with radiogenic and bleached grains often document lead loss (Rubatto et al. 1999;

Pidgeon et al. 2000). Indeed, modern lead loss on an inverse concordia diagram is indicated by a horizontal shift to the right along the 238U-206Pb axis, which is well documented in this sample (Fig. 11). The thin bright rims were too small to analyse. Discussion Petrogenesis Ultramafic and mafic rocks. Dzela ultramafic rocks represent oceanic lithospheric mantle residual after melt extraction; the mafic rocks apparently represent partial melt(s) derived from this lithospheric mantle. Arguments supporting this interpretation are developed below. (1) Dzela complex ultramafic rocks are compositionally diverse, from pyroxenite to dunite, including plagioclase Iherzolite. Their compositional variation in the field is gradational and not defined by tectonic and/or structural contacts, so we presume that they are somehow genetically related. These rocks are tectonized and generally lack cumulate fabrics, though layering in olivine clinopyroxenite provides evidence of some crystal segregation. (2) Despite the compositional variation of the ultramafic rocks, their chondrite normalized REE patterns are similar inasmuch as they are relatively flat (with the exception DR-522/2, LREE enriched harzburgite). They show weak middle REE enrichments or depletions (relative to LREEs or HREEs), generally accompanied by a negative Eu anomaly similar to mantle xenoliths (Fig. 5). (3) Mantle Iherzolites commonly have chondritic REE abundances (Ix chondrite), though there is a wide range of compositions from LREE depleted to LREE enriched. REE enrichments 10 times chondrite represent a maximum in mantle Iherzolites (Basaltic Volcanism Study Project 1981, pp. 292-297). Dzela ultramafic rocks have REE abundances and patterns similar to spinel Iherzolites (Fig. 5), and lack the depleted chondrite normalized LREE signature associated

Fig. 11. Inverse concordia diagram for quartz gabbro-norite, VP97-030. All analytical data are plotted with la error bars, whereas concordia ages (Ludwig 1998) are reported at the 2a confidence level and are indicated by grey ellispses (dashed lines represent analyses >5% discordant which were not included in the final analysis). The protolith crystallization age is defined.

NEOPROTEROZOIC DZELA ISLAND ARC

(4)

(5)

(6)

(7) (8)

with partial melts of shallow mantle lithosphere, e.g. NMORB or ophiolitic peridotites. The flat chondrite normalized REE patterns of the Dzela ultramafic and mafic rocks indicate a shallow lithospheric mantle source, rather than a deeper, garnet-bearing asthenospheric mantle source. A lithospheric mantle source is consistent with the stability fields of both plagioclase and spinel, and the observed presence of plagioclase Iherzolite and spinel Iherzolite in the Dzela complex. Dzela ultramafic and mafic rocks have predominantly negative Eu anomalies which also suggests that plagioclase was present in the source. Plagioclase Iherzolite is a component of the Dzela complex, so this is a logical conclusion and supports their derivation from a shallow mantle lithosphere. Of course, a variable oxidation state for Eu in the mantle source could also produce Eu anomalies, but this would be an ad hoc argument given the presence of plagioclase Iherzolite in the field. Dzela ultramafic and mafic rocks are intimately associated in the field. The mafic rocks have REE patterns similar to the ultramafic rocks, but with higher absolute concentrations. This suggests that the ultramafic rocks could be the source of the mafic melts via partial melting. The association of the ultramafic and mafic rocks with 'oceanic' metabasalts (see below) suggests that these rocks were generated in an oceanic setting. The Uralian platinum belt in the Kytlym region has middle REE enriched chondrite normalized patterns and the Palaeozoic Voikar ophiolite has light REE depleted chondrite normalized patterns (Savelieva et al. 2002). The negative Eu anomalies and flat REE patterns associated with Dzela ultramafic and mafic rocks distinguish them from the Palaeozoic Voikar ophiolite and from the intrusive rocks of the Uralian Plantinum Belt (Fig. 5).

Consequently, we interpret the ultramafic and mafic rocks of the Dzela complex to represent, respectively, a variably depleted residue of oceanic lithospheric mantle and the partial melt derived from it. This ultramafic residue is heterogeneous, reflecting either an originally heterogeneous source or variable degrees of partial melting of the original source. A likely environment for this intimate association of rock types might be a magma chamber. Coarse-grained hornblendite in the core Dzela antiform consists of primary brown hornblende that reflects a melt with an enriched volatile component/high partial pressure of water, e.g. the late stages of magma differentiation in a magma chamber. Blueschists and greenschists. We interpret these rocks to represent metamorphosed oceanic basalts, most probably generated in an intra-oceanic subduction zone environment. This interpretation is supported by the following: (1) Two distinct geochemical signatures are preserved within a continuous stratigraphic sequence of metabasalts found intercalated with jasper, indicating they formed in an oceanic environment. (2) LREE-depleted metabasalt has depleted LREE/HREE, high A12O3, high MgO, low CaO, extremely low TiO2, low LILE (e.g. Sr, Ba, Rb), and low HFSE (e.g. Ce, Ta, Th, Hf). Absolute concentrations are less than NMORB indicating that the mantle source for these rocks was NMORB-type mantle that had experienced a previous depletion event, i.e. previous melt extraction. Typical NMORB-type mantle for normal tholeiitic mid-ocean ridge basalt forms from shallow lithospheric mantle (Regelous et al 1999 and references therein). (3) LREE-enriched metabasalt has small enrichments of LREE/ HREE relative to NMORB, moderate A12O3, MgO, CaO, low TiO2, and low LILE. They also have NMORB-like (to slightly higher) concentrations of fractionated HREE. They

119

have low, NMORB-like (to slightly higher) concentrations of HFSE (but higher than the LREE-depleted group). Small variations in HREE concentrations in these rocks might be explained by fractional crystallization processes, i.e. fractionational crystallization of hornblende (Rollinson 1993), which increase REE concentrations without changing the REE pattern. Alternatively, the slightly enriched LREE/ HREE and fractionated HREE may indicate that garnet was a residual phase in the mantle source from which these basalts were derived. Consequently the LREE-enriched group could be derived from deeper, garnet-bearing asthenospheric mantle. (4) Dzela metabasalts lack some of the diagnostic geochemical indicators associated with the asthenospheric source of plume/ocean island basalt (OIB) genesis. One geochemical signature associated with OIB is the strong fractionation of LREE/HREE, due to both the presence of garnet in the source and the less depleted nature of asthenospheric (versus lithospheric, NMORB-type) mantle (Wilson 1989). Additionally, HFSE behave incompatibly and are preferentially concentrated (relative to NMORB) in OIB mantle sources: Zr/Nb is characteristically low in OIB (<10) and Nb and Ce higher than NMORB (Basaltic Volcanism Study Project 1981). Although Dzela metabasalt geochemistry permits them to be derived from garnet-bearing asthenospheric mantle, the small degree of LREE-enrichment (<10x NMORB) combined with low HFSE, suggests that they do not in fact represent melts derived from plume/OIB-type mantle. (5) Both the LREE-depleted and LREE-enriched groups have generally low HFSE, e.g. low Ce, Hf, and Ta. Ta can be regarded as a proxy for Nb, except in the presence of an aqueous fluid where Ta would be more abundant than Nb (Bau 1996; Kamber & Collerson 2000). Regardless, Ta concentrations in these metabasalts are so low that negative Ta (and Nb) anomalies must be pronounced (Fig. 5). Ta (and Nb) anomalies suggest the involvement of either mixing with partial melts of continental crust with a residual Nb, Ta, and Ti bearing mineral phase, or a subducted slab component (e.g. Drummond & Defant 1990; Drummond et al. 1996). Low Ce concentrations (average 2.5 ppm, LREE-depleted group; average 26 ppm, LREE-enriched group) preclude the involvement of continental crust in the genesis of these metabasalts (Wilson & Davidson 1984), consistent with point 1 above. Consequently, the low Ce and Ta (and Nb) concentrations favour the involvement of a subducted slab component in their genesis and suggest an ocean island arc setting for Dzela metavolcanic rocks. The evidence shows that these metabasalts formed in an oceanic environment. This, combined with the above distinctive geochemical features, favours the genesis of both the LREE-depleted and LREE-enriched metabasalts in an intra-oceanic island arc setting.

Timing Concentric, oscillatory zoning in zircon is evidence of an igneous origin, i.e. crystallization from a melt (Hanchar & Miller 1993). Continuous oscillatory zoning suggests a single growth phase, whereas truncated or embayed zoning indicates a resorption event that may be either magmatic or metamorphic. Textures and zoning in zircon associated with fluids, either late fluids associated with igneous processes or those associated with metamorphic processes, tend to obscure or erase igneous zoning and lead to the development of ghostly (relict), bleached, or amorphous zoning. Such textures are often associated with secondary processes, record later metamorphism/metasomatism, and can

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re-set U-Pb ages, e.g. Rubatto et al 1999; Pidgeon et al 2000; Corfu et al 2003; Tomaschek et al 2003. VP97-029 (metamorphosed dyke?). The older concordia age of 501 + 10 Ma (Fig. 8) from zircons with relict concentric oscillatory zoning (Fig. 7) probably represents the time of igneous crystallization. This is in agreement with a growing number of studies that document regional igneous activity at about 500 Ma and is thought to be associated with the initiation of post-orogenic rifting, e.g. Glodny et al. (2004 and references therein), Remisov (unpublished data). We interpret this sample to be a dyke because it contains metasedimentary xenoliths. An igneous age of 500 Ma would permit it to have intruded after obduction of the Dzela complex onto the margin of Baltica (sensu lato), which would be necessary to provide a source for the sedimentary xenoliths. Alternatively, an older protolith may have experienced Pb loss and slip along concordia, with the false 'concordia age' merely an artefact of the small number of analyses. This is not completely satisfactory, however, as there is no mechanism for the entrainment of the sedimentary xenoliths. We favour the former interpretation, but regardless of the interpretation the age of c. 500 Ma provides a minimum age for the Dzela complex. The younger age of 350 ± 11 from bleached and cracked grains documents a time of Pb loss consistent with a later fluidmetamorphic event. This younger, metasomatic event is documented only locally by the U-Pb isotopic system in samples from the Dzela complex, i.e. it is not recorded in all samples (see VP97-030 below). Pervasive, secondary zoisite in metamorphic equilibrium is abundant in this sample. Secondary zoisite is unstable at high temperatures, consequently, it probably formed under lowtemperature conditions at this time. A major low-temperature, high-pressure metamorphic event at c. 360 Ma is well defined in the Marun-Keu complex of the Polar Urals (Glodny et al. 2000, 2002, 2003; Scarrow et al 2002 and references therein). It reflects the time of subduction of the eastern margin of Baltica during collision with the Siberian craton after closure of the Uralian ocean (Brown et al 1996 and references therein). This is in good agreement with the age and low-temperature nature of metasomatism in the Dzela complex. Though the U-Pb zircon geochronometer is very robust, memory loss and recrystallization are facilitated by the presence of aqueous fluids (Villa 1998; Tomaschek et al 2003). Consequently, we presume that localized, late, low-temperature/ high-pressure fluid migration associated with Uralian orogenesis probably occurred throughout the Polar Ural region at this time and locally reset the U-Pb isotopic system. VP97-030 (quartz gabbro-norite). We interpret the oldest (1.09-3.14 Ga) concordant ages (Fig. 11) to represent inherited (xenocrystic) zircon grains. The main population of concordant analyses at 578 Ma from grains with igneous growth zoning (Fig. 10) are interpreted to represent the crystallization age of the quartz gabbro-norite. As part of the Dzela complex, this sample thus provides a crystallization age for the complex. Additionally, the preservation of spinel-pyroxene symplectite in gabbro of the Dzela complex (Fig. 4) indicates slow cooling at depth. The P-T regime at this time, documented by the spinelpry oxene symplectite, is between 1300 °C/18 kbar and 700 °C/ 5 kbar (Griffin 1972). A large number of discordant analyses represented by bleached and amorphous zircons reflects a disturbed isotopic system and Pb loss. The timing of this later event is poorly constrained to less than 500 Ma. Basalt genesis. The lack of obvious tectonic breaks between the metabasalts and the ultramafic and mafic rocks in the field, the structural position of the metabasalts in the footwall of the MUF, and their truncation by serpentinite and Silurian-Devonian island arc rocks of the hanging wall, suggest that they are

Neoproterozoic in age. Consequently, though the metabasalts have not been dated, we infer that the entire Dzela complex is Neoproterozoic in age. The time of blueschist and greenschist metamorphism of these basalts must post-date their protolith age. The spatial association of this low-temperature, high-pressure metamorphism within the MUF zone suggests that it may be related to Uralian orogenesis. Tectonic setting The tectonic setting most consistent with the above constraints from the Dzela complex is an island-arc environment. During the initial stages of intra-oceanic subduction, shallow oceanic lithospheric mantle is melted, generating an intimate association of basic magma and depleted mantle residue (Dzela ultramafic and mafic intrusive rocks). This magmatism begins to thicken the oceanic lithosphere. Coarse-grained hornblendite, with a high volatile component/high partial pressure of water, may record the late stages of magma differentiation within a magma chamber, consistent with the geochemistry of the LREE-enriched metabasalts and a crystallizing assemblage that includes hornblende. The low LREE, HREE, and HFSE concentrations of Dzela LREE-depleted metabasalts indicate their derivation from a previously depleted source, requiring previous melt extraction(s) from a NMORB-type, shallow lithospheric mantle (old spreading ridge?). The LREE-depleted group occurs at the base of the basaltic stratigraphic succession, which suggests they were 'early' magmatic products of the arc. As subduction proceeds, the down-going slab continues to dehydrate, metasomatizing the overlying mantle wedge. This facilitates enrichment of LREE and, with the exception of Ce, depletion of HFSE which are retained in the slab. Ce is enriched in the wedge, being mobile in an aqueous fluid as Ce + 4. Dehydration of the slab and metasomatism of the mantle wedge promotes melt generation. Melt genesis may involve deeper, garnet-bearing (asthenospheric) mantle or crystal fractionation of hornblende, and results in the production of basalt with slab related and enriched LREE (relative to NMORB) geochemical affinities, e.g. the Dzela LREE-enriched metabasalts. Quartz gabbro-norite, intruding at about 580 Ma, provides a crystallization age for the mafic intrusive rocks and a minimum age for the Neoproterozoic residual oceanic lithosphere of the Dzela complex. This documents the time of island arc genesis under generally high P-T conditions. Given the number of inherited zircon grains of various ages present in the quartz gabbro-norite and the probable lack of inheritable zircon from its ultramafic-mafic host, it is likely that these grains represent a continental sedimentary input. As far as we know, the presence of detrital zircon in the abyssal plain has never been documented. Therefore, the intra-oceanic subduction zone that generated the Dzela complex was likely proximal to the trench/continental margin at 580 Ma and the accretion of the complex to the Baltica (sensu lato) margin would have occurred shortly thereafter. A period of extension, subsidence, and marine transgression is documented along the length of the Urals in Late Cambrian to Early Ordovician time (Ivanov et al. 1975; PernTev 1979&; Puchkov 1997; Bogolepova & Gee 2004). The formation of Late Cambrian to Early Ordovician granites in the Polar Urals (Glodny et al 2004) is contemporaneous with alkaline dyke intrusion into Late Cambrian to Lower Ordovician sediments of the west Uralian foreland (Saveliev 1997 and references therein), and probably document incipient rifting of the eastern margin of Baltica. This may be recorded by the c. 500 Ma metamorphosed dyke (?) intruding the Dzela complex. Part of the Neoproterozoic Dzela complex was subducted, experiencing blueschist metamorphism that later retrogressed to

NEOPROTEROZOIC DZELA ISLAND ARC

greenschist facies. The age of this metamorphism is not known, but its spatial occurrence in the footwall of the MUF suggests it is related to Uralian Orogeny. The Dzela complex records local resetting of U-Pb zircon systematics at 340-360 Ma. Presumably this metamorphic event occurred locally throughout the Polar Urals (e.g. Glodny et al 2004) and was associated with Uralian Orogeny.

Timanian Orogeny The recognition of the 580 Ma Dzela island-arc complex in the southern part of the Polar Ural mountains reveals a pre-Uralian collisional history associated with the Neoproterozoic margin of northeast Baltica (sensu lato) that is little documented. In the Marun-Keu complex of the northern Polar Urals there is also evidence for late Neoproterozoic island arc magmatism (Glodny et al. 2004). The entire eastern margin of the East European Craton was characterized by active ocean-continent convergence in the late Neoproterozoic (Sengor et al. 1993; Scarrow et al. 2001), suggesting that accreted island arc material of late Neoproterozoic age, preserved as re-worked allochthons within Uralian thrust sheets and overprinted by Uralian metamorphism, may be more common than previously thought. Passive margin continental slope deposits of Neoproterozoic age are exposed in the Timan Range (Roberts et al. 2004 and references therein). The current extent of these deposits is inferred on the basis of drillcores and geophysical data (Belyakova & Stepanenko 1991; Kostyuchenko 1994) to lie c. 100km to the NE, beneath the Palaeozoic and younger sediments of the Pechora basin. These rocks define the pre-Timanian Neoproterozic passive margin of NE Baltica, now greatly shortened as a result of Timanian and younger tectonics. Geographically, the Neoproterozoic margin of eastern Baltica (sensu lato) coincides closely with its later Palaeozoic (Uralian) margin in the Northern, Central, and Southern Urals, as indicated by the distribution of Neoproterozoic age complexes (Zonenshain et al. 1990). Compressional deformation and metamorphism, including late Neoproterozoic to Early Cambrian blueschist- and eclogite-facies metamorphism, took place along the western Uralian foreland in the Middle and Southern Urals (e.g. Brown et al 1996; Puchkov 1996, 1997; Beckholmen & Glodny 2004). In the sub-Polar and Polar Ural region, however, the eastern margin of Baltica (sensu lato) diverges from its Palaeozoic (Uralian) margin (Zonenshain et al. 1990), with over 300 km of allochthonous material accreted between them (Pease & Gee 1999, 2003; Dovzhikova et al. 2004) prior to the intrusion of c. 550 Ma granitoids. The NE Palaeozoic (Uralian) margin of Baltica apparently follows this 'new' late Neoproterozoic accretionary margin. It is this fragmentary evidence of late Neoproterozoic collision, subduction, and island arc accretion that suggests that ocean-continent collision was the driving mechanism for Neoproterozoic Timanian orogenesis. The duration of Timanian convergence is poorly constrained, but is inferred to have lasted about 70 Ma. An alkaline complex intrudes Neoproterozoic passive margin deposits in the Timan Range at about 615 Ma (Larionov et al. 2004). This alkaline magmatism is thought to reflect extension of Baltica's (sensu lato) pre-Timanian Neoproterozoic passive margin. The tectonic environment then changed to a convergent setting with the onset of Timanian Orogeny. By c. 550 Ma, post-tectonic granitoids intruded across the basement to the Pechora Basin (Gee et al. 2000; Pease et al. 2004), marking the end of Timanian orogenesis. Accretion during this c. 70 Ma period would have been punctuated or episodic, rather than continuous, as oceanic highs (plateaus and arcs) approaching the subduction zone were forced onto the continental margin. Subduction polarity reversals were probably associated with such accretionary events.

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Consequently, the Timanian Orogeny along the entire eastern margin of Baltica (sensu lato) involved crustal growth via the accretion of oceanic crust and island arcs at this time. Along the northeastern margin of Baltica (sensu lato) this growth was either much more voluminous, or better preserved due to, for example, a more favourable geometry of the margin during subsequent convergence/rifting. Though much of this material was re-worked during Uralian Orogeny, information regarding the age and composition of its protolith and its metamorphic history is preserved. Future studies that seek to identify and describe Neoproterozoic rocks and their metamorphic history within Uralian thrust sheets further south will help to define Timanian orogenesis along the eastern margin of Baltica.

Conclusions 1. The Dzela complex comprises ultramafic and mafic rocks, and metabasalts. The complex represents residual oceanic lithosphere, the partial melt derived from it, and seafloor basalts derived from previously depleted, shallow NMORB-like mantle, as well as slightly LREE-enriched (relative to NMORB), deeper (garnet-bearing?) mantle sources. The tectonic setting most consistent with the geochemical data is an island arc. 2. Quartz gabbro-norite crystallized at 580 Ma under generally high P- T conditions and inherited zircon reflects the development of an island arc in proximity to Baltica (sensu lato). 3. Parts of the complex were subsequently subducted, metamorphosed under blueschist-facies conditions, and later retrogressed under greenschist-facies conditions. The age of this metamorphism is unknown, but is probably Uralian as it occurs in the footwall to the Main Uralian Fault. 4. Locally, metamorphism/metasomatism took place at c. 350 Ma and was probably related to Uralian orogenesis. 5. The Neoproterozoic Dzela complex was accreted to NE Baltica (sensu lato) during the Timanian Orogeny and was reworked during thrusting associated with younger Uralian Orogeny. 6. The inferred tectonic setting for the Neoproterozoic Dzela complex suggests that ocean-continent collision was the driving mechanism for the Timanian Orogeny. The authors gratefully acknowledge the support of N. Yushkin, the field work of N. Kostelov, support from the Swedish National Science Research Council, Uppsala University, the Swedish Museum of Natural History, and EU INT AS grant no. 97-1139. Thorough and constructive reviews by J. H. Scarrow and H. Austrhiem, as well as careful and constructive editorial comments by D. Gee, improved the quality of this manuscript. Thanks are also extended to the NORDSIM staff. The NORDSIM facility is funded by the research councils of Denmark, Norway, Sweden, the Geological Survey of Finland, and the Swedish Museum of Natural History. This paper is NORDSIM publication no. 95.

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Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia MONICA BECKHOLMEN1 & JOHANNES GLODNY2 1 Department of Earth Sciences, University of Uppsala, Villavdgen 16, S-752 36 Uppsala, Sweden (e-mail: [email protected]) GeoForschungsZentrum Potsdam, Telegrafenberg C2, D-14473 Potsdam, Germany

Abstract: The core of the Kvarkush Anticline, on the western side of the Northern Urals, contains pre-Ordovician metasediments and minor metabasites that have undergone multistage deformation and transitional greenschist- to blueschist-facies metamorphism. Metamorphic conditions for the blueschist-facies assemblages are estimated at 350-400 °C and 7-8 kbar. The metamorphic rocks are cut by mafic dykes, one of which yields a Sm-Nd mineral age of 398 ± 37 Ma (2a). The age of the high-pressure/low-temperature (HP/LT) metamorphism has been controversial. New Rb-Sr mineral data yield ages for the waning of blueschist-facies metamorphism at c. 536 Ma (535 + 6 Ma, 535.8 + 6.5 Ma, 536 ± 1 9 Ma, all 2a), demonstrating that the specific conditions required for blueschistfacies metamorphism were attained in latest Neoproterozoic to earliest Cambrian times. The new data are the first well-constrained documentation of HP/LT metamorphism within the Timanian Orogen. We suggest that easterly-directed subduction of the East European continental margin (present-day coordinates) during the Timanian Orogeny, involving compressional thickening of the margin sediments, is the most likely scenario for the formation of the Kvarkush blueschists. The Kvarkush HP/LT metasediments, together with lower greenschist-facies and subgreenschist-facies metasediments further west, form parts of a Timanian erogenic wedge on an east-dipping thrust ramp. Reworking during the Uralian orogenesis was of minor importance.

The East European platform is flanked to the east by the Uralian Orogen, which was formed by continent collision with an assemblage of Siberian-Kazakhian terranes in Devonian to Triassic time (Hamilton 1970; Zonenshain et al 1984). The Main Uralian Fault (MUF), inferred to be the principal tectonic suture of the Uralian Orogen, separates oceanic and island-arc lithologies in the east from rocks with East European continental affinity to the west. Along the 'European' flank of the Uralides, pre-Uralian metamorphic and magmatic complexes crop out in several nappes in the cores of antiforms below Early Palaeozoic platform cover. This preUralian basement shows evidence for at least one earlier orogenic event along the eastern edge of the East European Craton, which occurred in Late Neoproterozoic to Early Cambrian times. This event is termed the Timanian Orogeny; it has previously often been referred to as the 'Baikalian' (e.g. Shatskij 1963). Other terms used in the Russian literature are Pre-Uralides (Douralidy) or Rifeidy. This orogen has been correlated with the Cadomian of western Europe (Puchkov 1988, 1997; Scarrow et al 2001). Timanian structures are exposed along the entire eastern and northeastern margin (modern coordinates) of the East European Craton, particularly in the Bashkirian Anticlinorium in the southern Urals, in the Kvarkush-Kamennaya Gora Anticlinorium in the Middle to Northern Urals and in the Timan Range (Fig. 1). The Timanian orogenic belt continues northwestwards through the Kanin Peninsula (Southern Barents Sea) to the Varanger Peninsula in Northern Scandinavia. Other occurrences of Timanian rocks are known from drillcores in the Pechora Basin (Beliakova & Stepanenko 1991; Gee et al. 2000) and from outcrops along the western flank of the Polar Urals (Dushin 1997; Khain et al 1999; Scarrow et al 2001; Glodny et al 2004; Remizov & Pease 2004). The Timanian Orogeny mainly affected rocks that are Meso- to Neoproterozoic in age, spanning the 'Riphean' and 'Vendian' in Russian literature. The term 'Riphean' is derived from the Roman name for the Urals Mountains, the type section is the stratigraphic section in Bashkiria in the Southern Urals (see Maslov 2004 and references therein). In the Urals, these rocks occur at a tectono-stratigraphic level within the upper part of the Uralian foreland fold-and-thrust belt (e.g. Glasmacher et al 1999, 2001), beneath Palaeozoic strata, mainly shelf carbonates, of the Uralian passive margin (Fig. 1). The existence of the Timanian ('Baikalian') Orogeny has been questioned by Ivanov and Rusin (e.g. Ivanov 1977, 1981;

Rusin 1984, 1997; Ivanov & Rusin 1986, 2000). These authors have argued that the metamorphism and deformation commonly attributed to Timanian Orogeny are due instead to the increase in temperature and strain in the bedrock caused by slow, longterm rifting, and the final opening of the pre-Uralian ocean. They argue against a Timanian, Late Neoproterozoic Orogeny due to lack of Timanian-age oceanic lithologies and ophiolites, and the absence of unequivocal Timanian HP/LT metamorphic assemblages. Unconformities between Neoproterozoic and Palaeozoic rocks have been explained to be due to deposition on surfaces tilted due to listric faulting. The first of these two objections is no longer substantiated because, in the Engane-Pe Anticline (western foreland of the Polar Urals), fragmented ophiolites have been discovered (e.g. Dushin 1997) that mark a suture in the Timanian Orogen (Scarrow et al 2001) and undoubtedly have a Late Neoproterozoic age (Khain et al 1999). Pre-Palaeozoic basement containing blueschist is located within the Kvarkush Anticline, near the junction between the Timan fold belt and the Uralides in the Northern Urals (Fig. 1). The blueschist assemblages occur in pre-Uralian structures and are therefore potentially related to the Timanian Orogeny. The identification of HP/LT rocks such as blueschists is of key importance for understanding the evolution and the dynamics of the Timanian Orogen; generally, such rocks indicate processes of subduction, crustal thickening and formation of orogenic wedges (e.g. Ernst 1982; Platt 1986). The aim of this study has been to investigate the age and tectonic significance of the blueschist-facies high-pressure rocks reported from the Kvarkush hinge area (Figs 1 & 2).

Geological setting: the Kvarkush Anticline Within the framework of the Uralides, the Kvarkush Anticline (Fig. 2) forms the northern part of the c. 400 km long KvarkushKamennaya Gora Anticlinorium (Ablizin et al 1982). The Kvarkush hinge is situated in the 'triple point' between the southeastern extension of the Timan Range and the western flank of the Uralides, near the southern tip of the Pechora Basin (Fig. 1). Following the traditional division of the Urals into tectonic zones (e.g. Ivanov et al 1975), the Kvarkush-Kamennay a Gora

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125-134. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Schematic geological overview map, mainly based on work of Ivanov et al (1975) and Zonenshain et al (1990). The Timanides and Timanian age complexes within the framework of northeastern Europe (inset) and the Uralian Orogen. Numbers denote position of outcrops of Timanian complexes: (1) Varanger Peninsula; (2) Rybachij-Srednij Peninsula; (3) Kildin Island; (4) Kanin Peninsula; (5) Tsil'ma Kamen', Chetlas Kamen' and Vym Ridge; (6) Och-Parma and Dzhezhim-Parma; (7) Kvarkush-Kamennaya Gora Anticlinorium; (8) Bashkirian Anticlinorium. Letters refer to geological entities discussed in the text: (a) Kvarkush Anticline (study area); (b) Serebr'anka; (c) Ufaley complex; (d) Maksyutov complex.

Anticlinorium is located in the Central Uralian Zone, the eastern margin of which is the Main Uralian Fault (MUF). The main axis of the anticlinorium trends obliquely to the MUF, and the eastern limb of the anticlinorium is truncated by the MUF. The core of the Kvarkush Anticline is dominated by metasedimentary and mafic metavolcanic successions, cut by mafic dykes. These mostly greenschist-facies rocks are unconformably overlain by a thin Ordovician conglomerate and sandstone sequence followed by Palaeozoic limestones of the Uralian passive margin. The rocks above the unconformity have suffered less deformation and display a lower grade of metamorphism. The area has been investigated in detail by Esipov (1953), Starkov (1957, 1963, 1969) and L'vov (1957). A 1:200 000 map was published by Esipov (1960), providing the basis for Fig. 2. These authors describe the metamorphic rocks as argillaceous quartzites, albite - chlorite - sericite, chlorite - epidote - albite actinolite and epidote-glaucophane-garnet schists that have suffered some retrogression. The age of metamorphism, although

uncertain, was considered by most authors to be pre-Ordovician on geological evidence. The highest degree of metamorphism is recorded in the northern part of the Kvarkush Plateau, around the Verkhn'aya and Nizhn'aya Petelikha rivulets, tributaries to the Uls river (Fig. 2) with the occurrence of glaucophane and garnet assemblages (Starkov 1963) indicating blueschist-facies, high pressure-low temperature (HP/LT) metamorphism. Metamorphic temperatures and pressures for assemblages involving crossite are estimated at 350-400 °C and 7-8kbar (Rusin 1996). The rocks generally document different, prograde and retrograde stages of metamorphism and deformation, the significance and ages of which were for many years a matter of discussion (Starkov 1963; L'vov 1958; Minkin & Yakovleva 1974; Rusin et al. 1989; Rusin & Nikiforov 1991). Typical metamorphic assemblages (Voroshchuk et al 1999) are muscovite + biotite/chlorite + albite + quartz ± garnet (almandine-rich) ± Na, Na-Ca amphiboles (barroisite, Na-barroisite, Ca-crossite, crossite, Fe-glaucophane,

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limestones, implying post-Ordovician blueschist-facies metamorphism (based on a quoted 400 Ma K-Ar biotite age). In contrast, on a metamorphic map presented by Alekseev (1996), the glaucophane schist occurrences are shown in the eastern half of the northernmost Kvarkush hinge area, all underneath the Ordovician unconformity. Pre-Ordovician metamorphism is also inferred by previous K-Ar analyses of white mica from the Kvarkush area, giving ages in the interval 503-553 Ma (Nikiforov & Kaleganov 1991). The interpretation of the tectonometamorphic evolution of the Kvarkush Anticline is thus crucially dependent upon a well-defined age for the blueschist-facies metamorphism. Sampling and sample characterization

Fig. 2. Detailed map of the Kvarkush Anticline (based on the 1:200 000 map by Esipov 1960). Sample locations: (1) Mica schists B99:80, B99:81 and B99:82; (2) Chlorite schist 68-1; (3) Chlorite schist H152-2; (4) Mafic dyke B99:48.

winchite), which are present as blue prismatic, needle-shaped crystals, often as inclusions in plagioclase and as pseudomorphs after actinolite. Assemblages containing andalusite, cordierite and fibrolite may occur locally close to intrusions, and were interpreted as being late relative to the glaucophane-bearing assemblages (Minkin & Yakovleva 1974). The regional distribution of the blueschist-facies lithologies is also a matter of debate. Uncertainties mainly arise from bad exposure due to a thick soil cover. The area has not been glaciated and the boulders are therefore of local origin. Mapping is to a large extent based on the composition of loose blocks. However, on the eastern flank there is better exposure, with the rocks in the core of the Kvarkush Anticline rising steeply from the Uls valley. Rusin (1996) published a metamorphic map of the Kvarkush Anticline, with three north-south-trending glaucophane-bearing zones, a few hundred metres wide, cutting the pattern of the regional metamorphism. The central zone was tentatively shown to cut the Early Palaeozoic unconformity and overlying Ordovician quartzites and

During fieldwork on the Kvarkush Plateau and hinge area in 1999, various lithologies were collected for investigation of structures, metamorphic history and, in particular, timing of the HP/LT blueschist-facies metamorphism. The Ordovician unconformity was also examined. The samples (localities given in Fig. 2) for direct determination of the age of the blueschist-facies metamorphism were collected in the classical Nizhn'aya Petelikha rivulet (60°23.72/ N, 58°43.52/ E) (Fig. 2, samples B99:80, -81, -82, sample location 1). The sampled metapsammitic mica schists and quartzitic schists are composed of quartz, albite, blue amphibole (glaucophane), white mica (muscovite-phengite; sample B99:82 has paragonite in addition) and apatite. Tourmaline, pyrite, zircon and titanite are accessories. The rocks show polyphase deformation. An early schistosity (SO is defined by white mica and a fine-grained opaque mineral (hematite?). This schistosity is isoclinally folded with mineral growth (white mica, glaucophane) in the axial plane (S2), defining the main foliation in the rock. This texture is then overgrown by albite porphyroblasts (Fig. 3a). Strong asymmetric folding rotates albite porphyroblasts and gives rise to a prominent crenulation cleavage (S3) and intricate interference patterns in outcrop. A later cleavage (84), often at high angle to the metamorphic surfaces, is similar to that of the main cleavage in the overlying Palaeozoic sedimentary rocks. Additional information on the local tectono-metamorphic history can be found in Beckholmen (2000) and for the Serebr'anka section further south (cf. Fig. 1) in Beckholmen & Juhlin (1997) and Beckholmen (1998). Two samples of greenschist-facies metamorphic rocks (sample 68-1, chlorite-epidote schist and HI 52-2, chlorite-albite schist) come from the collection of the Institute of Geology and Geochemistry of the Academy of Sciences in Ekaterinburg (Fig. 2, sample locations 2 and 3, respectively). The deformed metamorphic rocks of the Kvarkush Anticline are cut by basaltic dykes which are clearly post-tectonic in relation to the pre-Ordovician deformation. At least two generations of mafic dykes are reported from the area, one of which crosscuts the overlying Palaeozoic sediments and therefore must be post-Ordovician (e.g. Esipov 1960). Age data for these dykes therefore provide a minimum age for both tectonism and metamorphism. In Skali (60° 05.1 r N; 58° 37.19' E, Fig. 2, sample location 4), c. 35 km south of the blueschist locality, we sampled such a dyke (sample B99:48) that is undeformed, largely unaltered and crosscuts deformed quartz-mica schists (Fig. 3b).

Analytical procedures For Rb-Sr and Sm-Nd isotope analyses, concentrates of all separable mineral phases were produced, including glaucophane, apatite, feldspar, white mica (phengite/paragonite) epidote, biotite, chlorite, and pyroxene. Since white mica is the main

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TIMS instrument at the GeoForschungsZentrum Postdam. The observed ratios were corrected for 0.25% per a.m.u. mass fractionation. Total procedural blanks were consistently below 0.10 ng for both Rb and Sr. Sample amounts between 5 and 50 mg lead to very high sample/blank ratios, and, consequently, no blank corrections were applied. For age calculations, an uncertainty of ±1%, as derived from replicate analyses of natural mica samples, is assigned to the 87Rb-86Sr ratios. 87Sr/86Sr ratios are reported with their 2 sigma internal precision plus uncertainties from spike correction. In the calculations of isochron parameters, a standard error of 0.005% for 87Sr/86Sr ratios was applied if individual errors were smaller than this value (cf. Kullerud 1991). This error estimate (2cr) was derived from reproducibility tests for Sr isotope ratios of spiked samples. A decay constant for Rb of I ^ - I O ^ V 1 (Steiger & Jaeger 1977) was used for all Rb-Sr age calculations. Sm and Nd isotope analyses were carried out on a Finnigan MAT 262 mass spectrometer at the GeoForschungsZentrum Potsdam. Nd was analysed in dynamic, Sm in static mode. For calculation, standard errors of ± 0.004% for 143Nd/144Nd ratios and ±0.3% for 147Sm-144Nd ratios were assigned to the results. Regression lines are calculated using Isoplot/Ex (Ludwig 1999). All results are reported with 2a errors if not indicated otherwise.

Results

Fig. 3. (a) Microphotograph of mica schist in the N Petelikha rivulet, sample B99:82. Blue amphibole in the main foliation overgrown by albite porphyroblasts, c. 1 mm across, (b) Photo of mafic dyke cutting folded quartzite at Skali, sample locality (4) in Fig. 2, sample B99:48, upper part mafic dyke, lower part folded quartzite. The visible part of the pencil is c. 10 cm.

Rb-bearing phase in the blueschist-facies rocks, its Rb-Sr data strongly influence the age determinations. Thin section observations suggest that different generations of white mica may be present. Consequently, in order to control the effects of this on the apparent ages, the white mica populations were separated into physically different fractions. This separation was possible on the basis of grain size (samples B99:80, B99:81), magnetic properties (sample B99:82) and density using heavy liquids (sample B99:80). With the exception of sample B99:48 (basaltic dyke), whole rocks were not analysed because of the partly weathered appearance of the schist samples. Mica sieve fractions (all >100 |Jim, to exclude small, possibly retrogression-related crystals from analysis) were ground under pure ethanol in a polished agate mortar and then sieved in ethanol in order to obtain clean, inclusion-free mica separates. The concentrates of the other minerals were finally purified by hand-picking under the binocular microscope. The schists were analysed for Rb and Sr, and the mafic dyke rock for Sm and Nd contents by isotope dilution methods. They were weighed into Savillex screw-top containers, spiked with a suitable mixed 87Rb-84Sr or 149Sm-150Nd spike solution, and dissolved in a mixture of HF and HNO3. Solutions were processed by standard cation-exchange techniques. Determinations of Sr isotope ratios were carried out on a VG Sector 54 multicollector thermal ionization mass spectrometer (GeoForschungsZentrum Potsdam) in dynamic mode. The values obtained for the NBS standard SRM 987 during the period of analytical work were 0.710250 ± 0.000010 (n = 16). All isotopic ratios were normalized to an 86Sr/88Sr ratio of 0.1194. Rb analyses were carried out on a VG Isomass 54 single collector

Thin sections reveal that blueschist-facies metamorphism was followed by a retrograde metamorphic stage involving crenulation and growth of albite + white mica + chlorite. During this retrograde stage the blueschist-facies phases partly altered to greenschist-facies assemblages. In particular, glaucophane commonly shows partial chloritization and/or transformation to green amphibole. It seems likely that not only a late-metamorphic stage but also post-metamorphic processes, such as lowtemperature alteration, played a role in the retrogression process. The results of the Rb-Sr and Sm-Nd analyses are summarized in Tables 1 and 2, and plotted as isochron diagrams in Fig. 4. The results from the three blueschist samples B99:80, -81 and -82 are quite similar; they share a common apparent age of about 536 Ma (Early Cambrian) and a common, fairly high, initial 87 Sr/86Sr isotopic composition of about 0.728. The datasets show Sr-isotopic disequilibria between apatite, chlorite, glaucophane, and albite, as evident from the mean squared weighted deviate (MSWD) values for the regression calculations (between 20 and 87, Table 1). It was not possible to purify the glaucophane in sample B99:82 from adherent chlorite; therefore, the analytical data from the separate was not considered for age calculation. In contrast to first expectations from thin section observation, the white mica population in the blueschist samples is very homogeneous in terms of its Rb-Sr systematics. There are no resolvable differences in apparent age between different grain size fractions or fractions defined by density or magnetic properties. The results from sample B99:82 indicate that two chemically distinct white-mica phases (phengite and paragonite) are present. However, these different white-mica phases generate concordant Rb-Sr apparent ages. The age obtained for the chlorite schist of sample H152-2 (526 ± 36 Ma) is, within the limits of error, similar to the Early Cambrian ages from the blueschist samples. However, there is considerable scatter of the data points around the regression line. In particular, the chlorite data point plots below the regression line defined by the white-mica concentrate, albite and apatite. The initial 87Sr/86Sr isotopic composition (0.7363, as defined by the low Rb-Sr phase apatite) is even higher than that of the three blueschist samples. For sample 68-1 (chlorite schist), only chlorite and epidote were separable, resulting in a two-point isochron with an apparent age of 481.5 ± 6.4 Ma.

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TIMANIAN BLUESCHISTS IN THE NORTHERN URALS Table 1. Rb-Sr analytical data 87

87

87

6.14 16.97 17.07 17.61 18.74

3.369 63.18 61.36 60.01 1.783

0.754460 1.207787 1.198222 1.185739 0.740777

0.0022 0.0010 0.0012 0.0016 0.0016

12.04 12.40 20.32 1023 19.81

1.724 1.874 55.01 0.007 56.01

0.740747 0.742474 1.149068 0.728952 1.155617

0.0018 0.0014 0.0016 0.0020 0.0018

± 19 Ma; MSWD =• 20, Sr, = 0.7295 ±0.0087 wm m = 0.36A, <500 (Jim 371 wm nm = 0.36A, <500 |jim 335 albite 2.56 apatite 1.20 blue amp + chlorite* 9.76

17.70 25.83 31.16 889.0 7.62

63.60 38.74 0.238 0.004 3.713

1.212134 1.028932 0.730617 0.728659 0.738521

0.0014 0.0014 0.0018 0.0014 0.0016

H152-2 Age: 526 ± 36 Ma; MSWD = 86, Sr/ = 0.7346 ± 0.0091 PS404 albite (+ quartz) 0.96 wm (+ chlorite) PS378 266 PS375 apatite 0.38 PS377 chlorite 0.80

12.57 21.91 941.5 3.20

0.221 36.11 0.001 8.608

0.737772 1.006240 0.736275 0.795002

0.0016 0.0014 0.0016 0.0016

1022 11.51

0.003 0.873

0.720529 0.726313

0.0016 0.0018

Sample no. Analysis no. 899:80 Age: 535 PS411 PS368 PS367 PS366 PS332

Material

Rb (ppm)

± 6 Ma; MSWD = 36, Sr, = 0.7279 ± 0.0021 albite 7.11 wmHBRF, <500 |xm 353 wm LBRF, <500 jjim 345 white mica >355 jjum 349 blue amphibole cc. 11.5

899:81 Age: 535.8 ± 6.5 Ma; MSWD = 87, Srf = 0.7282 ± 0.0010 PS405 albite 1 7.16 PS363 albite 2 8.01 PS362 white mica > 160 jjim 370 PS360 apatite 2.56 PS350 wm 125-1 60 |xm 367 899:82 Age: 536 PS370 PS369 PS346 PS345 PS371

Sr (ppm)

68-1 Age: 481.5 ± 6.4 Ma; Sr{ = 0.720325 ± 0.000037 PS154 epidote (impure) 10.5 PS 155 chlorite 3.46

Rb/86Sr

Sr/86Sr

Sr/86Sr 2
Errors are reported at the 2a level. *denotes analysis not used for age calculation. An uncertainty of ±1% has to be assigned to Rb-Sr ratios. Abbreviations: HBRF, sinks in bromoform; LBRF, floats in bromoform; cc., concentrate; w. mica, white mica; m, magnetic fraction (Frantz isodynamic separator, at 13° tilt as standard condition); nm, non-magnetic fraction. Mineral separate impurities are due to finescale intergrowths between blue amphibole and chlorite (blue amphibole of samples B99:80, B99:82), between albite and quartz (albite of sample H152-2), between paragonite and phengite (white-mica separates of B99:82) and between epidote and chlorite (epidote of sample 68-1).

The basaltic dyke (sample B99:48) which cross-cuts deformed quartz-mica schists, yielded a Sm-Nd two-point isochron age, defined by plagioclase and pyroxene, of 398 ± 37 Ma. Wholerock isotopic data infer Sm-Nd isotopic equilibrium within the rock and thus the validity of this age value. Discussion Age of blue schist-fades metamorphism The Rb-Sr isotope data on the blueschist samples (B99:80, -81, — 82) and the white mica-bearing chlorite schist (HI52-2) consistently yield Early Cambrian ages of c. 536 Ma, accepting the base of the Cambrian at 545 Ma (McKerrow & van Staal 2000). The glaucophane-bearing samples (B99-80, —81, 82) show minor isotopic disequilibria (as evident from the high MSWD

values), and no high quality Rb-Sr-isochrons could be generated (Fig. 4). This is most probably due to the complex metamorphic evolution of the rocks, involving late-/fP-metamorphic albite blastesis as well as slight, static greenschist- to sub-greenschistfacies retrogression, involving partial chloritization of glaucophane and biotite. Such processes are well known to disturb Rb-Sr systematics of the whole-rock system and of those mineral phases involved in the retrograde reactions. However, phases which form textural relicts with respect to retrogression (i.e. white mica, apatite, unaltered blue amphibole, and possibly most of the albite) will retain their isotopic signatures through the retrogression observed in our sample set, because these relict phases did not take part in the retrogression reactions, and retrogression temperatures were too low to allow for thermallydriven isotope exchange. Therefore, the Rb-Sr systematics of these mineral phases were defined by blueschist-facies metamorphic processes.

Table 2. Sm-Nd analytical data Sample no. Analysis no.

Material Sm (ppm) Nd (ppm) 147Sm/144Nd 143Nd/14/144Nd

B99:48 Age (plag +px): 398 ± 37 Ma; Ndt = 0.512403 ± 0.000053 PS478 pyroxene 1.2207 2.7338 plagioclase PS480 2.6870 10.733 PS481 whole rock 4.0152 13.515

2a m (%)

0.2700 0.1514 0.1796

An uncertainty of ±0.3% has to be assigned to the Sm-Nd ratios. Errors at 2a level.

0.513106 0.512797 0.512888

0.0014 0.0014 0.0014

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Fig. 4. Internal mineral isochrons (Rb-Sr and Sm-Nd) for blueschist-facies rocks, chlorite schists and a basaltic dyke from the Kvarkush Anticline, Northern Urals. Sample sites: see Fig. 2.

The apparent ages given by the, admittedly imperfect, isochrons for the samples B99-80, -81 and -82, are controlled largely by the data from the white-mica phases. The Rb-Sr systematics of white mica is known to be set during crystallization, by subsequent crystal-plastic deformation, or by diffusional, thermally-driven resetting. Thermal resetting can be ruled out for the Kvarkush blueschists since blueschist-facies metamorphic temperatures are below the 'closure temperature' for Rb-Sr in white mica (which is >550 °C, cf. Villa 1998). Furthermore, different grain size fractions of white mica yield identical apparent ages. If the white-mica

ages were cooling ages, smaller grain-size fractions should yield younger apparent ages; this is not observed. The white mica, together with glaucophane, define a strong synmetamorphic foliation in the rocks. The Sr isotope results indicate that, within limits of error, the isotopic closure of all white mica was contemporaneous, irrespective of their chemical composition (see paragonite-rich v. phengitic white mica in sample B99-82). We therefore interpret the apparent ages of c. 536 Ma to be deformation ages, reflecting the latest increments of pervasive, blueschist-facies metamorphism and thereby the age of development of the

TIMANIAN BLUESCHISTS IN THE NORTHERN URALS

dominant foliation (S2). Sr-isotopic equilibria set during this 536 Ma metamorphic stage were only partly and slightly disturbed later. Any later, greenschist- to sub-greenschist-facies thermal and/or deformational overprints were weak, because otherwise the blue amphibole would have been heavily altered, or even completely transformed to other phases. At c. 536 Ma, pervasive metamorphic deformation came to an end while the rocks still were at blueschist-facies conditions, i.e. in the stability field of the blue amphibole. This time constrains a late stage of the high-pressure metamorphic evolution. The subduction process that led to accretion of the metasedimentary pile and to the blueschist-facies HP/LT metamorphism may have started earlier, most likely in the Late Neoproterozoic. There are no time constraints on incipient subduction in our study area. Further south, in the southwestern Urals, single grain U-Pb zircon ages, in combination with heavy mineral spectra of sandstones, point to onset of convergence related to the Timanian Orogeny at around 620-610 Ma (Willner et al 2001, 2003). However, our dataset does not provide any information on the pre-536 Ma dynamics of the Timanian Orogen. Post-Timanian processes The age of 526 + 36 Ma for sample H 152-2, within the limits of error similar to the age data for the blueschist-facies metapsammites, is slightly affected by the retrogression process, namely by the chloritization of biotite. White mica in this sample is intergrown with chlorite, and thus the white-mica concentrate probably included some chlorite. The respective chlorite separate plots slightly below the regression line defined by albite, apatite and the white-mica (+chlorite) concentrate (cf. Fig. 4). The presence of some chlorite therefore 'lowers' the apparent age of the white mica-concentrate. Thus the main structural fabric of the rock was established prior to retrogression, during ductile deformation slightly before 526 Ma. The two-point chlorite-epidote isochron of the chlorite schist sample 68-1 corresponds to an apparent age of 481.5 ± 6.4 Ma. This date is more difficult to interpret. It may represent a stage of greenschist-facies deformation, a fluid-induced greenschistfacies recrystallization, or some degree of open-system behaviour of the chlorite after the latest stages of Early Cambrian metamorphism, perhaps related to Ordovician rifting. Since a twopoint isochron gives no control on the degree of possible resetting, and since an isochemical overprint would lead to apparent rejuvenation, this date of 481.5 ± 6.4 Ma is interpreted as a maximum age for a possible post-Cambrian overprint. It is therefore important for the interpretation of the post-Timanian tectonic history; chlorite schists are easy to deform, and the Rb-Sr isotopic system of chlorite will readily react on deformation. The clearly pre-Uralian apparent age indicates that Uralian internal deformation of the Kvarkush complex is most probably of minor importance. Any post-481 Ma event was very weak and did not reach biotite stability field conditions. Significance of the Early Devonian Sm-Nd age for the basaltic dyke The Sm-Nd isochron age of 398 + 37 Ma for the largely unaltered, undeformed, discordant basaltic dyke from Skali (sample B99-48, Fig. 2, sample location 4) is interpreted as a magmatic crystallization age. This is because pure separates of unaltered, magmatic plagioclase and pyroxene were used, and because the whole-rock data plots on the mineral isochron, indicating Ndisotopic equilibrium among the main Nd-bearing phases of the rock at the age of intrusion. The dyke thus belongs to the postOrdovician group of dykes, as defined by Esipov (1960). It cuts

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the S3 foliation in the host quartz-mica schists, which therefore must be older than 398 ± 37 Ma. This minimum age for deformation of the country rocks again confirms that the Uralian (Late Palaeozoic) tectonic reworking of the basement in the Kvarkush Anticline is of minor importance, and that tectonic and metamorphic fabrics of the Timanian Orogeny are largely preserved. Tectonic setting of the blueschist-facies metamorphism The complex structural configuration of the KvarkushKamennaya Gora Anticlinorium (cf. Knapp et al. 1998; Friberg et al. 2002) is the result of at least two orogenic events, namely the Timanian and Uralian Orogenies. The Uralian structural reorganization complicates detailed reassessment of the structural setting in which the blueschists were formed during the Timanian Orogeny. However, the geological and isotopic data provide some constraints for the Timanian evolution. We suggest that the very high initial 87Sr/86Sr isotopic compositions found for the blueschist samples, together with their quartzitic composition indicate a continental affinity for these metasediments. Most probably they represent material derived from the East European Craton. There are no known occurrences of volcanic-arc-related rocks west of the Kvarkush Anticline. Therefore, an eastward-directed subduction of the East European passive margin (present day coordinates) during the Timanian Orogeny, involving compressional thickening of the passive margin sediments, is the most likely scenario for the formation of the Kvarkush blueschists. The Kvarkush HP/LT metasediments, together with the metamorphosed sediments further west, form parts of a Timanian orogenic wedge. Seismic data support this view, where east-dipping reflections at mid-crustal depths underneath and east of the Kvarkush Anticlinorium have been interpreted as a major pre-Uralian, probably Late Proterozoic, thrust (Friberg et al. 2002; Knapp et al. 1998). With respect to the Uralian Orogen, the Timanian age Kvarkush HP/LT metamorphic rocks are located on the western flank of the Urals, in the footwall of the Main Uralian Fault (MUF). However, not all the occurrences of HP/LT rocks in this tectonic position are of Timanian age. Among these occurrences is the famous Maksyutov HP metamorphic complex (Southern Urals, Fig. 1), for which the age of eclogitization was dated by multiple Rb-Sr mineral isochrons at 375 + 2 Ma, i.e. in Devonian time (Glodny et al. 2002). Almost simultaneously an accretionary wedge was formed in the area west (modern coordinates) of the Maksyutov complex (e.g. Brown & Spadea 1999; Giese 2000). The Devonian orogenic evolution involved eastward-directed subduction of the leading edge of the East European continent (Hetzel 1999; Brown et al. 2000), as similarly inferred for the Timanian evolution (see above). Apparently, along the western flank of the Urals, two strikingly similar orogenic evolutions are spatially related and partly superimposed, one of Vendian to earliest Cambrian age and the other Devonian. Regional implications The latest Neoproterozoic to Early Cambrian age for the blueschist-facies metamorphism has implications for the interpretation of the regional geology of the Kvarkush area. As outlined above, very poor exposure led to debate about the regional distribution and time of formation of the blueschist-facies rocks. Blueschists are found in pre-Ordovician rocks in the northern hinge of the Kvarkush Anticline (Starkov 1963; Alekseev 1996), but also in other localities further to the south (Alekseev 1996) and SW (Rusin 1996, and observations by one of us (MB)). This regional distribution suggests the presence of a major Timanian-age accretionary complex with a regional blueschist-facies metamorphic

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overprint. The isotopic age of the blueschists presented here is at variance with previously presented ideas of narrow post-Ordovician blueschist-facies zones cross-cutting the regional metamorphic pattern (Rusin 1996).

and references therein). Late- to post-tectonic granites, apparently of calc-alkaline affinity (Pease et al 2004) intrude the basement beneath the Pechora Basin at 550-560 Ma (Gee et al 2000).

Kvarkush blueschist-facies metamorphism in the framework of Timanian Orogeny

Conclusions

Within the framework of the entire Timanian orogenic belt, along the northeastern and eastern margins of Baltica, the occurrence of compression-related, high-pressure metamorphic rocks of latest Neoproterozoic to Early Cambrian age is not unique. The new Rb-Sr mineral ages reported here, together with previous K-Ar ages, indicate that the Timanian metamorphic imprint is dominant over the entire Kvarkush area, although metamorphism in places may only have reached greenschist-facies conditions in the present level of exposure. Another example of subductionrelated high-pressure rocks most probably of Timanian age are the eclogites of the Ufaley Complex (Fig. 1), Middle Urals (Kazak 1970; Keilman 1974; Belkovski 1989). Similar to the HP/LT rocks of the Kvarkush area, these eclogites are structurally located in the footwall of the Main Uralian Fault, indicating affinity of the complex to Baltica. In the Ufaley Complex, HP/LT rocks are only locally preserved. This may be due to intense Uralian tectono-metamorphic overprint, as described by Echtler et al. (1997). Although there is no reliable isotopic age determination for this eclogite-facies metamorphism, Puchkov (pers. com. 2001) has provided structural arguments for inferring a Timanian age. Glaucophane-bearing, possibly subduction-related, highpressure rocks in the Timanides have also been reported from the Kanin Peninsula (Kazak et al. 1989) and detrital glaucophane has been identified in Devonian limestones from the Tsil'ma area, Timan Range (Kochetkov, pers. com. 2001). From the structural position of these occurrences within the framework of the Timanides (Fig. 1), an age comparable to the age of blueschist-facies metamorphism in Kvarkush is inferred. Apart from our data for Timanian metamorphic rocks with clear HP/LT signature, there are many other age data for Timanian orogenic, metamorphic and magmatic processes in a similar time frame. Abundant 40Ar/39Ar white-mica ages roughly around 550 Ma from metasediments of the Bashkirian Anticlinorium, southwestern Urals (Fig. 1) and the style and timing of the Late Neoproterozoic Orogeny here show striking similarities with the Timan Range (Glasmacher et al 1999, 2001). From the Beloretsk metamorphic complex, in the eastern limb of the Bashkirian Anticlinorium, amphibolite facies rocks with an age of c. 550 Ma are clearly pre-Ordovician (Matte et al 1993). These are associated with eclogites, the precise age of which remains unknown (Glasmacher et al 2001). The style of deformation of the Precambrian rocks in the Bashkirian Anticlinorium as reported by Glasmacher et al (1999) closely resembles what has been described from similar rocks of the Middle Urals (Beckholmen & Juhlin 1997; Beckholmen 1998) and in the KvarkushKamennaya Gora Anticlinorium of the Northern Urals (Beckholmen 2000). Studies of heavy mineral spectra and determinations of U-Pb ages of single zircons from sandstones of the southwestern Urals revealed a conversion of the eastern margin of the Baltica protocontinent to an active margin at about 620-610 Ma (Willner et al 2001; Willner et al 2003) and magmatic activity at c. 580 Ma (Willner et al 2003). Evidence for compressional deformation along the Timanide belt is known from the Varanger, Srednij and Rybachij Peninsulas and Kildin Island (see Fig. 1) with illite ages dating diagenesis at 635 Ma (Dallmeyer & Reuter 1989) and 610-620 Ma (Gorokhov et al 1995). The main deformational event occurred around 550-570 Ma (Roberts 1995, 1996; Olovyanishnikov et al 2000

The waning stage of blueschist-facies deformation in the Kvarkush Anticline is dated by Rb-Sr mineral isochrons at c. 536 Ma, i.e. earliest Cambrian time. The lithological characteristics and Sr-isotopic signatures of the Kvarkush metasediments indicate affinity of these rocks to the pre-Neoproterozoic East European continent. The most likely scenario for formation of the blueschists is HP/LT metamorphism in an orogenic wedge in a Timanian age, east-dipping subduction system. The Sm-Nd mineral age for a discordant mafic dyke (398 ± 37 Ma) confirms the pre-Uralian timing of blueschist-facies metamorphism. Since the age of 536 Ma relates to a late stage of the high-pressure metamorphic evolution, the subduction process which led to accretion of the Kvarkush metasedimentary pile probably started earlier, in the Vendian. A review of the literature on HP/LT metamorphic rocks in the Urals suggests that relics of Timanian age HP rocks occur as a dismembered belt along the western flank of the Southern, Middle and Northern Urals. In the footwall of the Main Uralian Fault, two very similar cycles of subduction and formation of accretionary wedges are spatially related or even superimposed, firstly in the Vendian-Early Cambrian (Timanian) and subsequently in the Devonian-Carboniferous (Uralian). This repetition of very similar orogenic evolutions along the eastern margin of the East European continent complicates the interpretation of structural geology in the region. Although some of the Timanian structures are obscured by Uralian processes, the orogen as a whole is unambiguously comparable with modern orogens, involving subduction, tectonic wedging and HP/LT metamorphism. However, for a detailed reconstruction of the Timanian orogenic processes, far more geochronological, structural, geophysical and petrological data are needed. The fieldwork in the Kvarkush Anticline was organized by the Department of Geology and Geochemistry of the Russian Academy of Sciences, Ekaterinburg; the efforts of A. Rusin and his team are warmly acknowledged. The work in Uppsala has been funded via a grant to D. G. Gee by the Swedish Research Council, the Royal Society (KVA) and Uppsala University. J. G. wishes to acknowledge the support provided by the GeoForschungsZentrum Potsdam for completion of this work. We thank J. Herwig, M. Langanke, V. Kuntz and M. Dziggel for help with figure design and sample processing. S. Sindern and M. Tichomirova provided careful and constructive reviews. V. Pease is thanked for editorial handling. This is a contribution to EUROPROBE's Uralides Project.

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Northern Urals.] Ezhegodnik-1990, Institut geologii i geokhimii, Sverdlovsk UrO AN SSSR, 64-66. RUSIN, A. L, NIKIFOROV, O. V. & YAKOVLEVA, O. M. 1989. Problemy metamorfizma zony sochlenenia Urala i Timana [Metamorphic problems in the junction between the Urals and the Timans.] Ezhegodnik-1988, Institut geologii i geokhimii, UNC AN SSSR, Sverdlovsk, 56-58. ScARROW, J. H., PEASE V., FLEUTELOT, C. & DUSHIN, V. 2001. The late Neoproterozoic Enganepe Ophiolite, Polar Urals, Russia: An extension of the Cadomian arc? Precambrian Research, 110, 255-275. SHATSKIJ, N. S. 1963. Rifejskaya era i Baikalskaya skladchatost' [Riphean era and the Baikalian folding.] In: Izbrannye trudy, 1, AN SSSR, 600-619. STARKOV, N. P. 1957. Khlorit iz kvartsevykh zhil plato Kvarkush na Urale [Chlorite from quartz veins of the Kvarkush Plateau in the Urals.] AN SSSR, Zapiski vsesoyuznogo mineralogicheskogo obshchestva, 2 ser, 86, 4, 505-508. STARKOV, N. P. 1963. Kvoprosu o metamorfizme drevnikh svit zapadnogo sklona Severnogo Urala. [To the question on the metamorphism of the ancient suites of the western slope of the Northern Urals.] Trudy pervogo Ural'skogo petrograficheskogo soveshchania. UFAN SSSR, Sverdlovsk, 111, 223-233. STARKOV, N. P. 1969. Magmatism, petrologicheskie i geokhimicheskie osobennosti porod bazaVtoidnoj formatsii zapadnogo sklona Severnogo Urala [Magmatism, petrologic and geochemical characteristics of the rocks of the basaltoid formations of the western flank of the Northern Urals.] Trudy vtorogo Ural'skogo petrograficheskogo soveshchania. UFAN SSSR, Sverdlovsk. STEIGER, R. H. & JAEGER, E. 1977. Subcommission on geochronology; convention on the use of decay constants in geo- and cosrnochronology. Earth and Planetary Science Letters, 36, 359-362. VILLA, I. M. 1998. Isotopic closure. Terra Nova, 10, 42-47. VOROSHCHUK, D. V., BECKHOLMEN, M., RUSIN, A. I. & GEE, D. G. 1999. Mineral chemistry of Glaucophane-bearing Zones in the Kvarkush Anticline, Northern Urals. Abstract for EUROPROBE Uralides workshop, Miinchen. WILLNER, A. P., ERMOLAEVA, T., STROINK, L., GLASMACHER, U. A., GIESE, U., PUCHKOV, V. N., KOZLOV, V. L & WALTER, R. 2001. Contrasting provenance signals in Riphean and Vendian sandstones in the SW Urals (Russia): constraints for a change from passive to active continental margin conditions in the Neoproterozoic. Precambrian Research, 110, 215-239. WILLNER, A. P., SINDERN, S., METZGER, R., ERMOLAEVA, T., KRAMM, U., PUCHKOV, V. & KRONZ, A. 2003. Typology and single grain U/Pb ages of detrital zircons from Proterozoic sandstones in the SW Urals (Russia): early time marks at the eastern margin of Baltica. Precambrian Research, 124, 1-20. ZONENSHAIN,

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PECHERSKY, D. M., KHAIN, V. V. & MATVEENKOV, V. V. 1984. Plate tectonic model of the South Urals development. Tectonophysics, 109(1-2), 95-135. ZONENSHAIN, L. P., KUZMIN, M. I. & NATAPOV, L. M. 1990. Uralian Foldbelt. In: Geology of the USSR: a plate-tectonic synthesis. Geodynamics series, Vol. 21, American Geophysical Union, Washington, D. C., 27-54.

The Precambrian rocks of Novaya Zemlya EUGENE A. KORAGO, GALINA N. KOVALEVA, BORIS G. LOPATIN* & VLADIMIR V. ORGO VNIIOkeangeologia, 1, Angliysky pr.,190121, St Petersburg, Russia (""e-mail: [email protected])

Abstract: Neoproterozoic and perhaps older metamorphic and igneous rocks in the Novaya Zemlya Archipelago are exposed in four main localities. On the South Island of Novaya Zemlya (the Southern Domain) largely metasedimentary, low greenschist facies, pelitic and flyschoidal formations, dated by acritarchs as late Riphean-Vendian occur; they are intruded by metagabbro-dolerites and lamprophyres. These rocks are overlain by Ordovician strata with sharp, angular unconformity and are therefore considered to be products of Timanian (Baikalian) tectonogenesis. On the southern part of the North Island (the Central Domain), metamorphic rocks of high greenschist to amphibolite facies are preserved which are partly migmatized and comprise metasedimentary and igneous rocks; their contact with Palaeozoic strata here is tectonic. Isotopic age determinations on metasediments and igneous rocks vary from 1550 to 598 Ma and are not yet conclusive. These relatively highly metamorphosed formations may be older than Neoproterozoic. A Neoproterozoic (late Riphean) age was determined for the Mitushev Kamen'Granite Massif, which is located within the Main Novozemel'sky thrust zone. On the northern part of North Island, (the Northern Domain) Neoproterozoic strata (mainly turbidites) of low greenschist facies pass conformably up into Cambrian and younger strata. A major NW-trending fault (the Baidaratsky fault zone) is believed to separate the Southern Domain from Central and Northern Domains. It is concluded that the southern area is a peripheral part of the Neoproterozoic Timanian (Baikalian) fold belt. Further to the NE across the Baidaratsky fault zone, lies a broad area of possibly older basement, including the Central and Northern Domains of Novaya Zemlya and possibly the Franz Josef Land Archipelago.

The Pai-Khoi-Novaya Zemlya fold belt is the youngest (early Cimmerian) segment of the Uralides. In plan, it has an arcuate form, convex to the west, and comprises several synclinoria and anticlinoria, changing their trend from WNW in the south to NE in the north. This fold belt separates the Barents and Pechora basins in the west from the West Siberian-South Kara Basin in the east. Along with the Polar Urals structures and the Byrranga (Taimyr) structures, the Pai-Khoi-Novaya Zemlya fold belt composes a gigantic sigmoidal fold belt bordering the North KaraBarents and South Kara (West Siberian) basins of younger sedimentation. Neoproterozoic outcrops in the Novaya Zemlya Archipelago are sparse and small, but their significance for understanding the structure and evolution of the Pai-Khoi-Novaya Zemlya fold belt is critical. Neoproterozoic rocks are known in the extreme south and north of the archipelago, and also in the central part within the Main Novozemel'sky thrust zone (Fig. 1). The Precambrian rocks differ in age, composition and structure in the different areas. More importantly, in relation to the younger overlying strata, the archipelago is considered to be an assemblage of three tectonic domains: (1) a Southern Domain with Neoproterozoic basement uncomformably overlain by Ordovician cover; (2) a Central Domain (separated from the Southern by the Baidaratsky fault zone) with probably older basement, inferred to be unconformably overlain by Cambrian rocks; and (3) a Northern Domain, separated from the Central by the Sporonavoloksky fault zone, with unknown basement (probably of same age as in the Central Domain) and with a continuous sedimentary sequence from the Neoproterozoic into the Early Palaeozoic. The Baidaratsky fault zone is expressed by a series of strike-slip faults, which can be seen on the seismic records in the Barents Sea (Lopatin et al. 2001). The main movement on this fault zone is thought to be Timanian (Baikalian) in age. The Southern Domain Metasedimentary formations and associated intrusions of Neoproterozoic age are exposed in the core of Yuzhnonovozemel'sky (Southern Novaya Zemlya) Anticlinorium on the Piritovy and Rusanova peninsulas (Fig. 2) and neighbouring islands (Korago & Chukhonin 1988; Kovaleva et al. 1984). The metasediments include two main units, a lower one of Late Riphean and an upper one of Vendian age. Both are unconformably overlain by

Ordovician formations. The Riphean unit, in its lower part (Loginovskaya Series), is composed of polymictic sandstones with layers of siltstones, slates and carbonate concretions. Its upper part (Piritovskaya Series) comprises black pyritized chlorite-sericite phyllites with subordinate thin layers of siltstones and rare sandstones. The thickness of these Riphean strata are about 1 km and the base is not seen. Acritarchs: Synsphaeridium sorediforme Tim., Protosphaeridium nervatum Tim., P. densum Tim., P. duricorum Andr., Strictosphaeridium sinapticuliferum Tim., S. implexum Tim. have been reported in these sediments. Together with filaments of the algae Trachythorichoides ovalis Herm. and other forms (Kovaleva et al. 1984), they demonstrate a Neoproterozoic (late Riphean) age. In the lower part (Krivenerskaya Series) of the overlying Vendian succession polymictic volcaniclastic sandstones are interbedded with siltstones and phyllites; some gravelites and conglomerates are also present. Its upper part (Reinekskaya Series) is composed of dark-grey and greenish banded sericite-chlorite slates and phyllites with beds of metasiltstones and metasandstones. Secondary carbonates and gradational bedding are usual. The thickness of this unit exceeds 1 km; its age is inferred to be Vendian, based on the presence of acritarchs and algae: Origmatosphaeridium rubiginosum Andr., Leiosphaeridia gigantea Schep., Litosphaeridia minor Schep., and other forms (Kovaleva et al. 1984). The Precambrian rocks are deformed by isoclinal folds of different orders with thickening in the fold hinges and other evidence of material movement. The folds are broken by numerous faults and reworked by a younger NW-oriented deformation. Ordovician siliciclastic and carbonate sediments overlie the Neoproterozoic rocks with sharp, angular disconformity (Fig. 3). Cambrian strata are apparently absent within the Yuzhnonovozemel'sky Anticlinorium. The Neoproterozoic rocks were metamorphosed under low greenschist facies conditions. Judging by their initial composition (greywacke) and sedimentary structures (grading, etc), they are predominantly turbidites, which are also characteristic of the sedimentary complexes of Kanin, Timan, Pai-Khoi and Severnaya Zemlya of apparently similar age. Detrital minerals (garnet, zircon, titanite, rutile, pyroxene, hornblende, spinel, epidote) and clasts in the sandstone indicate that the source area was heterogeneous and probably included volcanic arcs, ophiolite belts and projections of older crystalline basement (Kovaleva etal 1984).

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 135-143. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Simplified geological map of the Novaya Zemlya Archipelago and adjacent offshore areas (from unpublished Geological Map of the Novaya Zemlya Archipelago at 1:500 000 scale).

THE PRECAMBRIAN ROCKS OF NOVAYA ZEMLYA

137

Fig. 2. Geological map of the core of the Yuhznonovozemelsky anticlinorium. (From Kovaleva et al. 1984; see Fig. 1 for location.)

Within the Neoproterozoic successions, in the area of the Rusanov and Piritovy peninsulas and Olenii island, are small intrusions grouped compositionally into two associations: the Rusanovsky gabbro-dolerites and the Rahkovsky lamprophyres

(Table 1). The Rusanovsky Suite includes sills, laccoliths and dykes. The sills (up to 30 m thick and 6 km long) are differentiated: their inner parts are composed of medium- to coarsegrained dolerites and gabbro-dolerites, their upper parts are

138

E. A. KORAGO CTAL.

to the original magma and that little subsequent fractional crystallization has occurred. These intrusions have undergone considerable metamorphic alteration including albitization of plagioclase and amphibolization of pyroxene; secondary chlorite and epidote are present and carbonate and leucoxene are widely distributed. Hydrothermal processes have generated quartz, quartz carbonate and hematite veins (or lenses) up to 0.5 m in width and 20 m in length. In the host metasediments, contact alteration is expressed as thin zones of sericite-chlorite-quartz hornfels. Both the metasediments and the meta-igneous rocks were influenced by the late Neoproterozoic (?Cambrian) deformation. Lamprophyre (and possibly wehrlite) dykes of the Rahkovsky Suite cut the Neoproterozoic succession and Rusanovskaya gabbro-dolerites in the area of the Piritovy Peninsula. The thickness of the lamprophyres is l-3m, and their length 3-4 km. Amphibole and olivine monchiquites predominate. Their chemical composition varies, and is characterized by high contents of TiO2 (3-4%) and P2O5 (0.8-1.6%) due to considerable amounts of ilmenite, kaersutite and apatite. High alkali and MgO contents with low SiO2 are also typical. The rocks have undergone intensive alteration, as evidenced by the carbonitization and chloritization of the both groundmass and the phenocrysts of biotite and amphibole (Korago et al 1982). The geochemical affinities of the Rahkovsky and Rusanovsky suites are similar, which may indicate a co-genetic relationship. On AFM and Na2O-K2O-CaO diagrams their fields are very close. Both suites have a SI index of 40. Nevertheless, the Rahkovsky Suite has a higher SiO2/K2O + Na2O ratio and higher potassium, titanium and phosphorus contents. The latter, along with the mineral composition of lamprophyres, indicates that the initial magma was rich in volatile components. The lamprophyres also contain large (up to 5 cm) crystals of ore minerals (solid solutions of ilmenite and hematite) which might have been generated at mantle depths (Korago et al. 1982). Thus, it seems that the lamprophyres may have originated from a deeper magmatic source. The Rusanovsky and Rahkovsky igneous suites, intruding the metamorphic rocks of late Riphean-Vendian age, are deformed and metamorphosed together with them. They are completely absent in nearby Ordovician rocks. K-Ar whole rock and amphibole analyses from both suites yield ages in the interval 405 ± 28 to 445 ± 22 Ma (Korago et al 1992). This probably represents the time of a superimposed metamorphic event,

Fig. 3. Contact between Neoproterozoic and Ordovician strata on the Pyritovy Peninsula (Southern Novaya Zemlya). (From Kovaleva et al 1984.)

fine-grained quartz-bearing dolerites and gabbro-dolerites and their lower parts are olivine-bearing gabbros. The dykes are l-3m thick and reach up to 5km in length. Small bodies are mainly composed of fine-grained dolerites and porphyritic basalts. Sometimes their margins are chilled. Dolerites of the Rusanovsky Suite generally have small variations of SiO2 (46-50%), Ti02 (1.70-3.15%) and Na2O + K2O (2.34-3.0%). On discrimination diagrams, they group within the ocean island basalt and abyssal ocean basalt fields (Pearce 1968). The Rusanovsky dolerites have an average solidification index (SI) of 40 (Kuno 1964), suggesting that they are close in composition Table 1. Chemical composition of intrusive rocks of Novaya Zemlya Yakorninsky Suite

Oxides

1 Si02 TiO2 A12O3 Fe203

FeO MnO MgO CaO

Na20

K20 P205 Ignition loss

40.99 0.02 1.58 4.03 4.75 0.15 37.10

— 0.07 0.03 0.03 11.33

2

3

50.04 1.48 14.70 3.03

8.0 0.19 6.57 8.43 2.31 1.63 0.27 3.38

47.08 0.70 19.48 2.74 6.39 0.15 6.77 9.15 2.03 2.19 0.04 3.50

Sulmenevsky Suite

4 48.99 0.57 12.34 1.78 6.88 0.18 12.41 11.44 1.06 1.38 0.08 2.95

5

6

78.06 0.03 12.47 0.24 0.50 0.01 0.62 1.68 5.49 0.63

tr 0.70

74.06 0.03 15.02 0.38 0.47 0.03 0.17 0.42 3.80 4.82 0.09 0.87

Rusanovsky Suite

7 70.82 0.26 15.01 0.92 1.69 0.07 1.35 2.59 3.00 2.67 0.11 1.70

9

8 49.75 3.15 11.53 1.96 9.72 0.19 9.91

7.0 2.16 0.18 0.22 4.34

46.04 1.07 15.13 6.60 5.43 0.20 9.92 8.52 2.52 0.47 0.13 4.26

Rahkovsky Suite

10

11

34.88 4.17 9.52 6.91 8.62 0.35 11.08 14.04 1.52 0.68 1.55 6.47

42.20 3.24 10.45 1.80 9.48 0.18 10.82 8.69

1.0 0.56 0.83 10.70

Mitushevsky Suite

12 77.10 0.10 11.70 0.84 0.71 0.20 0.15 0.50 3.66 4.84

—

13 74.36 0.23 12.47 0.78 1.36 0.03 0.38 0.82 3.73 4.96

— 0.42

0.88

14

15

16

75.29 0.20 13.13 1.37 0.22 0.02 0.45 1.40 6.09 0.40

66.67 0.36 15.21 1.87 1.85 0.06 1.57 2.98 3.94 3.75

65.08 0.46 14.84 2.56 1.70 0.09 2.48 2.45 2.57 4.92

—

—

—

0.96

1.57

2.63

Sum 100.08 100.03 100.22 100.06 100.43 100.16 100.19 100.11 100.29 99.79 99.95 100.22 100.00 99.53 99.83 99.78 Number of analyses 1 7 2 7 1 1 1 3 723 612131 The samples are from O. Baklund, N. N. Mutafi, T. N. Timofeeva collections, analysed by G. S. Danilova, L. G., Syrnikova, L. G. Finshina (VNIIOkeangeologia). 1, serpentinite (Mitushihka bay); 2, contaminated diabase (Sulmenev Bay); 3, metadiabase (Sulmenev Bay); 4, amphibolites and contaminated amphibolite (the same); 5, pegmatitic plagiogranite (the same); 6-7, muscovite granite (the same); 8, diabase with high titanium content (Rusanov Peninsula); 9, diabase (Rusanov Peninsula); 10, amphibole monchikite (Pyritivy Peninsula); 11, olivine monchikite (the same); 12, pegmatitic granite (Mitushihka Bay); 13, leucocratic and muscovite granite; 14, albitized granite (Litke Mt); 15, plagiogranite-granodiorite (the same); 16, granosyenite (the same).

THE PRECAMBRIAN ROCKS OF NOVAYA ZEMLYA

because the geological evidence favours a pre-Ordovician age for these rocks.

The Central Domain The basement of the Central Domain crops out in the southern part of North Island (Fig. 1) within the Main Novozemel'sky thrustfault zone. These Precambrian rocks occur in two main areas in the vicinity of Sulmenev Bay and Mitushev Bay. On the northern coast of Sulmenev Bay (Fig. 4), the rocks comprise a small (about 5 km2) tectonic body bounded by faults within middle Palaeozoic sequences. Two metamorphic units, together composing the Sulmenevskaya Series, can be distinguished: a lower unit of marbles, mica schists and quartzites and an overlying unit of quartz-plagioclase-amphibole-biotite schists, biotite plagiogneisses and amphibolites, with rare layers of marbles and mica quartzites, all intruded by numerous granitic veins. The exposed thickness of the lower marble sequence is about 450 m and of the upper schist is about 700 m. The metamorphic grade of the Sulmenevskaya Series is epidote-amphibolite facies, based on the usual mineral paragenesis: quartz-plagioclase (albite-oligoclase)-biotite-muscovite hornblende. Locally, there is intensive diaphthoresis, expressed by actinolitization of hornblende, sericitization of plagioclase and chloritization of biotite. Veins of aplite, pegmatite and muscovite granite intrude the metasediments. They vary in thickness up to 4 m, and strike for distances up to 30 m. The dominant composition of the granite veins is quartz (25-35%), plagioclase (35-40%), microclineperthite (5-25%), and muscovite (2-10%). Accessory minerals include tourmaline, zircon, apatite, tantalo-niobate, cassiterite and scheelite. Intruded into the Sulmenevskaya Series are small metabasic bodies, exposed on the northern coasts of Sulmeneva and Mitushikha bays. These intrusions are named the Yakorninsky Suite. They cross-cut the metasediments in several locations. One of the largest bodies, with tabular form, comprises metagabbro, metadolerite and amphibolite. The rocks are banded, mottled and injected by the granite veins, similar to those penetrating the Sulmenevskaya Series. Chemical analyses of the Yakorninsky Suite are grouped by petrographic affinities (Table 1). Considerable compositional variation can be explained partly by hybridization effects resulting from granitic intrusions. The average composition of the veined granites (Sulmenevsky Suite) intruding the Sulmenevskaya Series corresponds to plagiogranites-alaskites (Table 1, analyses 5-7), varying from oversaturated to saturated by SiO2. A12O3 content is high, due to the presence of muscovite, total alkalis vary from 6.5 up to 8%, and the amounts of K2O and Na2O are similar. Ages of 1550 ± 80 and 1490 ± 100 Ma were obtained from the metamorphic rocks of the Sulmenevskaya Series, using the Pb/Pb evaporation multigrain method on zircons (Korago & Chukhonin 1988); ages of c. 1300 Ma were obtained on the Sulmenevskaya granites. Recently, new data (Kaplan et al 2001), using the U-Pb method on single zircons from the Sulmenev veined granites, yielded 598 ± 26, 618 ± 18 and 604 ± 9 Ma ages. The Pb/Pb data are thought to represent the age of the source rocks of the Sulmenevskaya Series, and the U-Pb show the age of the granite intrusion. Further to the south on North Island, also in the Main Novozemel'sky fault zone is another igneous intrusion—the Mitushev Kamen' Granite Massif (Fig. 5). There are four separate exposures with a general northwestern extension of over 40 km and a width of 3-8 km. The main intrusive body is 50 km2 and has an oval shape in plan view. It is composed predominantly of coarse-grained alaskite granite with cataclastic and blastohypidiomorphic textures, large deformed grains of quartz,

Fig. 4. Geological map of the Sulmenev Bay area. (From Korago 1984; see Fig. 1 for location.)

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E. A. KORAGO £7 AL.

Fig. 5. Geological Map of the Neoproterozoic Mitushikha Kamen' granitoid Massif. (From Korago 1984; see Fig. 1 for location.)

microcline-perthite and albitic plagioclase are cemented by a finely fragmented matrix of quartz, feldspar and mica. The preliminary composition of the alaskite is quartz (20-35%), microclme-perthite (15-60%), plagioclase (5-15%) and chloritized biotite and muscovite (3-5%). Accessory minerals include apatite, zircon, titanite, orthite and fluorite. The marginal parts of the body are composed of foliated and more melanocratic biotite and biotite-amphibole plagiogranites. This intrusion usually has tectonic contacts with Upper Silurian and Lower Devonian strata, but in an exposure near the Zapasova Mountain (Fig. 6), it is overlain by coarse-grained Upper Silurian rocks, containing granite pebbles in the basal strata (Korago 1984). One of the bodies on the coast of Serebryanka Bay is made up of similar porphyroblastic amphibole-biotite

Fig. 6. Contact between the Neoproterozoic Mutushev granitoides and unconformably overlying Silurian sedimentary strata on Mt Zapasova. (From Korago 1984.)

THE PRECAMBRIAN ROCKS OF NOVAYA ZEMLYA

plagiogranites-granodiorites (Figs 5 and 6). They consist of quartz (20-30%), sericitized oligoclase (40-60%), microcrineperthite (15%), biotite (5-20%) and amphibole (0-15%). Isotope age studies by the Pb evaporation method on zircon populations yielded ages of 680 ± 50 and 730 ± 50 Ma, corresponding to the late Riphean (Korago et al 1984, 1993). These data are close to ages obtained by Kaplan et al. (2001), using the U-Pb single grain method: 609 ± 4, 587 + 7 and 717 + 4 Ma. The Northern Domain In the Northern Domain, located in the northern part of North Island, Neoproterozoic strata pass conformably up into the Cambrian, Ordovician and Silurian, without any unconformities or gaps in sedimentation. In the lowermost parts of the sequence (base not seen), the Makovskaya Series is exposed along both shores of the Maka Bay in the core of an anticline within the hinge of the Severozemel'sky Anticlinorium (Fig. 7). This unit

141

is composed of metasandstones, metasiltstones and phyllites, forming turbiditic cycles 0.5-15 m thick. Metasandstones (sometimes including conglomerates) give way to fine-grained sediments. In the upper part of the group, the microfossils Bavinella faveolata Schep. of Neoproterozoic (Late Riphean) age are present. The group is more than 1300 m thick. The Makovskaya Series is conformably overlain by the Lomonosovskaya Series, reported from the Shokal'sky Glacier to the Inostrantseva Bay. The Lomonosovskaya Series is composed of dark-grey, highly pyritized phyllites with beds of quartz-feldspar metasandstones and metasiltstones. The strata are cyclic and have a total thickness of 1700 to 2430 m. Cambrian trilobites are present in the upper part of the sequence. In its middle part, the Late Riphean microfossils Bavlinella faveolata Schep., Symplassosphaeridium sp., Leiosphaeridia sp., Pterosphermopsimorpha sp. and Trachystrichosphaera sp. occur. The geological position and fauna indicate a Late Riphean to Early Cambrian age for the Lomonosovskaya Group (Korago et al 1984).

Fig. 7. Geological map of the Maka Bay-Inostrantsev Bay area of North Island. (From Korago et al 1984; see Fig. 1 for location.)

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Conclusions The Sulmenevskaya Series, exposed on North Island in the Central area within the Main Novozemel'sky fault zone, may be the oldest (Mezoproterozoic?) formation in the Novaya Zemlya Archipelago, but isotopic age determinations are not yet conclusive. This terrigenous-carbonate sequence suffered a relatively high grade of metamorphism, up to epidote-amphibolite to amphibolite-migmatite facies. Its final stage of consolidation was manifested by granitoid veins of Vendian age. The Sulmenevskaya Series has tectonic boundaries with Palaeozoic (Devonian) rocks. However, further south on North Island, c. 600 Ma granites of the Mitushev Kamen' Massif are overlain unconformably by Silurian strata. Cambrian strata are widespread in the Central area, in contrast to the Southern area, where they are absent. The Upper Riphean and Vendian sequences in the Southern Domain, exposed in the hinge of Yuzhnonovozemersky Anticlinorium, represent siliceous greywacke (turbidite) formations with gradational and rhythmic bedding, metamorphosed in low greenschist facies. Their age was determined by acritarchs. Together with the associated metagabbro-dolerite Rusanovsky Suite and lamprophyre Rakhovsky Suite, they are considered

to be products of Timanian (Baikalian) tectogenesis. The pre-Ordovician age of tectonogenesis is demonstrated by a major unconformity separating this basement from the overlying Ordovician cover and the absence of Cambrian strata. In the Northern domain, the Upper Riphean Makovskaya and the Vendian to Lower Cambrian Lomonosovskaya Series, exposed within the core of the Severozemel'sky Anticlinorium, are conformably overlain by Lower-Middle Cambrian strata. The Neoproterozoic and Lower Palaeozoic strata comprise a uniform unit of little metamorphosed sandstones, siltstones, and phyllites of predominantly quartzfeldspar composition and rhythmically-bedded turbidite. This implies continuous sedimentation through Neoproterozoic up into the Early Palaeozoic, probably in a continental slope-rise environment; important Neoproterozoic tectonic events are lacking. Basement is not exposed in the Northern Domain and may be similar to that in the Central Domain. The Northern Domain had an independent evolution up to the late Palaeozoic, when it joined the Novaya Zemlya fold system as a result of early Cimmerian orogenesis. Comparison of the Precambrian geology of Novaya Zemlya with that of other areas (Timan, Kanin, Svalbard and Pai-Khoi) around the Barents Sea, along with results of geophysical investigations of the Barents Sea area (Gramberg 1988), suggests the following Precambrian tectonic evolution. The pre-Riphean crystalline basement of the Kola Peninsula and the East European Craton gives way northwards to a wide heterogeneous belt (Fig. 8), bordered by fault zones and Mesoproterozoic massifs (similar to the Sulmenev Series or their source rocks), separated by Neoproterozoic mobile zones, all trending NW. These mobile zones are mainly composed of Neoproterozoic (Upper Riphean-Vendian) formations, folded and metamorphosed during Timanian (Baikalian) tectonogenesis. The Southern Domain of Novaya Zemlya is a peripheral part of this belt. Further to the NE, beyond the Neoproterozoic fold belt, across the Baidaratsky fault and its supposed northwesterly continuation (Lopatin et al. 2001), lies a broad area of older basement. It may include Nordaustlandet, possibly the Franz Josef Land Archipelago and the Central and Northern Domains of Novaya Zemlya. The juxtaposition of the Southern and the Central Domains along the Baidaratsky fault zone presumably took place in pre-Ordovician time during Timanian (Baikalian) orogenesis. The authors are grateful to the Polar Marine Geological Expedition (Lomonosov - St Petersburg) for recently acquired data on Novaya Zemlya, to David Gee (Uppsala) for discussions of the geology, and both David G. Gee and Olga K. Bogolepova (Uppsala), for generous assistance with linguistics in preparing the manuscript for publication.

References

Fig. 8. Precambrian basement of the western sector of the Russian Arctic.

GRAMBERG, I. S. (ed.) 1988. [Barents shelf plate]. VNIIOkeangeologia, Nedra, Leningrad, 1-262 [in Russian]. KAPLAN, A. A., COPELAND, P., BRO, E. G., KORAGO, E. A., PROSKURNIN, V. F., VINOGRADOV, V. A., VROLIJK, P. J. & WALKER, J. D. 2001. New radiometric ages of igneous and metamorphic rocks from the Russain Arctic. In: VNIIGRI/AAPG International Conference 'Exploration and production operations in difficult and sensitive areas', St Petersburg, 06-2. KORAGO, E. A. 1984. [Peculiarities of the structure of the Mitushev Kamen' granitoid massif (Novaya Zemlya)]. In: BONDAREV, V. I. (ed.) [Novaya Zemlya on the earliest stages of its geological evolution]. Sevmorgeo, Leningrad, 126-145 [in Russian]. KORAGO, E. A. & CHUKHONIN, A. P. 1984. [New data on the geological structure and the age of granitoids of the Mitushev Kamen' Massif (Novaya Zemlya)]. Doklady Academii Nauk USSR, 277, 445-448 [in Russian].

THE PRECAMBRIAN ROCKS OF NOVAYA ZEMLYA

KORAGO, E. A. & CHUKHONIN, A. P. 1988. [Granite formations of Novaya Zemlya]. Izvestiya Academii Nauk USSR, seriya geologicheskaya, 110, 28-36 [in Russian]. KORAGO, E. A., BYEVA, E. P., MAKSIMOVSKY, V. A. & TIMOFEEVA, T. N. 1982. [Lamprophyre Complex of the southern of Novaya Zemlya]. In: BONDAREV, V. I. (ed.) [Geology of Novaya Zemlya], 78-88 [in Russian]. KORAGO, E. A., ANDREEVA, I. A. & ERSHOV, Y. I. 1984. [Precambrian formations of the North island of Novaya Zemlya]. In: BONDAREV, V. I. (ed.) [Novaya Zemlya on the earliest stages of its geological evolution]. Sevmorgeo, Leningrad, 5-19 [in Russian]. KORAGO, E. A., KOVALEVA, G. N., IL'IN, V. F. & PAVLOV, L. G. 1992. [Tectonics and Metallogeny of the Early Kimmerides of Novaya Zemlya]. Leningrad, Nedra, 1-196 [in Russian].

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KORAGO, E. A., KOVALEVA, G. N., IL'IN, V. F. & PLATONOV, E. G. 1993. [The Pre-Cambrian of Novaya Zemlya]. Otechestvennaya geologia, 12, 36-48 [in Russian]. KOVALEVA, G. N., KORAGO, E. A. & SMIRNOVA, L. N. 1984. [Stratigraphy and tectonic position of the oldest formations of the South Island of Novaya Zemlya]. Bulleten Moskovskogo obschestva ispytatelej prirody, otdelenie geologicheskoe, 59, 80-88 [in Russian]. KUNO, H. 1964. [Series of igneous rocks]. In: Chemistry of the Earth Crust, 2, Moscow, 107-121 [in Russian]. LOPATIN, B. G., PAVLOV, L. G., ORGO, V. V. & SHKARUBO, S. I. 2001. Tectonic structures of Novaya Zemlya. Polarforschung, 69, 131-137. PEARCE, T. H. 1968. A contribution to the theory of variation diagrams. Contributions to Mineralogy and Petrology, 19, 142-157.

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Early Palaeozoic unconformity across the Timanides, NW Russia O. K. BOGOLEPOVA & D. G. GEE Department of Earth Sciences, Uppsala University, Villavdgen 16, SE-75236, Uppsala, Sweden (e-mail: [email protected], [email protected])

Abstract: The Lower Palaeozoic sequences, unconformably overlying the Timanides of Timan, Pechora, Pai-Khoi, Vaigach, Novaya Zemlya and the Polar, Northern, Middle and Southern Urals are described and interpreted with regard to their stratigraphy, sedimentation, structure and biogeography. These mainly shallow marine sedimentary successions, with associated igneous rocks (largely alkaline), are Ordovician in age, reaching back into the Late Cambrian in some areas, particularly in the east, along the front of the Urals. They were deposited during rifting of Baltica's northeastern margin and subsequent development of a passive continental margin. The underlying, mainly Neoproterozoic basement of turbidites in the west and calc-alkaline volcanites and ophiolites further east are briefly referred to, with particular emphasis on the age of the youngest rocks. Based on these data, a Timanian orogenic belt can be traced along the northeastern and eastern margin of the East European Craton. The timing of orogeny can be constrained to the Vendian, perhaps extending into the Early Cambrian.

A major unconformity transgresses the Timanides from the western front of the orogen in the Timan Range to the foreland fold-and-thrust belt of the Uralides (Fig. 1). It oversteps a wide range of pre-Palaeozoic rocks, from Vendian platform successions in the west to Neoproterozoic ophiolites in the far east. The character and age of the unconformity are key components for defining the regional tectonics of the eastern margin of the East European Craton. The unconformity can be followed from Novaya Zemlya and the Polar Urals southwards through the Late Palaeozoic fold-and-thrust belt nearly to the Aral Sea. The strata directly overlying the unconformity range in age from Late Cambrian to Devonian, and are generally older in the east than the west. Lower and Middle Cambrian units have been reported locally, but their occurrence is not well constrained. The lithologies and structures in the rock complexes underlying the unconformity are not a prime concern of this paper. They have been variously referred to as Timanides' (Shatsky 1935), 4 pre-Uralides' (Kheraskov 1948,1967), 'Baikalides' (Shatsky 1935, 1958), 'Ripheides' (Bekker 1976), and more recently 'Cadomian' (Puchkov 1979, 1997, 2000). We prefer a terminology that is unique to the eastern margin of the East European Craton and therefore use only the term Timanides'. Where lithological and palaeontological information is available concerning the Timanide units below the unconformity, these are summarized in order to provide an overview of the variability of this basement and the age and significance of the unconformity. However, the main objective of this paper is to provide a synthesis of data, not all of which is published, on the stratigraphy of the formations immediately overlying the unconformity. Some of the sources cited here provide palaeontological data that we have reported on face value, because we have been unable to locate the original descriptions of the fossils. Most authors accept the regional significance of Timanian Orogeny (Puchkov 2002; Roberts & Siedlecka 2002) recognizing the presence of a foreland molasse, continental margin deep-water siliciclastic successions, calc-alkaline igneous associations and, in the hinterland, ophiolites; however an alternative hypothesis is preferred by some authors (Ivanov 1977, 1981; Ivanov & Rusin 2000). They conclude that the Timanian deformation was not orogenic and that the Neoproterozoic lithologies can best be ascribed to an intracontinental rift basin, their deformation and metamorphism resulting from basin inversion. Descriptions of the unconformity The post-Timanian unconformity and overlying Palaeozoic successions are best known from the Timan Range and the foreland

fold-and-thrust belt of the Polar Urals. In the Timan Range, the sedimentary cover is dominated by Devonian sandstones, locally underlain by Silurian limestones. Older Palaeozoic strata are present further east. In the Polar Urals, sandstone-dominated successions of Late Cambrian and Early Ordovician age comprise a rift-related facies that is overlain by younger passive margin, drift-related Ordovician and Silurian shelf carbonates. Further east, these sandstones and carbonates give way to basinal shaledominated siliciclastic formations. This facies change has been documented (Voinovsky-Kriger 1967; Puchkov 1997) throughout the western front of the Uralide Orogen, defining a transition from the shallow continental shelf (the Eletskaya Facies) to deeper water continental slope environments (the Lemvinskaya Facies). The descriptions that follow begin with the Timan-PechoraPolar Urals transect and then proceed southwards through the Northern, Middle and Southern Urals. The unconformity is then followed northwards into the high Arctic of Pai-Khoi, Vaigach and Novaya Zemlya, where it appears to diminish in importance or die out in the far north. With regard to the stratigraphy below the unconformity, we refer to Neoproterozoic (locally Vendian) when it is demonstrable, and to Riphean where the successions may include Mesoproterozoic rocks. The boundary between the Neo- and Mesoproterozoic is taken at 1000 Ma (Remane et al. 1996). Timan Range and Kanin Peninsula Mid-Palaeozoic formations rest unconformably on folded and metamorphosed Neoproterozoic basement of the Timan Range, reaching from the front of the Northern Urals in the SE to the Kanin Peninsula in the NW. Devonian continental sandstones dominate the cover, often with basalts; locally, they are underlain by Silurian limestones. A major angular unconformity has been described from several localities along the Timan Range (location 1 on Fig. IB), where Neoproterozoic turbidites, metamorphosed in greenschist facies and intruded by a Vendian igneous suite, are overlain by Lower Silurian limestones and Devonian continental sandstones and basalts (Olovyanishnikov 1998). To the NW, on the Kanin Peninsula (Fig. 2), similar Silurian and Devonian rocks rest unconformably on metasedimentary units which occur in a major NW-trending anticline. The metasediments are composed mainly of turbidites that vary in metamorphic grade from low greenschist facies in upper structural levels to high amphibolite facies in the deepest exposed units (Lorenz et al. 2004). These metamorphic rocks are thrust southwestwards onto dolomites and limestones with stromatolites and oncolites (Ramsay 1911) of the Gnil'skaya

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 145-157. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. The main geographical elements (A) and generalized tectonic map of the Urals and Timan-Pechora Region (B), showing localities with the post-Timanian, Palaeozoic angular unconformity referred to in the text (based on Peive et al 1977).

Formation. The latter enable correlation with the Neoproterozoic (Upper Riphean) strata of the southern Urals (Olovyanishnikov 1998). To the SW of the Timanide Range, in the Arkhangelsk area of the eastern White Sea, Vendian successions have been described

by Grazhdankin (2004). He has demonstrated a progression through the Vendian from passive margin successions to those of a foreland basin, sourced from the NE. Towards the NW, the Timanian complex can be followed along the margin of the Barents Sea coast of the Kola Peninsula

TIMANIAN UNCONFORMITY, NORTHWESTERN RUSSIA

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Fig. 2. A geological map of the Kanin Peninsula and northern part of the Timan Range (based on Peiveetal. 1980).

via the Sredny-Rybachy Peninsula, to northernmost Norway, where it is overthrust by the nappes of the Scandinavian Caledonides (Roberts & Siedlecka 2002).

Pechora Basin Relationships between the sedimentary cover and underlying Timanian basement are known from deep drillholes through the Pechora Basin successions, summarized by Nikonov et al (2000). The metaturbidites of the Timan Range extend northeastwards beneath the Pechora Basin, and are intruded by c. 560 Ma granites (Gee et al. 1998). They are flanked to the NE by a greenschistfacies volcano-sedimentary association of calc-alkaline affinities (Belyakova & Stepanenko 1991). A lower Palaeozoic succession transgresses this basement, with Lower Ordovician quartzites usually separating an Ordovician-Silurian carbonate platform facies from the underlying basement. Locally, red beds separate the Ordovician basal siliciclastic formations from the volcano-sedimentary basement; these have been interpreted to be of Late Vendian to Cambrian age (Bogatsky et al. 1996). Similar lithologies have also been reported from east of the Timanides in the Izhma-Omra area (location 2 on Fig. IB), where polymict red sandstones and conglomerates, varying in thickness from 200 to 1000 m (Belyakov 1991) occur above the unconformity. This basal succession passes up into the (?) Middle Cambrian-Lower Ordovician Sed'iol'skaya Formation (100200 m), which comprises light quartz sandstones with interbeds of shales (Timonin 1998). The ages of the rocks above the unconformity are poorly constrained. At one locality in the Pechora Basin (the Malaya Pera-11 drill-hole, location 3 on Fig. IB) shales unconformably overlying granites at 3169m depth, have yielded the acritarchs Celtiberium dedalinum, Cristallinum cambriense, Eliasum llaniscum, Leiosphaeridia div, Leiovalia sp., Micrhystridium (=Heliosphaeridium) obscurum, Ovulum lanceolatum, Solisphaeridium flexipilosum and Timofeevia phosporitica (Jankauskas in Belyakova 1988). The acritarchs indicate a Cambrian age. Higher in the drillhole, Tremadocian acrotretid brachiopods have been found (Belyakova 1988). The Bol'shepul'skaya-l drillhole (location 4 on Fig. IB) penetrated coarse-grained sandstones

unconformably overlying Riphean metamorphic rocks at a depth of 1587 m. About 50 m above these unfossiliferous sandstones, grey mudstones yield the monoplacophoran mollusc Kirengella (Gubanov & Bogolepova 2003), which was described for the first time from the Upper Cambrian of Siberia (Rozov 1968). This mollusc is known from the Upper Cambrian-Lower Ordovician Kidryasovskaya Formation of the southern Urals (Doguzhaeva 1972); also from the Upper Cambrian-Lower Ordovician of the Ozark Uplift of Missouri (Stinchcomb & Angeli 2002). Higher in the drillhole (interval of 1246-1252 m) are siltstones intercalated with limestones with rare acrotretid brachiopods and the conodonts Acodina aff. lirata, Drepanodus ex. gr. lineatus, D. aff. simplex, Oneotodus aff. gracilis, O. aff. erectus, and Paltodus (?) aff. bassleri, that suggest a late Tremadoc-early Arenig age (Belyakova 1988). Further south, in the Northern Urals, similar lithologies, most probably coeval with the Izhma-Omra complex have been described from the Polyudov Range (location 8 on Fig. IB) above the Ordovician unconformity (Timonin 1998). Urals mountains Within the Uralide Orogen, the Timanide complexes are exposed locally beneath the strata in the western Uralian Palaeozoic to Early Mesozoic foredeep basin and also in the allochthonous Uralian oceanic and island-arc terranes of the hinterland (location 5 on Fig IB). Riphean (Neoproterozoic, reaching back into the Mesoproterozoic) successions occur in the cores of major anticlines (Peive et al 1977) that dominate the Uralian foreland fold-andthrust belt (e.g. the Bashkirian and Kvarkush-Kamennogorsk anticlinoriums, locations 9 and 11 on Fig. IB). The Timanide complexes, which locally include older crystalline rocks of the East European Craton margin, are unconformably overlain by the Early Palaeozoic successions that are the main subject of this paper. Polar Urals Beneath the Uralian foredeep basin of the Polar Urals, the basement is too deep to be sampled by drilling; hence the basal

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Fig. 3. The Polar Urals, showing the western limb of Engane-Pe (E) and other anticlines, cored by Timanide complexes with Early Palaezoic cover, overthrust by Uralian allochthons (based on Peive et al 1977; Didenko et al 2001).

Palaeozoic successions between the central part of the Pechora Basin and the Urals are unknown. However, in the Urals foreland fold-and-thrust belt, Timanian basement and overlying Early Palaeozoic cover are well exposed in anticlines (e.g. Engane-Pe and Manita-Nyrd), where a major unconformity is readily demonstrable. Further east, the Palaeozoic successions occur in thrust sheets in the footwall to the main Uralian allochthons of highpressure rocks and ophiolites (Zoloev et al 2002). The unconformity below the Lower Palaeozoic conglomerates and quartzites, which overlies the Neoproterozoic volcaniclastic sediments and volcanics, is well exposed in the western limb of Engane-Pe Anticline (location 5 on Fig. IB, Figs 3 & 4). In the core of this inlier, the greenschist facies Neoproterozoic rocks strike NW, strongly discordant to the overlying Palaeozoic strata. A fragmented ophiolite (Dushin 1997; Scarrow et al. 2001) has been described and associated plagiogranites have been dated to 670 Ma by the U-Pb method on zircons (Khain et al 1999). Other localities with a well-exposed angular unconformity between the Neoproterozoic and Lower Palaeozoic rocks are situated on the Usa River in the vicinity of Vorkuta (Fig. 5). The units below the unconformity belong to the Upper Riphean-Lower Vendian Enganepeiskaya Group (up to 700 m), which is composed of turbidites, tuffs, basic effusive rocks and tuff breccias. The lower part of this group contains black shales, siltstones and subordinate sandstones; acritarchs have been reported, but appear not to be diagnostic. The upper part of the group is represented by light siltstones and sandstones. They yield the acritarchs Botuobia wernadskii (=Oscillariopsis wernadskii), Leiosphaeridia sp., Nucellosphaeridium sp., the algae Arctacellularia doliiformis, Eomycetopsis tipicus, Polythrichoides lineatus and Tortunema sibirica (Dembovsky et al 1988). These Polar Urals Vendian fossils and successions are closely comparable with those of the East European Platform (IFshenko in Dembovsky et al 1988), the southern part of Novaya Zemlya (Smirnova et al 1982), Vaigach Island and northwestern Pai-Khoi (Dembovsky etal 1988). The strata unconformably overlying the Timanian formations in the Usa River section are composed of conglomerates, gritstones, quartz sandstones and siltstones of the Manitanyrdskaya Group. These strata are typical of the Cambrian-Ordovician successions in the Engane-Pe and Manita-Nyrd areas (Fig. 5), where the relationships between these units have been studied in detail (Varganov et al 1973; Dembovsky et al 1990). On the banks of the Iz'ya-Vozh, Niyua-Yu, Smerti and Khoidyshor creeks, at the base of the Manitanyrdskaya Group, the Upper Cambrian Bad'yashorskaya Formation (up to 3000 m), comprises thick units (up to

100 m) of acid and basic volcanics and volcaniclastic rocks, which grade laterally and upwards into sandstones and siltstones. They are succeeded by the Schugorskaya Formation, composed of Middle Ordovician sandy limestones and overlying limestones and dolomites. Cambrian microfossils have been found from the lower part of the Manitanyrdskaya Group in the Khoidyshor Creek (Fig. 5, section A) about 30 km north of the type section. They are represented by the acritarchs Archaeodiscina umbonulata, Leiomarginata simplex, Skiagia compressa (=Baltisphaeridium compressum), S. orbiculare (=Baltisphaeridium orbiculare), Simia sinica (=Granomarginata squamacea), and Tasmanites tenellus (Varganov et al 1973). In the uppermost part of the Manitanyrdskaya Group, the problematic molluscs Angarella lopatini and A. obrutchevi have been recorded in sandstone lenses. These forms also occur in the Arenig-Llanvirn of Siberia and Taimyr. Further east in the Polar Urals, in the Kharbei Complex and immediately underlying allochthons, fragmented ophiolites occur, overlain by quartzites and limestones in association with basalts and rhyelites. The sedimentary rocks have not yielded fossils. This association may be an eastern extension of the Early Palaeozoic platform facies. Further SE in the Polar Urals, the basal succession of conglomerates and quartzitic sandstones is overlain by a 1200 m thick sequence of shales with cherty beds and volcanic and volcaniclastic units. The volcanics are dominated by subalkaline basalts, tuffs and turfites. This basinal facies is referred to as the Lemvinskaya Facies. Outcrops along the Lemva River, situated in the southern part of the Polar Urals (Fig. 5, Sections F, E) show the Upper Cambrian to Tremadoc Pogureiskaya Formation unconformably overlying the Vendian Man'inskaya Formation (Saranin 1972; Vodolazsky 1983; Dembovsky et al 1988). The Man'inskaya Formation is composed of grey shales and metatuffs, and is intruded by porphyritic diabases. Along the neighbouring Malaya Bad'ya-Yu River, these Timanian rocks occur in the cores of Uralian anticlines. They include thin-bedded sandstones, siltstones and clayey limestones with subordinate lenses of tuffaceous sandstones. The sandstones and shales contain the microphytofossils Arctacellularia doliiformis, Chuaria circularis, Leiosphaeridia sp., Leiominuscula sp., Macroptycha biplicata, M. uniflicata, Protosphaeridium densum, P. holtedahlii, Satka granulosa, Stictosphaeridium implexum, S. pectinate, and Synsphaeridium sorediforme (Dembovsky et al 1988). Chuaria shows a global distribution and is most abundant in the Neoproterozoic, around the Late Riphean-Vendian boundary, where it is considered to be stratigraphically important (Vidal et al 1993).

TIMANIAN UNCONFORMITY, NORTHWESTERN RUSSIA

Fig. 4. The major angular unconformity between the Neoproterozoic Engane-Pe complex and overlying Early Ordovician (perhaps Late Cambrian) quartzites (A) a general view, photo by G. Savelieva 2002; (B) an outcrop, photo by D. G. Gee 1996.

The lowermost part of the unconformably overlying Pogureiskaya Formation (600 m thick) consists of polymict conglomerates, gritstones, and quartzitic sandstones. This part of the succession is well exposed on the Pogurei-Egarta Creek (Fig. 5, Section E), where the basal sandy beds above the unconformity yield the brachiopod Eoorthis plenus, and the trilobites Micragnostus sp. and Olenidae gen. indet (Varganov et al 1973; Dembovsky et al 1990), indicative of a Late Cambrian age. The upper part of the Pogureiskaya Formation consists of sandstones, siltstones and shales. The most complete section is known from the Pokoinitsa River (Varganov et al 1973), where glauconitic sandstones and siltstones with interbeds of gravelly sandstones, calcareous sandstones and glauconitic limestones occur. Fossils

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recorded from these strata include the brachiopods Alimbella armata pagaensis, Altorthis kinderlensis, Finkelnburgia (?) sp., Lermontella pagensis, Medessia uralica, Plectotrophia (?) alata, Tritoechia lermontovae, Synbasic trophopsis, the trilobites Dolgedola (Dolgeuloma) multicava, Hystricuridae, Leimitzia bavarica, L. pagica, Neoaldanaspis sibiricus, the conodonts Cordylodus proavus, Oneotodus singularis, graptolites and gastropods (Varganov et al. 1973). Conodonts suggest a Late Cambrian age, based on the presence of Cordylodus proavus (Cooper et al 2001). From the overlying limestones, the brachiopods Alimbella sp., Apheoorthis vicini, Obolus sp., Orthidae sp. indet, Tritoechia aff. quebecensis, T. lermontovae, the trilobites Apatokephalus (?) sp. indet., Geragnostus adductus, Jdyia tuberosa, Niobe sp., and the conodonts Acodus erectus, Acontiodus latus, Oneotodus variabilis and Scolopodus peselephantus have been reported (Dembovsky et al 1990; Antsy gin 2001). They indicate an early Tremadocian age for the beds, most probably the upper part of the C. angulatus zone. A lateral equivalent of the upper part of the Pogureiskaya Formation is the c. 900 m thick Grubeinskaya Formation, described from the Grube-Yu River type section (Voinovsky-Kriger 1967). These strata comprise green shales with interbeds of siltstones, limestones and volcanic rocks. Limestones have provided a few fossils, for example, the brachiopods Bobinella ex gr. kolumbensis, Finkelnburgia sp. and Eoorthis sp. (Varganov et al 1973), indicating a Late Cambrian age. Dark-grey and green-grey siltstones of the upper part of the Grubeinskaya Formation yield the Tremadocian trilobites Niobe aff. laviceps, the conodonts Trinodus cf. sedenbladhi, the brachiopods Altorthis kinderlensis, Alimbella armata pagaensis and Lingullela sp. In the silico- and tuffitic shales of the section on the Pokoinitsa-Shor River, the trilobites Apatokephalus sp., Asaphidae, Ceratopyge ex. gr.forficula, Geragnostus abductus, Euloma sp., Niobe (?) sp., Pliomeroides sp. and Shumardia sp. have been found (Varganov et al 1973). Shales exposed along the Paga and Pokoinitsa rivers yield Agnostidae, Apatokephalus sp., Asaphidae, Ceratopyge ex gr. forficula, Euloma sp., Shumardia sp. and the graptolite Dictyonema sp. Within the limestone interbeds among the shales, Cyclopyge (?) sp., Euloma sp., Harpides (?) sp., Niobe sp. and Pleiomegalaspis (?) sp. have been collected. These strata have been correlated with the upper Tremadocian of the southern Urals and Pai-Khoi (Dembovsky et al 1988). Ordovician acritarchs, probably of Tremadocian age, have been reported from the upper part of the Grubeinskaya Formation. In the uppermost part of the Grubeinskaya Formation, the graptolites Phyllograptus densus and P. elongatus have been found, indicating an Arenigian age (Dembovsky et al 1990). Locally, the Grubeinskaya strata replace the Pogureiskaya Formation, resting with angular unconformity on the Vendian Man'inskaya Formation. This contact with the underlying rocks has been studied (Dembovsky et al 1990) along the Malaya Bad'ya-Yu River (Fig. 5, section F), where Vendian porphyritic diabase dykes are unconformably overlain by conglomerates of the Grubeinskaya Series. Sub-Polar Urals Within the Sub-Polar Urals (Fig. 1), the Timanian Complex is deformed by Uralian folds and thrusts, with deformation increasing from west to east. In the vicinity of Kozhim, southwards from the Polar Urals, in the so-called Sub-Polar region (location 7 on Fig. IB), Neoproterozoic, low greenschist-facies metamorphosed sedimentary and volcanic rocks are unconformably overlain by Upper Cambrian— Lower Ordovician clastic and volcanic successions, up to 3000 m thick. The Timanian rocks are represented by Upper Riphean marbles, basalts, and tuffs, up to 2000m in thickness. They are intruded by Late Riphean-Vendian gabbros and diorites of

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Fig. 5. The Neoproterozoic-Palaeozoic boundary interval in the Sub Polar and Polar Urals (based on Varganov et al. 1973; Dembovsky et al. 1990, modified).

the Khatalambinsky and Parnuksky suites and Vendian granites of the Kozhim suite (Pystin 1994; Zaporozhtseva & Pystin 1994). The overlying Upper Cambrian deposits (Tel'posskaya Formation) correspond to the basal part of the siliciclastic Manitanyrdskaya Group (see the Polar Urals, above). Along the Bol'shaya Katalambiya River, siliceous gravelstones and conglomerates occur at the base of the group. Upwards, they are gradually replaced by quartz sandstones, the thickness of which increases westwards to c. 650 m. The Lower Ordovician part of the Manitanyrdskaya Group, in this area, is subdivided into the Obeizskaya (450 m) and Saledskaya formations (500 m), composed of intercalations of cross-bedded sandstones, siltstones and subordinate shales. The precise age of the Obeizskaya Formation is not known. However, as the overlying Saledskaya Formation contains Arenigian fossils (brachiopods, lingulids, and bryozoa), a Tremadocian age is considered probable (Antsygin 1993). On the Rossomakha Ridge (location 6 on Fig. IB), a contact between Upper Cambrian conglomerates and sandstones and underlying Upper Neoproterozoic sericite-quartz schists is well defined (Fig. 5, section G). An angular unconformity between the conglomerates of the Grubeinskaya Formation and the underlying volcanics is well represented on the Molyudzh-Vozh River, crossing the Rossomakha Ridge. Here, a few metres of basal conglomerates occur as 'pockets' overlying different horizons of basement rocks (Dembovsky et al. 1990). The age of the Manitanyrdskaya and Grubeinskaya formations in the Sub-Polar Urals is thought to be Late Cambrian-Early Ordovician (Fig. 6). In the upper part of the Manitanyrdskaya Group (the Usinskaya Formation) layers of calcareous sandstones yield the problematic molluscs Angarella cf. lopatini, A. cf. laevis, the brachiopods Lycophoria sp. and Ranorthis sp., the trilobites

Niobe sp., and crinoids and conodonts that together indicate an Arenigian age (Dembovsky et al. 1988). In another river crossing the Rossomakha Ridge, the Lower Ordovician sandy limestones yield the brachiopod Rhysostrophia vorkutaensis (Andreeva 1977). Silty shales developed on the Srednyaya Kokpela River (Rossomakha Ridge) yield the Lower Ordovician acritarchs Baltisphaeridium bifurcation, Leiosphaeridia sp., Lophosphaeridium obtusatum, Trapezochitina (?) sp., Trashysphaeridium atteniatum and Tylosphaeridium unduratum (Dembovsky et al. 1990). Northern Urals A Lower Ordovician angular unconformity on the Polyudov Ridge (location 8 on Fig. IB) has been described by Chechia (1955). Below the unconformity, Upper Vendian conglomerates, sandstones and siltstones of the Kocheshorskaya Formation occur. K-Ar isotope dating of authigenic glauconite from these strata has yielded an age of 560-590 Ma (Sokolov & Fedonkin 1990). Unconformably overlying these Timanide units are quartz sandstones, gritstones, and conglomerates of the Polyudovskaya Formation. Although fossils have not been found, these strata may be of the same (?) Middle Cambrian-Tremadocian age as in the neighbouring Pechora Basin (Timonin 1998, see above). In the upper reaches of the Pechora River, similar strata yielded the Tremadocian brachiopods Dinorthis sp., Tritoechia lermontovae and trilobites (Antsygin 2001), but the descriptions of these sections are incomplete and the relationships with the underlying units are unknown. Figure 6 shows a profile across the Polar and Northern Urals, summarizing the Cambrian-Ordovician, west-east lithofacies development.

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Fig. 6. A lithofacies profile of the Upper Cambrian and Ordovician sediments across the northern Urals (based on Dembovsky et al. 1990, modified).

Middle Urals The Palaeozoic foreland fold and thrust belt of the Middle Urals (location 9 on Fig. IB) is dominated by the Kvarkush-Kamennogorsk Anticlinorium. This Uralian structure separates the islandarc volcanics and fragmented ophiolites of the Tagil Zone from the Uralian foredeep to the west (Fig. IB). The KvarkushKamennogorsk Anticlinorium is cored by a greenschist-facies (locally with blueschists, see Beckholmen & Glodny 2004) sedimentary complex, unconformably overlain by Palaeozoic cover. In the hinge and eastern limb of the Kvarkush-Kamennogorsk Anticlinorium, the latter is Ordovician in age; further west, Silurian strata overlap onto the Neoproterozoic basement and these, in turn, give way to Devonian sandstones that rest directly on Vendian in the western limb of the anticline. The Timanide Complex, in the core of the KvarkushKamennogorsk Anticlinorium, consists of phyllites, psammites, subordinate marbles, trachybasalts and trachyandesites of probable Neoproterozoic age, passing up into Vendian tillites. Gabbro-diabase dykes with K-Ar ages from 630 to 590 Ma are known within this succession (Raaben 1994). The overlying Upper Vendian rocks are less deformed and metamorphosed than the underlying rocks and do not contain the mafic intrusions; they may be separated from them by an unconformity. The uppermost part of the pre-Palaeozoic succession is interbedded polymict and quartz sandstones, siltstones and shales of the Ust'sylvitskaya Formation (Antsygin 1993). The Upper Vendian Ediacarian Metazoa Nemiana simplex and Tirasiana concentralis have been recorded from the upper part of this formation (Bekker 1980). The basal strata above the Palaeozoic unconformity, represented by the Lower Ordovician Kozinskaya and Kolpakovskaya formations, are composed of conglomerates, quartz sandstones, tuffs, trachybasalts and rhyolites (Antsygin 1993). The Kozinskaya sandstones (location 10 on Fig. IB) yield the brachiopods Tritoechia lermontovae, T. kodymi, T. subkolichai and the trilobites Jdyia tuberosa (Antsygin 2001) indicating a Tremadocian age. Southern Urals In the southern Urals (location 11 & 12 on Fig. IB), another major fold, the Bashkirian Anticlinorium (Fig. 7), comparable to the

Kvarkush-Kamennogorsk Anticlinorium of the Middle Urals, dominates the bedrock geology to the west of the Uralian ophiolites and island-arc complexes of the Magnitogorsk Zone. This major structure provides the type area for Riphean successions (Keller & Chumakov 1983); it also includes metamorphosed units (Beloretsk Complex) with amphibolites and eclogites (Matennar et al 1999; Glasmacher et al 2001). As in the KvarkushKamennogorsk Anticlinorium, Ordovician successions unconformably overlie the pre-Palaeozoic basement in the eastern limb of the Bashkirian Anticlinorium and there are Silurian and Devonian strata immediately above the unconformity further to the west. The Riphean and Vendian of the Southern Urals are well defined by means of fossils: stromatolites, microphytolites, acritarchs, algae and an Ediacarian fauna (Bekker 1977, 1980). Riphean sequences are composed of carbonate, siliciclastic and volcanic rocks. Rich assemblages of microfossils characterize the whole succession. Vendian deposits (the Ashinskaya Group) are represented by arkoses and quartz sandstones, breccias, and conglomerates, shales, and limestones (Maslov et al 1996), and are well characterized by microfossils (Jankauskas 1982). In the Basinskaya and Ziganskaya formations acritarchs and the algal filaments Kapitophyma ovalis, Omalophyma angusta, and O. gracilis, typical of the Vendian of the East European Platform have been reported (Kozlov et al 1996). Sandstones of the Uryukskaya Formation yield the acritarchs Kildinella hyperboreica, K. nordia and Synsphaeridium sp. Within the eastern limb of the Bashkirian Anticlinorium (1 on Fig. 7), Middle Ordovician rocks rest on different levels of the Upper Riphean and Vendian strata with an angular unconformity (sections on the Belaya River, close to Mindibaevo). Brachiopod-bearing quartz sandstones, sandy dolomites and conglomerates represent the Ordovician strata (Kamaletdinov & Kazantseva 1977; Puchkov 1977). South of Beloretsk, in a section on the Buganak River (2 on Fig. 7), Late Ordovician Lower Silurian limestones unconformably overlie Riphean marbles (Mamaev 1977). South of the Bashkirian Anticlinorium, apparently within the Lemvinskaya Facies zone in the Sakmara area, Upper Cambrian and Lower Ordovician successions are well documented. An Upper Cambrian fauna has been reported from the Khmelevsky member (200-450 m thick) in the lower part of the Kidrysovskaya Formation (Antsygin 1993), exposed in the Ural River tributaries (the Khmelevka, Kayala, Blyava, Kidryasovo, Chaushka, Medes

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Fig. 7. A geological sketch map of the southwestern Urals showing relationships between the Neoproterozoic and Palaeozoic rocks (based on Sobolev et al 1986; Willner et al 2001).

and Alimbet rivers). Conglomerates are at the base of the succession that passes into sandstones and siltstones. Sandstones yield the brachiopods Billingsella akbulakensis, Lingulella sp., Obolus sp., and the trilobites Eoshumardia pustulata, Jdyia sp., Kujandaspis aff. kujandensis, Micragnostus porosus, and Olentella sp. indet. (Antsygin 2001). From the overlying upper part of the Kidryasovskaya Formation (150m thick), composed of sandstones and siltstones with limestone and shale lenses, the brachiopods Acrothyra sp., Alimbella armata, Altorthis kinderlensis, Finkelnburgia sp., Imbricatia lesnikovae, Ingria clouda, Lingula sp., Lingulella sp., Medessia uralica, Obolus rasumovskii, Triseptata nelidovi, Tritoechia lermontovae, and the trilobites Acerocarina keisaranica, Alimbetaspsis kelleri, Apatokephalus arduus, A. serratus, Dolgedola obunca, D. (Dolgeuloma) multicava, Hystricuridae sp., Geragnostus sidenblandhi, Jduia exussa, J. kosagachica, J. tuberosa, Jujuaspis keideli, Kainella sp., Leimitzia bavarica, L. furculosa, L. pagica, Macropyge foliacea, Micragnostus kidrasensis, Medeselaspis ampus, Niobe sp., Peltocare recta, Promegalaspides kasachstanensis, Pseudokaniella pustulata, the graptolites Dictyonema inexpectatum and D. ex. gr. flabelliformis (Varganov et al 1973; Antsygin 2001), and the monoplacophoran molluscs Kirengella kultavasaensis, Romaniella aebitensis, R. getlingi and R. zwerevi (Doguzhaeva 1972) have been reported. These strata are correlated with the Upper Cambrian-Tremadocian interval (Antsygin 1993). A remarkable and enigmatic association of Lower Cambrian strata occurs in the Sakmara area. About 25 km west of Orsk (location 12 on Fig. IB), massive archaeocyathid limestones have been described in the banks of the Kidryasovo River (Lermontova & Razumovsky 1933); these apparently occur as olistoliths (Puchkov 2000) in a largely siliciclastic and volcaniclastic unit,

the Tereklinskaya Formation. The latter is described as resting unconformably on Precambrian schists. The age of the Tereklinskaya Formation is not well constrained. Both late Cambrian conodonts (Nasedkina in Puchkov 2000) and Tremadocian acritarchs (Chibrikova 1977, 1997; Chibrikova & Olli 1999) have been described from these strata, but these may be exotic, and the host rocks may be as young as Devonian (Ryazantsev et al 2000). The Early Cambrian assemblage is represented by the algae Acanthina multiformis, Batinevia bayancolica, Chabakovia nodosa, Gordonophyton distinctum and Epiphyton penicillatum. They have been described by Korde (1973), who also identified the same forms in the Lower Cambrian Kameshkovskaya Formation of Altay-Sayan in southern Siberia. The latter overlies Riphean formations that are strikingly similar to those of the Bashkirian Anticlinorium. It has been suggested (Didenko & Pechersky 1993) that these areas of southeastern Baltica and southern Siberia were united until the Early Cambrian, and separated during the opening of the Early Palaeozoic Uralian ocean (Maslov et al 1996). Resolution of this enigma awaits comprehensive investigations of the Sakmara successions.

North of the Polar Urals and Pechora Basin A major anticline extends northwestwards from the northern part of the Polar Urals, through Pai-Khoi and Vaigach Island to southern Novaya Zemlya (Fig. 1A). Pre-Palaeozoic successions are exposed in the core of this structure. A major angular unconformity between Proterozoic basement and overlying Lower Palaeozoic cover, the latter represented mainly by siliciclastic

TIMANIAN UNCONFORMITY, NORTHWESTERN RUSSIA

and carbonate successions, has been reported in these areas (Bondarev 1964, 1970; Bondarev et al 1973). Pai-Khoi The Timanide Complex of Pai-Khoi (Fig. 1 A) is represented by a c. 6000 m thick Riphean-Vendian sequence, composed of three formations. A lower unit, the Upper Riphean, Amderminskaya Formation, is made up of limestones and dolomites. A middle unit, the Upper Riphean-Vendian Morozovskaya Formation, is dominated by tuffs, and tuffitic sandstones, and an upper unit, the Vendian Sokol'nikskaya Formation, by tuffs, dark shales, sandstones and conglomerates, intercalated with acid volcanics and volcaniclastic strata. Magmatism on both Pai-Khoi and Vaigach is reported to be calc-alkaline (Korago & Stolbov 2002). On the Bol'shaya Edunei Ridge of western Pai-Khoi, Ordovician strata of Eletskaya Facies rest unconformably on the Timanian Complex. These Ordovician strata comprise light quartzites, quartz sandstones and shales that pass up into sandy limestones with Angarella. Further east, the unconformably overlying strata are basal sandstones, passing up into limestones with subordinate shales with the brachiopods Alimbella cf. armata, Altorthis kinderlensis, Aphaeorthis sp., Clarkella sokolyna, Ingria sp., Finkelnbergia sp., Medessia sp., Syntrophopsis ? sp., Tritoechia lermontovae, the trilobites Apatokephalus heterosulcatus, Asaphellus sp., Carolinites ex gr. genacicana, Ceratopyge ? ex gr. forficuloides, Cybelurus sokoliensis, Dikelokephalina aff. discraeura, Geragnostus karskensis, GeragnoStella sp., Harpides sp., Hystricurus sp., Lakorsalina limbata, Megalaspides sp., Megistaspis sp., Nileus limbatus, Niobella parvula, Nyaya novozemelica, Pliomeroides primigenus, Remopleuridiella ? sp., Shumardia sp., Tersella ? magnaoculus, the graptolites Tetragraptus (Tetragraptus) sp., and the monoplacophorans Pseudoscinella aff. sibirica (Bursky 1970; Antsygin 2001). The age of these strata is Arenig (Cocks & Fortey 1998). The brachiopod and trilobite faunas of Pai-Khoi are mainly endemic and show taxonomic similarities to the other parts of Baltica. However, the trilobites Nileus and Shumardia, known from Nevada and California (Fortey & Droser 1999), also occur. Vaigach Island On Vaigach Island (Fig. 1A), the Proterozoic rocks, c. 500m thick, are represented by basic volcanics and tuffs, overlain by shales and limestones. Vendian microfossils have been reported in this unit (Zaporozhtseva & Pystin 1994). A 90° unconformity separates these Neoproterozoic rocks from the overlying Palaeozoic strata of gritstones, quartz sandstones and siltstones. In the uppermost part of this succession, the brachiopods Altorthis sp., Imbricatia sp. and the trilobites Apatokephalus sp. and Pliomeroides defensus have been reported (Bondarev 1970) and correlated with the Arenig of Pai-Khoi. Novaya Zemlya The Proterozoic rocks of Novaya Zemlya (Fig. 1A, Fig. 8) occur in southern, central and northern areas (Bondarev 1970; Korago et al 2004). In the southernmost part of the island, the Timanide unconformity and an overlapping Lower Ordovician sequence has been described (Bondarev 1964; Korago et al. 1992) from the Piritovy Peninsula (Fig. 8). Siliciclastic rocks (up to 2000 m thick) of the Timanide Complex, were metamorphosed under low greenschist facies. These strata contain well-preserved sedimentary features (grading and ripple-drift lamination, characteristic of turbidites). The whole Riphean-Vendian sequence is subdivided into four

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units that have been well defined by Riphean and Vendian microphytofossils (Smirnova et al. 1982). The unconformably overlying Lower Ordovician strata are composed of conglomerates and quartz and polymict cross-bedded sandstones, overlain by limestones (Fig. 8). The latter have yielded the brachiopods Apheoorthis vicina, Syntrophopsis vetusta, the problematic molluscs Angarella sp., and bryozoa (Bondarev 1964) indicating an Arenigian age; and the underlying siliciclastics may be of Tremadocian age. In central Novaya Zemlya, Riphean rocks are intruded by plagiogranites and gabbros. A major intrusion of Vendian granites forms the Mityushev Kamen' Massif, directly overlain by Silurian sandstones (Korago et al. 1992). Cambrian rocks occur along the Matochkin Strait and contain trilobites, brachiopods and hyoliths in variegated sandstones and shales. These fossils were also found along the northwestern coast of the Southern Island, on the Karpinsky Peninsula (Fig. 8), in a fine-grained dark calcareous sandstone and, further north, in a light-coloured sandstone with thin layers of limestone (Holtedahl 1922). Walcott (1924) and Walcott & Resser (1924) reported the assemblage to include the brachiopods Acrotreta sp., Billingsella holtedahli, B. opius, Eoorthis sabus, Huenella triplicata, Lingulella arctica, Obolus (Westonia) sp., and the trilobites Acrocephalites vigilans, Agnostus holtedahli, A. pisiformis, A. septentrionalis, Dolgania megalops, Irvingella arctica, L septentrionalis, Kaninia crassimarginata, K. divaricans, K. lata, Kazelia speciosa, Koldinia typa, Orlovia arctica, Pesaia exsculpta and P. latifrons. Walcott and Resser (1924) proposed a Late Cambrian-Early Ordovician age for these strata. Popov (1985) reported the following Upper Cambrian inarticulate brachiopods from the northwestern part of the Southern Island: Acrothele cf. coriace, Angulotreta postapicalis, Anobolotreta sp., Dictyonina sp. and Prototreta gribovensis. This assemblage shows a taxonomic similarity with those of Estonia (Volkova 1982) and Laurentia (Palmer 1954). Investigations by Ermolaev between 1931-1933 showed the presence of Middle and probably Lower Cambrian rocks in the Karpinsky area (Bondarev 1970). From the lowermost part of the section (Astaf'evskaya Formation), represented predominantly by green-grey shales, alternating with grey siltstones, sandstones and limestone lenses, the Toyonian-Amginian trilobites Ellipsocephalus sp., Protagraulos priscus, P. priscus novozemelica, P. proscus gribovae, P. subpriscus, Strenuella sp. and Solenopleurella sp. have been described (Solov'ev et al. 1986). Higher up, the sequence is characterized by the Amginian trilobites Chondragraulos sp., Ellipsocephalus polytomus, E. gurichi, Strenuella (Comluelld) sp., Paradoxides (Eccaparadoxides) oelandicus pinus, P. (Paradoxides) paradoxissimus, P. (Acadoparadoxides) sacheri, Solenopleura sp., the brachiopods Acrothele sp., Acrothyra sp., Hadrotreta sp., trace fossils and hydrozoid imprints (Solov'ev et al. 1986). In northern Novaya Zemlya, an apparently continuous Late Neoproterozoic-Palaeozoic succession occurs. The RipheanVendian Lomonosovskaya Formation is represented by a siliciclastic and carbonate successions. The Vendian part is well defined by acritarchs (Korago et al. 1993). The overlying shales contain Paradoxides aurora and P. hicksi of the Middle Cambrian (Mendeleevskaya Formation) and the succession passes up into Upper Cambrian-Lower Ordovician sandstones with Billingsella holtedahli, Irvingella septentrionalis and Orlovia arctica (Solov'ev et al. 1986). Higher in the succession, green shales and sandstones of the Middle Ordovician Vel'kenskaya Formation occur. They are characterized by Arenigian-Llandeilo graptolites (Bondarev et al. 1985). Conclusions The data presented above show the presence of a major angular unconformity between Neoproterozoic complexes and Lower

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Fig. 8. A geological sketch map of Novaya Zemlya, showing the location of the Neoproterozoic-Palaeozoic unconformity (based on Romanovich 1970).

Palaeozoic successions in the Timan Range, beneath the Pechora Basin and along the western front of the Urals, from the far south to southern Novaya Zemlya. This unconformity provides regional evidence of an erogenic process that took place along the entire northern and eastern margins of the East European Craton during the latest Precambrian. The Riphean-Vendian history of the eastern margin of the East European Craton is characterized by the development of a wide pericratonic basin, giving way eastwards, in the Neoproterozoic, to continental slope and rise environments that accumulated thick turbidite-dominated successions. Outboard of this continental margin there is evidence of island-arc magmatism and ophiolites; also, at least locally, of subduction-related (high P/T) metamorphism. Calc-alkaline igneous rocks beneath the Pechora Basin and a variety of granitoids, ranging in age from c. 630 Ma on Novaya Zemlya to c. 560 Ma beneath the Pechora Basin, provide evidence of Timanian accretion and associated magmatism. The metamorphism is largely of Vendian age, but may have continued into the earliest Cambrian.

During much of Cambrian time, the Timanide area was uplifted and eroded. In general, rocks of the Lower, Middle and lowermost parts of the Upper Cambrian are missing. Only in central Novaya Zemlya are there are reliable records of Middle and Upper Lower Cambrian strata, suggesting that Timanian Orogeny may have ended earlier in northernmost areas. The apparent lack of a Timanian unconformity in northern Novaya Zemlya is enigmatic. Beneath the Lower Palaeozoic unconformity there is local evidence of extensive weathering. Locally, in the basement several kilometres beneath the Pechora Basin, red sandstones and acid volcanic rocks have been found beneath the Palaeozoic unconformity. Their low-grade metamorphism and deformation have suggested that they may represent a hinterland-basin molasse facies, deposited on underlying more altered Timanide complexes. Decisive evidence is lacking. Sedimentation started in the Ordovician or, particularly in eastern areas along the front of the Urals, in the Late Cambrian. Upper Cambrian and Tremadoc strata were apparently deposited during early rifting of the Baltica margin and a passive margin was

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established by the Arenig. Volcanism occurs throughout the Ordovician and, at least locally, in the Late Cambrian, in general being characterized by alkaline basalts, porphyrites and tuffs. The Early Palaeozoic rifting of northeastern Baltica was discordant to the Timanide Orogen in the Timan-Pechora-Polar Urals region; further south, it is nearly parallel and apparently controlled the deformation front of the Urals during the subsequent Palaeozoic orogeny. The Upper Cambrian to Ordovician successions throughout most of the area underlain by Timanide basement have a similar facies development, dominated by shallow marine environments, known as the Eletskaya Facies. The basal beds that unconformably overlie the Timanide units, are composed of siliciclastic, poorly sorted continental and shallow marine deposits. Overlying these basal elastics in the east (Uralian front) is a monotonous Upper Cambrian and/or Tremadocian succession of sandstones, siltstones, limestones and shales. The sandstones and sandy shales are cross-bedded, rippled and they generally lack macrofossils (or the latter have extremely low diversity) and the facies is interpreted to represent coastal marine, shallow water, high energy environments. These occur in Pechora, Kozhim, Vaigach, the western Pai-Khoi (Bol'shaya Edunei), Novaya Zemlya, and the Middle Urals. Only in the most easterly outcrops in the front of the Urals is a basinal succession of Late Cambrian and Ordovician age represented, dominated by shales and termed the Lemvinskaya Facies. This facies was deposited in somewhat deeper water. The Late Cambrian-Tremadocian strata of Lemva, Pogurei and Grube-Yu in the Polar Urals, the eastern Pai-Khoi (Lakorsale), and the Southern Urals are represented by limestones and shales with brachiopods, trilobites, gastropods, conodonts, acritarchs and graptolites. A deepening of the basin from west to east (in modern coordinates) was accompanied by increasing sediment thickness. These two parallel facies zones, the Eletskaya and Lemvinskaya, are traced along the whole eastern margin of Baltica. The fauna show no evidence of barriers for distribution and migration along this wide, long (2000 km) platform margin. Thus common species such as Alimbella armata, Altorthis kinderlensis, Apatokephalus sp., Medessia uralica and Tritoechia lermontovae occur throughout the Polar, Middle, Southern Urals and PaiKhoi. The graptolite Dictyonema flabelliformis, widely known from the Baltoscandian platform is also found in the Polar and Southern Urals. The Cambrian-Ordovician acritarch association shows a taxomonic similarity with the other parts of Baltica, e.g. the Moscow Syneclise. Some of the taxa are geographically widely distributed. The brachiopod and trilobite faunas listed above are mainly endemic to Baltica, including lycophoriid brachiopods (Lycophoria is known from the Sub-Polar Urals) and asaphid and megistaspid trilobites (e.g. Megistaspis, with ocurrences on Pai-Khoi) or have a wide geographic distribution (Eoorthis, Finkelnburgia, Geragnostus, Irvingella, Lingulella). Interestingly, Novaya Zemlya and the Polar Urals show cosmopolitan elements in their fossil assemblages (Apheoorthis, Syntrophopsis and Tritoechia), whereas the Southern Urals share the occurrence of the trilobites Kujandaspis kujandensis with Kazakhstan, and Leimitzia bavarica is known from the Polar Urals, the southern Urals and Bohemia. The monoplacophoran mollusc Kirengella is found beneath the Pechora Basin and in the southern Urals. These shells were described from the southern (the Kirenga River) and northwestern (the Kulyumbe River) parts of the Siberian platform, and are also recorded in Laurentia. Kirengella has been reported from Severnaya Zemlya as well, but our studies do not support this assignment. Among the inarticulate brachiopods, Acrothele cf. coriacea, known from Novaya Zemlya, has been described before from Scandinavia, and Angulotreta postapicalis, from Laurentia and Estonia. Angarella, widely distributed in the Arenigian-Llanvirnian of Siberia and Taimyr, is typical of the Polar Urals, Pai-Khoi, Novaya Zemlya and the SubPolar Urals, but unknown so far from the Middle and Southern

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Urals. The presence of the deep-water trilobites Nileus and Shumardia in the Pai-Khoi and the Polar Urals successions might imply that these areas were located along the edge of the continent in the early Ordovician. This paper is a contribution to the INT AS projects HALE and NEMLOR. We are grateful to Dennis Brown (Barcelona) and Victor Puchkov (Ufa) for their constructive reviews, and in particular, to Malgorzata Moczydlowska-Vidal (Uppsala) and Robin Cocks (London) for their positive criticism and help with the fossil assemblages. We also thank Georgy Petrov (Ekaterinburg) for his help with the Russian literature.

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SHATSKY, N. S. 1935. [On tectonics of the Arctics. Geology and prospective minerals in northern USSR], 1, 476-509 [in Russian]. SHATSKY, N. S. 1958. Les relations du Cambrian avec le Proterozoigne et les plissements Baikaliens. Les relations entre Precambrian et Cambrian. CNRS Coll Institute, Paris, 91-101, 1957. SMIRNOVA, L. N., KORAGO, E. A. & KOVALEVA, G. A. 1982. [A subdivision of the Precambrian rocks of the Southern Island of the Novaya Zemlya Archipelago according to plant microfossils.] In: BONDAREV, V. I. (ed.) [Microfossils of the polar areas and their Stratigraphic significance] Sevmorgeologiya, Leningrad, 84-91 [in Russian]. SOBOLEV, I. D., AVTONEEV, S. V., BELKOVSKAYA, R. P., et al. 1986. [A tectonic map of the Urals, scale 1:1000000. Explanatory notes] Uralgeologiya, Sverdlovsk, 1-168 [in Russian]. SOBOLEV, N. V., DOBRETSOV, N. L., BAKIROV, A. B., SHATSKY, V. S. 1986. Eclogites from various metamorphic complexes in the USSR and the problems of their origin. Geological Society of America Memoir, 164, 349-363. SOKOLOV, B. S. & FEDONKIN, M. A. 1990. The Vendian System. SpringerVerlag, 2, 1-273. SOLOVEV, I. A., TRUFANOV, G. V. & KORAGO, E. A. 1986. [The Cambrian deposits of Novaya Zemlya.] Sovetskaya Geologia, 3, 65-76 [in Russian]. STINCHCOMB, B. L. & ANGELI, N. A. 2002. New Cambrian and Lower Ordovician monoplacophorans from the Ozark Uplift, Missouri. Journal of Paleontology, 76 (6), 965-974. TIMONIN, N. I. 1998. [The Pechora Plate: the history of geologic development in the Phanerozoic.] Ekaterinburg, 1-239 [in Russian]. VARGANOV, V. G., ANTSYGIN, N. Y. & NASEDKINA, V. A. 1973. [Stratigraphy and fauna of the Ordovician of the Middle Urals.] Nedra, Moscow, 1-228 [in Russian]. VIDAL, G., MODZYDLOWSKA, M. & RUDAVSKAYA, V. A. 1993. Biostratigraphic implications of a Chuaria-Tawuia assemblage and associated arcritarchs from the Neoproterozoic of Yakutia. Palaeontology, 36, 387-402. VODOLAZSKY, A. I. 1983. [Palaeozoic sediments and a development of the Lemva structural-facial zone of the Polar Urals.] Short PhD thesis, Leningrad, 1-17 [in Russian]. VOINOVSKY-KRIGER, K. G. 1967. [A review on the Lemva Zone tectonics.] Moskovskoe obschestvo ispytatelej prirody (geologiya), 42 (3), 3-27 [in Russian]. VOLKOVA, N. A. 1982. [Age constraints of the Yul'gazes suite and the Cambrian-Ordovician boundary in Estonia.] Sovetskaya Geologiya, 9, 85-88 [in Russian]. WALCOTT, C. D. 1924. Ozarkian brachiopods from Novaya Zemlya. In: Report of the scientific results of the Norwegian expedition to Novaya Zemlya 1921, 25. Kristiania. A.W. Br0ggers Boktrykkeri, 3-10. WALCOTT, C. D. & RESSER, C. E. 1924. Trilobites from the Ozarkian sandstones of the island of Novaya Zemlya. In: Report of the scientific results of the Norwegian expedition to Novaya Zemlya 1921, 24. Kristiania. A.W. Br0ggers Boktrykkeri. 3-15. WILLNER, A. P., ERMOLAEVA, T., STROINK, L., GLASMACHER, U. A., GIESE, U., PUCHKOV, V. N., KOZLOV, V. I. & WALTER, R. 2001. Contrasting provenence signals in Riphean and Vendian sandstones in the SW Urals (Russia): constraints for a change from passive to active continental margin conditions in the Neoproterozoic. Precambrian Research, 110, 215-239. ZAPOROZHTSEVA, I. V. & PYSTIN, A. M. 1994. [Composition and structure of the pre-Phanerozoic lithosphere of the European northeast of Russia.] Nauka, St'Petersburg, 1-110 [in Russian]. ZOLOEV, K. K., KOROTEEV, V. A., DUSHIN, V. A., RAPOPORT, M. S., SAVELIEVA, G. N. & SMIRNOVA, T. A. 2002. [Geology and mineragenesis of the Polar Urals and adjacent margin of the East European platform.] In: DODIN, D. A. & SURKOV, V. S. (eds) [The Russian Arctic: geological history, mineragenesis, environmental geology.} Vniiokeangeologiya, St Petersburg, 328-346 [in Russian].

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Late Cambrian relative age constraints by acritarchs on the post-Timanian deposition on Kolguev Island, Arctic Russia M. MOCZYDLOWSKA1, M. STOCKFORS1 & L. POPOV2 Uppsala University, Department of Earth Sciences, Palaeobiology, Norbyvdgen 22, SE-75236 Uppsala, Sweden (e-mail: [email protected]) ^Department of Geology, National Museum of Wales, Cathays Park, Cardiff CF1 3NP, UK 1

Abstract: Platformal sedimentary successions, inferred to overlie folded Neoproterozoic basement complexes and the post-Timanian unconformity, were studied in the Bugrino 1 and North-Western 202 boreholes on Kolguev Island, southern Barents Sea. They yielded diverse, age-diagnostic acritarchs and some shelly fossils. The Upper Cambrian strata, referred to the Peltura and Acerocare zones, are recognized by means of acritarch taxa known from the East European Platform and other areas in Baltica, Avalonia, and Gondwana. Shelly faunas of brachiopods, phyllocarid crustaceans and problematic molluscs are indicative of the Tremadoc and Arenig stages in the upper part of the succession. The boreholes did not reach the Timanian unconformity and older Cambrian strata may be present. Within the area of the Pechora Basin, Cambrian strata have been identified with certainty only on Kolguev Island. Further east in the Polar Urals, Upper Cambrian strata are present and Lower or Middle Cambrian and Vendian strata have been claimed, though not yet proven. They are reported from several localities, but the fossils are poorly documented. The biostratigraphic evidence and facies development indicate that Upper Cambrian deposits may have been distributed widely along the northeastern margin of Baltica, as they were on the East European Craton and in Baltoscandia.

Palaeozoic siliciclastic and carbonate sediments, several hundreds to thousands of metres thick, are widespread all over the Arctic Russia in northeastern Europe; they form the platformal cover, overlying an unconformity that separates them from the Timanian basement complexes. The post-Timanian unconformity is defined by an erosion surface that transgresses Neoproterozoic deformed and metamorphosed basement rocks and is overlain by nearly horizontal platformal sediments (Bogatsky et al. 1996; Gee et al. 2000). The Lower Palaeozoic strata are exposed and also known from subsurface occurrences in the Pechora Basin, which stretches between the Timan Range and the Polar Urals (Fig. 1). They accumulated in shallow marine and deep environments of the passive margin of the Baltica palaeocontinent (Nikishin et al. 1996). The Neoproterozoic metasedimentary and volcanic rocks, cut by granites and other intrusions dated to c. 550-620 Ma (Gee et al. 2000), are separated by faults into several crustal blocks. The latter were accreted to the edge of the East European Craton during the Timanian Orogeny (previously also referred to as Baikalian), which occurred c. 600-570 Ma (Gee et al. 2000). Sub sequently, the elevated Timanide Orogen was eroded for a substantial period of time to form a peneplained surface prior to the Early Palaeozoic transgression. The age of the basal sediments, including variously dated Cambrian and Ordovician strata, is essential for constraining the onset of the post-Timanian depositional cycle. The oldest sedimentary successions are known in two boreholes on Kolguev Island, allowing the definition of an accurate biostratigraphy. Kolguev Island is situated in the southern Barents Sea. The sedi mentary succession is inferred to overlie a basement that is the northwestern extension of the Pechora Zone and is a part of the Pechora Basin (Fig. 1). On Kolguev Island, the lowermost undeformed strata, known only from boreholes, were interpreted to be Early Ordovician in age (Preobrazhenskaya et al. 1995); however, the occurrence of Cambrian-Vendian sediments deeper in the successions was postulated from geophysical profiles (Olovyanishnikov 1998). The sedimentary successions in the Bugrino 1 and North-Western 202 boreholes (Fig. 2) yielded well-preserved and age-diagnostic acritarch microfossils and a few shelly fossils. Acritarchs are marine, phytoplanktonic, mostly algal, microorganisms, that lived in the photic zone and were dispersed as a marine snow and accumulated in sediments from suspension. Therefore, they are largely facies independent. The invertebrate shelly faunas are benthic brachiopods and

problematic molluscs, that are strongly influenced by the facies distribution; free-swimming phyllocarid crustaceans are also present. In this paper, we summarize the recently established acritarch data on the age of the lowermost sedimentary succession sampled on Kolguev Island (Moczydlowska & Stockfors 2004), re-evaluate the microfossil record by Rudavskaya (in Preobrazhenskaya et al. 1995) from the same sections and review the invertebrate faunas. The inferred Late Cambrian age of the strata constrains the timing of the Palaeozoic deposition on Kolguev Island. We also summarize and critically review the previous assessments, based on acritarchs, of the relative age of strata overlying the post-Timanian unconformity in the context of the regional occurrence of Cambrian rocks. Geological background The structural geology and tectonic history of the crustal blocks beneath the Pechora Basin are complex and the provenance of some of the units is not established with certainty. It is commonly accepted that the basement below Kolguev Island belongs to the Pechora Zone (Bogatsky et al. 1996, p. 24, 26; Gee et al. 2000), but other interpretations favour its affiliation with the Bolshezemelskiy Microcontinent (Gee & Ziegler 1996; Bogatsky et al 1996, p. 27; Olovyanishnikov et al. 1996, 1997, 2000; Olovyanishnikov 1998). The sedimentary cover overlying the folded and metamorphosed basement is, by contrast, lithologically and sedimentologically uniform, conforming to several regionally recognized depositional cycles. Differences between the cover successions on the individual blocks mostly involve the preservation of particular stratigraphic units and the duration of depositional hiatuses. The sediments immediately overlying the post-Timanian unconformity are of diverse ages on the discrete blocks of the northeastern margin of Baltica. In the Timan Range they are late Vendian in age; in the Pechora Basin, above the Izhma Zone, Early Ordovician with an exception locally of being Middle Cambrian, and further east above both the Pechora Zone and the Bolshezemelskiy Microcontinent Early Ordovician (Bogatsky et al 1996). However, the Middle Cambrian strata in the Izhma Zone are, neither documented in a published palaeontological record nor evident from the stratigraphic succession (see below).

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 159-168. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Simplified map of Arctic Russia in NE Europe showing (a) major tectonic zones of the Timan-Pechora plate and the location of Kolguev Island (modified after Gee et al. 2000) and (b) map of Kolguev Island with position of the North-Western 202 and Burginol boreholes.

In summary, the sedimentation began in most areas in the Early Ordovician or latest Cambrian times, the latter demonstrated unambiguously only on the Kolguev Island. The Palaeozoic sedimentary successions on Kolguev Island are among the best known in the Pechora Basin and their well-dated unconformities are important for regional tectonic and geodynamic reconstructions. The depositional environments are consistent with an open shelf along the margin of the Palaeozoic platform that extended across the Pechora Basin and passed into slope and deeper marine environments towards the Barents Sea, Novaya Zemlya and the Urals. During the Palaeozoic Era, deposition along this passive margin of the Baltica palaeocontinent was interrupted by uplift and regression, as evident from the regional unconformities. The Palaeozoic-Mesozoic succession on the Kolguev Island, which is nearly 5000 m thick, is known from about fifty hydrocarbon-prospecting boreholes; none reached the pre-Palaeozoic basement, which has been interpreted from seismic profiles (Bro et al. 1988; Preobrazhenskaya et al 1995). The lowermost rocks that have been core-sampled and contain fossils have been attributed to the Lower Ordovician (Preobrazhenskaya et al. 1995) or Upper Cambrian (Moczydlowska & Stockfors 2004). The Upper Cambrian strata are recognized only in two boreholes, Bugrinol and North-Western 202, by age-diagnostic microfossils (Moczydlowska & Stockfors 2004). In other parts of the island, unfossiliferous, but probably Middle Ordovician, quartzitic sandstones and siltstones occur in the base of some boreholes (Peschanoozerskie 4 to 46 boreholes; Preobrazhenskaya et al. 1995). The Palaeozoic strata in the boreholes are almost horizontal, unmetamorphosed and have a maximum thickness of 2700m (in the Bugrinol borehole), comprising Ordovician, Devonian, Carboniferous and Permian siliciclastic and carbonate strata and Upper Devonian basalts (Preobrazhenskaya et al. 1995). The succession is interrupted by four regional unconformities with stratigraphic breaks spanning the Late Ordovician-Silurian, Middle Devonian, and some intervals between the Late Devonian-Early Carboniferous and Early-Middle Carboniferous. The lowermost known Upper Cambrian and Lower-Middle Ordovician deposits accumulated in an offshore marine environment and form a single sequence (see description below), that is paraconformably overlain by the Lower Devonian sandstones of Old Red Sandstone facies. The pre-Devonian sedimentary cycle

was ended by uplift causing erosion and a significant stratigraphic hiatus. The spatial distribution and thickness of the CambrianOrdovician strata indicate that nearshore environments dominated towards the east and the fault bordering with the Bolshezemelskiy Microcontinent, whereas outer shelf environments prevailed towards the west and south of the island. The pre-Lower Devonian hiatus is substantially larger here than in the mainland Pechora Zone, where it spans only a part of the Middle Ordovician (Melnikov 1999). This depositional break is in sharp contrast to most areas in the Pechora Basin, where the Middle Ordovician to Early Devonian carbonate deposition was nearly continuous (Bogatsky et al 1996; Melnikov 1999). Cambrian rocks in the Pechora Basin and adjacent areas Cambrian strata have been reported in several parts of the Pechora Basin, in addition to Kolguev Island, but diagnostic fossils have not been unambiguously described. The occurrence of Upper Cambrian strata in deep troughs north and NE of the Kolguev Island, i.e. in the Pechora Sea, has been inferred from the deep seismic profiles (compilation in Bogatsky et al. 1996, p. 5, 78). In view of the evidence presented here, the extension of such rocks in the vicinity of the island seems likely. However, they pinch out towards the SW where, on the Kanin Peninsula and in the northern Timan, Devonian (locally Silurian) strata rest directly on basement. The occurrence of a pocket of Middle Cambrian strata above the Izhma Zone (Gee et al 2000), based on unpublished data on microfossils in shales unconformably overlying granite in the Mala Pera 11 borehole, suggests the possibility that Cambrian strata may have covered the region. However, this record awaits confirmation because of some inconsistencies between the stratigraphic ranges of the taxa reported. The acritarch assemblage from the Mala Pera 11 borehole, cited in more detail by Bogolepova & Gee (2004), is in general of Cambrian age, but the ranges of some species exclude each other, being restricted either to the Early or to the Middle-Late Cambrian. These discrepancies may depend on erroneous taxonomic determination of microfossils and/or imply that their stratigraphic ranges should be clarified. Undifferentiated Cambrian-Ordovician deposits may occur in the Malozemelskay a-Kolguev monocline (above the Pechora

LATE CAMBRIAN ACRITARCHS, KOLGUEV ISLAND

Fig. 2. The Upper Cambrian-Ordovician geological sections in the Bugrinol and North-Western 202 boreholes on Kolguev Island with the stratigraphic levels of the fossiliferous acritarch samples and occurrence of shelly faunas.

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Zone) and over the Bolshezemelskiy Microcontinent, extending into the Urals, and the Vendian-Cambrian deposits immediately beneath the Lower Ordovician in the Izhma-Pechora syneclise (map-sheet 2 and 57 in Bogatsky et al. 1996). Inferred Vendian-Cambrian terrigenous red beds, which include tuffaceous and acidic effusive rocks, are widespread underneath the Ordovician strata in the subsurface succession above the Bolshezemelskiy Microcontinent (Bogatsky et al 1996) and the Kolguev massif (Olovyanishnikov 1998). These rocks have been interpreted as products of washout and redeposition of the Riphean volcanics, forming a molasse related to the Timanian (Baikalian) Orogeny (Bogatsky et al. 1996, p. 23; Olovyanishnikov 1998, p. 151). In the Polar Urals, adjacent to the east of the Pechora Basin and the Bolshezemelskiy Microcontinent, Upper Cambrian deposits are thought to form the lowermost section of the terrigenous Manitanyrd 'Superformation' (or Group). The unit is exposed in the Kozhym River section and comprises also Lower Ordovician strata and initiates the Lower Palaeozoic sedimentary cycle (Antoshkina & Shishkin 2000, p. 15). The Cambrian deposits range up to 640 m in thickness and their age is established from acritarchs (unpublished report by Puchkov & Chuvaskov 1988, cited by Antoshkina & Shishkin 2000). In the report by Puchkov & Chuvaskov (1988, p. 30-35), the acritarch species identified by L.N. Ilchenko from the Manitanyrd Group in the Miniseysk and Baidaratsk regions, such as Tasmanites tenellus Volkova 1968, T. bobrowskae Wazyriska 1967 (cited under invalid name T. variabilis'), Granomarginata squamacea Volkova 1968, G. prima Naumova 1960, Skiagia orbiculare (Volkova 1968) Downie 1982 (cited as "Baltisphaeridium orbiculare'), Skiagia ornata (Volkova 1968) Downie 1982 (='B. ornatum), Asteridium tornatum (Volkova 1968) Moczydlowska 1991 (='Micrhystridium tornatum'), were referred to the Upper Cambrian. These taxa, if correctly identified (and this is the assumption on which they are taxonomically transferred to valid species herein), are known elsewhere exclusively from the Lower Cambrian and basal Middle Cambrian, with the exception of G. squamacea which extends into the lower Upper Cambrian (Volkova et al. 1983; Moczydlowska 1991, 1998). The identification and/or age assessment of the Manitanyrd microfossils require revision and therefore the age of the group remains uncertain. Cambrian acritarchs are also reported from the Khoidyshor Formation in the Khoidy-Shor rivulet section in the western Polar Urals, including Skiagia orbiculare (cited as invalid '#. orbiculare'), Skiagia compressa (Volkova 1968) Downie, 1982 (='B. compressum'), Granomarginata squamacea, Tasmanites tenellus and Archaeodiscina umbonulata Volkova, 1968 (Puchkov & Chuvaskov 1988, p. 15-16). These taxa also suggest the presence of Lower-Middle Cambrian strata. Besides these reports, there is no age control on the pre-Arenigian deposits in the Polar Urals, and the Cambrian-Tremadocian ages were mostly inferred from the stratigraphic position underneath strata with Arenigian brachiopods and bryozoans (Antoshkina & Shiashkin 2000, p. 15). In summary, the various Cambrian ages of marine terrigenous and partly molasse-like rocks underlying the Lower Ordovician deposits in the Timan-Pechora area and the Polar Urals are poorly established. In most cases, they are indirectly inferred from their position in base of the geological successions above the Timanian unconformity and beneath Ordovician strata. In general, the palaeontological evidence for such ages is too poorly documented to be accepted without reservation. Unambiguous Cambrian deposits that are geographically closest to Kolguev Island occur in the subsurface of the White Sea eastern coast, near to Arkhangelsk. Upper Cambrian clays, containing age-diagnostic brachiopods and acritarchs, are preserved as xenoliths in volcanic rocks in four drillcores (Popov & Gorjansky 1994). These occurrences are isolated from other Cambrian

strata on the East European Craton (St Petersburg area) by a distance of 500km. The preservation of this local outlier near Arkhangelsk is explained as a remnant of denudation of once widely distributed Upper Cambrian rocks in the White Sea area (Popov & Gorjansky 1994) as well as in Baltoscandia (e.g. on the island of Aland; Martinsson 1974). However, the records from the Izma Zone at Mala Pera and from the Polar Urals appear to indicate that even the possibility of Lower-Middle Cambrian strata should not be excluded. Although without exact determination of the series, the occurrence at Mala Pera adds to the known extension of Cambrian rocks around or across the Timan Range, stretching between the East European Craton (Moscow syneclise to the Arkhangelsk area), Pechora Basin and up to the Polar Urals. All these Cambrian rocks seem to belong to the same biogeographic province and contain taxonomically comparable acritarch assemblages and some benthic shelly fossils in common; they need further investigation.

Cambrian-Ordovician successions studied on Kolguev Island The oldest part of the sedimentary succession recognized on Kolguev Island is reached in the Bugrino 1 and NorthWestern 202 boreholes and referred to the Lower Ordovician (Preobrazhenskaya et al. 1995) and now to the Upper Cambrian (Fig. 2); yet none of the boreholes reached the Timanian basement. In one additional borehole, the East Peschanoozerskaya, quartz sandstones attributed to undetermined Ordovician are unfossiliferous and their age is inferred from lithostratigraphic correlation elsewhere and unconformably overlying Lower Devonian strata. The thickness of the Upper Cambrian and Lower Ordovician strata in the Bugrino 1 and North-Western 202 boreholes is 1280.0m and 361.0m, respectively. Formations have not been established, only chronostratigraphic series, and various informal lithological units have been described within the depth intervals measured in the cored successions or interpolated from the seismic profiles (Fig. 2). These strata underlie the Devonian unconformity and occur in the southern and northwestern areas of Kolguev Island; they consist of fine-grained siliciclastic rocks, well-bedded claystones and siltstones with subordinate intercalations of conglomerates up to c. 10m thick. In the Bugrino 1 borehole, the lowest section, 496 m of claystones and siltstones within the interval of 3660.0-4156.0 m, may be treated as a formation and subdivided into a few informal members. In ascending order these are: (1) monotonous, parallellaminated, dark to almost black claystones (3980.0-4156.0 m); (2) thin-bedded siltstones (3922.0-3980.0 m); (3) laminated claystones interbedded with siltstones (3710.0-3922.0 m); (4) thin-bedded siltstones (as in the unit 2) (3660.0-3710.0 m). The succeeding interval of 175m (between 3485.0-3660.0 m) also can be treated as a formation; it comprises an alternation of grey, greenish and reddish claystones and siltstones with some thin beds of conglomerate. The overlying strata (609 m) up to the depth of 2876.0 m, are predominantly siltstones, which are cut by the unconformity separating them from Upper Devonian pink and greenish-grey quartz and polymictic sandstones (Preobrazhenskaya et al 1995). In the North-Western 202 borehole, the entire succession below the Upper Devonian unconformity is made of monotonous dark grey claystones with thin intercalations of siltstones, approximately 361 m thick. It is truncated by a red, kaolinite-hematite bearing weathering layer, which is unconformably overlain by a thin bed of Frasnian sandstones and then basalt (Fig. 2). The whole sequence on Kolguev Island was interpreted to form a single transgressive-regressive cycle (Preobrazhenskaya et al. 1995). The grey clay stone-siltstone formation at the base was

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considered to reflect 'lagoonal-marine' and the overlying greenish-red and less frequently grey claystone, siltstone and sandstone unit as alternating marine and 'continental' conditions (Preobrazhenskaya et al 1995). There are only a few sedimentary structures reported in this sequence, and there is no convincing evidence to identify lagoonal or continental depositional environments. The grey, laminated clay stone-siltstone formation (3660.0-4156.0 m) of Bugrino 1 and the entire pre-Devonian sequence in North-Western 202 is inferred here to have accumulated under dysaerobic conditions in marine shelf environments. The deposition was mostly from suspension (parallel laminated clays) with periodic increase in grain size (thin-bedded silts) and occasional input of poorly sorted sand and grit (or sedimentary breccia) that may represent tempestites. The fossils embedded in the sediments are typically marine organisms, benthic and freeswimming invertebrates and microscopic planktonic algae. The uniform lithological facies and their large thickness (approximately 500 m) are consistent with continuous and fast subsidence of the basin; the calculated rate of sedimentation is in the range of 30 m/Ma. Such sedimentation rates are comparable to the average rates of 17-25 m/Ma estimated for the Pechora Basin and 23-30 m/Ma for the Pechora-Kolva Zone (=Pechora Zone) during Ordovician times (Malyshev 2000). The greenish-red sediments above the depth of 3660 m in the Bugrino 1 borehole show a higher content of siltstone and sandstone and interbeds of conglomerate. Small-scale cross-stratification, scoured surfaces and slumping structures are reported from this succession, along with interbeds of polymictic sandstone with grains and fragments of metamorphic slates and acidic effusive rocks (Preobrazhenskaya et al. 1995). These features indicate a shallowingup sequence with enhanced circulation and aerobic conditions established in the bottom zone (reddish oxidized silts and sands). The increased content of coarser siliciclastics and more polymictic than quartzitic sands are evidence of a regressive phase, but there is no sign of continental deposition. The depositional conditions are those of an inner shelf, and the input of polymictic sand and formation of conglomerate may indicate local faulting and/or uplift of the adjacent continental margin and probable close proximity to volcanic islands (clasts of acidic effusive rocks). The biostratigraphic interpretation of the successions as Lower Ordovician (Tremadoc-Arenig) was based on the occurrences of acritarchs studied by Rudavskaya, whereas invertebrate faunas (brachiopods, phyllocarids and some problematic molluscs) also of Ordovician age were described by Popov (in Preobrazhenskaya et al. 1995). The new studies of the acritarchs and shelly fossils reported here, provide additional evidence of Upper Cambrian strata in the lowermost Palaeozoic part of the succession on Kolguev Island.

Previous fossil records The basal portion of the sedimentary succession reached in the Bugrino 1 and North-Western 202 boreholes was only fragmentarily cored. Invertebrate faunas are represented by lingulid brachiopods, problematic molluscs and arthropod phyllocarids (crustaceans), whereas acritarchs are taxonomically disparate and assigned to numerous taxa. Brachiopods attributed to Angarella ex. gr. lopatini Assatkin have been reported from the Bugrino 1 borehole (depth of 3781.0m) and Lingulinella aff. davissi (McCoy) from the Bugrino 1 (4082.5 and 4154.0m) and the North-Western 202 (interval 4225.0-4492.0 m) boreholes (Preobrazhenskaya et al 1995). The identification of these taxa is revised below. The initial investigation of acritarchs in the Bugrino 1 and North-Western 202 successions, and in general on Kolguev Island, was carried out by the late Valeria Rudavskaya and reported by Preobrazhenskaya et al. (1995). The acritarch

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assemblages were rich, but they were neither illustrated nor described in this paper. They include some species that we have not found. These are Nellia longiscula Golub and Volkova 1985, Ladogiella rotundiformis Golub and Volkova 1985, Stellechinatum uncinatum (Downie 1958) Molyneux 1987, Stelliferidium aff. cortinulum (Deunff 1961) Deunff et al. 1974, Cymatiogalea cuvillierii (Deunff 1961) Deunff 1964, Acanthodiacrodium 'invictum' Rasul 1979 = Actinotodissus ubuii Martin 1969, Acanthodiacrodium comptulum Rasul 1979, and Micrhystridium shinetonense Downie 1958 (the taxonomic names and synonymous species are revised herein). The cited taxa are biostratigraphically complementary to the assemblages in the present study and the entire acritarch association recovered by Rudavskaya is taxonomically less diverse in comparison to that described here. Rudavskaya assigned the acritarch assemblages to the Upper Cambrian-Lower Tremadoc (North-Western 202, at depths of 4418.0m and 4492.0m), Tremadoc (Bugrino 1, at 3783.03859.0m, 4082.0-4150.0 m, and North-Western 202, at 4275.0m), and Arenig (Bugrino 1, at 2802.0m). She considered one species, Pirea sp., from the North-Western 202 borehole at the depth of 4225.0 m, to be diagnostic for the Upper Tremadoc Stage. However, Preobrazhenskaya et al. (1995) attributed the whole geological succession to the Lower Ordovician, neglecting Rudavskaya's evidence of Cambrian strata. The Rudavskaya's collection was not available to us for examination and new samples were studied from different depths although some came from the same core intervals (Fig. 2).

Materials and methods Organic-walled microfossils (acritarchs) were extracted from dark grey, thin-bedded mudstone samples, 50 g in weight, by the standard palynological method of chemical digestion in inorganic acids (hydrofluoric, hydrochloric and nitric), followed by controlled oxidation to remove organic debris and then filtration in a closed apparatus (see the method by Vidal 1988). The microfossils were mounted in a plastic medium (epoxy resin Epotek 301) on glass microscope slides and these permanent strewn slides were studied using a transmitted light microscope with interference contrast, complemented by observations in a high definition real-time 3D microscope, Edge R400. The collection of microfossils and all illustrated specimens are stored in the Museum of Evolution, Palaeontological Section, Uppsala University (MEU-Ru-95-1 to 13). The photomicrographs (Figs 3 & 4) were taken with oil immersion and interference contrast. The position of the specimens is marked by England Finder Co-ordinates with the microscope slide label oriented to the left. Of the thirteen processed samples, three from the Bugrino 1 borehole at the depths of 3604.8 m, 3607.3 m and 3607.8 m appeared to be barren. Ten fossiliferous samples in both successions from the stratigraphic intervals shown in Fig. 2 yielded abundant specimens in a relatively good state of preservation. The thermal alteration of organic matter building the vesicle wall of the microfossils is low and corresponds to the mesocatagenesis stage of lithogenesis, which is within the oil window. More than 500 specimens of morphologically complex and ornamented acritarchs were identified and numerous spheromorphs were observed. The macrofaunal collection of brachiopods, molluscs and phyllocarids is housed in the National Museum of Wales, Cardiff, and the illustrated specimens are assigned with the pre-fix NMGW.

Acritarchs Acritarchs in this study are from ten stratigraphic horizons in the Bugrino 1 and North-Western 202 successions. The microfossil

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Fig. 3. Late Cambrian-Tremadocian acritarchs from Kolguev Island, (a, b) Vulcanisphaera britannica Rasul 1976. (a) Specimen MEU-Ru-95-09/2. N40. (b) MEURu-95-04/1. S34-1. (c, d) Saharidia fragilis (Downie 1958) Combaz 1967. (c) Specimen MEU-Ru-95-04/1. X47. (d) MEU-Ru-95-11/1. C29. (e) Eliasum llaniscum Fombella 1977. Specimen MEU-Ru-95-04/2.M34-2. (f) Actinotodissus burmannae (Burmann 1968) Fensome et al 1990. Specimen MEU-Ru-9507/2. J39. Specimens on micrographs from the Bugrino 1 borehole, at the following depths (a) 4152.9 m, (b, c, e) 4083.8 m; and (f) 4152.3 m. Specimen (d) is from the North-Western 202 borehole, at a depth of 4274.4 m. Scale bar in (a) is equal to 7 (Jim for (a, b); 14 jjim for (c); 18 |xm for (d, f); and 12 jjim for (e).

assemblages in the North-Western 202 borehole are from the depths of 4227.0 m, 4274.4 m and 4496.8 m (Fig. 2) and they include a few biostratigraphically useful taxa. The basal portion of the succession (sample at 4496.8 m, that is about 3.3 m above the borehole bottom) yielded the species that are diagnostic for the Late Cambrian, thus defining the maximum age of the strata in the borehole and regionally on Kolguev Island. In the Bugrino 1 borehole, the fossiliferous horizons are at a depth of 4083.8 m (one sample) and within the interval of 2 m thick beds at depths of 4151.6m, 4152.0m, 4152.3m, 4152.6 m, 4152.9 m and 4153.6 m, yielding six samples (Fig. 2). The basal portion of the strata in this borehole is thought to be of latest Cambrian age (see below). The acritarch association (all taxa recorded) consists of twenty-seven species and ten additional taxa left under an open

nomenclature. One species is new, but it is based on a synonymous taxon described from China and it will be formally established (Moczydlowska & Stockfors 2003). The dominant species in the Kolguev association, both in their taxonomic diversity and abundance, are within the acanthodiacrodian-dasydiacrodian group of acritarchs and of the genera Polygonium and Vulcanisphaera, which are characteristic for the Cambrian-Ordovician transition. The taxonomic diversity of the acritarchs at each stratigraphic horizon is remarkable and the number of species may vary from a few to twenty or so, but several species occur commonly throughout the successions. Most of the species have welldefined stratigraphic ranges and are known from a wide palaeogeographic distribution, rendering them useful for recognizing of the Cambrian-Ordovician boundary. Recently, the Global Stratotype Section and Point of this boundary has been ratified in the

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Fig. 4. (a) Actinotodissus spinutisus (Timofeev 1959) Moczydlowska and Stockfors 2004. Specimen MEU-Ru-95-04/1. X27. (b) Actinotodissus achrasii (Martin 1973) Yin Leiming 1986. Specimen MEU-Ru-95-07/2. G24-2. (c, d) Solisphaeridium lucidum (Deunff 1959) Turner 1985. (c) Specimen MEU-Ru-95-07/2. N28. (d) MEU-Ru-95-07/2. N26-4. (e) Actinotodissus polimorphus (Timofeev 1959) Moczydlowska and Stockfors 2004. Specimen MEU-Ru-95-07/2. D26-4. (f) Actinotodissus secundarius (Timofeev 1959) Moczydlowska and Stockfors 2004. Specimen MEU-Ru-95-07/2. Z32-3. All specimens are from the Bugrinol borehole, at a depth of 4083.8 m (a) and 4152.3 m (b-f). Scale bar in (a) is equal to 7 jjim for (a-e), and 8 jxm for (f).

section at Green Point, western Newfoundland, and it coincides with the first appearance of the conodont lapetognathus fluctivagus Nicoll et al 1999 (Cooper & Nowlan 1999). The newly accepted Ordovician lower boundary is slightly below the base of the traditional Tremadoc stage and the Rhabdinopora flabelliformis graptolite zone in the British subdivision scheme (Cooper 1999). The acritarchs are not yet known from the stratotype section at Green Point, but are well recognized in the Cambrian-Ordovician interval in other localities in Newfoundland. Their taxonomic turnover is detected, within an acceptable degree of stratigraphic resolution and correlation, at the boundary between the Upper Cambrian Acerocare trilobite Zone and the Tremadoc Stage, in previously applied stratigraphic divisions. For a review of Cambrian-Ordovician boundary studies, see

Basset & Dean (1982), Barnes & Williams (1991), Cooper & Nowlan (1998), and Fortey (2000), and for acritarch studies in Newfoundland see Parsons & Anderson (2000). Taxonomic descriptions and detailed evaluation of the acritarchs from Kolguev Island are provided by Moczydlowska & Stockfors (2004) and only biostratigraphically significant species are considered herein. Most species in the Kolguev association have ranges overlapping the uppermost Cambrian and lowermost Ordovician, as established in the global records (Dean & Martin 1982; Martin 1982, 1992; Yin 1986; Volkova 1990, 1993; Paalits 1992, 1995; Vecoli 1996; Parsons & Anderson 2000). Such cosmopolitan species are Saharidiafragilis, Polygonium pellicidum, P. pungens, P. gracile, P. sexradiatum, Solisphaeridium akrochordum,

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S. lucidum, Actinotodissus achrasii, A. spinutisus, A. ubuii, and Vulcanisphaera africana. Some other species, including Polygonium martinae, Poikilofusa squama, Impluviculus stellaris, Dasydiacrodium tricorne and Actinotodissus formosus, have a more limited geographic distribution, but their range also embraces the Cambrian-Ordovician boundary. The species Vulcanisphaera britannica and Acanthodiacrodium angustum had been thought to appear at the base of Tremadoc (Combaz 1967; Downie 1958; Rasul 1976; Vecoli 1996; Parsons & Anderson 2000), but their ranges seem to extend into the uppermost Cambrian Acerocare trilobite Zone in the Baltica palaeocontinent (Volkova 1990; Moczydlowska & Stockfors 2004). A few species in the association are restricted to the Upper Cambrian and these are Actinotodissus polimorphus, A. secundarium, Impluviculus villosiusculus and Dasydiacrodium obsonum (Yin 1986; Martin & Dean 1988; Volkova 1990; Paalits 1992; Ribecai & Vanguestaine 1993; Vecoli 1996; Moczydlowska & Stockfors 2004). Only two species, Actinotodissus crinitus and A. burmannae are known exclusively from the lowermost Ordovician (Rasul 1979; Moczydlowska & Stockfors 2004). The vertical distribution of acritarch species in the Bugrino 1 succession and their ranges established from the global records indicate that the strata below the depth of 4152.0 m are uppermost Cambrian (Fig. 2). The acritarch assemblage previously reported by Rudavskaya (in Preobrazhenskaya et al. 1995) from the depth of 4150.0 m in this succession is convincingly Tremadocian. Based on the above evidence, the Cambrian-Ordovician boundary is recognized within the interval of 4150.0-4152.0 m in the Bugrino 1 borehole. Only a few acritarch species were recovered in the NorthWestern 202 succession in the present study, including Eliasum llaniscum, Saharidia fragilis and Polygonium sexradiatum, and their ranges span the Upper Cambrian and Tremadocian transition (Moczydlowska & Stockfors 2004). Some additional species of similar stratigraphic ranges, and in particular Nellia longiuscula Golub & Volkova 1985, which is restricted to the Upper Cambrian Peltura zone (Volkova & Golub 1985; Volkova 1990), have been

recovered by Rudavskaya at the depth of 4492.0 and 4418.0m (in Preobrazhenskaya et al 1995). Acritarch assemblages of Tremadocian age were also recorded at the depth of 4275.0 m and higher up in the North-Western 202 succession and therefore we suggest that the base of the Ordovician may be placed approximately at this horizon (Fig. 2).

Shelly fossils Among the shelly fossils preserved in the studied intervals of cores there are brachiopods, phyllocarids and a mollusc. The lingulid brachiopod Lingulella nicholsoni (Calloway) occurs throughout the succession in abundance (Figs 2 & 5). The stratigraphic range of the species has been defined in the lower Tremadoc Stage of England and Wales, where it is common, but it is also known in Scandinavia associated with different species of the graptolite Rhabdinopora (Sutton et al. 1999). However, the record from the uppermost Cambrian strata in the NorthWestern 202 borehole indicates that the range is longer. Specimens previously assigned to as Angarella ex. gr. lopatini Assatkin (Preobrazhenskaya et al. 1995) are transferred to Solandangella! sp. The external morphology and muscle scars of these fossils are more similar to Solandangella vizcainoi Horny, the taxon recorded from the upper Foulon Formation (Arenig) of the Montagne Noire, France (Horny 1995), than to the problematic mollusc Angarella from the Arenig to lower Llanvirn of Siberia. The present record suggests that the upper part of the pre-Devonian succession in the Bugrino 1 borehole is actually of Arenigian age (Fig. 2). Some fossils described from the Indysei beds (Arenig) in the Kosju River section, North Urals and referred to Angarella lopatini Assatkin by Bogoyavlenskaya (1991) are similar in their general conical shell shape to the present specimens, but they differ by being bivalved. The latter feature excludes a relationship to Solandangella.

Fig. 5. Lingulid brachiopods Lingulella nicholsoni (Calloway) from the NorthWest-202 borehole, (a) Specimen NMW 98.62.13, dorsal valve exterior, x 7, at a depth of 4274.4 m; (b) NMW 98.62.10, dorsal valve exterior, x8, depth interval 4225.04235.0 m. Scale bar in (b) is equal to 2.5 mm.

LATE CAMBRIAN ACRITARCHS, KOLGUEV ISLAND

Numerous phyllocarid crustaceans, taxonomic affiliations of which remain to be determined, occur in the upper part of the North-Western 202 succession (interval of 4225.0-4280.8 m) and at the base of the Bugrino 1 succession (4149.6-4156.0). Their possible stratigraphic significance has not yet been established. Biostratigraphic implications and conclusions The age of the lowermost known portion of the sedimentary cover overlying the Neoproterozoic basement complexes and the postTimanian unconformity on Kolguev Island is established by means of acritarchs as Late Cambrian, equivalent to the Peltura and Acerocare trilobite zones of Baltica (Moczydlowska & Stockfors 2004; Fig. 2). Currently, this age may be accepted as the only one that is palaeontologically documented in the region and thus as the minimum age of the post-Timanian unconformity. The stratigraphic hiatus between the two structural complexes, defining the unconformity, is also constrained by the age of the Neoproterozoic metasedimentary successions cut by late Vendian intrusive rocks in the basement (c. 550-560 Ma; Nikishin et al 1996; Gee et al 2000). By Late Cambrian-Tremadocian times, outer shelf marine sedimentation, evident from the successions studied here, was proceeding on the Pechora shelf. The predominantly fine-grained siliciclastic deposits and facies distribution are consistent with a stable platform along the passive margin of the Baltica craton (Nikishin et al. 1996). This margin might not have been directly connected with the epicontinental basin on the East European Platform, to the SW across the Timan Range, but the existence of a seaway along the shelves surrounding the range is highly probable because of the common occurrence of shelly faunas (brachiopods) and acritarchs. Acritarchs are planktonic biota, passively dispersed by surface currents, and their uniform taxonomic distribution indicates the lack of environmental barriers and certainly the presence of a connection between the Pechora shelf and the epicontinental basin covering the East European Craton in the present-day areas of Arkhangelsk, St Petersburg and up to the Moscow syneclise. The evidence is provided by numerous acritarch taxa recorded there in the Upper Cambrian-Tremadoc strata (Volkova 1990; Popov & Gorjansky 1994; Preobrazhenskaya et al. 1995; Moczydlowska & Stockfors 2004). Brachiopods are strong indicators of marine channels for migration and their record from the Arkhangelsk area shows such a connection with the Moscow syneclise during the Late Cambrian Peltura biochron (Popov & Gorjansky 1994). The Upper Cambrian sediments might have been widely distributed, perhaps up to the White Sea area, where similar sediments also occur, as this new record from Kolguev Island suggests. The known outliers of the Upper Cambrian sediments suggest that these strata may originally have extended across most of the post-Timanian peneplained terranes that were accreted to the margin of Baltica in the Vendian. Marine biota thrived, freely migrated and were uniformly dispersed around the margins of the Baltica palaeocontinent, including the Pechora shelf. We thank David Gee, Anna Siedlecka and Dima Grazhdankin for their thorough comments and suggestions on the manuscript. The research by M. Moczydlowska and M. Stockfors was supported by grant from the former Swedish Natural Science Research Council (NFR) to M. Moczydlowska (Project G-AA/GU 09939-319). Studies by L. Popov were funded by the Royal Society of London and the National Museum of Wales, and are a contribution to the IGCP Project 410.

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MARTIN, F. 1992. Uppermost Cambrian and Lower Ordovician acritarchs and Lower Ordovician chitinozoans from Wilcox Pass, Alberta. Geological Survey of Canada, Bulletin, 420, 57 pp. MARTIN, F. & DEAN, W. T. 1988. Middle and Upper Cambrian acritarch and trilobite zonation at Manuels River and Random Island, eastern Newfoundland. Geological Survey of Canada, Bulletin, 381, 91 pp. MARTINSSON, A. 1974. The Cambrian of Norden. In: HOLLAND, C. H. (ed.) Lower Palaeozoic rocks of the World, Vol. 2, Cambrian of the British Isles, Norden, and Spitsbergen. Wiley-Interscience, London, 185-283. MELNIKOV, S. V. 1999. Konodonty ordovika i silura TimanoSeverouralskogo regiona [The Ordovician and Silurian conodonts of the Timan-Northern Urals region.] VSEGEI, St Petersburg, 136 pp [in Russian]. MOCZYDLOWSKA, M. 1991. Acritarch biostratigraphy of the Lower Cambrian and the Precambrian-Cambrian boundary in southeastern Poland. Fossils and Strata, 29, 127 pp. MOCZYDLOWSKA, M. 1998. Cambrian acritarchs from Upper Silesia, Poland—biochronology and tectonic implications. Fossils & Strata, 46, 121 pp. MOCZYDLOWSKA, M. & STOCKFORS, M. 2004. Acritarchs from the Cambrian-Ordovician boundary interval on Kolguev Island, Arctic Russia. Palynology, 27, 110 pp (in press). NIKISHIN, A. M., ZIEGLER, P. A., STEPHENSON, R. A., et al 1996. Late Cambrian to Triassic history of the East European Craton: dynamics of sedimentary basin evolution. Tectonophysics, 268, 23-63. OLOVYANISHNIKOV, V. G. 1998. Verkhniy dokembriy Timana i poluostrova Kanin [Upper Precambrian of the Timan and Kanin Peninsula.] Russian Academy of Sciences Uralian Division, Ekaterinburg, 163 pp [in Russian]. OLOVYANISHNIKOV, V. G., BUSHUEV, A. S. & DOKHSANYANTS, E. P. 1996. The structure of the conjugation zone of the Russian and Pechora plates from geological and geophysical data. Transactions (Doklady) of the Russian Academy of Sciences, Earth Sciences Section, 351, 1228-1232. OLOVYANISHNIKOV, V. G., SIEDLECKA, A. & ROBERTS, D. 1997. Aspects of the geology of the Timans, Russia, and linkages with Varanger Peninsula, NE Norway. Norges geologiske unders0kelse, Bulletin, 433, 28-29. OLOVYANISHNIKOV, V. G., ROBERTS, D. & SIEDLECKA, A. 2000. Tectonics and sedimentation of the Meso- to Neoproterozoic Timan-Varanger Belt along the northeastern margin of Baltica. Polarforschung, 68, 267-274. PAALITS, I. 1992. Upper Cambrian acritarchs from boring core M-72 of North Estonia. Proceedings of Estonian Academy of Sciences Geological Section, 41, 29-37. PAALITS, I. 1995. Acritarchs from the Cambrian-Ordovician boundary beds at Tonismagi, Tallinn, North Estonia. Proceedings of Estonian Academy of Sciences Geological Section, 44, 87-96. PARSONS, M. G. & ANDERSON, M. M. 2000. Acritarch microfloral succession from the Late Cambrian and Ordovician (Early Tremadoc) of Random Island, Eastern Newfoundland, and its comparison to coeval microfloras, particularly those of the East European Platform.

American Association for Stratigraphic Palynologists Contribution Series, Number 38, 129 pp. POPOV, L. & GORJANSKY, V. 1994. First record of Upper Cambrian from the eastern White Sea coast: new evidence from obolids (Brachiopoda). Geologiska Foreningens i Stockholm Forhandlingar, 116,31-35. PREOBRAZHENSKAYA, V. I., USTRITSKIY, V. I. & BRO, E. G. 1995. Paleozojskie otlozheniya ostrova Kolguev (Barentsevo morie) [The Palaeozoic deposits of the Kolguev Island (Barents Sea).] Stratigrafiya, Geologicheskaya korrelatsiya, 3, 75-85 [in Russian]. PUCHKOV, V. N. & CHUVASHOV, B. I. (eds) 1988. Novye dannye po stratigrafii verkhnego proterozoiya i nezhego paleozoya zapadnogo sklona severa Urala [New data on stratigraphy of the Upper Proterozoic and Lower Palaeozoic of the western slope of the North Urals.] AN SSSR, Urals Division, Sverdlovsk, 62 pp (Unpublished report) [in Russian]. RASUL, S. M. 1976. New species of the genus Vulcanisphaera (Acritarcha) from the Tremadocian of England. Micropalaeontology, 22, 479_484. RASUL, S. M. 1979. Acritarch zonation of the Tremadocian Series of the Shineton Shales, Wrekin, Shropshire, England. Palynology, 3, 53-72. RIBECAI, C. & VANGUESTAINE, M. 1993. Latest Middle-Late Cambrian acritarchs from Belgium and northern France. Special Papers in Palaeontology, 48, 45-55. SUTTON, M. D., BASSETT, M. G. & CHERNS, L. 1999. Ungulate brachiopods from the lower Ordovician of the Anglo-Welsh Basin, Part L Monograph of the Palaeontographical Society London, 1-60. VECOLI, M. 1996. Stratigraphic significance of acritarchs in CambroOrdovician strata, hassi-Rmel area, Algerian Sahara. Bollettino della Societa Paleontologica Italiana, 35, 3-58. VIDAL, G. 1988. A palynological preparation method. Palynology, 12, 215-20. VOLKOVA, N. A. 1990. Akritarkhi srednego i verkhnego kembriya vostochnoevropejskoj platformy [Middle and Upper Cambrian acritarchs from the East European Platform.] Nauka, Moscow, 115 pp [in Russian]. VOLKOVA, N. A. 1993. Akritakhi pogranichnykh otlozheniy kembriya i ordovika proglintovoy polosy Estonii (skvazhina M-56) [Acritarchs of the Cambrian-Ordovician transitional deposits of the glint zone in Estonia (borehole M-56).] Eesti Teaduste Akadeemia Toimetised, Geoloogoa, 41, 15-22 [in Russian]. VOLKOVA, N. A. & GOLUB, I. N. 1985. Novye acritarkhi verkhnevo kembriya leningradskoy oblasti (Ladozkaya svita) [New acritarchs from the Upper Cambrian of the Leningrad district (Ladoga Formation).] Paleontologicheskiy Zhurnal, 4, 90-98 [in Russian], VOLKOVA, N. A., KIRJANOV, V. V., PISCUN, L. V., PASHKYAVICHENE, L. T. & JANKAUSKAS, T. V. 1983. Plant microfossils. In: URBANEK, A., ROZANOV, A. Y. (eds) Upper Precambrian and Cambrian palaeontology of the East-European Platform. Wydawnictwa Geologiczne, Warszawa, 5-46. YIN, L.-M. 1986. Acritarchs. In: CHEN, J.-Y. (ed.) Aspects of CambrianOrdovician boundary in Dayangcha, China. China Prospect Publishing House, Beijing, 314-373.

Northeastern and northwestern margins of Baltica in Neoproterozoic time: evidence from the Timanian and Caledonian Orogens ANNA SIEDLECKA1, D. ROBERTS1, J. P. NYSTUEN2 & V. G. OLOVYANISHNIKOV3 1 Geological Survey of Norway, N-7491 Trondheim, Norway (e-mail: [email protected]; [email protected]) 2 Department of Geology, University of Oslo, NO-0316 Oslo, Norway (e-mail: j.p. nystuen @ geologi. uio. no) 3 Institute of Geology, Komi Research Centre, Ural Division, Russian Academy of Sciences, Pervomaiskaya 54, Syktyvkar, 167000 Russia (e-mail: [email protected])

Abstract: The Neoproterozoic depositional histories of the Timanian and Baltoscandian, orthogonal margins of Baltica show several important differences but also some similarities. The Timanian margin comprises mainly low-grade, terrigenous sedimentary successions with a distinctive, margin-parallel fault zone separating pericratonic and basinal domains. Magmatic rocks are comparatively rare on land, but are common in deep drillcores recovered from beneath the Pechora Basin. Conversion to an active margin occurred in latest Riphean time, ultimately leading to the accretionary and transpressional regime of the Vendian-age, Timanian Orogeny. Along the Baltoscandian margin, successions of low to high metamorphic grade are preserved in diverse Caledonian nappe complexes. Three main types of palaeobasin are distinguished, based largely on sedimentary facies and basin geometry. Magmatic rocks are more common than in the Timanides, ranging from mafic dyke swarms to the voluminous Seiland Igneous Province. This margin remained passive throughout the Neoproterozoic era. The Vendian-dated dyke swarms signify the onset of lapetus/^gir ocean opening at precisely the time when the orthogonal Timan margin was being deformed and telescoped during the Timanian Orogeny.

The palaeocontinent Baltica is generally considered to have been derived from the break-up of the supercontinent Rodinia (e.g. Dalziel 1997; Meert & Powell 2001). Heralding a long period of continental crustal extension in the Rodinian domain, rupture and rifting began in the Mid to Late Neoproterozoic, earlier in some regions than in others (Harland & Gayer 1972; Kumpulainen & Nystuen 1985). As the outline of Baltica gradually took form, its separate, rift-basinal margins developed in different ways. (Present-day geography is referred to throughout this paper.) Along the southern margin of Baltica, the foundations of a proto-Tethys seaway were delineated. The western, Baltoscandian margin of Baltica, on the other hand, is characterized by disparate, rift-basinal successions that ultimately gave way to a developing oceanic realm—the JEgu Sea (Torsvik & Rehnstrom 2001) that developed through time into the lapetus Ocean (Roberts & Gale 1978). Along its northeastern, Timanian margin, pericratonic and basinal sedimentation regimes, also typifying a passive margin, were succeeded by a Palaeo-Asian ocean (Olovyanishnikov et al. 1995). Each of these margins was subsequently converted from passive to active and affected by major, compressional to transpressional orogenic deformations; Early Caledonian in the case of the western and southern margins (Ramsay et al. 1985; Dallmeyer & Gee 1986), and Timanian along the northeastern periphery of Baltica (Roberts & Siedlecka 1999, 2002; Olovyanishnikov et a 2000; Scarrow et al 2001). Most of the Late Neoproterozoic Timanian Orogen occurs beneath the Palaeozoic and younger cover of the Pechora Basin (Fig. 1) (Olovyanishnikov et al 1995; Gee et al 2000). Its southwestern marginal part, in the Timan Range, extends from the junction with the Urals northwestwards to the Kanin Peninsula. It then continues offshore just off the Kola Peninsula to the onshore occurrences on Kildin Island, the Rybachi and Sredni Peninsulas and the Varanger Peninsula in northern Norway (Fig. 1). The northwesternmost exposed part of this 1800km long Timan-Varanger Belt (TVB) terminates beneath the northeastern tip of the Norwegian Caledonides (Roberts 1996). In the region of interference between the two orogens in northern Norway, several features of the Neoproterozoic basinal development show similarities with the main part of the Timanian Basin in the Timans. Some features, however, are

different and have more in common with the western Baltoscandian Basins and the Neoproterozoic occurrences in East Greenland and Spitsbergen. In this paper we will focus firstly on the development of the Neoproterozoic Timanian margin successions, pointing out the similarities and differences observed along and across strike based mainly on our own research but also on data from the Russian literature. We will then describe the development of the Neoproterozoic Baltoscandian basinal successions and their history in latest Neoproterozoic and Early Palaeozoic time. Sedimentary successions occurring in extensional basins in more central locations on Baltica, such as the Neoproterozoic Vattern Basin in southern Sweden (Andreasson 1994) and basins in the region of the Gulf of Bothnia are outside the scope of this paper.

The Timanian margin of Baltica Northern Norway and NW Russia: Varanger Peninsula, Rybachi-Sredni Peninsula and Kildin Island The geology of Varanger Peninsula is a suitable starting point for understanding the Neoproterozoic development of the northern margin of Baltica (Figs 1 & 2). A major, NW-SE trending fault zone, the Trollfjorden-Komagelva Fault Zone (TKFZ), is one of the most striking features of the geology of Varanger Peninsula (Fig. 2) and has played an important role in the Neoproterozoic and later history of this region. The fault zone was active at several times in pre- and post-Caledonian time and is still active today. Geological mapping, and stratigraphic and sedimentological studies have shown that the Neoproterozoic successions occurring to the NE and SW of the fault zone are time-correlative but exhibit considerable differences both in thickness and in sedimentary facies development (e.g. Siedlecka 1985). To the NE, in the Barents Sea Region (Fig. 2), there are two major successions, c. 9000m and c. 6000m thick, respectively, separated from each other by a low-angle angular unconformity. The nature of their substratum is unknown. The 9000 m thick lower succession (the Barents Sea Group) commences with submarine-fan turbidites

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 169-190. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. The northeastern (Timanian) and northwestern (Caledonian) margins of Baltica. NE margin', VP, Varanger Peninsula; RP, Rybachi Peninsula; SP, Sredni Peninsula; KI, Kildin Island; KP, Kanin Peninsula; NT, Northern Timan; VR, Vymskaya Ridge; ETF, East Timan Fault; CTF, Central Timan Fault; WTF, Western Timan Fault. Exposed areas of Neoproterozoic successions along the Timanian margin are indicated by a diagonal-lined pattern. Westward continuation of the Neoproterozoic successions into the Laksefjorden (L), Porsangerfjorden (P) and possibly even into the Kvaenangsfjorden (KV) area are indicated by the dotted ornament. White Sea coast: TC, Tury Cape; CR, Chapoma River; A, Arkhangelsk area. Other occurrences of Proterozoic rocks along the Kola coast are indicated by dots. NW margin; H, Hedmark; V, Valdres; H, H0yvik; He, Heidal; S, Sarv; T, Tossasfjallet. Other occurrences are indicated by the dotted ornament. Based mainly on Roberts & Gee (1985), Kumpulainen & Nystuen (1985) and Olovyanishnikov et al (2000).

Fig. 2. Simplified geological sketch map of the Varanger, Rybachi and Sredni Peninsulas and Kildin Island, NE Norway and NW Russia, showing the main divisions to the NE (basinal domain, dotted ornament) and SW (pericratonic domain, grey) of the Trollfjorden-Komagelva Fault Zone. Crosses, Archaean and Palaeoproterozoic crystalline rocks: oblique-line ornament, Tanahorn Nappe.

(Fig. 3b,c), grading upwards into prodelta and delta front deposits and terminating with coastal and fluvial beds (Fig. 3a,d) (Siedlecka & Edwards 1980; Pickering 1981, 1982; Siedlecka et al. 1989). The succession is predominantly terrigenous with subordinate carbonates only in its upper part. Acritarchs in the upper part of the succession (the Batsfjord Formation) testify to a Late Riphean age (Vidal & Siedlecka 1983). The upper, entirely

terrigenous succession (the L0kvikfjellet Group) is c. 6000 m thick and consists predominantly of shallow-marine deposits with subordinate fluvial accumulations. At most, about half of the lower succession has been removed by erosion at the unconformity surface. There is no fossil record in the upper succession and its possible Vendian age has been suggested on the basis of various comparisons and correlations. There are no glacial deposits in this upper succession. The Neoproterozoic succession preserved SW of the TKFZ, in the Tanafjorden-Varangerfjorden Region (Fig. 2), is only c. 3000 m thick and rests in its southern part (south of the Varangerfjord) with erosional contact on the older Precambrian metamorphic basement. It is predominantly terrigenous and contains several gaps, identified by mapping or by biostratigraphic work on microfossils (Siedlecki 1980; Vidal 1981). There is one, major, low-angle angular unconformity between the Upper Riphean and the Vendian parts of the succession. The Upper Riphean part is predominantly fluvial in its lower half (Fig. 4d) and shallow marine higher up, and terminates with a stromatolitebearing dolomite (Fig. 4e,f). The upper, Vendian succession is entirely terrigenous, starts with Varangerian glacial deposits (Fig. 4a-c) (two tillites are present), passes upwards into terminal Proterozoic strata and terminates with fossiliferous Lower Palaeozoic beds (Cambrian to Tremadoc) just west of Tanafjorden, on the Digermulen Peninsula (Edwards 1984; Reading 1965). At most, and progressively to the south, about 2000 m of the lower Neoproterozoic succession has been removed by erosion prior to

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Fig. 3. Lithologies of the Barents Sea Region, Varanger Peninsula, (a) Interbedded mudstone, siltstone and carbonate rocks with non-columnar stromatolites, lower part of the Batsfjord Formation, Barents Sea Group; Veinesodden, inner Syltefjorden. (Photo: A. Siedlecka.) (b) Turbidites of the Kongsfjord Formation, lower Barents Sea Group, with synsedimentary deformation structures; coastal section, Kongsfjorden. (Photo: A. Siedlecka.) (c) Thin-bedded turbidites of the Kongsfjord Formation, coastal section, Kongsfjorden; normal younging, from bottom to top. (Photo: A. Siedlecka.) (d) Sandstones and mudstones of the Batsfjord Formation, and overlying Tyvjofjell Formation, outer part of Batsfjorden. The succession is here intruded by vertical dolerite dykes. (Photo: A. Siedlecka.)

the Early Vendian (Fig. 5). The chronostratigraphy of the succession SW of the TKFZ is fairly well documented by acritarchs, stromatolites, Ediacara and, in its Palaeozoic part, by macrofossils (Vidal 1981; Bertrand-Sarfati & Siedlecka 1980; Raaben et al 1995; Farmer et al 1992). Rb-Sr isotopic work on illite subfractions suggests that the Varanger Ice Age, at least in this type area, started after 630 Ma (Gorokhov et al. 2001). This same study also documented the Cryogenian (Late Riphean) age of the sub-glacial succession. The contrasting development of the Neoproterozoic record NE and SW of the TKFZ led to the conclusion that the TKFZ originated in pre-Neoproterozoic time as a fault bordering an extensional sedimentary basin that developed along the northern/northeastern margin of Baltica. Thick basinal successions accumulated in the hanging wall while slightly later and in part contemporaneously, pericratonic sediments were deposited in the coastal and shelf areas SW of the faulted margin (e.g. Siedlecka 1975, 1985). The precise stratigraphic relationship between the basinal and the pericratonic realms was uncertain until Rice (1994) described an unconformity between the lower-middle part of the basinal succession (below) and the middle part of the pericratonic succession from an area close to the TKFZ in the western part of Varanger Peninsula. The southeastward continuation of the TKFZ has been identified on land on the isthmus between the Rybachi and Sredni Peninsulas and by geophysical methods north of Kildin Island (Fig. 2). Here termed the Sredni-Rybachi Fault Zone (SRFZ), it separates

basinal and pericratonic successions, as on the Varanger Peninsula. Only the lower, c. 4000 m thick part of the lower basinal succession is preserved on Rybachi Peninsula. It comprises an olistostrome (Fig. 6c), submarine-fan turbidites (Fig. 6a,b) and prodelta deposits and, as on Varanger Peninsula, the nature of the substratum is unknown (Siedlecka et al. 1995a). On the basis of the character of the olistoliths, however, it is inferred that the older Precambrian basement was exposed and eroded at the time of deposition of the olistostrome and turbidites. The pericratonic accumulations are preserved mostly on Sredni Peninsula and Kildin Island, and include only the Upper Riphean part of the record. The c. 2000 m thick succession in these areas is composed mainly of fluvial, coastal and shallow-marine deposits (Fig. 6d) with minor occurrences of stromatolitic carbonate rocks (Fig. 5) (Krylov & Lyubtsov 1976; Lyubtsov et al. 1989, 2000). The biostratigraphy of these strata is established on the basis of microfossils and stromatolites. There are considerable differences in lithology between the SW Varanger and the Sredni-Kildin successions; this might be expected in pericratonic areas, and has been explained by differences in composition and topographies of the hinterland (Siedlecka et al. 1995£). On Sredni, the pericratonic succession starts with fluvial beds that rest with erosional contact on Archaean crystalline rocks. The contact is not exposed on Kildin Island. Evidence judged to be important in the interpretation of the sedimentary basins development in Neoproterozoic time in northern Norway and NW Russia is listed in Fig. 5.

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Fig. 4. Lithologies of the Tanafjorden-Varangerfjorden Region, Varanger Peninsula, (a) Contact between sandstones of the Veinesbotn Formation (Vads0 Group) below and tillite of the Smalfjord Formation (Vestertana Group), Kvalneset, SW Varangerfjorden; looking north. (Photo: D. Roberts.) (b) Sharp erosional contact between sandstone of the Veinesbotn Formation and tillite of the Smalfjord Formation, Bigganjar'ga (Oaibaccanjar'ga), Varangerfjorden; looking approximately west. Note the glacially scoured top surface of the sandstone. (Photo: D. Roberts.) (c) Close-up of part of the Smalfjord tillite, Bigganjar'ga locality. (Photo: D. Roberts.) (d) Sandstone and intercalated mudstone of the Fugleberget Formation (Vads0 Group) showing cross bedding with overturned foresets in the upper part of a bed. Southern coastal section of Vads0 Island, Varangerfjorden. (Photo: D. Roberts.) (e) Part of the eastern coastal section of Tanafjorden showing upper parts of the Tanafjorden Group to the right, the dolomite of the Grasdalen Formation in the centre, and Vestertana Group to the left. (Photo: A. Siedlecka.) (f) Columnar stromatolites (Linella acaniella) of Unit F of the Porsanger Formation, Tanafjorden Group, Porsangerfjorden. (Photo: A. Siedlecka.)

Westward extension of the Neoproterozoic successions and sedimentary basins The Neoproterozoic deposits of the pericratonic domain of the Varanger Peninsula continue west of Tanafjorden to the inner Porsangerfjorden area, a distance of c. 200km (unrestored) within the Lower Allochthon (Gaissa Nappe Complex) of the Norwegian Caledonides. They appear to thin westwards and exhibit minor lateral sedimentary facies changes. If one accepts

suggested correlations farther to the west, based on lithological grounds, the pericratonic succession may continue as far west as the Kvaenangsfjorden area, an additional c. 200 km along strike (e.g. Reusch 1900; Gayer & Rice 1989) (Fig. 1). This inferred correlation was the reason for suggesting that a Neoproterozoic basin, the Gaissa Basin, extended westwards from the Varanger Peninsula, accumulating fairly shallow-water, fluvial and marine strata. This Gaissa Basin was separated from the basinal domain by a NE-SW-trending, basement-cored Finnmark Ridge

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

C

Late Vendian to Early Cambrian, c. 590-520 Ma

Northwestern Baltoscandian Basins:

Varanger Peninsula:

Rybachi-Sredni Pen.:

Kanin P. and Timan:

Stable tectonic conditions: Continuous Upper Vendian Lower Cambrian deposition

Stable tectonic conditions: Continuous Upper Vendian Lower Cambrian deposition

No sedimentary record

Stable tectonic Conditions ? Early Cambrian transgression (terrigenous Middle Cambrian deposits)

Glacioeustatic sea-level rise, Extension, subsidence, some block-faulting, intrusions of mafic dykes Crustal depression: Widespread terrigenous fluvial and shallow -marine deposits

Basinal realm: Trace of Timanian Orogeny Transgression, mafic dykes Mostly shallowmarine deposits

Compression, Timanian Orogeny No sedimentary record

Compression, Timanian Orogeny, Erosion Major angular unconformity,

Pericratonic realm: Subsidence transgression Fluvial to shallowmarine deposits Ediacara

B

Early Vendian, Varanger Ice Age c. 630-c. 590 Ma

NW Baltoscandian Basins:

Varanger Peninsula:

Rybachi-Sredni Pen.:

Kanin Peninsula and Timan:

Glaciation Upper Varangerian Tillite

Basinal realm: Shallow-marine deposits Pericratonic realm: Glaciation Upper Varangerian Tillite Post-glacial fluvial deposits

No record

? No sedimentation ? Erosion No evidence of glaciation

? No sedimentation

Basinal realm: Shallow -marine deposits Pericratonic realm: Terrigenous deposits, subordinate carbonates

No record

? No sedimentation ? Erosion No evidence of glaciation

No evidence of glaciation

Basinal realm: Shallow -marine deposits Pericratonic realm: Lower Varangerian Tillite

No record

? No sedimentation Erosion No evidence of glaciation

A

?Mid to Late Riphean, c. 1000-c. 630 Ma

Northwestern Baltoscandian Basins:

Varanger Pen. & Rybachi-Sredni Pen.:

Kanin Peninsula and Timan:

Uplift and erosion, unconformity

Uplift and erosion, low-angle angular unconformity

Erosion ? Continuous sedimentation

Decreasing relief and subsidence, Terrigenous deposits and carbonates

Decreasing relief and subsidence, Basinal realm: Shallow terrigenous deposition. Pericratonic realm: Shallow terrigenous deposits with intermittent erosion, stromatolitic carbonates

Decreasing relief and subsidence, Basinal realm: starvation Pericratonic realm: shallow terrigenous deposits with intermittent erosion, Extensive stromatolitic carbonates

Extension, rifting, several grabens Thick fluvial conglomerates and sandstones, proximal to distal turbidites

Extension, origin of the TKFZ and SRFZ forming the SW margin of the Timanian Basin Basinal realm: Olistostrome, turbidites -prodelta muds Pericratonic realm: Delta, shallow -marine

Extension, origin of WTF, CTF, ETF, forming the SW margin of the Timanian Basin Basinal realm: Turbidites-to-prodelta and associated deposits Pericratonic realm: erosion

Fig. 5. Neoproterozoic tectonics of the northeastern, northern and western margins of Baltica as reflected by the sedimentary record.

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Fig. 6. Lithologies of the Rybachi and Sredni Peninsulas, (a) Parallel-laminated, thinly bedded, silty sandstone and mudstone/shale in synsedimentary, slide-scar contact with slightly thicker bedded turbidites below. Tsypnavolokskaya Formation, NE Rybachi. (Photo: A. Siedlecka.) (b) Turbiditic sandstones and shales of the Zubovskaya Formation (upper Bargoutnaya Group), northeastern coast of Rybachi. (Photo: A. Siedlecka.) (c) Coarse, unsorted conglomerate with both angular and rounded, extrabasinal clasts, Motovskaya Formation (Einovskaya group), southern Rybachi Peninsula. Note the sharp contact between two beds. (Photo: A. Siedlecka.) (d) Sandstones of the Kuyakanskaya Formation showing herring-bone cross bedding, Sredni Peninsula. (Photo: A. Siedlecka.)

(Gayer et al. 1987). However, the more recent discovery of the stratigraphic relationship between the basinal and pericratonic successions mentioned above (Rice 1994) testifies to an interconnection between the basinal and pericratonic domains in this northeastern area, at least in Late Riphean time. The lower succession of the basinal realm (9000m thick), however, has not been identified in the lithostratigraphic successions of the Caledonian nappes. The upper L0kvikfjellet Group succession (6000 m thick) was compared on lithological grounds with the Laksefjorden Group in the Laksefjord Nappe Complex (F0yn 1969; Laird 1972) (Fig. 1). The lowermost part of the Laksefjorden Group contains a diamictite that has been interpreted as a tillite (Laird 1972), and this interpretation, combined with the angular unconformities identified on Varanger Peninsula on both sides of the TKFZ, was the reason for suggesting a Vendian age for the upper succession of the basinal realm and correlating it across the fault zone with the Varangerian and post-Varangerian deposits. Since then, however, the suggested glacial origin of the diamictites in the Laksefjorden Group has been rejected in favour of a debris flow alluvial deposit (Chapman 1980; F0yn et al 1983). Nevertheless, the upper basinal, L0kvikfjellet Group succession continues to be correlated with the Neoproterozoic tillite-bearing succession across the TKFZ, even though lithostratigraphically it is quite different and contains no positively identified tillites. This correlation, along with the palaeomagnetic data (Kj0de et al. 1978; Torsvik et al. 1995), was explained by a dextral strike-slip translation of the basinal succession along the fault zone separating the two sedimentation domains, probably in Early Palaeozoic time.

Summing up, the important evidence for discussion and interpretation of the likely westward extension and palaeogeography of the Neoproterozoic basins include: (1) correlation of the pericratonic successions with successions occurring 200-400 km SW of Varanger Peninsula; and (2) the absence of the lower basinal succession in the Caledonides west of Tanafjorden.

Coastal Kola Peninsula and the White Sea Along the northeastern coastal areas of the Kola Peninsula, c. 200-400 km SE of Kildin Island, there are several isolated occurrences of ?Middle Riphean conglomerates, sandstones and siltstones resting unconformably on older Precambrian crystalline rocks (Koistinen et al. 2001). The area is poorly accessible but based on published descriptions, the rocks probably represent what we here designate as pericratonic deposits. These rocks were correlated with the Terskaya and Turyinskaya Formations that occur on the northern side of the White Sea (Lyubtsov et al. 1988,2000). Two of us (DR & AS) examined some of these occurrences briefly in 1992. Polymict conglomerates and feldspathic sandstones with large-scale cross-bedding occur at Tury Cape, close to Umba (TC in Fig. 1). In addition, rippled mudstones with desiccation cracks and salt pseudomorphs crop out at Cape Korablovyi. In the section along the Chapoma river, farther to the east, there are mostly purple or grey, rippled mudstones with desiccation cracks, subordinate lenses of coarse-grained arkose

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

and thin carbonate beds with bulbous stromatolites. In addition, Lyubtsov et al. (1988, 2000) have recorded numerous, isolated occurrences of Neoproterozoic deposits along the southwestern and southeastern coastline of the Kola Peninsula at the mouths of several rivers, here schematically indicated in Fig. 1. These deposits comprise alluvial, coastal marine and possibly lacustrine accumulations and probably represent the remains of a once extensive, pericratonic sedimentary cover. On the southern side of the White Sea there are Upper Vendian deposits 1.5km thick that are famous for their abundant and diversified trace fossils, non-skeletal fossils and microfossils (e.g. Fedonkin 1981; Sokolov & Fedonkin 1990). The basal formation of this succession either unconformably overlies Riphean sedimentary rocks or rests directly on the older Precambrian crystalline basement. The Vendian succession is terrigenous and has been interpreted as a shallow-water accumulation (Sokolov & Fedonkin 1990). Riphean faulting, in an extensional regime, resulted in a horst and graben topography which is reflected in differences of thickness of the Vendian succession (Stankovsky et al. 1981). A tuff horizon in approximately the middle part of the Vendian succession (Ust Pinega Formation) has been dated at 555 + 0.3 Ma (Martin et al 2000). No glacigene deposits have been reported. There are no indications that the northern and southern occurrences of the Neoproterozoic rocks of the Kola Peninsula have ever been interconnected, although the Vendian transgression may have come from the north and covered the Kola Peninsula (Sokolov & Fedonkin 1990). We therefore assume that all the above-mentioned, shallow-water, Neoproterozoic sedimentary rocks of the coastal areas of the White Sea could have represented accumulations of either a shallow embayment connected to the Timanian Basin or of a separate shallow basin. Recently, Grazhdankin (2004) provided more documentation on the Neoproterozoic successions of the White Sea area and proposed a new interpretation of the palaeotectonic setting of the Vendian accumulations of this area and of the Mezen and Moscow Basins. In this model, the Vendian successions are seen as representing the infilling of a foreland basin that developed ahead of the evolving Timanian Orogen.

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The Timans and Kanin Peninsula A distance of c. 500 km separates the well documented, northern Kola, Neoproterozoic successions described above and those exposed in the Timans and on the Kanin Peninsula (Figs 1 & 7). Several, extensive, NW-SE oriented fault zones are present in these areas and there are geophysical indications that at least one of them, the Central Timan Fault, continues northwestwards just off the Kola Peninsula and then extends farther to the NW as the SRFZ and TKFZ (Figs 1 & 2) (Olovyanishnikov et al 2000). Most of the faults in the Timans and Kanin are identified primarily by geophysics. Surface exposure of Neoproterozoic strata is poor and the area is not easily accessible. There are data from drillcores, however, that supplement the restricted surface observations. Neoproterozoic successions exposed on both sides of the Central Timan Fault (CTF) compare readily with the basinal and pericratonic domains identified in northwestern areas of the Timan-Varanger Belt. This points to the existence of a large, elongate Timanian Basin, the southwestern extensional faulted margin of which can be traced from Varanger Peninsula in the NW to the junction with the Urals in the SE. In addition, a major, laterally extensive, stromatolite-bearing dolomite unit occurs between the pericratonic and the basinal successions (Figs 5 & 8). Both the dolomite on Kanin Peninsula (AS, DR, 1992, unpubl.) and the siliciclastic successions of Kanin and northern Timan (DR, VO, 2000 unpubl.) have been studied. The basinal succession of the Vymsky Ridge in Central Timan was studied in 1995 (Siedlecka & Roberts 1995). These comparatively recent observations, along with the earlier work of Getsen (1975, 1987) and Olovyanishnikov (1998), constitute the basis of our current understanding and modelling of the development of the Timanian Basin (see e.g. Olovyanishnikov et al 2000; Roberts & Siedlecka 2002). The successions exposed on Kanin Peninsula and in northern Timan (Fig. 7) are predominantly siliciclastic, consist in part of turbidites and are several thousand metres in thickness; their substratum is unknown. The Mikulkinskaya Group, exposed in the core of an anticline in SE Kanin, is considered to be the

Fig. 7. Geological maps of (A) Kanin Peninsula and (B) Northern Timan showing the exposed Neoproterozoic successions. (Modified from Roberts et al. 2004.)

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A. SIEDLECKA ETAL

Fig. 8. Tentative model for the development of the Timanian Basin; modified from Roberts & Siedlecka (2002). (A) Initial rifting, erosion of footwall shoulders and accumulation of slope-to-basin deposits. (B) Shallow-marine, pericratonic sedimentation and build up of stromatolitic, reefal, shelf-margin carbonate rocks. (C) Timanian deformation arising from SW-directed compression (open arrow), with resultant inversion of the basin-marginal faults. CTF, Central Timan Fault. Time intervals-Mid (R2), Late (R3) and Terminal (R4) Riphean. V-C, Vendian to Cambrian.

oldest unit and is metamorphosed in amphibolite facies, whereas the remaining (younger) successions of both Kanin and northern Timan have undergone greenschist-facies parageneses. Recent studies have shown that the oldest, Mikulkinskaya, succession was buried to a depth of c. 25 km to 28 km (Lorenz et al 2004). Looking at the overall picture of the Timans, the Kanin-northern Timan segment belongs to a deeper part of the basinal domain and is not bordered by the Central Timan Fault. The succession on Kanin continues southwestwards in the subsurface and meets the dolomite unit along a fault contact (Olovyanishnikov et al. 2000). The thickness of the succession in northern Timan is over 5000 m and on the Kanin Peninsula more than 10 000 m (Getsen 1975, 1987). The successions are not continuous, and have been described as separate formations with uncertain stratigraphic interrelationships, in particular in northern Timan. The successions of the Kanin and northern Timan differ in sedimentological development; the northern Timan succession consists entirely of turbidites (Fig. 9a-e) whereas the Kanin succession is composed only partly of turbiditic lithologies (see Roberts et al. 2004). In summary, the successions of Kanin and Timan are several kilometres in thickness and represent a basinal, partly turbiditic deposit, proximal to the faulted slope of, an extensional basin. Both isotopic dates and microfossils give no conclusive evidence of age except that: (1) the successions are Neoproterozoic; (2) the dolomite (Fig. 9f) is Upper Riphean (Raaben et al. 1995); and (3) the Vendian is

probably not represented in these particular lithostratigraphical successions (Fig. 5). It has also been suggested (Roberts & Siedlecka 2002) that transverse faults or accommodation zones may have resulted in some form of segmentation of the elongate basin, with fault-block rotations, which would account for unusually steep dips in some areas. River sections across the Vymskaya Ridge in Central Timan (Fig. 10) expose a succession totalling c. 6500 m in thickness. In spite of the presence of minor folds and later faults, this estimated thickness is probably close to the true stratigraphic thickness. Neither the bottom nor the top of this succession is exposed. Geophysical data indicate the presence of tectonic contacts against adjacent units (Getsen 1987; Olovyanishnikov 1998). The succession is monotonous, dominated by fine-grained sediments and is terrigenous with a few examples of diagenetic carbonates. Just one unconformity has been recorded, overlain by somewhat coarser, sandy, sedimentary rocks; the stratigraphic range of the hiatus is unknown. The unconformity is interpreted to represent an uplift and erosional event followed by a new transgression. The pre-unconformity section starts with turbidites followed by mud-clay sediments deposited by dilute turbidity currents and from hemipelagic suspensions. The succession has been tentatively interpreted as representing submarine-fan and prodelta basinal sediments. The environment of accumulation of the upper part of the section is uncertain (Siedlecka & Roberts 1995; Roberts et al. 2004). Although a Vendian age has been suggested for a part of the succession (Olovyanishnikov 1998), the microfossil content is not conclusive (Olovyanishnikov et al. 2000). Parts of the pericratonic succession of the Central Timan are discontinuously exposed in the Chetlasky Kamen (Fig. 1). The succession is terrigenous and 2000-3000 m in thickness (Getsen 1987; Olovyanishnikov 1998) and is tentatively interpreted as a shallow-marine to coastal deposit (Siedlecka & Roberts 1995). A terminal Riphean age is suggested by microfossils (Getsen & Pykhova 1977). Sections at Cylmensky Kamen in the Central Timan (Fig. 1) expose both Upper Riphean stromatolitic dolomite and fine-grained terrigenous deposits comparable to the basinal succession of the Vymskaya Ridge (Getsen & Pykhova 1977). The contacts are tectonic and the stratigraphic relationships uncertain (Getsen 1987, 1991; Dedeyev & Getsen 1987). Dedeyev & Getsen (1987) suggested a Vendian age for parts of the pericratonic successions of the Central Timans. Summing up the Neoproterozoic geology of the Timans, it appears that the contacts between the pericratonic succession, the dolomite and the basinal succession are tectonic and mostly detected in the subsurface by geophysical data. The immediate substratum remains unknown. Although the biostratigraphic evidence is restricted, it may be tentatively concluded that the terrigenous successions are mostly Neoproterozoic, probably in part even late Mesoproterozoic. The dolomite comprises stromatolites of Late Riphean (Cryogenian) age. The occurrence of strata of definite Vendian age is uncertain. Although their presence has been assumed, based on a reassessment of the microfossil content (Olovyanishnikov 1998), we believe that it is not conclusive enough for suggesting a Vendian age for the uppermost parts of the succession of the Central Timan. The Neoproterozoic strata marginal to the northeastern periphery of Baltica, and exposed in the Timans and on Kanin Peninsula, were affected by SW-vergent folds and thrusts during the Timanian Orogeny (Roberts & Siedlecka 1999, 2002; Olovyanishnikov et al. 2000; Roberts et al 2004). The Mid to Late Vendian age of the orogeny derives from a variety of isotopic ages obtained both from rocks exposed at the surface and from deep drillcores in the Pechora Basin (Malkov 1992; Gee et al. 2000). Moreover, there is a major angular unconformity at the Neoproterozoic-Palaeozoic interface. In northern Timan and on Kanin, Silurian limestones locally define and overlie this unconformity, but the oldest

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

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Fig. 9. Lithologies of northern Timan and Kanin Peninsula, (a) Sandy and muddy, turbiditic graded-bedded unit of the Rumyanichnaya Formation, coastal section, N. Timan. The lenticular, starved-ripple beds are here partly modified by open folding; north to the left. (Photo: D. Roberts.) (b) Turbiditic, silty sandstone and pelite with climbing ripples clearly defined in the paler sandier units, Rumyanichnaya Formation, coastal section, N. Timan. Top is to the NNW; palaeocurrent flow towards ENE. (Photo: D. Roberts.) (c) Interbedded silty sandstone and pelite, with a prominent, early-diagenetic, carbonate-rich layer just below centre, Rumyanichnaya Formation; coastal section, N. Timan, looking NE. In higher-grade rocks on Kanin Peninsula, such concretionary lenses are called 'scarnoids'. (Photo: D. Roberts.) (d) Thin-bedded and laminated, silty and pelitic turbidites of the Rumyanichnaya Formation, with a well-developed spaced cleavage (axial planar to local folds); looking approximately north. (Photo: D. Roberts.) (e) Laminated turbidites of the uppermost part of the Rumyanichnaya Formation, with evidence of synsedimentary sliding of semi-consolidated packages of multilayered sediment. Chorny River, N. Timan; inverted sequence. (Photo: D. Roberts.) (f) Columnar stromatolites (Gymnosolenides) from the Lyudovatovskaya Formation, Cape Lyudovatyi, near Shoina, SW Kanin Peninsula. (Photo: A. Siedlecka.)

reported, exposed Palaeozoic beds are of Cambrian age from the northeastern slope of the southern Timans (Getsen 1987). However, the oldest, biostratigraphically well documented, Middle Cambrian-Ordovician strata have recently been reported from drillcores on Kolguev Island (Moczydlowska-Vidal & Popov 2001; Moczydlowska et al 2004).

In summary, some of the most important evidence relevant to any interpretation of the basinal and tectonic evolution of the Kanin-Timan region is: (1) The presence of basinal and pericratonic deposits throughout the Timan Range;

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REGIONS Main Tectonostratigraphic Divisions Upper Allochthon (lower part only)

NORTHERN NORWAY

SWEDEN

S0R0Y-SEILAND NAPPE (top of Kalak Nappe Cpx)

SOUTHWESTERN NORWAY

SOUTHERN & SOUTHEASTERN NORWAY & SWEDEN

SEVE NAPPES

KALAK & SEVE NAPPE CPXS.

SARV NAPPES (SjETRA, LEKSDAL, REMSKLEPP)

Fig. 10. Geological sketch map of part of Vymskaya Ridge, Central Timan (based on Olovyanishnikov 1995).

(2) the occurrence of an Upper Riphean stromatolitic dolomite along the footwall shoulder of the fault contact between the pericratonic and basinal successions; (3) the absence of glacial deposits; and (4) the regional, angular unconformity at the Neoproterozoic Palaeozoic interface.

The Baltoscandian margin of Baltica

Middle Allochthon

Lower Allochthon

Parautochthon & Autochthon

LAKSEFJORD NAPPE CPX.

GAISSA NAPPE CPX. BLAIK, RAUTAS NAPPES

JOTUN-VALDRES NAPPE COMPLEX

JOTUN-VALDRES NAPPE COMPLEX

KVITVOLA NAPPE COMPLEX

HARDANGERRYFYLKE NAPPE COMPLEX

OSEN-R0A NAPPE CPX.

SYNNFJELL, AURDAL NAPPES

VARIOUS MINOR THRUST SHEETS

The western Baltoscandian basins The northwestern margin of Neoproterozoic Baltica, also termed the Baltoscandian margin, was strongly disrupted and imbricated during the compressional and extensional stages of the protracted Caledonian orogenic cycle (Fig. 1). Consequently, the original character and geometry of the Baltoscandian margin has to be restored from a series of polyphasally deformed thrust sheets now present in different parts of the Scandinavian Caledonides. Remnants of the margin are identified as Neoproterozoic rock successions and igneous rocks in a number of nappe complexes within the Lower, Middle and Upper Allochthons (Fig. 11), that can be traced from the southwestern parts of Norway to western Finnmark in the far north (Kumpulainen & Nystuen 1985; Nystuen & Siedlecka 1988; Andreasson 1994; Andersen et al. 1998; Andreasson et al. 1998) (Fig. 1). Several of these nappe units consist of Neoproterozoic to Cambrian successions up to 3-4 km thick that have been displaced cratonward from their original sites of deposition over distances of at least 150km, and probably as much as 300-400 km (Gee 1975, 1978; Nystuen 1981; Morley 1986; Gayer et al. 1987). Most of these allochthonous sedimentary successions include attached slices of their original Baltican, crystalline basement, consisting of granites, diorites and mafic rocks with Proterozoic ages of c. 1450-1700 Ma. The sedimentary rock successsions and associated igneous rocks were metamorphosed in greenschist to amphibolite facies during Caledonian tectono-thermal events. Based on sediment infill facies, architectural style and content of syndepositional igneous rocks, the western Baltoscandian palaeobasins can be subdivided into three major types. Type I basins are dominated by very coarse-grained, fluvial to deep-marine, feldspathic sandstones and conglomerates, and developed as rift basins in a cratonic setting. Type II basins are characterized by successions up to 3-4 km thick of dominantly medium- to finegrained sandstone of wide lateral extent. The structural setting of the type II basins is likely to have been pericratonic in coastal plain and shallow-marine environments. The type III basins

Fig. 11. Simplified tectonostratigraphy of the Scandinavian Caledonides, showing the diverse nappes and nappe complexes mentioned in the text. The higher, outboard, oceanic terranes of the Upper Allochthon (Koli Nappes), and the nappe complexes of the Uppermost allochthon, comprising rocks of inferred Laurentian origin, are purposely omitted from this figure. (Modified and simplified from Roberts & Gee 1985).

contain varied sedimentary facies and are characterized by the emplacement of large volumes of magmatic rocks during the stage of active basin formation. By comparison, volcanic rocks are either absent or present in only small amounts in type I basinal successions. Mafic sills or dykes are either poorly represented or locally common, in places occurring in swarms, in the basinal successions of types II and III. One diamictite unit has been recorded in several of the western Baltoscandian basins of types I and II. The diamictite has been identified as a Varangerian tillite, most likely corresponding to the younger of two tillite formations in the pericratonic succession of East Finnmark (Nystuen 1976, 1985; Bj0rlykke & Nystuen 1981; Edwards 1984; Kumpulainen & Nystuen 1985). According to Gorokhov et al. (2001), the Varangerian glaciation occurred between 630 and 560 Ma. Ages of other Neoproterozoic-Early Cambrian successions within the marginal regions of Baltica have been constrained by the distribution of acritarchs (Vidal & Nystuen 19900, b). Type I Basins in the Lower and Middle Allochthons. The most characteristic, rift-type basin succession derived from the western Baltoscandian margin is recorded in the Osen-R0a Nappe Complex of the Lower Allochthon (Bj0rlykke et al. 1976; Nystuen 1982, 1987) (Fig. 11). This is the classical 'sparagmite region' of the Scandinavian Caledonides. The ancient basin, earlier informally called the 'sparagmite basin', was termed the Hedmark Basin by Nystuen (1982) and identified

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

as an allochthonous, Neoproterozoic, western Baltoscandian basin (Nystuen 1981; Kumpulainen & Nystuen 1985). A similar, coarsegrained, rift basin succession, the Valdres Group, occurs within the Middle Allochthon in the Jotun-Valdres Nappe Complex in South Norway (Figs 1 & 11). This Valdres Basin succession contains several fault-related, alluvial conglomerate units (Nickelsen et al 1985; Kumpulainen & Nystuen 1985). The Valdres Group was deposited on Baltica-type crystalline rocks similar to those in the Jotun Nappe Complex (Bryhni & Sturt 1985). Type I basinal successions also occur in several nappe units in Jamtland and Vasterbotten, Sweden, as the Risback, Sjoutalven and Offerdal Groups in the Lower and Middle Allochthons (Gee et al. 1974, 1985; Gee & Kumpulainen 1980; Stephens et al 1985). The Risback Group and the overlying Sjoutalven Group were deposited in the western Baltoscandian Risback Basin (Kumpulainen 1982; Kumpulainen & Nystuen 1985). These

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basinal successions are also dominated by coarse-grained sandstone and conglomerate units. The infill history of the Hedmark Basin best illustrates the Neoproterozoic rift-type basins of the western Baltoscandian margin. The basin was formed by rifting of a Mesoproterozoic granitic basement in Late Riphean time (Bj0rlykke et al. 1976; Nystuen 1982, 1987, 1999; Kumpulainen & Nystuen 1985; Vidal & Nystuen 199(k) (Fig. 12A), but the precise age of the initial rifting is not known. The pre-rift, crustal extensional stage or initial rift stage is reflected by a fluvial sandstone formation (Fig. 13b,c) (Nystuen 1987, 1999). Basin widening by lateral fault propagation occurred during the early synrift stage, accompanied by progradation of alluvial fans and infill of deep marine troughs with turbidite sandstone (Fig. 13a), submarine conglomerate fans and black, organic-rich shale. This stage was succeeded by a rise in relative sea-level, with deposition of

Fig. 12. Conceptual diagram of Late Riphean to Early Cambrian basin evolution along the western Baltoscandian margin. (A) Late Riphean (R3) situation showing the embryonic stage of type I basins and the pericratonic type II basins. (B) Late to Terminal Riphean (R3-R4) stage, coinciding with peak magmatic activity for the type I basins. The type II basinal sedimentation proceeded in a continuing extensional regime. (C) Vendian - Cambrian situation V-C) showing maximum magmatic activity (mafic dykes, gabbros, basaltic lavas and sills) in the type II and type III basins. Farther out (to the left), lapetan oceanic crust was being generated. Inland, crustal submergence led to the widespread Early Cambrian marine transgression. (Modified from Kumpulainen & Nystuen 1985.)

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Fig. 13. Lithologies of the Hedmark Basin, SE Norway, (a) Turbiditic sandstone of the Br0ttum Formation, lower part of the Hedmark Group, Lillehammer. Stratigraphic way up is to the right. (Photo: J. P. Nystuen.) (b) Massive and thick-bedded fluvial sandstone of the Rendal Formation, lower part of the Hedmark Group, R0a. (Photo: J. P. Nystuen.) (c) Cross-bedded fluvial sandstone, Rendal Formation, Rendalen. (Photo: J. P. Nystuen.) (d) Intraformational limestone breccia of the Biri Formation, Hedmark Group, Hatten, Trysil. (Photo: J. P. Nystuen.) (e) Laminated micritic limestone with mudstone laminae, Biri Formation, Bin, lake Mj0sa. (Photo: J. P. Nystuen.) (f) Moelv Tillite, lithified basal till with clasts of granite, porphyry and red sandstone; Slemdalen. (Photo: J. P. Nystuen.)

shallow-marine limestone and organic-rich shale (Biri Formation) (Fig. 13d,e) onto continental alluvial facies. Renewed faultinduced, coarse-clastic, fan-delta sedimentation took place in the late synrift stage (Fig. 12B) prior to the onset of the Varangerian glaciation and deposition of the Moelv Tillite formation (Fig. 13f). Erosional remnants of the tillite lying on autochthonous crystalline basement along the present front of the Caledonides prove the wide regional extent of the Varangerian ice sheet. The basin was finally filled by deltaic, fluvial and shallow-marine sediments during Vendian to Early Cambrian time (Fig. 12C). In the

Early Cambrian, the sea transgressed over large parts of Baltica. Quartz-arenite sandstone and associated greyish-green shales covered most of peneplained Baltica, as well as the Neoproterozoic successions of the Hedmark Basin and other western Baltoscandian basins. Local tholeiitic basalt flows occur in the pre-Varangerian, synrift part of the Hedmark Group and have been suggested to represent fissure eruptions dating from the peak rifting period (Saether & Nystuen 1981; Nystuen 1987, 1999). The Varangerian Moelv Tillite rests with primary depositional contact on some of these local basalt flows, which have

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

a geochemical signature that is typical of continental tholeiites (Furnes et al. 1983). No reliable isotopic age has yet been obtained on these basalts. Lithostratigraphy, facies and architecture vary appreciably from one type I rift basin to another. However, two chronostratigraphically significant marker horizons are recorded in several of the basins. The pre-Varangerian carbonate level is present in the Valdres and Risback Basins as well as in the Hedmark Basin (Kumpulainen & Nystuen 1985). This level represents a marked regional transgression, the Bin transgression (Bj0rlykke et al. 1976), a marine incursion with associated sedimentation which probably filled erosional or structural depressions in the Baltican craton, as indicated by the correlation of acritarchs between the Hedmark Basin and the Vattern Basin in southern Sweden (Vidal 1985; Andreasson 1994; Vidal & Moczydlowska 1995). The other marker bed is the Varangerian tillite formation in the Hedmark, Valdres and Risback Basins. Type II Basins in the Middle Allochthon and Lower Part of the Upper Allochthon. Type II basins are present in the Kvitvola Nappe Complex in South Norway as the Engerdalen Basin (Nystuen 1980; Kumpulainen & Nystuen 1985), and in the Sarv Nappe in southern Sweden and Norway as the Tossasfjallet Basin and other equivalent palaeobasins in the Middle Allochthon (Gee 1977, 1980; Gee & Kumpulainen 1980; Krill 1980; Greiling 1989; Greiling & Kumpulainen 1989; Greiling et al 1999) (Fig. 11). They also occur as several other, Neoproterozoic, basinal successions in the Seve, Kalak and Laksefjord Nappe Complexes (Fig. 11) throughout northern Sweden and Norway (Laird 1972; Andreasson 1994; St01en 1997; Andreasson et al 1998). In the southwestern part of the Norwegian Caledonides, in the Sunnfjord area, type II basinal sedimentation is probably represented by the H0yvik Group within an outlier of the JotunValdres Nappe Complex (Brekke & Solberg 1987; Corfu & Andersen 2002). Stratigraphies and depositional environments of type II basins are particularly well displayed by the Tossasfjallet and Engerdalen Basins. The pre-Varangerian succession in the Tossasfjallet Basin is dominated by up to 4000 m of alluvial or coastal plain sandstone with some marine or lacustrine mudstone intercalations, overlain by a dolomite formation (Gee & Kumpulainen 1980; Kumpulainen 1980; Kumpulainen & Nystuen 1985). A similar pre-Varangerian stratigraphy is present in the Engerdalen Basin (Nystuen 1980), and in type II successions reported from several thrust sheets in the Seve, Kalak and Laksefjord Nappe Complexes of northern Sweden and Norway (Williams et al 1976; F0yn et al 1983; Ramsay et al 1985; Andreasson et al 1998). The stratigraphic level of the dolomites, commonly associated with black shales and evaporite-facies magnesite, also with chert nodules (Nystuen 1969), forms an important marker unit in many of these type II basinal successions (Arnbom 1980; Sjostrom 1983; Andreasson 1986; Kullerud^a/. 1990; Svenningsen 1994a, b, 1995; St01en 1994, 1997). One single, Varangerian tillite formation rests on the dolomite marker unit in both the Engerdalen and the Tossasfjallet Basins; but the glacial horizon has not been identified in the type II basinal successions in the Seve and Kalak Nappe Complexes. The Varangerian tillite and the dolomite marker are succeeded by a several hundred metres thickness of Vendian to ?Cambrian marine mudstone, and shallow-marine and fluvial sandstone in all of the type II basinal successions mentioned above (Kumpulainen & Nystuen 1985; Andreasson et al 1998). The sedimentary record of the Engerdalen Basin is interrupted by thrust sheets consisting of augen gneiss, granite and gabbroic rocks, inferred as tectonic slices derived from the western Baltoscandian margin (Nystuen 1980). The Engerdalen Basin succession is devoid of both syndepositional volcanic and intrusive rocks, and post-depositional intrusions. Rift-related mafic dykes and sills with intrusive ages of less than 650 Ma characterize the

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Sarv Nappes (the Ottfjallet dolerite dyke swarm) (Fig. 14f) and other thrust units in the overlying Seve and Kalak Nappe Complexes. Andreasson (1987) referred to these dyke swarms as the Baltoscandian Dyke Swarm. Sheeted dyke complexes and gabbro intrusions occupy several of the type II basinal successions in the Seve thrust sheets of northern Sweden and in the Laksefjord and Kalak thrust sheets of northern Norway. These dykes and sills are generally tholeiitic with MORB affinity, and range in age from about 615 to 550 Ma (Andreasson et al 1998). The 616Ma Egersund mafic dyke swarm, of similar character in the autochthonous basement of SW Norway, is considered to be related to the opening of the lapetus Ocean (Bingen et al 1998; Bingen & Demaiffe 1999). A type II basinal succession consisting of sandstone units several thousand metres in total thickness occurs in the Rondane area, central southern Norway. The arkosic sandstone package occurs in thrust sheets of the Kvitvola Nappe Complex, though the thrust sheets of the Rondane sandstones are located tectonostratigraphically above the Kvitvola Nappe sensu stricto (Siedlecka et al 1987). A conglomerate (Rosten Formation) with granite and gneiss clasts carrying pre-sedimentary deformation structures (Strand 1951) occurs in the lower part of the thick sandstone succession, that is structurally associated with a thrust sheet of crystalline rocks of Baltican provenance. The units of arkosic sandstone were probably deposited within fluvial and/or shallow-marine environments in a basin possessing a high rate of accumulation. The succession may correlate laterally with the pre-Varangerian sandstone units in the Engerdalen and Tossasfjallet Basins of the Kvitvola and Sarv Nappes, respectively (Bockelie & Nystuen 1985). Although the Rondane succession represents a type II basin that originated along the western Baltoscandian margin, it is uncertain whether it has been derived from a separate basin or from the same basin as the Engerdal Group in the Kvitvola Nappe. In the Sunnfjord area of western Norway, the H0yvik Group is the southernmost representative sandstone succession of established Neoproterozoic age in the Scandinavian Caledonides. The basal part of the succession rests with primary depositional contact on 1500 Ma old, Proterozoic basement rocks that are correlated with the Precambrian crystalline rocks of the JotunValdres Nappe Complex (Brekke & Solberg 1987; Corfu & Andersen 2002). The H0yvik Group has an exposed thickness of more than 1500m and is dominated by fine- to mediumgrained, shallow-marine quartz arenites and arkosic sandstones, with a carbonate member in the lower part of the group. The succession is penetrated by a mafic dyke swarm of similar geochemical signature to that of the Ottfjallet dolerite dyke swarm in the Sarv Nappe (Daehlin & Andersen 1988), and may be part of the Baltoscandian Dyke Swarm. Brekke & Solberg (1987) correlated the H0yvik Group with the 'sparagmites' of south-central Norway. The H0yvik Group is here interpreted as a type II western Baltoscandian basinal succession of pericratonic origin, cut by post-depositional, rift-related, mafic dykes, a palaeogeographic position also suggested by Brekke & Solberg (1987) and Osmundsen & Andersen (1994). Other Neoproterozoic sandstone successions may be present in the Caledonian nappe pile of southwestern Norway. In central-south Norway, the psammitic rocks of the Heidal Group (the 'Heidal Series' of Strand 1951), in the southern part of the Upper Allochthon, probably represent a succession derived from a distal part of the western Baltoscandian margin or possibly an outboard microcontinental block of Baltican derivation. This can also be inferred from its gneissic basement, which is of Baltican affinity (Sturt & Ramsay 1999). The feldspathic sandstones and flaggy psammites of the Hummelfjell mountain tract south of R0ros and NE of the Rondane area, now part of the Remsklepp Nappe Complex, were probably also derived from the western Baltoscandian margin. This succession is intruded by abundant metadolerite dykes (Gee et al 1985). Comparable flaggy

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Fig. 14. Lithologies in Seve, Kalak and Sarv Nappes, Scandinavian Caledonides. (a) Arkosic metasandstones with cross bedding (inverted sequence), Kalak Nappe Complex, North Troms, Norway. (Photo: K. B. Zwaan.) (b) Flagstone quarry in strongly tectonized, feldspathic metasandstones of the Klubben Formation, S0r0y Group, Kalak Nappe Complex; Alta, Finnmark, north Norway. (Photo: T. Heldal.) (c) Commercially exploited, thin-banded, tectonic flagstone with pelitic interbeds of the Leksdal Nappe (Sarv Nappe), Imsdalen, near Snasa, central Norway. (Photo: D. Roberts.) (d) Typical association of feldspathic metasandtone, felsic mica schist and amphibolite of the Seve Nappe; road-cut east of Are, west-central Sweden. Boudinage is quite common in this region, and farther to the west into Norway. (Photo: R. Kumpulainen.) (e) Mafic dyke cutting obliquely through multilayered feldspathic metasandstones and garnet-mica schists of the Klubben Formation, Kalak Nappe Complex, S0r0ya, west Finnmark, north Norway. (Photo: D. Roberts.) (f) Steeply dipping arkosic sandstones transected at high angle by a dolerite dyke, which itself is offset along a fracture or shear surface. Tossasfjall Formation, Sarv Nappe, Ottfjallet, west-central Sweden. (Photo: R. Kumpulainen.)

metasandstone successions are found in several lensoid thrust sheets in central Norway, e.g. Saetra Nappe (Krill 1983) and Leksdal Nappe (Andreasson et al 1979; Beckholmen & Roberts 1999) (Fig. 14c), belonging mostly to the Sarv Nappes of the Middle Allochthon (Fig. 11). These successions are also intruded by metadolerite dykes. In Finnmark, the bulk of the c. 8 km thick succession of the Laksefjord Nappe Complex (Laird 1972; F0yn et al 1983) con-

sists of basal, unsorted, conglomeratic sandstones followed by tabular and trough cross-bedded, arkosic sandstones. These psammitic rocks are interpreted as representing alluvial fan, braided stream and flood-plain alluvial deposits deriving from a highland area of basement gneisses to the SE (F0yn et al 1983), known as the Finnmark Ridge (Gayer et al 1987; Gayer & Rice 1989). The sandstones are succeeded by shallow to deeper marine siltstones and mudstones. A feature of this thick succession

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

is its intrusion by moderately abundant metadolerite dykes of generally tholeiitic composition (Gayer et al. 1985) and, as yet, unknown age. Type III Basins, Intruded by Large Igneous Plutonic Bodies. Basinal successions of this third type are similar in their general depositional setting to those of type II, but occur at a higher tectono-stratigraphic level and are consequently representative of more outboard locations along the Baltoscandian margin. Moreover, they are characterized by the occurrence of voluminous magmatic rocks in the highest nappes, varying widely from extensive, sheeted mafic dyke swarms to major complexes of ultramafic, gabbroic and alkaline compositions, and are considered to have originated in the Baltican continent-lapetus Ocean (TEgir Sea) transition zone (Andreasson 1994; Andreasson et al. 1998). Palinspastic restoration based on balanced cross-sections has indicated that the outermost parts of the Baltoscandian continental margin extended c. 600 km W-NW of their present-day locations in western Finnmark (Gayer et al. 1987). The principal structural units hosting these outermost, coastal plain to marine, Neoproterozoic successions are the Seve Nappes of northern Sweden and Kalak Nappe Complex of northern Norway (Fig. 11). These laterally extensive units were united by Andreasson et al. (1998) into what they have termed the SeveKalak Superterrane (SKS), with the magmatic rocks interpreted as the fragmented remains of a large igneous province. In the Swedish areas, there is a general succession of psammites, quartzites and diverse schists (Fig. 14d) with thin marbles and graphitic pelites, and mafic volcanites and rare diamictite horizons, but this is dissected into many thrust sheets, variably mylonitized and, in places, almost totally engulfed by sheeted dykes and gabbroic plutons (Andreasson 1986; Svenningsen I994a,b). A minimum age for the succession, and the approximate age of rifting immediately prior to ocean development, is given by a U-Pb zircon date of 608 ± 1 Ma from a sheeted dyke unit in one of the Seve thrust sheets (Svenningsen 2001). The Seve Nappes were initially deformed and metamorphosed, in eclogite facies, in Late Cambrian/Early Ordovician time (M0rk et al. 1988; Essex et al. 1997) during a major tectono-thermal event known as the Finnmarkian (Sturt et al 1978; Roberts 2003). Farther north, in Troms and Finnmark, northern Norway, the typical Seve rocks pass into the Kalak Nappe Complex (Sturt et al. 1975; Zwaan & Roberts 1978; Ramsay et al. 1985). Although subdivided into 9 or 10 separate nappes or thrust sheets, the Kalak as a whole preserves a distinctive lithostratigraphical succession known as the S0r0y Group. This mainly amphibolite-facies succession (c. 4.5-5 km thick) consists largely of cross-bedded arkosic psammites (Fig. 14a,b) and intercalated pelitic schists, followed by specific schist, marble and graphitic schist formations with a turbiditic metagreywacke and schist formation at the top (Roberts 1968a, b, 1974; Ramsay 1971). Greenstone (metabasalt) horizons are present in this turbiditic unit in parts of Troms (Padget 1955). Mafic dykes are common in many of the thrust slices that compose the Kalak Nappe Complex (Fig. 14e). A feature of the various thrust sheets is that they comprise a pairing of Palaeoproterozoic ortho- and paragneisses or Karelian Raipas Supergroup rocks, and parts of the Neoproterozoic S0r0y Group succession, in places with an unconformable contact preserved (Ramsay & Sturt 1977). This particular basement-cover relationship is also recorded in the highest, and most distal-allochthonous, S0r0y-Seiland Nappe, on Seiland (Akselsen 1982). The S0r0y-Seiland Nappe is also host to the voluminous Seiland Igneous Province (SIP) (Robins & Gardner 1975; Robins 1996) which covers an exposed area of close to 5000 km2. The magmatic evolution of the SIP involved the intrusion of many gabbroic to dioritic plutons, ultramafic complexes, alkaline rocks, carbonatites and mafic dykes. The oldest gabbroic plutons, based

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on field relationships, are of calc-alkaline character and considered to be syndeformational; but only one isotopic dating (Rb-Sr) has been reported, giving an inferred intrusion age of c. 830 Ma (Krogh & Elvevold 1990). The precise age of the bulk of the magmatic activity, mostly gabbro and ultramafic complexes, and dyke swarms, is uncertain, although Sm-Nd isochron dates favour the period c. 705-530 Ma, during a major phase of crustal extension and associated foliation development. Based on field relationships, five separate generations of mafic dykes have been reported from the 0ksfjord district, four of which are constrained to the time interval c. 600 to less than 550 Ma (Reginiussen et al 1995). Recent U-Pb zircon dating of gabbro and diorite from 0ksfjord and S0r0ya also falls within this range (R. J. Roberts, pers. comm. 2003). The youngest group of igneous rocks in the SIP is represented by alkaline rocks and carbonatites, intruded at shallow crustal levels in the period c. 530-520 Ma (U-Pb zircon dates; Pedersen et al. 1989). This was followed by a highgrade metamorphic event, the Finnmarkian, in Late Cambrian time (Robins 1996). Before 1990, the age of the S0r0y Group had been inferred as latest Riphean to Vendian, and possibly Cambrian. Since glacigene diamictites were absent, the maximum age was unknown (Zwaan & Roberts 1978, p. 57), though not older than a granite dyke Rb-Sr-dated to c. 1470 Ma in subjacent basement gneisses (Ramsay & Sturt 1977). The principal polyphase deformation of the succession was then generally considered to be Finnmarkian (Sturt et al 1978), thus correlating with the eclogite-facies accretionary event recognized later in the Seve Nappes; and the SIP was interpreted as broadly syn-Finnmarkian. Andreasson (1987) and Krill & Zwaan (1987) suggested, instead, that the SIP could be largely pre-Finnmarkian. This idea was supported by subsequent isotopic work, noted above. In addition, Daly et al (1991) provided U-Pb, Sm-Nd and Rb-Sr dating evidence indicating that a major phase of orogenic deformation and metamorphism affected the S0r0y Group before c. 800 Ma, and Nd isotopic data suggested that deposition of the S0r0y Group occurred in the period c. 1300-800 Ma (Aitcheson 1990). Reginiussen et al. (1995), on the other hand, interpreted the penetrative foliation in the S0r0y Group and in the oldest intrusions in the SIP to be an extensional fabric related to rifting and crustal thinning. Robins (1996) accepted this possibility, but also noted that emplacement and later foundering of the large mafic and ultramafic plutons could have generated widespread local foliations in the country rocks. Summing up, the outermost parts of the Baltoscandian margin representing the type III basins are characterized by mainly mafic-ultramafic magmatism with several generations of mafic dyke swarms. In the Seve Nappes, intrusive activity appears to have been largely confined to Vendian time. Farther north, in the Kalak Nappes, the very earliest gabbroic magmatism in the SIP appears to have started in the Late Riphean, and mantle diapirism may have been a contributory factor (cf. Reginiussen et al. 1995; Andreasson et al 1998). Whatever the fundamental cause of the magmatism (which is not the subject of this paper), crustal extension (Krill & Zwaan 1987) or transtension (Roberts 1988) clearly facilitated the episodic ingress of magma, culminating in the Vendian mafic dyke swarms and other mafic bodies, and carbonatites and related alkaline rocks, that immediately preceded the opening of the seaway separating western Baltica and Siberia (Torsvik & Rehnstrom 2001).

Discussion Basinal, stratigraphic and palaeoclimatic features From the foregoing accounts, the two complementary, northeastern and northwestern margins of Baltica show both similarities

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and contrasts in their developments in Neoproterozoic time (Fig. 5). In the case of the less well exposed Timanian margin, a long-lived passive margin was ultimately converted to a convergent or transpressional one during Mid to Late Vendian time. In contrast, in precisely this same time interval the northwestern, Baltoscandian margin remained essentially passive, and had just reached its rifting climax, heralding the rift-drift transition into a developing oceanic realm (Fig. 5). Thus, approximately orthogonal margins of Baltica record coeval compression and extension at the termination of the Neoproterozoic era. The northeastern and northwestern, rifted passive margins show pronounced differences in the geometries of the developing sedimentary basins in response to the extensional regime. The Timanian margin is characterized by disparate, elongate, pericratonic and basinal domains separated by a prominent fault zone along which there are developments of stromatolitic carbonate rocks (Olovyanishnikov et al. 2000). Evidence of the basin deepening to oceanic status outboard of NE Baltica, with both primitive and later evolved arc development, comes largely from the many deep drillholes penetrating the Early Palaeozoic and younger cover of the Pechora Basin (Getsen 1991), and from the geology of the Polar Urals (Scarrow et al 2001). However, synextensional magmatism in the exposed pericratonic and basinal domains is restricted to uncommon, thin basalt and tuff layers. The northwestern, Baltoscandian margin, on the other hand, is segmented and characterized by specific rift basins based on the infill facies, basinal architecture and the character of syndepositional magmatism. Basalt flows of continental tholeiite character are present in one of the most inboard rift basins, and elsewhere dolerite dyke swarms of Vendian age are a prominent feature in the thick, arkosic sandstones of many thrust sheets. Farther north, large volumes of mainly mafic-ultramafic magmatic rocks and mafic dykes, of imprecise latest Riphean to mainly Vendian age and of inferred mantle-diapiric origin, characterize the most distal thrust-sheets of the magmato-sedimentary successions. Another aspect of the Neoproterozoic history of these two marginal areas of Baltica is the similarity in the development of the Riphean record from predominantly siliciclastic sedimentary to carbonate rocks in the uppermost part, in all areas that we have described (Fig. 5). The belt of stromatolitic dolomites in the Timans is the most extensive, and the carbonates of Kildin Island and Sredni Peninsula are their distant correlatives. In the latter area there are distinct indications of evaporitic conditions, exemplified by tepee structures and nodules, interpreted as having been sulphates (Siedlecka et al. 1995Z?). On Varanger, the Riphean pericratonic record terminates with a stromatolitic dolomite (the Grasdalen Formation), just beneath the overlying Varangerian glacial deposits (Smalfjord Tillite) (BertrandSarfati & Siedlecka 1980). The Porsanger Dolomite, a western stratigraphic correlative of the Grasdalen Dolomite, is a much more extensive unit with abundant columnar stromatolites and evidence of evaporitic conditions (silicified sulphate nodules, pseudomorphs of gypsum crystals) in the dolomite itself and in the underlying red shale (Siedlecka 1976; Tucker 1976). The dolomite on Kanin Peninsula with well documented Upper Riphean stromatolites (Raaben et al 1995) contains irregular or, in places, nodular silica concentrations reminiscent of evaporite nodules (A. Siedlecka, unpubl. data). The only sign of carbonate rocks with non-columnar stromatolites within the basinal succession is in the upper Barents Sea Group (Siedlecka 1982) far beneath the angular unconformity that is inferred, by correlation, to represent the RipheanVendian interface. There too, evidence of evaporitic conditions is documented (Siedlecka et al 1998), notably in the red beds just above the carbonate unit, where salt pseudomorphs, including 'hopper' crystals, have been observed. Another interesting aspect (discussed below) is that nowhere in the basinal domain of the

Varanger Peninsula, or in the entire basinal TVB, are there any glacial deposits. The stratigraphic record of the northwestern Baltoscandian basins includes a carbonate formation in the upper part of the Riphean succession. Like the carbonate-facies associations of the Late Riphean along the northeastern Timanian margin of Baltica, the carbonate formations of the type I and II basins of the western Baltoscandian margin are characterized by evaporite-facies rocks that are suggestive of a prevailing arid and warm climate and local sabhka environments at the time of deposition. The wide lateral occurrence of carbonate rocks at this particular stratigraphic level may indicate an extensive regional transgression, or a series of transgressions, onto Baltica. This has been called the 'Biri transgression' in the Hedmark Basin; below, we refer to it more generally as the 'carbonate transgression'. In the western Baltoscandian basins the carbonate-facies association is succeeded by thin to very thick, siliciclastic deposits overlain unconformably by the only tillite known in this region. As we have described, evidence of the Early Vendian, Varangerian glaciation in northern Norway is provided by two, stratigraphically separate, tillite formations in the peri-cratonic succession. On the other hand, tillites appear to be absent in the Rybachi-Sredni and Kildin areas, on the White Sea coast of the Kola Peninsula, and also on the southern side of the White Sea. One, widespread tillite unit, the Moelv Tillite, is present on the Baltoscandian margin of Baltica. It is correlated with the upper tillite of northern Norway; the Mortensnes Tillite (Fig. 5). The lower (Smalfjord) tillite is known only in northeastern Norway. The succession of the Smalfjord Formation comprises diverse sedimentary facies where moraine is a subordinate constituent. The sedimentary facies also show marked lateral variability, and variability in thicknesses (Edwards 1984). In addition, parts of the diamictite have more recently been reinterpreted as probable debris flows associated with alluvial fan accumulations (Jensen & Wulff-Pedersen 1996), which makes the distribution of the Smalfjord moraine even more restricted. Therefore, we assume that the Smalfjord glaciation was not as widespread, either geographically or in time, as the later Mortensnes event. In general, siliciclastic deposition ceased towards the end of Late Riphean time. The 'carbonate transgression' must have been widespread, at least embracing northeastern and northwestern parts of Baltica. At that time, stromatolites were flourishing in many regions, in a warm, arid/semi-arid climate. Retrogression, erosion and some tectonic movements (angular unconformity in the north) were followed by a marked deterioration of climate that heralded the Varangerian glaciation, initially areally restricted to the Varanger region and later widespread over large parts of Baltica. However, nowhere in postglacial time are there any thick 'cap dolomites' that immediately postdate the Varangerian glaciation. The only comparable lithology is a thin dolomite with stromatolite-like structures, present in the Varanger area on top of the lower tillite, the Smalfjord Formation, where there is a gradual transition into overlying shaly deposits. In this respect, the Varangerian glacial successions of northern Baltica differ from many other Late Neoproterozoic glacial units around the world that are terminated by carbonateevaporite rocks, a feature of critical significance for the 'snowball Earth' hypothesis (Kirschvink 1992; Hoffman et al 1998). On the contrary, the situation on Baltica was described by Roberts (1971, 1976) as an 'anti-greenhouse effect', the locking up of atmospheric carbon dioxide in the dolomites leading to climate deterioration and extensive glaciation. However, it is difficult to say how extensive or continuous the ice sheet actually was. No evidence of glaciation has been recorded along the Timanian margin, and none has been reported from the adjacent hinterland and foreland areas. The possible presence of Vendian strata in the Timans has been suggested

NE AND NW MARGINS OF NEOPROTEROZOIC BALTICA

(e.g. Dedeev & Getsen 1987, pp. 91-93; Olovyanishnikov 1998, pp. 117-119) but it is not well documented, in contrast to the White Sea area where the biostratigraphic record is excellent. There was probably a period of uplift and non-deposition during Early Vendian time along the Timanian cratonic margin heralding the tectonic inversion and Mid-Late Vendian Timanian Orogeny, and the syn- to post-tectonic sedimentation in the adjacent foreland basin (Grazhdankin 2004). Plate-tectonic aspects of the developing margins The birth of the Neoproterozoic continent Baltica derives from the break-up of Rodinia. According to traditional 'Wilson Cycle' concepts, the Baltoscandian margin of Baltica is considered to have rifted away from eastern (Appalachian) Laurentia (Dewey 1969; Gee 1975; Roberts & Gale 1978). This notion was questioned first by palaeomagnetic data and interpretations requiring that Baltica lay in an inverted position during Late Riphean to Vendian and even Cambrian time (Torsvik et al. 1991). Other

Fig. 15. Distribution of the palaeocontinents and oceanic areas in Late Riphean (above) and latest Vendian time, based on the palaeomagnetic data, reconstructions and interpretations of Torsvik & Rehnstrom (2001) and Hartz & Torsvik (2002), with small modifications. The Seiland Igneous Province (S) of the S0r0y Nappe (highest part of the Kalak Nappe Complex) is here shown as a microcontinental sliver that originally detached itself from Baltica, and was eventually reunited with the parent continent during the circa 500 Ma Finnmarkian orogeny. BM, Baltoscandian margin; TM, Timanian margin; PNZ, Pechora and Novaya Zemlya blocks; PGACT, Peri-Gondwanan Avalonian-Cadomian terranes. The black star marks the approximate location of the inferred Balto-Timanian triple junction. Double lines, sea-floor spreading axes; Lines with ticks, intercontinental, ocean-margin rifting; Line with triangular tags, inferred subduction polarity, Avalonian-Cadomian arc.

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features, such as basin asymmetry and differences in sedimentary successions and rift magmatism, led Andreasson (1994) and Soper (1994) to suggest alternative fits for Baltica. In more recent years the Siberian plate has entered the arena (Torsvik et al 1995), with a 'new' oceanic tract, the JEgir Sea (Torsvik & Rehnstrom 2001) separating Siberia from Baltica (Fig. 15) in Late Riphean to Cambrian time (Cocks & Torsvik 2002; Hartz & Torsvik 2002). An inverted Baltica has also been shown to facilitate reconstructions of the peri-Gondwanan, Avalonian-Cadomian terranes and their linkage with the Timanian Orogen (Roberts & Siedlecka 1999, 2002; Roberts & Torsvik 2000; Hartz & Torsvik 2002). Most plate-tectonic reconstructions of Baltica include the Pechora, Polar Urals and Novaya Zemlya lithospheric segments as part of this plate (e.g. Torsvik et al. 1991, 1996; Dalziel 1997; Dalziel et al. 2000; Torsvik & Rehnstrom 2001; Rehnstrom & Corfu 2004). These segments represent different blocks or subterranes that were later juxtaposed against the Neoproterozoic Baltica sensu stricto, which itself developed as a result of initial break-up of Rodinia. The Neoproterozoic Baltica thus differs from the Fennosarmatian Craton as outlined by Nikishin et al. (1996) and the Early Palaeozoic Baltica as defined by Cocks & Fortey (1998), by excluding the Pechora and Barents Sea regions, parts of Novaya Zemlya, Taimyr and adjacent Arctic islands. A more likely representation of Baltica in latest Riphean time is shown in Hartz & Torsvik (2002; Fig. 1). The welding of these diverse blocks to the northeastern, Timanian margin of Baltica thus probably occurred at the termination of the Neoproterozoic (Nikishin et al. 1996) in connection with Timanian orogenesis and accretion (Roberts & Siedlecka 2002; Roberts & Olovyanishnikov 2004). The original northeastern margin was broadly defined by the array of faults and major fault zones extending from the Timans in the SE to Varanger Peninsula and beyond, below the Caledonian nappes, in the NW. A critical problem in plate-tectonic reconstructions of the Neoproterozoic lithospheric segments in this region is the structural relationship between the Timanian and the Baltoscandian margins. These two near-orthogonal margins are inferred to have intersected in an area situated several hundred kilometres WNW of present-day Finnmark (Gayer et al. 1987). It has also been suggested (Siedlecka 1975; Roberts & Gale 1978; Rice & Reiz 1994) that these margins merged at a common triple junction, although in these particular models Baltica is assumed to have rifted away from Greenland. Accepting the more recent palaeomagnetic evidence of an inverted Baltica in Late Riphean time (Torsvik et al. 1996; Hartz & Torsvik 2002), with the Siberian plate forming the opposite, conjugate margin, we suggest that a triple junction did exist in Late Riphean-Vendian time, which we refer to, informally, as the Balto-Timanian triple junction (BTTJ) (Fig. 15). We consider that this orthogonal marginstriple junction configuration may explain both differences and similarities in the successions that accumulated along these adjacent margins of Baltica. The northwestern Baltoscandian margin hosted pericratonic type-II basins, typical of rifted, passive continental margins, disrupted by type-I rift basins that indented this passive margin. The northeastern Timanian margin, on the other hand, with its progressively subsiding, outboard basinal domain, is likely to have formed by a combination of normal extension and strike-slip movements. A developing BTTJ would have been an area of crustal thinning and abnormal heat flow, thus creating a situation conducive to the generation of a mantle plume with associated, voluminous magmatic activity. It is in this northwesternmost corner of Neoproterozoic Baltica that the most outboard, type-Ill, S0r0y Group succession accumulated, in Mid to Late Riphean time, together with the latest Riphean to Vendian magmatic rocks of the Seiland Igneous Province, before their ultimate incorporation into the highest levels of the Kalak Nappe Complex (Gayer et al. 1987). An allied feature is the tendency for alkaline rocks and carbonatites to have clustered at or close to

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triple-junction sites just prior to the opening of lapetus (Roberts & Gale 1978). Doig (1970) showed that in the region of North America and NW Europe, many such centres of alkaline magmatism, or failed arms containing alkaline rocks, are of Vendian age. Parts of the type-Ill basinal successions of the Seve Nappes of northern Sweden are also inferred to have derived from this outermost, northwestern corner of Baltica, according to the studies of Rehnstrom et al (2002) and Paulsson & Andreasson (2002). In the former case, a metamorphic age of 637 ± 3 Ma from a late Palaeoproterozoic gneissose granite is interpreted as relating either to early stages of Timanian accretion or to Vendian extension. The latter study reported a c. 850 Ma anatectic granite, and a c. 605 Ma metamorphic age associated with the emplacement of a dolerite dyke swarm.

Conclusions (1) The Neoproterozoic depositional histories of these two orthogonal margins of Baltica—the northeastern Timanian and the northwestern Baltoscandian margin—show several important differences but also some similarities. Both started life in ?Mid to Late Riphean time during crustal-extensional regimes arising from break-up of the supercontinent Rodinia. The two margins are believed to have intersected at a common, triple junction, the Balto-Timanian triple junction, situated some several hundred kilometres to the NW of present-day Finnmark. Ultimately, these long-lived passive margins were converted to active, compressional/ transpressional scenarios, although within the time-frame of the Neoproterozoic era the Baltoscandian margin remained passive throughout. (2) The principal features of the two margins during the Neoproterozoic era can be summarised as follows: (a) The northeastern Timanian margin comprises mainly terrigenous, low- to medium-grade, sedimentary successions with a prominent, margin-parallel, extensional fault zone separating distinctive pericratonic and basinal domains. Syndepositional magmatic rocks are rarely seen on land, but they are reported to increase dramatically in volume in the deeper oceanic realm detected in many deep drillholes in the Pechora Basin. Conversion to an active margin occurred towards the end of Neoproterozoic time just prior to reversal to a compressional and partly transpressional regime which gave rise to the Mid-Late Vendian, Timanian Orogeny. During this period of accretion, diverse microcontinental blocks in the Timan-Pechora-Polar Ural-Novaya Zemlya region were united with the northeastern, Timanian margin, thus forming the familiar palaeogeographical outline of Early Palaeozoic Baltica. (b) The Baltoscandian margin successions are of variable, low to high metamorphic grade and are preserved in distinctive nappe complexes recognized throughout the Caledonides. Three main types of palaeobasin are distinguished based on sedimentary facies, basin geometry, and the character and volume of magmatic products. Mafic dyke swarms are ubiquitous, continental tholeiitic basalts occur in one, inboard, type-I rift basin, and voluminous mafic-ultramafic plutons cut the type-Ill basinal successions in parts of northern Norway. The type-Ill basin of western Finnmark and the associated SIP are considered to have originated in the vicinity of the Balto-Timanian triple junction, near the confluence of the two margins. This Baltoscandian margin remained passive until the middle to late stages of the Cambrian period.

(3) The Baltoscandian mafic dyke swarms marking the major rift phase just prior to lapetus/^Egir ocean opening are isotopically dated to the Vendian period, which is precisely when the orthogonal, northeastern Timanian margin was being subjected to compression, accretion and telescoping of its pericratonic and basinal/oceanic successions. These opposed, coeval, stress regimes along the two, orthogonal margins fit neatly into a common, large-scale, plate-tectonic framework. (4) The general features of facies development through Late Riphean to Early Cambrian time in the basins of these two margins of Baltica reflect a common history of continental denudation and climate change. A major deterioration of climate affected the northern and western parts of Baltica at the transition between Riphean and Vendian time. A warm and arid to semi-arid, Late Riphean climate with carbonate and red-bed deposition was followed by glaciation that was probably first restricted to northern areas and spread quite rapidly to central and western parts of the palaeocontinent. As thick, post-Varangerian, 'cap carbonate' successions are absent in these parts of Baltica, the general situation here is the opposite of that described by advocates of the 'snowball Earth' hypothesis for the Late Neoproterozoic. (5) Lower Vendian glacial deposits are absent in the Timanides and adjacent White Sea area and it is conceivable that they were never present. Instead, the Early Vendian was probably a time of uplift and erosion in these areas, immediately preceding the Mid-Late Vendian, Timanian Orogeny. Much of the fieldwork upon which this review is based was carried out by the four authors, and many other colleagues, over a period of more than three decades. Field studies by AS and DR on Kola Peninsula during the years 1990-1993 were part of a collaborative research programme between the Geological Survey of Norway and the former USSR (now Russian) Academy of Sciences. Fieldwork for DR and VO in northern Timan and on Kanin Peninsula, in the year 2000, was supported largely by the Swedish Polar Research Secretariat. We are grateful to Ian Dalziel and Maurits Lindstrom for their constructive reviews, which led to improvements in the final manuscript. Valuable and pertinent editorial comments from Vicky Pease are much appreciated. Thanks also go to Tom Heldal, Risto Kumpulainen and Bouke Zwaan for the loan of photographic slides, and to Irene Lundquist for preparation of the final figures.

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Svalbard: a fragment of the Laurentian margin 1

D. G. GEE1 & A. M. TEBEN'KO

Department of Earth Sciences, Uppsala University, Villavdgen 16, Uppsala, SE-752 36, Sweden (e-mail: [email protected]) 2 Polar Marine Geological Research Expedition (PMGRE), Pobeda Street 24, Lomonosov - St Petersburg, 189510, Russia

Abstract: The last decade of structural and isotope-age dating studies in Svalbard and East Greenland has provided strong support for the close correlation of these segments of the Caledonide Orogen, as had previously been inferred from stratigraphic evidence. Prior to Tertiary opening of the Norwegian-Greenland Sea, Svalbard's Caledonian terranes were an essential part of the Laurentian margin, as witnessed not only by the Early Palaeozoic depositional environments and fauna, but also by the character of the Palaeoproterozoic basement, the Meso- to Neoproterozoic cover, the evidence of late Grenvillian tectono-thermal activity, Caledonian structural style and timing of movements, Caledonian granitic magmatism and Old Red Sandstone (ORS) deposition. Recently published maps of East Greenland show the hinterland allochthons of central East Greenland to strike out obliquely into the continental shelf. The hypothesis promoted here requires that they continue offshore northwards, extending to the northern edge of the NE Greenland shelf and that most of the Svalbard terranes were northerly continuations of the East Greenland Caledonides. Only along the west coast of central Spitsbergen are 'foreign' terranes exposed that have affinity with Pearya, having been located north of the North Greenland foldbelt, apparently unrelated to Laurentia, prior to Ellesmerian Orogeny. The unambiguous affinity of the Svalbardian and Greenlandian (Laurentian) Caledonides contrasts markedly with the Timanide evolution of northeastern Baltica. It confirms previous interpretations that an important Caledonian suture-zone transgresses northeastwards across the Barents Sea, separating Laurentian domains in the NW from the Timanides of Baltica in the SE. The Timanides of northeastern Europe are truncated by, and terminate in the Barentsian Caledonides of the Barents Shelf.

In the late Palaeozoic and Mesozoic, prior to the opening of the Eurasian Basin and North Atlantic Ocean, Svalbard was located north of Greenland (Fig. la). Svalbard's Caledonian bedrock includes major complexes that are closely comparable with rock units in the East Greenland Caledonides. Most of Svalbard's Caledonian bedrock shares a tectonic history with the margins of Laurentia that can be traced far back into the Proterozoic; perhaps earlier. However, some units, in particular those occurring along the central part of the west coast of the island of Spitsbergen, appear to have close affinities with northernmost Ellesmere Island (Pearya) and the North Greenland Fold Belt. This paper compares the tectonic history of the Svalbard Caledonides with that of East and North Greenland and northern Ellesmere Island and discusses alternative hypotheses for their Proterozoic and Palaeozoic evolution. It also considers the relationships between the Svalbardian development and that of neighbouring terranes to the east and south, now composing the Barents Shelf and the Timanide Orogen, marginal to the Kola Peninsula. The sutures in the Scandinavian Caledonides (Gee & Sturt 1985) that separate allochthonous Laurentian margin terranes from Baltica (Gee 1975), continue northeastwards from northernmost Norway into the Barents Sea. Their locations are of importance for understanding the evolution of both the Barents Sea and the Arctic Basin since the Palaeozoic. Some authors (e.g. Ohta 1994; Verba & Sakoulina 2001) have argued that the Timanide Orogen extends northwestwards across the Barents Sea to Svalbard and into the North Greenland foldbelt; this hypothesis neglects the unambiguous evidence of Svalbard's affinity with Laurentia and the East Greenland Caledonides. Caledonian research on both the East Greenland margin and Svalbard, with the combined application of tectonic and isotope age studies in both areas, is providing new insight into the tectonic evolution of this part of the North Atlantic-Arctic Caledonides and the various terrane hypotheses (e.g. Fig. Ib, c). Transcurrent displacements of Svalbard terranes over distances of many hundreds (even thousands) of kilometres have been favoured by most authors since the terrane hypothesis for the North Atlantic Caledonides was presented in the late 1960s (Harland 1969). The new data favour a less complex hypothesis.

Svalbard The pre-Carboniferous basement of Svalbard, summarized in Harland (1997), is composed of Caledonian-deformed Early Palaeozoic and Proterozoic complexes and an Old Red Sandstone (ORS) graben (Fig. 2). The Caledonian rocks are separated into different provinces by major north-trending fault-zones that have been shown to have both dip-slip and strike-slip movement. Important differences in stratigraphy, structure and tectonothermal history across these faults led to the proposal that the different Caledonian provinces represent independent tectonostratigraphic terranes (Harland 1972; Harland & Gayer 1972; Gee & Page 1994). many hundreds of kilometres of lateral (strike-slip) movement have been claimed to explain the juxtaposition of these terranes, much of the evidence being based on comparisons of Svalbard's Caledonian bedrock with parts of the East Greenland margin of Laurentia. Harland (1985, 1997) divided the Svalbard Caledonides into three main provinces or terranes: eastern, central and western. Others (e.g. Gee 1986; Gee & Page 1994) have found the boundaries between the provinces less easy to define and have preferred a subdivision into eastern, northwestern and southwestern terranes. For convenience, particularly when relating to previous literature, the Svalbard Caledonides are summarized below in these categories and with reference also to the island of Bj0rn0ya, located 250 km to the south of Spitsbergen (Fig. la).

Eastern terranes Svalbard's composite Eastern Terrane is readily subdivided into two parts, referred to as the Nordaustlandet Terrane in the east (Fig. 3) and the West Ny Friesland Terrane further west (Fig. 4). The former extends eastwards across the Barents Shelf via Kvit0ya towards Franz Josef Land, but is thought not to reach this archipelago (Gee 2001, 2003). The West Ny Friesland Terrane has been considered to be bounded in the west by a major sinistral strike-slip fault running north-south from

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 191-206. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Late Mesozoic reconstruction of the North Atlantic Caledonides (A) and alternative hypotheses for the tectonic evolution (B) from Harland (1997), and (C) from Gee & Page (1994).

Fig. 2. Svalbard's Caledonian terranes.

Wijdefjorden to Billefjorden and called the Billefjorden Faultzone (Harland et al 1974), separating the Caledonian rocks from the ORS of the Andreeland-Dicksonland Graben. This fault was defined as a terrane boundary when the main movements were thought to have occurred in the late Devonian (Svalbardian) and not in the late Silurian to early Devonian, as has been inferred to be more probable (Gee & Page 1994; Witt-Nilsson et al 1998). With the Caledonian basement beneath the Andreeland ORS unexposed, it remains possible (indeed probable, see below) that these crystalline rocks are also related to Svalbard's West Ny Friesland Terrane. Nordaustlandet terrane. The Caledonian rocks of Nordaustlandet (Fig. 3) are exposed in northern and eastern areas, and are overlain to the south by Carboniferous and younger strata. Some Caledonian granites (c. 415 Ma) in this basement have related magnetic anomalies that extend far southwards towards

Edge0ya, beneath the platform cover (Gee et al. 1999; Johansson et al. 2002). Caledonian structures are dominated by upright to west-verging folds, generally associated with low greenschistfacies metamorphism (Flood et al. 1969). However, the Caledonian deformation and metamorphic grade increases eastwards, with migmatization in the Duvefjorden area and further to the east (Tebenkov et al. 2002; Johansson et al 2004). The Nordaustlandet Terrane is composed of a Grenville-age basement (Gee et al 1995) overlain by a Neoproterozoic platform cover (Murchisonfjorden Supergroup). The latter passes upwards into the Hinlopenstretet Supergroup, which stretches through the Vendian and Cambrian into the mid Ordovician (Llanvirn). The westernmost parts of the Nordaustlandet Terrane occur in eastern Ny Friesland, and are largely composed of Neoproterozoic successions (Lomfjorden Supergroup) that are very similar to formations of the Murchisonfjorden Supergroup of Nordaustlandet (Sandelin et al. 2001). The oldest rocks so far identified in the Nordaustlandet Terrane are metasedimentary units of the Brennevinsfjorden Group (western areas) and Helvetesflya Formation (central areas). These dark shales (phyllites), turbidites and subordinate quartz sandstones, generally metamorphosed in greenschist facies, are extensively intruded by c. 950 Ma granites (Gee et al 1995; Johansson et al 2000) and locally migmatized. No basement to the Brennevinsfjorden-Helvetesflya metasedimentary units has been recognized. Studies of detrital zircons (A. Larionov, pers.comm. 2000) in these formations have shown the dominance of Mesoproterozoic source areas, with little evidence of older provenance. The youngest zircons in these sedimentary rocks are c. 1200 Ma, implying deposition in the latest Mesoproterozoic or earliest Neoproterozoic, before intrusion of the 950 Ma granites. A major angular unconformity separates the Brennevinsfjorden (Ohta 19820) and Helvetesflya (Gee & Tebenkov 1996) metasedimentary units from overlying volcanic rocks of the Kapp Hansteen Group (in the west) and Svartrabbane Formation (central areas). Both the Helvetesflya and Brennevinsfjorden units are, at least locally, inverted beneath this unconformity and had apparently been folded, with high angle axial surfaces, prior to uplift, erosion and deposition of the overlying volcanic and volcaniclastic rocks. The latter are dominated by andesites and rhyolites, and intruded by quartz porphyries (Tebenkov 1983; Ohta 1985), all these rocks yielding ages similar to or slightly older (c. 960 Ma) than those of the granites intruding the underlying Brennevinsfjorden Group (Johansson et al 2000). Both the sedimentary (Brennevinsfjorden-Helvetesflya) and volcanic (Kapp Hansteen-Svartrabbane) formations were folded together during intrusion of the 950 Ma granites and metamorphosed in

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Fig. 3. Geology of Nordaustlandet, based on Sandelin et al (2001).

greenschist facies (Tebenkov et al. 2002). Uplift and erosion apparently occurred shortly thereafter, prior to the deposition of the unconformably overlying Neoproterozoic Murchisonfjorden Supergroup (Gee & Tebenkov 1996; Sandelin et al 2001). The Neoproterozoic Murchisonfjorden and Lomfjorden successions of the Nordaustlandet Terrane are closely comparable and readily divisible into two parts: an upper section dominated by limestones and dolomites and a lower siliciclastic section, mainly of shales and quartzites. The unconformable base of the Murchisonfjorden Supergroup has been found only in central Nordaustlandet (Sandelin et al 2001); further west, this junction is faulted, as is probably also the case in Ny Friesland with the base of the Lomfjorden Supergroup. These mostly shallow marine, little metamorphosed, strata are about 5 km thick and have been described in some detail from Nordaustlandet by Kulling (1934), Flood et al (1969), Ohta (1982Z>) and Sandelin et al (2001), and from Ny Friesland by Wilson (1958), Harland et al (1992) and Harland (1997). Studies of acritarchs have indicated that the Murchisonfjorden-Lomfjorden successions reach back to c. 800 Ma (Knoll 1982; Butterfield et al 1988). The Lomfjorden Supergroup in eastern Ny Friesland is underlain towards the west by metasedimentary units of the Planetfjella Group (Wallis 1969). The contact has been interpreted as a stratigraphic transition (Wallis 1969; Harland 1997), or a major fault (Nathorst 1910) with either strike-slip (Manby 1990) or extensional movement (Gee et al 1994). The Planetfjella Group has been correlated both with the Kapp Hansteen Group (Harland 1985) on the basis of an inferred volcanic component, and with the Brennevinsfjorden Group on the basis of general lithological comparability. However, Larionov et al (1998) have shown that

the Planetfjella detrital zircons include a c. 950 Ma population, indicating that the group is younger than, and probably sourced by, the Grenville-age basement of the Nordaustlandet Terrane. The Murchisonfjorden and Lomfjorden supergroups are overlain conformably by Vendian and Early Palaeozoic strata of the Hinlopenstretet Supergroup. Basal tillites (Kulling 1934; Harland et al 1993) are overlain by Cambrian sandstones and then limestones that pass up into the Early Ordovician. The youngest part of the succession is exposed along the western coast of Hinlopenstretet and includes a passage from platform limestones up into dark shales and limestones of late Arenig to early Llanvirn age (Fortey & Bruton 1973). Younger strata may occur beneath the waters of Hinlopenstretet. These Cambrian and Early Ordovician successions contain a fauna of unambiguous Laurentian affinities; only in the uppermost, deeper water Llanvirn strata are a small proportion of Baltica-related fossils reported (Fortey & Barnes 1977). Throughout western Nordaustlandet, the Hinlopenstretet and Murchisonfjorden supergroups are little metamorphosed, the grade reaching low greenschist facies only in the lower siliciclastic formations of the succession. Further east, in central Nordaustlandet, the metamorphic grade of these lower units increases to amphibolite facies near the contact with the migmatites of northern Austfonna (Fig. 3). Migmatization of the lower units of the Murchisonfjorden Supergroup, along with the Helvetesflya Formation, provides unambiguous evidence of the increasing Caledonian tectono-thermal regime towards the east. Isotopic dating has demonstrated the widespread crystallization of Caledonian (c. 430-450 Ma) zircons in these migmatites, both in the Duvefjorden area (Tebenkov et al 2002) and further east on Nordaustlandet

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Fig. 4. Geology of Ny Friesland, based on Gee et al. (2001).

and as far as Kvit0ya (Fig. 6) on the northern Barents Shelf (Johansson et al 2004). West Ny Friesland terrane. The Caledonian bedrock of western Ny Friesland (Harland et al 1992; Witt-Nilsson et al 1998) contrasts markedly with that of Nordaustlandet. The rocks of western Ny Friesland (Fig. 4) are notable for their more intense penetrative Caledonian deformation and generally higher grade (amphibolite facies) of regional metamorphism. The rocks units in the two terranes are of different lithology and age and the West Ny Friesland Terrane lacks evidence of Grenville-age tectono-thermal activity.

The West Ny Friesland crystalline rocks are dominated by orthoand paragneisses, schists and quartzites—the Atomfjella Complex (Krasil'shikov 1973). This succession is c. 8 km thick and occurs in a large upright to west-vergent fold, the Atomfjella Antiform, that is superimposed on earlier phases of isoclinal folding and thrusting (Harland 1959; Witt-Nilsson 1998). The complex is flanked to the west by the Billefjorden Fault-zone and the Andreeland ORS and unconformably overlain by Early Carboniferous sandstones. Within the Atomfjella succession, granitic gneisses occur at four main levels, overlain by metasediments. Both the igneous and

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sedimentary units are highly strained and primary structures are preserved only locally in low strain zones. However, basal sedimentary facies have been located in a few areas and consist of conglomerates overlying the granitic basement and containing pebbles derived from the latter (Hellman et al 1997; Witt-Nilsson et al 1998). The granites, occurring at different levels in the Atomfjella Complex are extensively deformed to orthogneisses; they have been dated by the U-Pb method on zircons (also, in some cases, titanites) and have yielded consistent c. 1750 ± 20 Ma ages (Johansson et al. 1995; Larionov et al. 1995; Johansson & Gee 1999; Johansson 2001). Only one orthogneiss has yielded an older (Late Archaean) age (Hellmann et al. 2001). Zircons have also been extracted from the intervening formations of mostly quartzite-dominated metasedimentary rocks (Hellman 2000). These yield single crystal Pb/Pb and ion microprobe ages closely similar to those of the orthogneisses; however, both older (Palaeoproterozoic and Late Archaean) and some younger populations occur, the former dominating. The quartzites (e.g. Polhem Formation) have detrital zircons as young as c. 1500 Ma (also one crystal of 1320 Ma) and a schist and marble unit (Smutsbreen Formation) contains zircons as young as c. 1190 Ma (Gee & Hellman 1996). Both the orthogneisses and the quartzites contain an abundance of mafic sheets, many of which can be shown to be of intrusive origin. One of these amphibolitized dolerites has yielded a zircon age of c. 1300 Ma (Hellman & Witt-Nilsson 1999). The Smutsbreen Formation, in the core of the Atomfjella Antiform, is notable for both younger detrital zircons and a general lack of the mafic rocks, that are so conspicuous in the overlying units. In combination, the structural and isotopic age studies of the Atomfjella Complex have demonstrated that this thick succession is tectono-stratigraphic (Gee et al. 1994; Johansson et al. 1995; however, see Harland 1997, for another interpretation), and composed of at least four large thrust-sheets (Witt-Nilsson et al. 1998; Gee et al. 2001). The thrusting was west-vergent and the thrust stack was assembled in the Silurian to Early Devonian (Gee & Page 1994; Johansson et al. 1995). A single late Ordovician Ar/Ar age points to the possibility of an earlier start to orogenesis (Gee & Page 1994). The Atomfjella Antiform is characterized by extreme axial north-south elongation, with ubiquitous boudinage (Harland 1959; Witt-Nilsson 1998). Along the west coast of Ny Friesland, retrogression from amphibolite to greenschist facies is accompanied by sinistral strike-slip faulting. The contact with the ORS is faulted (Harland et al. 1974) and the movements involve a late Devonian dip-slip (reverse) component of a few kilometres (Lamar & Douglas 1995; McCann & Dallmann 1996). As mentioned above, this major fault (the Billefjorden Fault-zone) has been claimed to be the western boundary of Svalbard's Eastern Terrane (Harland et al. 1974). Evidence from further west suggests that igneous rock units related to the West Ny Friesland Terrane may be present beneath the Andreeland-Dicksonland Graben and the boundary, therefore, should be placed further west. The contact between the Nordaustlandet and West Ny Friesland terranes is inferred to be located at the base of the Planetfjella Group. Garnet-mica schists (± staurolite and kyanite) of the latter overlie quartzites of the Atomfjella Complex and the thrust (Gayer 1969; Gee et al. 1994) that separates these two units in northern Ny Friesland, is marked by lenticular ultramafites. Other authors (Manby 1990; Lyberis Manby 1999) have inferred that the upper contact of the Planetfjella Group, towards the overlying Lomfjorden Supergroup, is a major strike-slip fault, defining the terrane boundary. We regard this deformation zone as an extensional phenomenon, related to late orogenic uplift (Gee et al 1994). Other authors (e.g. Harland 1997) regard both the upper and lower contacts of the Planetfjella Group to be primary (i.e. with normal stratigraphic relationships preserved), though somewhat disturbed by faulting.

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A large Caledonian granite batholith intrudes the Lomfjorden Supergroup and Planetfjella Group in southern Ny Friesland; it has yielded a Rb-Sr whole-rock and mineral (apatite) age of 432 ± 10 Ma (Tebenkov et al 1996).

Northwestern terranes The Andreeland-Dicksonland Graben (Fig. 2), is bounded to the west by a major fracture, the Breibogen-Bockfjorden Fault (BBF). There is volcanic activity with hot-springs in the vicinity of Bockfjorden. These Pleistocene volcanic rocks contain abundant mantle and lower crustal xenoliths (Amundsen et al 1987), providing evidence for the deep penetration of this fault system today. To the west of the BBF, Caledonian crystalline rocks are preserved in two north-trending horsts, separated by the narrow Raudfjorden Graben (Fig. 2). In both these horsts, thick schist and marble successions pass down into migmatites intruded by granites. Caledonian metamorphism is widespread, as are Caledonian granites at deeper structural levels. In the eastern, Biscayerhalv0ya-Holtedahlfonna Horst (a narrow belt, c. 20 km wide, flanking the Andreeland-Dicksonland Graben), the schist and marble succession dominates the southern parts, occurring in a large antiform cored by migmatite. Locally (e.g. Liefdefjorden), Ohta & Larionov (1998) have shown that some granites intruded at c. 960 Ma, using the single-grain zircon-Pb-evaporation method. In northern parts (Biscayerhalv0ya) of the horst, the schistmarble association is overthrust, apparently from the east, by an eclogite-bearing association (Richarddalen Complex) of hornblendic gneisses, marbles and other sedimentary rocks (Gee 1972). Augen gneisses, derived from c. 960 Ma porphyritic granites, locally in association with gabbros of similar age (Peucat et al 1989), are intercalated with these gneisses. Isotope dating (U-Pb method on zircons and titanites) of the eclogite-facies metamorphism (Peucat et al 1989; Gromet & Gee 1998) has proved difficult, but it has been concluded that this high-grade metamorphism was probably of late Ordovician (c. 455 Ma) age (Gromet & Gee 1998), the mafic protoliths being derived from Neoproterozoic intrusions. The higher pressure of the Richarddalen metamorphism contrasts markedly with the regional metamorphism of most of Svalbard's northwestern Caledonides, where high greenschist to amphibolite facies dominates, with relatively high 7/low P parageneses and the development of migmatites at deeper structural levels. The combination of major differences in both protolith and metamorphism between the Richarddalen Complex and all the other Caledonian units in northwestern Spitsbergen provide the basis for treating the northwestern province as representing at least two terranes. The character of the basement beneath the AndreelandDicksonland Graben is of importance for interpreting Svalbard's Caledonian tectonics. The inferred west-vergent thrusting of the eclogite-bearing Richarddalen Complex, indicates that this allochthon may be present to the east of the BBF. Another line of evidence was provided by Hellman et al (1998), who dated quartz porphyry boulders in conglomerates, which unconformably overlie the crystalline rocks of the Biscayerhalv0ya Horst. These clasts were selected for isotope-age studies because they were deposited close to the BBF and are totally foreign to Svalbard's northwestern terranes, both in composition and also very low metamorphic grade. An age of c. 1735-1740 Ma was obtained from these boulders of felsic volcanites that are inferred to have been derived from the Caledonian basement beneath the ORS, close to the east of the BBF. Thus, there are two lines of evidence concerning the basement of the Andreeland-Dicksonland Graben. These indicate that it is composed, at least in part, of a high P/ high T allochthon of Meso-Neoproterozoic crystalline rocks,

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and also low-grade Palaeoproterozoic volcanites, the latter being similar in age to the granitic orthogneisses of the West Ny Friesland Terrane. To the west of the Raudfjorden Graben (Fig. 2) and flanking its western margin, northwestern Spitsbergen is dominated by migmatites, schists and marbles (Gee & Hjelle 1966; Hjelle & Ohta 1974). major upright to west-vergent folds plunge gently to the south and deform the metasedimentary units and the migmatites; the latter and associated granites dominate the northern areas. Some of the metasedimentary rocks, such as the marbles and quartzites, can be followed deep into the migmatites, leaving no doubt that the migmatization influenced the entire sedimentary succession. All metamorphic transitions have been reported from pelitic schists to neosome-bearing paragneisses to migmatites with variably segregated granitic neosome. Granites occur as discrete intrusions and have been dated by the single-zircon, Pbevaporation method to c. 420-430 Ma (Peucat in Balashov et al. 19960; Ohta et al 2002). A younger batholith dominates the Caledonian intrusions—the Hornemantoppen Granite, and has yielded a Rb-Sr isochron age of c. 415 Ma (Hjelle et al 1979; Balashov et al 19960). As in the case of the Biscayerhalv0ya-Holtedahlfonna Horst, referred to above, granitic intrusions of late-Grenvillian age (c. 960 Ma) have been detected (Ohta et al 2002) locally within the migmatites. The age of the metasedimentary rocks is probably late Mesoproterozoic, but not well defined. Single-zircon, Pbevaporation studies of detrital zircons (Larionov, in Ohta et al 2002) provide evidence of late Archaean and Palaeoproterozoic source areas for the sedimentary formations; a few Mesoproterozoic crystals occur, but none as young as 960 Ma. Thus, although the widespread migmatization of northwestern Spitsbergen is very probably of Caledonian age, the possibility remains that this tectono-thermal event was superimposed on a Grenville-age episode of similar P and T. A variety of Kr/Ar and Ar/Ar ages on micas and hornblendes and Pb-evaporation zircon data provide evidence of the widespread Caledonian, c. 430-410 Ma, metamorphism of Svalbard's Northwestern terranes (Gayer et al 1966; Dallmeyer et al 1990a). Unconformably overlying these are ORS units (Siktefjellet Group) of earliest Devonian or late Silurian age (Gee & Moody-Stuart 1966; Friend et al 1997). If the Silurian-Devonian boundary is dated at 418 Ma (McKerrow & van Staal 2000), then metamorphism of this terrane occurred in the Silurian. Only in the Richarddalen Complex is there unambiguous evidence of an earlier, probably mid-late Ordovician event.

Southwestern terranes Whereas the Caledonian terranes of northwestern Svalbard and further to the east are little influenced by post-Devonian deformation, all the areas south of Kongsfjorden (Fig. 2) were involved in the West Spitsbergen Tertiary fold-and-thrust belt (Harland & Horsfield 1974; Dallmann et al 1993). Interpreting the Palaeozoic tectonic history is, therefore, more difficult than in other areas. Nevertheless, despite the superimposed Tertiary deformation, it has been possible to reconstruct Proterozoic and Palaeozoic histories some of which are similar to those of terranes further NE, and others that are clearly unrelated and 'exotic'. In this account, two major complexes are distinguished, the one occurring north of Isfjorden and including a blueschist-eclogite allochthon and the other further south and apparently part of the Laurentian platform. In the area of western Spitsbergen, south of Isfjorden to S0rkapp, a Proterozoic succession can be reconstructed (Birkenmajer 1975, 1981; Bj0rnerud 1990; Ohta and Dallmann 1999) that involves both a Grenville-age basement complex and Neoproterozoic cover. The latter is overlain by Vendian tillites,

except in southernmost areas, where Cambro-Ordovician platform successions unconformably overlie the Neoproterozoic rocks (Birkenmajer 1991). Although the Vendian and Early Palaeozoic units are, in general, comparable with those of the Nordaustlandet Terrane, there are some significant differences that throw light on Svalbard's connections with northeastern Greenland. The complexity of the tectono-thermal histories of the Caledonian rocks south of Isfjorden, with some rock units having experienced Precambrian, Caledonian and Tertiary deformations, has implied that this part of Svalbard has been particularly problematical both for basic geological mapping and for terrane interpretations. Caledonian terrane boundaries projected through the area (e.g. Harland 1972, 1985) have been rejected by those involved in subsequent detailed mapping (e.g. Dallmann et al 1990; Bj0rnerud et al 1991). The view that Wedel Jarlsbergland (Fig. 2) is divisible into western and eastern parts, separated by a major fault zone (Harland 1997) is not accepted here, and hence the southwestern terranes south of Isfjorden are treated as a single province. The oldest Precambrian units are exposed in southern Spitsbergen where a three-fold subdivision was described (Birkenmayer 1981, 1991), composed of metasedimentary rocks (Isbjornhamna Group, largely garnet-mica schists) overlain (perhaps overthrust) by various metamorphosed volcanites, gabbro-diorites and schists (Eimfjellet Group), and then siliciclastic metasediments (Deilegga Group). An attempt has been made (Balashov et al I996b) to date detrital zircons in the amphibolite facies Isbjornhamna schists, using the Pb-evaporation method on small populations (not single crystals); this imprecise method yielded ages of c. 1500 Ma. The overlying metavolcanites and intrusive rocks are a suite of mafic and felsic rocks, metamorphosed in greenschist facies, that have provided a consistent group of single-zircon Pb-evaporation ages of c. 1160 ± 40 Ma (Peucat in Balashov et al I996b). Other volcanic rocks have yielded c. 1200 Ma U-Pb zircon ages and a lower intercept age c. 930 Ma, thought to be the time of regional metamorphism (Balashov et al 1995). Relationships between the Deilegga Group and the inferred underlying units are not seen in the field, but the Deilegga phyllites have been assumed to be younger because of their lower metamorphic grade. Unconformably overlying the Deilegga Group and presumably also the Isbj0rnhamna and Eimfjellet groups is a c. 3 km thick Neoproterozoic succession (Sofiebogen Group) of conglomerates passing up into stromatolite-bearing limestones and dolomites and overlain by a fine-grained siliciclastic formation. The basal Sofiebogen facies of conglomerates and sandstones thickens southwestwards, suggesting a source from that direction. Tillites (Kapp Lyell Group) of inferred Vendian age overlie the Sofiebogen Group, except in the southernmost areas. The tillites of SW Spitsbergen are inferred to include a thick succession (3-4 km) of diamictites (Harland et al 1993). Their relationship to the overlying Cambro-Ordovician successions (Major & Winsnes 1955) is not unambiguously defined, the latter generally overlying Neoproterozoic formations with a tectonic contact. Birkenmajer (1960, 1991) has inferred primary unconformable relationships and a significant break (Jarlsbergian diastrophism), with phyllitic clasts of the underlying units occurring in the basal Cambrian conglomerates. Early Cambrian quartzites, dolomites, sandy limestones and dark shales compose the basal formation, which is overlain by Early Ordovician limestones and dolomites. A low-angle unconformity marks a Middle and Upper Cambrian hiatus. The fossiliferous Ordovician limestones, with unambiguous Laurentian faunas, are reported to be overlain by an unfossiliferous siliciclastic sequence (Dallmann et al 1993) and these Early Palaeozoic formations were strongly folded prior to the deposition of Devonian ORS. Whereas the rock units in areas of SW Svalbard described above can be broadly correlated with those of Nordaustlandet, the northern part of Svalbard's southwestern terrane (from Isfjorden to

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Kongsfjorden and including the island of Prins Karls Forland) has not yielded a coherent pre-Devonian history. Harland (1997) has emphasized the possibility that the Early Palaeozoic and older rocks may not have been influenced by Silurian deformation, but rather by a late Devonian episode which he correlated with the Ellesmerian Orogeny. One complex in western Spitsbergen (Motalafjellet, Fig. 2) is clearly exotic (Gee 1986). A blueschist-eclogite association referred to as the Vestgotabreen Complex (Kanat & Morris 1988; Ohta 1979), yielding white mica K-Ar (Horsfield 1972) and Ar/Ar (Dallmeyer et al 1990/?) uplift ages of c. 470 Ma, is unconformably overlain by polymict conglomerates and limestones with Middle Ordovician conodonts (Armstrong et al. 1986) and then turbidites of Late Ordovician or Early Silurian age (Scrutton et al. 1976). The high-P basement, with its Palaeozoic cover, is thrust over a thick diamictite of inferred Vendian age. The Vestgotabreen Complex is mainly composed of mica schists, with some amphibolitized eclogites; it also includes both fine-grained mafic units (perhaps sheeted dykes) and occasional serpentinites and gabbros. This subduction-related assemblage, probably including fragments of oceanic crust, is foreign to the dominating platform facies of Svalbard's Early Palaeozoic rocks, as is the timing of its deformation and metamorphism. Connection with the M'Clintock Orogen of Pearya (see below) has been proposed (Harland & Wright 1979; Ohta et al. 1989; Trettin 1989). Bj0rn0ya Two hundred and fifty kilometres south of Spitsbergen, on the western edge of the Barents Shelf (Fig. la), the small island of Bj0rn0ya provides further evidence for understanding the Svalbardian evolution. On the southern tip of the island, which is mostly dominated by ORS and younger strata, Neoproterozic and Ordovician successions (Holtedahl 1920; Krasil'schikov & Livshits 1974; Dallman & Krasil'schikov 1996) are exposed in west-vergent folds (Braathen et al. 1999). Late Neoproterozoic dolomites (Russehamna Formation) are overlain by sandstones and shales, with rare clasts, perhaps of glacial origin (Harland et al. 1993). Ordovician dolomites and limestones (Arenig to Llanvirn) of Laurentian affinities overlie these Neoproterozoic formations unconformably and both Cambrian and Tremadoc units are absent. Smith (2000) has drawn attention to the close similarities between the pre-Devonian stratigraphy of Bj0rn0ya and that of NE Greenland, both in terms of lithofacies and fauna, placing tight constraints on the location of Bj0rn0ya in the early Palaeozoic. Comments on interrelationships between the Svalbard terranes Although marked contrasts between Svalbard's Caledonian provinces have been recognized for many years, defining the boundaries between the different parts has proved controversial. Interpretation of the various provinces as far-transported, transcurrently displaced terranes intensified the debate. Extrapolation of the on-shore geology to neighbouring shelf areas has also been problematic, but helped by the release of aeromagnetic data (AMAROK 1994). Highly magnetic Caledonian granitoids have been identified on Nordaustlandet (Gee et al. 1999) that probably extend southwards to Edge0ya. Equally relevant has been the correlation of strong positive linear magnetic anomalies with the c. 1750 Ma magnetite-rich granites and orthogneisses of the Atomfjella Complex; these anomalies can be followed from Ny Friesland, far south beneath their post-Caledonian cover of

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southern Spitsbergen into the shallow Barents Shelf (Skilbrae 1992). Thus the western and eastern provinces, with their preCaledonian basement of Grenville-age complexes, are probably separated throughout Svalbard by a basement complex (the West Ny Friesland Terrane) that was untouched by Grenville-age tectono-thermal activity. Of the Svalbard terranes with Early Palaeozoic successions, three are clearly interrelated by their shelf facies and Laurentian faunas: Nordaustlandet, the southern part of the southwestern terranes, and Bj0rn0ya. As Smith (2000) has emphasized, the absence on Bj0rn0ya, of Cambrian and Tremadocian strata is a characteristic that binds this terrane more closely to northeastern Greenland than to Nordaustlandet. Southernmost Spitsbergen also has a Middle-Upper Cambrian to Tremadocian hiatus, but it is shorter (and less well constrained) than that on Bj0rn0ya; nevertheless it indicates similar affinities. In addition, in the case of southern Spitsbergen, there is a Neoproterozoic succession that can be correlated with both Nordaustlandet and northern Greenland, but is more similar to the latter (see below). The Vastgotabreen Complex (southwestern terranes) in central west Spitsbergen provides strong evidence for 'foreign' terranes on Svalbard (Gee 1986). Also in this area, north of Isfjorden and on Prins Karls Forland, there are a number of thick low-grade metasedimentary units, so strongly influenced by Tertiary and older faulting that it has proved impossible to define an agreed stratigraphy or correlate elsewhere on Svalbard (Harland et al. 1979; Hjelle et al. 1979). These complexes may have affinities with Pearya. Comparison with the East Greenland Caledonides The East Greenland Caledonides (Fig. 5) compose the northeastern part of the Laurentian continental margin, which extends southwards to the Appalachian foreland of North America. Early Palaeozoic platform carbonate-dominated successions are overthrust by extensive allochthons that were originally located tens to hundreds of kilometres to the east (Haller 1971; Henriksen 1985; Henriksen et al. 2000; Higgins & Leslie 2000). These hinterland-derived thrust-sheets are generally composed of Palaeoproterozoic and older crystalline rocks, overlain by MesoNeoproterozoic and Early Palaeozoic sedimentary successions. All these strata were originally deposited on the Laurentian craton, inboard of the Early Palaeozoic rifted margin. Outboard terranes of oceanic affinities have not been found in the Greenland Caledonides. The thrust-sheets have been categorized by Higgins etal. (2001) as 'thick-skinned' in the hinterland and 'thin-skinned' towards the foreland and transport to the WNW has been estimated to exceed 200km. These foreland and hinterland allochthons strike NNE at a small oblique angle to the coast of East Greenland and the eastern parts of the 'thick-skinned' units are developed only in the central parts of the orogen as far as about 76°N, from where they strike offshore into northeastern Greenland's wide continental shelf. Strong similarities between the Caledonian geology of eastern Greenland and that of Svalbard have been recognized since the early exploration of these areas in the 1920s and 1930s. Kulling's work on the Neoproterozoic and Early Palaeozoic successions of both central East Greenland (Kulling 1930) and Nordaustlandet (Kulling 1934) drew attention to the close comparability of their depositional environments. Subsequently, many authors (particularly Harland 1969 and thereafter) have drawn attention to and amplified these correlations. During the last decade, a wide range of new data on northeastern Greenland have further substantiated the similarity of the two areas. Whereas in the past this comparibility led to terrane hypotheses involving long distances of strike-slip displacement of the Svalbard terranes, on and offshore mapping of northeastern Greenland has shown that less complex hypotheses are possible (Fig. 6) and to be preferred (Gee 2001).

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ig. 5. Geology of East Greenland, based on Higgins & Leslie (2000).

Until recently, the most striking similarities between the Caledonides of Svalbard and East Greenland concerned the stratigraphical and sedimentological correlation of three major units: the Neoproterozoic Eleonore Bay Supergroup in Greenland (S0derholm & Jepsen 1991; S0derholm & Tirsgaard 1993; Jepsen & Kalsbeek 2000) with the Murchisonfjorden and Lomfjorden supergroups of Svalbard's Nordaustlandet Terrane; the Vendian tillite-bearing successions in both areas (Hambrey 1983); and the characteristic Cambrian sandstone to limestone-dolomite formations, the latter reaching into the Early Ordovician, that are

typical of the entire platform margin of Laurentia from East Greenland, via Scotland to eastern Canada (Swett & Smith 1972; Swett 1981). More detailed analyses of both Nordaustlandet and central East Greenland have shown that the similar Neoproterozoic successions in both these areas are underlain by nearly identical late Mesoproterozoic to earliest Neoproterozoic basement complexes. The latter are composed of thick siliciclastic units—the Krummedal and Smallefjord groups of East Greenland and the Brennevinsfjorden and Helvetesflya units of Nordaustlandet;

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Fig. 6. Laurentia-Baltica relationships in the Mesozoic (c. 250 Ma), with inferred correlation of Svalbard's Caledonian terranes with the Greenlandmargins and Pearya (reconstruction based on Roberts et al. (in press) and Scott & Turton 2001, with minor adjustments).

in both areas, these are intruded by megacrystic granites (extensively deformed to augen gneisses) of similar age, 930-940 Ma in Greenland (Watt et al. 2000) and 950 Ma in Nordaustlandet (Gee et al 1995; Johansson et al 2000, 2004). Remarkably, the detrital zircons in the metasedimentary host rocks in Greenland (Watt & Thrane 2001; Kalsbeek et al 2000) as on Nordaustlandet, are dominated by Mesoproterozoic ages as young as c. 10501100 Ma and a few as old as c. 1700 Ma; these sediments were largely derived from a Grenville-age source, little influenced by older Proterozoic and Archaean basement; they may have shared the same basin of deposition. North of 76° in the East Greenland Caledonides, the hinterland allochthons are dominated by highly deformed and metamorphosed Palaeoproterozoic basement. Between 76° and 78°N, the Caledonian metamorphism reaches eclogite facies (Gilotti 1993; Brueckner et al 1998; Elvevold & Gilotti 2000). Further north, amphibolite facies prevails and the Palaeoproterozoic crystalline rocks are thrust together with metasandstone successions that are thought to be of Mesoproterozoic age and correlatable with the Independence Fjord Group (Collinson 1980) of the western foreland and the lower 'thin-skinned' thrust sheets (Fig. 5). The Palaeoproterozoic gneisses of the hinterland allochthon have yielded a range of late Palaeoproterozoic ages (c. 1900-2100 Ma), and these older rocks are cut by c. 1750 ± 20 Ma granites (Kalsbeek et al 1999). The Independence Fjord Group is extensively intruded by dolerites of the Midsommers0 Suite that have been dated to c. 1250 Ma. The latter are probably feeders to overlying, c. 1200m thick basalts of the Ziz-Zag Dal Basalt Formation. In the hinterland allochthon, Kalsbeek et al (1999) reported metasandstones with

intercalated basalts and rhyolites, the latter yielding c. 1740 Ma U-Pb zircon ages at one locality (Pedersen et al 2002). Kalsbeek et al (1999) have suggested that the Independence Fjord Group may be Palaeoproterozoic in age. Another interpretation is given below that favours the existence of an older sandstone-basaltrhyolite association and a younger 'classical' Independence Fjord Group of Mesoproterozoic age. There are some remarkable similarities between the hinterland thrust-sheets of NE Greenland and the Atomfjella Complex of the West Ny Friesland Terrane. Although the Palaeoproterozoic thrust intercalations in Ny Friesland are mostly c. 1750 Ma in age (only one exceptional unit has been detected of Late Archaean age; Hellmann et al 2001), the intercalated, younger metasandstone units contain a large population of somewhat older Palaeoproterozoic detrital zircons (Hellmann et al 1997) that are closely similar to those in the NE Greenland hinterland basement (Kalsbeek et al 1999). In Ny Friesland, the c. 1750 Ma granites contain local xenoliths and large rafts of sandstone, that are older than the dominating Mesoproterozoic cover. Thus, there are close similarities between the West Ny Friesland Terrane and the hinterland allochthons of northern East Greenland. As in western Ny Friesland, the hinterland allochthon of NE Greenland (north of 76°N) has provided no evidence of Grenville-age tectono-thermal activity. In the same way as the West Ny Friesland Terrane is influenced by sinistral strike-slip faulting along Wijdefjorden and Billefjorden, so in NE Greenland the 'thick-skinned' hinterland thrust sheets are separated from the 'foreland' allochthon by a wide (locally 5 km) intense zone of sinistral shearing (the Storstr0mmen shear-zone of Holds worth & Strachan 1991). This

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shear-zone and related structures in the adjacent basement complex has many structural similarities with the western flank of the Atomfjella Complex in Ny Freisland and the Billefjorden Fault zone (Witt-Nilsson 1998). Whereas the folded and thrust Caledonian bedrock to the west of the hinterland 'thick-skinned' allochthons in northeasternmost Greenland is well exposed and documented, the potential equivalents on Svalbard are hidden beneath the ORS of the AndreelandDicksonland Graben. The evidence on Svalbard (referred to above) that the basement beneath this ORS cover includes c. 1740 Ma rhyolites suggests correlation with the rock units now known in NE Greenland. However, the proposed presence of an eclogite-bearing allochthon with Neoproterozoic protoliths in the basement of the Andreeland Graben, remains an enigma that is discussed further below. In the Caledonian foreland of northeasternmost Greenland, the Independence Fjord sandstones and Zig-Zag Dal basalts are unconformably overlain by a Neoproterozoic succession, the Hagen Fjord Group, that can be correlated with the Sofiebogen Group of SW Spitsbergen; the latter is comparable in general, but not in detail, with that of the Eleonore Bay Supergroup of East Greenland (S0nderholm & Jepsen 1991) and Nordaustlandet's Muchisonfjorden Supergroup. The Sofiebogen Group was deposited on a basement of Mesoproterozoic metasedimentary and metavolcanic rocks that were strongly influenced by late Mesoproterozoic deformation and regional high greenschist- to amphibolite-facies metamorphism. This basement complex is apparently foreign to the pre-Caledonian basement of northeasternmost Greenland, but is overlain by Cambrian and Early Ordovician strata of Laurentian platform affinities.

Comparison with the North Greenland fold belt and Pearya The North Greenland fold belt trends E-W along the northern margin of Greenland (Higgins et al. 1985) and is an eastern extension of the Ellesmerian orogenic belt of high Arctic Canada. In northernmost Greenland's deep Franklinian Basin (Higgins et al. 1991), deposition continued from the earliest Palaeozoic until the Early Devonian; the Ellesmerian deformation is inferred to have occurred in the Late Devonian. However, in Arctic Canada, the Ellesmerian orogen provides evidence of both 'classical' Caledonian-age orogeny (Late Silurian to Early Devonian) with ORS deposition, and also later pulses of deformation through the Devonian, during the final phases of the orogeny (Trettin 1989). In the most relevant area for this paper, the northeastern corner of Greenland, there is strong evidence that advance of the Caledonian nappes was providing a source of turbidites into the North Greenland trough from the Middle Silurian onwards (Peel & S0nderholm 1991). Higgins et al. (2001) draw attention to the lack of evidence of Precambrian basement along the outermost margin of North Greenland fold belt; most previous workers have accepted that the Franklinian trough defines the southern margin of an Early-Mid Palaeozoic ocean. The Franklinian Basin of North Greenland is a spectacularly well preserved trough with an early extensional history in the Cambrian, a starved basinal period in the Ordovician to Early Silurian and then a reactivation of subsidence during the rest of the Silurian and Early Devonian. The Laurentian platform margin rifted and collapsed in the Early Cambrian, feeding several kilometres of turbidites into the basin. After an Ordovician interval of mainly shale accumulation, turbidite-deposition again dominated, in response to further collapse of the outer shelf. The basinal successions were thrust southwards onto the southern shelf in the Late Devonian. Northerly vergence characterizes the outermost northern margin of the North Greenland fold belt, where Tertiary thrust systems are also prominent.

From northern Greenland, the Franklinian Basin passed westwards into Ellesmere Island, where the sediments are overthrust in the north by the Pearya Terrane (Trettin 1987). Much of the pre-Carboniferous history of Arctic Canada is buried beneath the vast Sverdrup Basin, but on northern Axel Heiberg Island, Old Red Sandstones have been shown to overlie older Palaeozoic strata unconformably. From northern Ellesmere Island, Trettin (1987, 1989) has described the Pearya Terrane to be composed of a Mesoproterozoic basement complex, overlain by Neoproterozoic shallow marine successions and overthrust by an Early Ordovician fragmented ophiolite, with some associated Ordovician volcanites of calcalkaline affinity. Younger Ordovician and Silurian siliciclastics, carbonates and volcanites overlie the older components of the Pearya Terrane with major unconformity. Reliable age control is limited and reconstructions of the terrane evolution are therefore provisional. Pearya's crystalline basement is composed of granitic gneisses, with subordinate amphibolites and minor quartzites, schists and marbles, the entire association apparently having been subject to amphibolite facies metamorphism. The granitic rocks, in one area, have yielded a Rb-Sr isochron of c. 1060 Ma (Sinha & Frisch 1976) and in another location, a c. 1040 Ma U-Pb zircon upper intercept age (Trettin et al. 1987). Late Mesoproterozoic, upper intercept, U-Pb zircon ages on Palaeozoic and younger igneous rocks support the interpretation that the Pearya basement is Mesoproterozoic in age, i.e. it is foreign to the crystalline basement of most of northern Laurentia, but not to Svalbard's southwestern terranes, which were also a part Laurentia (see above). Overlying the Pearya basement with inferred unconformity (the contact is usually faulted) is a largely metasedimentary succession. Although the details of this are little known, it includes arkoses, thick diamictites (perhaps glacial and Vendian), carbonates, schists and quartzites, in all c. 6-8 km thick, and mostly metamorphosed in greenschist facies. A rhyolite near the top of this shelf succession has yielded a U-Pb zircon age of c. 500 Ma (Trettin et al. 1987). The Mesoproterozoic Pearya basement, with its shallow marine cover, is overlain technically by a deeper marine assemblage of inferred Early Ordovician age. This includes both basalts and andesites of calc-alkaline affinity and slices of serpentinites and gabbros with associated trondhjemites and other acid-intermediate intrusions that may, in part, represent a fragmented ophiolite; zircons in one of the acid intrusions yielded a U-Pb zircon age of c. 480 Ma (Trettin et al. 1987). Emplacement of this complex of probable oceanic and island arc affinities on to the Pearya shelf occurred late in the Early Ordovician, prior to c. 460 Ma. This event is referred to in Canada as the M'Clintock Orogeny and was followed by the rapid accumulation of a thick (7-8 km) siliciclastic (including turbidites), carbonate and volcanic succession of Middle Ordovician to Late Silurian age. This mid Palaeozoic succession, resting with major unconformity on the Early Ordovician and older rocks, appears to have accumulated in isolation from the Franklinian Trough and Trettin (1989) has suggested that the Pearya terrane assembly did not dock with Laurentia before the Early Devonian; he proposed that east-west trending strike-slip sinistral displacements controlled the docking of Pearya with the Laurentian margin. Pearya is clearly a composite terrane. The comparability of Pearya's Ordovician history with that of Svalbard's exotic, subduction-related assemblage in western Spitsbergen (southwestern terranes) is striking (Harland & Wright 1979; Ohta et al. 1989; Trettin 1989; Ohta 1994). In both Pearya and western Spitsbergen, allochthons derived from oceanic or subduction-related environments were emplaced onto Neoproterozoic shelf assemblages in the Early Ordovician (c. 470 Ma), prior to deposition of siliciclastic and carbonate units as old as Caradoc. Thus an episode of deformation, late in the Early Ordovician (the M'Clintock

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Orogeny) is important in both areas and clearly foreign to both the North Greenland margin (Franklinian Basin) and Svalbard's Laurentia-related terranes. The earlier history of Pearya is less easily correlated with Svalbard's Caledonian geology, but has some resemblance to the southwestern terranes. Some of the inferred Neoproterozoic to Cambrian successions on northern Ellesmere Island, that compose the footwall to Pearya's ocean-derived allochthons, have similarities with those of central-west Spitsbergen, e.g. the thick diamictites of inferred Vendian age and the underlying Neoproterozoic succession. The underlying Mesoproterozoic basement on Pearya has a clear Grenville-age signature, with granite intrusions at c. 1050 Ma. This is a hundred million years older than the widespread granitic intrusions of Svalbard's eastern and northwestern terranes, and a hundred million years younger than the magmatic suites (in the Eimfjellet Group) of southwestern Spitsbergen. Thus, although a general 'Grenville' signature unites these areas, closer correlation will require a more detailed knowledge of the bedrock geology.

Assembly of Svalbard's terranes The evidence presented above and summarized in Figure 7, provides a basis for reconsidering the different hypotheses for Caledonian terrane assembly on Svalbard. Previous hypotheses (e.g. Harland 1985), involving very large (about 1000km) strike-slip displacements of eastern terranes from the central East Greenland margin, involve great complexity (Fig. Ib). The recent mapping of eastern Greenland and correlations with Svalbard's eastern terranes suggest instead that the latter extended directly northwards from the shelf-edge of northeastern Greenland into the high Arctic (Fig. 6). Nevertheless, the evidence on both Svalbard and in eastern Greenland for sinistral strike-slip movements remains relevant to the discussions of terrane assembly, which may have occurred partly by sinistral transpression. The summaries of the Proterozoic and Early Palaeozoic histories of Svalbard, east and north Greenland and northern Ellesmere Island, presented above (Fig. 7), have drawn attention both to remarkable similarities and also some differences; they raise questions, the more notable of which are considered below. Fundamental to the discussions that follow is the concept that Laurentia and Baltica were independent palaeocontinents, at least by the mid Cambrian (perhaps as early as latest Vendian) that were distinguished by characteristic Cambro-Ordovician platform lithofacies and related endemic faunas (Swett 1981; Cocks & Fortey 1998). With regard to the correlation of Svalbard's eastern terranes and the hinterland, 'thick-skinned' allochthons of East Greenland, not only are the details of the rock units, their structure and geochronology similar, but also the contrast in both areas between their eastern and western parts. Whereas on Svalbard this contrast has been regarded as profound and requiring independent terrane identifications (the West Ny Friesland and Nordaustlandet terranes), in east Greenland comparable juxtapositions have been described as tectonic and the question left open. In east Greenland, evidence of late-Grenvillian Orogeny has been questioned. Thus, there may be some differences between these areas and those of Nordaustlandet, where the c. 950 Ma granite magmatism was syntectonic and followed immediately after calc-alkaline volcanicity (apparently absent in Greenland). An enigma of the mid-late Proterozoic history is the evidence in southwestern Spitsbergen of a Grenville-age basement overlain by shallow marine successions that can be correlated with the Hagen Fjord Group of NE Greenland. The fauna in the overlying Cambro-Ordovician strata define Laurentian affinity. However, evidence of Grenville-age tectono-thermal activity in the foreland

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and allochthons of northeasternmost Greenland is lacking. The southwestern Spitsbergen record implies that a late Mesoproterozoic complex was accreted to the Palaeoproterozoic craton of NE Greenland very early in the Neoproterozoic, probably during late Grenvillian orogeny. The inferred location of Svalbard's characteristic Palaeo-, Meso-, and Neoproterozoic rock units to the NE of Greenland during Caledonian Orogeny, raises the question as to whether this northerly extension of the Caledonides existed in the late Neoproterozoic or was the result of Caledonian Orogeny. If only the latter, it would imply that transpression was an important component of the Caledonian collision in this part of the orogen and resulted in sinistral strike-slip displacements of the hinterland terranes. With the evidence both on Spitsbergen (Billefjorden and Breibogen-Bockfjorden faults) and East Greenland (Storstr0mmen shear zone) of important zones of transcurrent sinistral shearing, this remains a possibility. A simpler alternative is that the Svalbard units existed as a northeastern (present coordinates) projection of the Greenland margin, prior to Caledonian Orogeny, as a result of the pre-Caledonian rifting and development of the Early Palaeozoic oceanic basins. A tectonic scenario for the Early Cambrian that appears to be generally accepted is that rifting and collapse of both the eastern and northern shelves of Greenland led to development of ocean basins flanking these margins (e.g. Torsvik et al. 2001). By the Early Ordovician, these oceans were substantial global features. Closure of the lapetus Ocean marginal to eastern Greenland, occurred during the latter part of the Ordovician and Silurian and culminated in Scandian collision (Gee 1975) between Baltica and Laurentia. In northeasternmost Greenland, this collision started in the Early Silurian, with late Llandovery flysch deposition in the foreland basin and extensive turbidite influx into the Franklinian Basin. Svalbard's present position, trending NW at nearly a right angle to the NE Greenland foldbelt, requires anticlockwise rotation of these Caledonian terranes during the mid Palaeozoic, prior to collision with the Pearya-related complexes of western Spitsbergen and closure of the Franklinian Basin.

Relationship of the Svalbard Caledonides to the Timanides The evidence provided above demonstrates that Svalbard's Caledonian lithologies and structure correlate in remarkable detail with the Palaeozoic-deformed outer margins of northern Laurentia. Unambiguous evidence (ophiolites, etc) throughout the Scandinavian Caledonides that the Baltoscandian margin of Baltica was separated from Laurentia by oceanic domains through at least the latter part of the Cambrian and all the Ordovician supports the faunal evidence that Baltica and Laurentia were isolated continents for much of the Early Palaeozoic. The Timanide Orogen developed along the northeastern margin of the Fennoscandian Shield during the Vendian and is overlain by Late Cambrian and younger Palaeozoic successions of unquestioned Baltica affinities—an eastern extension of the Baltica platform facies (Bogolepova & Gee 2004). In combination with the evidence of sutures in the Scandes, these lines of evidence eliminate the possibility of an extension of the Timanide Orogen across the Barents Sea into the Svalbard Caledonides. The absence of evidence of Vendian Orogeny on Svalbard (only in southernmost Spitsbergen is there evidence of a significant Vendian unconformity) likewise testifies against correlations with the Timanides. Thus, the Mesozoic and younger strata of the Barents Sea cover an important NNE-trending suture zone dividing the pre-Devonian basement into two major domains, Laurentia in the NW and the Timanide Orogen of Baltica in the SE.

Fig. 7. Correlation of the Svalbard terranes with northern Laurentia and Pearya.

SVALBARD: LAURENTIAN CALEDONIAN MARGIN

Conclusions For the last thirty years, Svalbard's Caledonian provinces have usually been interpreted in terms of far-transported terranes, with most having affinities to the East Greenland Caledonides, but some to the North Greenland Foldbelt and Pearya (northern Ellesmere Island). The reassessment of existing information presented here confirms the comparability with Laurentia; indeed some of the previous correlations find remarkable support in the new databases published over the last decade from Svalbard and East Greenland. However, the terrane hypothesis is not favoured by the new work. Mapping of East Greenland has shown that the Caledonian allochthons strike obliquely at a small angle to the coastline, implying that the Central East Greenland correlatives of Svalbard's Nordaustlandet Terrane trend offshore into the shallow continental shelf of northeasternmost Greenland. It is proposed here that Svalbard's eastern terranes (Nordaustlandet and West Ny Friesland terranes), northwestern terranes and southern part of the southwestern terranes, comprised a direct northern continuation of East Greenland's hinterland, 'thick- and thin-skinned' allochthons (Higgins & Leslie 2000). The geometry of this Svalbardian 'extension' of northeastern Laurentia, prior to Caledonian Orogeny, has yet to be defined precisely. WittNilsson et al (1998) showed that the west-vergent Caledonian thrusting of Svalbard's eastern terranes (comparable with that of East Greenland) initiated with orthogonal shortening and progressed into transpressive transcurrence. Thus the geometry of the Svalbardian 'extension' may have been significantly modified by sinistral transcurrent displacements, resulting from late Scandian transpression. The Caledonides of northeasternmost Laurentia may not have terminated north of Svalbard in an oceanic domain; indeed, the orogen probably continued (with the Laurentian part as a narrow prong) across the Lomonosov Ridge into what are now the continental margins of the Amerasia Basin (Embry 2000). Only the bedrock of central western Spitsbergen and particularly the Vestg0tabreen high-pressure subduction complex, thrust onto poorly defined Neoproterozoic and Early Palaeozoic shelf successions, appears to have direct affinities with Pearya and the northern margin of the Franklinian Basin. The presence of the Franklinian Basin, flanking the orogen to the west would have allowed, indeed controlled, the progressive anticlockwise rotation of the Svalbard terranes, from NNE-trending during Caledonian collision to NW-trending in the Late Devonian; amalgamation with Pearya occurred with the final closure of the basin during the Ellesmerian Orogeny. Our work over the last decade on Svalbard has been supported by our home institutions in Uppsala and Lomonosov and also by the Swedish Polar Research Secretariat. Funding from the Swedish Research Council and EUROPROBE's INT AS projects HALE and NEMLOR has also been important. Early versions of this manuscript have been improved by comments from Robin Cocks, Tony Higgins, Niels Henriksen, Ake Johansson, Robert Scott, Rob Strachan and Vicky Pease.

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OHTA, Y., LARIONOV, A. N., TEBENKOV, A. M., LEPVRIER, C., MALUSKI, H., LANGE, M. & HELLIBRANT, B. 2002. Single zircon Pb-evaporation and 40Ar/39Ar dating of the metamorphic and granitic rocks in north-west Spitsbergen. Polar Research, 21, 73-89. PEDERSEN, S. A. S., GRAIG, L. E., UPTON, B. G. J., RAMO, O. T., JEPSEN, H. & KALSBEEK, F. 2002. Palaeoproterozoic (1740 Ma) rift-related volcanism in the Hekla Sund region, eastern North Greenland: field occurrence, geochemistry and tectonic setting. Precambrian Research, 114, 327-346. PEEL, J. S. & S0NDERHOLM, M. 1991. Sedimentary basins of North Greenland. Bulletin Gr0nlands Geologiske Unders0gelse, 160,1-164. PEUCAT, J. J., OHTA, Y., GEE, D. G. & Bernard-GRiFFiTHS, J. 1989. U-Pb, Sr and Nd evidence for Grenvillian tectonothermal activity in the Spitsbergen Caledonides, Arctic Ocean. Lithos, 22, 275-285. ROBERTS, R. J., TORSVIK, T. H., ANDERSEN, T. B. & REHNSTROM, E. F. 2003. The Early Carboniferous Mager0y Dykes, Northern Norway: Palaeomagnetism and Palaeogeography. Geological Magazine, 140,443-451. SANDELIN, S., TEBENKOV, A. M. & GEE, D. G. 2001. The stratigraphy of the lower part of the Neoproterozoic Murchisonfjorden Supergroup in Nordaustlandet, Svalbard. Geologiska Foreningens Stockholm Forhandlingar, 123, 113-127. SCOTT, R. A. & TURTON, M. A. K. 2001. Mesozoic tectonic events in the North Atlantic and Arctic: stratigraphic response in an adjacent riftflank basin (Sverdrup Basin, Arctic Canada). Polarforschung 69, Jahrgang 1999, 73-84. SCRUTTON, C. T., HORSFIELD, W. T. & HARLAND, W. B. 1976. Silurian fossils from western Spitsbergen. Geological Magazine, 113, (6), 519-523. SINHA, A. K. & FRISCH, T. 1976. Whole-rock Rb/Sr and zircon ages of metamorphic rocks from northern Ellesmere Island, Canadian Arctic Archipelago. II. The Cape Columbia Complex. Canadian Journal of Earth Sciences, 13, 774-780. SKILBREI, J. R. 1992. Preliminary interpretation if aeromagnetic data from Spitsbergen, Svalbard Archipelago (76°-79°N): Implications for structure of the basement. Marine Geology, 106, 53-68. SMITH, M. P. 2000. Cambro-Ordovician stratigraphy of Bj0rn0ya and North Greenland: constraints on tectonic models for the Arctic Caledonides and Tertiary opening of the Greenland sea. Journal of the Geological Society, 157, 459-470. S0NDERHOLM, M. & JEPSEN, H. F. 1991. Proterozoic basin of North Greenland. In: PEEL, J. S. & SONDERHOLM, M. (eds) Sedimentary basins of North Greenland. Bulletin Gr0nlands Geologiske Unders0gelse, 160, 49-69. S0NDERHOLM, M. & TiRSGAARD, H. 1993. Lithostratigraphic framework of the Upper Proterozoic Eleonore Bay Supergroup of East and North-East Greenland. Bulletin Gr0nlands Geologiske Unders0gelse, 167, 1-38. SWETT, K. 1981. Cambro-Ordovician strata in Ny Friesland, Spitsbergen and their palaeotectonic significance. Geological Magazine, 118, 225-250.

SWETT, K. & SMITH, D. E. 1972. Paleogeography and depositional environments of the Cambro-Ordovician shallow marine facies of the North Atlantic. Bulletin of Geological Society of America, 83, 3223-3248. TEBENKOV, A. M. 1983. [Late Precambrian magmatic formations of Nordaustlandet.] In: KRASILSHIKOV, A. A. & BASOV, V. A. (eds) Geologiya Spitsbergena: sbornik nauchnykh trudov [The Geology of Spitsbergen: a collection of papers.] 'Sevmorgeo', Leningrad, 74-86 [in Russian]. TEBENKOV, A. M., OHTA, Y., BALASHOV, Y. A. & SIROTKIN, A. N. 1996. Newtontoppen granitoid rocks; their geology, chemistry and Rb-Sr age. Polar Research, 15, 67-80. TEBENKOV, A., SANDELIN, S., GEE, D. G. & JOHANSSON, A. 2002. Caledonian migmatization in central Nordaustlandet, Svalbard. Norsk Geologisk Tidskrift, 82, 15-28. TORSVIK, T. H., VAN DER Voo, R., MEERT, J. G., MOSAR, J. & WALDERHAUG, H. J. 2001. Reconstructions of the Continents Around the North Atlantic at About 60th Parallel. Earth and Planetary Science Letters, 187, 55-69. TRETTIN, H. P. 1987. Pearya: a composite terrane with Caledonian affinities in northern Ellesmere Island. Canadian Journal of Earth Sciences, 24, 224-245. TRETTIN, H. P. 1989. The Arctic Islands. In: BALLY, A. W. & PALMER, A. R. (eds) The Geology of North America—An overview. Boulder, Colarado, Geological Society of America, The Geology of North America, A, 349-370. TRETTIN, H. P., PARRISH, R. & LOVERIDGE, W. D. 1987. U-Pb age determinations on Proterozoic to Devonian rocks from the northern Ellesmere Island, Arctic Canada. Canadian Journal of Earth Sciences, 24, 246-256. VERBA, M. L. & SAKOULINA, T. S. 2001. The reconstruction of the Early Paleozoic structure of the Barents Sea sedimentary Basin inferred from geophysical surveys along Profile I-AR. Polarforschung 69, Jahrgang 1999. II, 85-94. WALLIS, R. H. 1969. The Planetfjella Group of the lower Hecla Hoek of Ny Friesland. Norsk Polarinstitutt Arbok 1967, 80-108. WATT, G. R. & THRANE, K. 2001. Early Neoproterozoic events in East Greenland. Precambrian Research, 110, 165-184. WATT, G. R., KINNY, P. D. & FREDERICHSEN, J. D. 2000. U-Pb geochronology of Neoproterozoic and Caledonian tectonothermal events in the East Greenland Caledonides. Journal of Geological Society, London, 157, 1031-1048. WILSON, C. B. 1958. The Lower Middle Hecla Hoek rocks of Ny Friesland, Spitsbergen. Geological Magazine, 98, 89-116. WITT-NILSSON, P. 1998. The West Ny Friesland Terrane. An Exhumed Mid-Crustal Obliquely Convergent Orogen. Acta Universitatis Upsaliensis. Comprehensive Summaries of Uppsala Dissertations from the Faculty of Science and Technology 415, 1-28. WITT-NILSSON, P., GEE, D. G. & HELLMAN, F. 1998. Tectonostratigraphy of the Caledonian Atomfjella Antiform of northern Ny Friesland, Svalbard. Norsk Geologisk Tidsskrift, 78, 67-80.

Grenvillian and Caledonian tectono-magmatic activity in northeasternmost Svalbard AKE JOHANSSON1, ALEXANDER N. LARIONOV1'2, DAVID G. GEE3, YOSHIHIDE OHTA4, ALEXANDER M. TEBENKOV5 & STEFAN SANDELIN3 1 Laboratory for Isotope Geology, Swedish Museum of Natural History, Box 50 007, SE-104 05 Stockholm, Sweden (e-mail: [email protected]) 2 Now at Centre oflsotopic Research, Russian Geological Institute (USEGEI), St Petersburg, Russia 3 Department of Geophysics, Uppsala University, Villavdgen 16, SE-752 36 Uppsala, Sweden ^Norwegian Polar Research Institute, c/o IASC, Postboks 5156 Majorstua, N-0302 Oslo, Norway 5 Polar Marine Geological Expedition, Pobeda street 24, Lomonosov 189 510, Russia

Abstract: The Nordaustlandet Terrane of NE Svalbard forms an exposed part of the Barentsia microcontinent. Augen gneisses, migmatites, granites and gabbros dominate the scattered outcrops along the northeastern coast of Nordaustlandet, and on the smaller islands to the north and east, as far as Kvit0ya. These outcrops probably represent the deepest exposed crustal levels within the folded Caledonian basement of the Nordaustlandet Terrane. In the present study, a variety of rock types have been analysed by different U-Pb dating techniques (conventional, Pb-evaporation and ion microprobe) on zircon, titanite and monazite The major and trace element compositions and Sm-Nd isotope geochemistry of these rocks have also been investigated. The augen gneisses yield U-Pb ages of c. 950 Ma, indicating that they are deformed late Grenvillian granites, similar to the Grenville-age granites and augen gneisses of northwestern and central Nordaustlandet. Migmatites, grey granites, aplitic dykes and a syenite (boulder) yield U-Pb ages mainly falling in the 430-450 Ma range, slightly older than the 410-420 Ma late-tectonic Caledonian granites further west. Both the Grenvillian and Caledonian granites are of crustal anatectic origin, and the Caledonian granites and migmatites may have formed largely by remelting of Grenvillian crust. The ages of the mafic rocks are uncertain, but Sm-Nd data indicate a possible emplacement age of c. 700 Ma for two of the gabbros, suggesting that they may be the result of rift-related magmatism in connection with the opening of the lapetus Ocean. A few enigmatic inherited zircons of similar late Neoproterozoic age found in younger granites may possibly be related to this event. No evidence for late Neoproterozoic orogenic activity, similar to that in the Timanides of northern Russia, is seen in eastern Svalbard. At this time, eastern Svalbard (Barentsia) was probably part of the Laurentian margin, and probably located far away from northern Baltica.

The island of Nordaustlandet, together with adjacent smaller islands and the eastern part of Ny Friesland, form the eastern part of Svalbard's Eastern Caledonian Terrane (e.g. Harland 1972, 1997; Gee 1986; inset in Fig. 1). Research in recent years has led to the subdivision of the Eastern Terrane into a separate Nordaustlandet Terrane, encompassing the above areas, and a West Ny Friesland Terrane (Gee et al 1995; Witt-Nilsson 1998). The Nordaustlandet Terrane forms a part of the Barentsia microcontinent, which otherwise encompasses the northwestern part of the Barents Shelf to the south and east of Svalbard (cf. Gee & Ziegler 1996; Gudlaugsson et al 1998). Recent structural, stratigraphic and isotopic studies in western and central Nordaustlandet have shown that the basement is composed of a Grenville-age (c. 950 Ma) complex of metasedimentary, metavolcanic and intrusive granitic rocks (Gee et al. 1995; Gee & Tebenkov 1996; Johansson et al 2000), overlain by Neoproterozoic and Early Palaeozoic platformal sediments (Kulling 1934; Flood et al 1969; Ohta 1982; Sandelin et al 2001), and intruded by Caledonian granitoids (Gee et al 1999; Johansson et al 2002). Caledonian metamorphism generally is greenschist facies, but in central Nordaustlandet, evidence for Caledonian migmatization has been found (Tebenkov et al 2002). Further east along the north coast of Nordaustlandet into Orvin Land east of Duvefjorden (Fig. 1), various granitic gneisses and migmatites dominate the bedrock and have been grouped together as the 'Duvefjorden Complex' by Tebenkov (in Dallmann & M0rk 1991). Similar rocks (augen gneisses, migmatitic paragneisses, red and grey granites, aplitic dykes, rare amphibolites), as well as massive gabbros, also occur in the scattered outcrops along the east coast of Nordaustlandet (Nordmarka, S0rmarka, Isispynten) and on the islands to the north (Sju0yane, Foyn0yane) and east (Stor0ya, Kvit0ya; Fig. 1). Detailed descriptions have been made of some of these outcrops (Sandford 1950, 1954; Hjelle et al 1978; Hjelle 1978a, b) and rock types (Ohta 1978), as well

as early reconnaissance Rb-Sr dating (Hamilton & Sandford 1964) generally yielding Caledonian ages. However, the remoteness of the area, the sparsity of outcrops separated by vast expanses of glaciers and water, and the lack of modern geochronological data, have all contributed to a lack of comprehensive maps and descriptions of the area as a whole. Two alternative hypotheses have been suggested for the geotectonic setting of northeasternmost Svalbard in Caledonian time: it may either have been a stable foreland area to the Caledonian Orogen, composed of Grenvillian or older basement rocks, with Caledonian deformation and metamorphism waning towards the east, or alternatively, it may represent the strongly metamorphosed and migmatized core zone of the Caledonian Orogen. In the present paper, samples of various rock types from northeasternmost Svalbard, collected during Russian and Swedish fieldwork in recent years, have been used for U-Pb geochronological studies. During earlier conventional U-Pb multigrain zircon studies in western and central Nordaustlandet, the presence of inherited zircons both in the Grenvillian and Caledonian magmatic rocks, often has resulted in ages that were erroneous or uninterpretable (Johansson et al 2000, 2002). Therefore, the emphasis in this study has been on single-grain Pb-evaporation and ion microprobe spot analyses of individual grains. These studies have been complemented with major and trace element geochemistry and Sm-Nd isotope investigations of whole rock samples, in order to compare these rocks with the established igneous suites of western and central Nordaustlandet. Structure of the Nordaustlandet Terrane The overall structure of the Nordaustlandet Terrane is one of north-south-trending upright to west-vergent Caledonian anticlines and synclines, with rocks of the Grenvillian complex as well

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 207-232. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Simplified geological map of Nordaustlandet and adjacents islands (based on Flood et al 1969; Hjelle & Lauritzen 1982; Lauritzen & Ohta 1984), with sample locations. Inset map of Svalbard shows Caledonian terranes (from Gee 1986), main Caledonian fault lines (RF, Raudfjorden Fault; BBF, Breibogen-Bockfjorden Fault; BF, Billefjorden Fault) and major Caledonian granitoids (Hornemantoppen batholith on NW Spitsbergen, Newtontoppen granite in Ny Friesland, and Rijpfjorden granite on Nordaustlandet).

as Caledonian granites being exposed in the anticlinal areas, and Neoproterozoic and younger sediments in the intervening synclines (Fig. 2). The folds plunge gently southwards, perhaps as a result of east-west crossfolding (Hjelle 1978£), or southwards tilting of the whole basement complex during Tertiary uplift of Svalbard, related to the opening of the Arctic and North Atlantic Oceans. As a result of the latter, the Caledonian basement complex disappears under Carboniferous and younger sediments in southern Nordaustlandet, whereas progressively deeper basement levels are exposed towards the north. At the same time, a slight westward tilt of the Nordaustlandet basement complex may be suggested, perhaps related to extension along the Eolussletta shear zone in eastern Ny Friesland in late Caledonian time (cf. Witt-Nilsson et al. 1998). In the westernmost syncline, the Hinlopenstretet Syncline, the whole of the Neoproterozoic Murchisonfjorden Supergroup as well as the overlying Vendian to Ordovician Hinlopenstretet Supergroup are preserved, whereas in central Nordaustlandet only the basal parts of the Murchisonfjorden Supergroup remain. (Fig. 2). Further east, no Neoproterozoic or younger sediments have been observed with certainty; instead deeper basement levels with granites, augen gneisses and migmatites as well as mafic rocks are exposed. According to Sandford (1954), the Carboniferous cover probably overlies the igneous and metamorphic rocks in eastern Nordaustlandet directly, without any intervening 'Hecla Hoek' rocks (i.e. Neoproterozoic metasediments), indicating that the latter had been eroded away prior to the Carboniferous transgression. The Nordaustlandet structure is illustrated schematically in Figure 2. The overall result of this 'double tilting' of the Caledonian basement, is that the deepest, most highly metamorphosed levels of basement are exposed in easternmost Nordaustlandet and on the smaller islands to the north and east. It now remains to

establish whether this deep basement represents Grenvillian basement with little superposed Caledonian metamorphism and deformation, or a highly metamorphosed root zone to the Caledonian Orogen. The western boundary of the Nordaustlandet Terrane extends north-south through central Ny Friesland, probably being defined by the thrust at the base of the Planetfjella Group (cf. Gee et al 1995, 2001). In northern Ny Friesland (Mosselhalv0ya) ultramafic lenses are associated with this thrust (Witt-Nilsson et al. 1998). The Nordaustlandet Terrane extends southwards and eastwards from Nordaustlandet, and probably also northwards to the shelf-edge of the Barents Platform. Regional aeromagnetic anomalies (AMAROK A.S. 1994) provide a basis for relating onshore and offshore geology, and recent work (Gee et al. 1999) has shown that a north-trending belt of conspicious positive anomalies located in central Nordaustlandet is related to late Caledonian intrusions, the strongest anomaly being identified on land to be of quartz -syenitic to quartz-monzonitic composition (the Djupkilsodden pluton). This belt of anomalies can be followed southwards from Nordaustlandet to Barents0ya and perhaps to Edge0ya (see fig. 2 in Gee et al. 1999). Easternmost Nordaustlandet lacks major magnetic anomalies, but another north-trending belt of strong positive anomalies occurs offshore north and south of Stor0ya, and similar anomalies also occur further east towards Kvit0ya (AMAROK A.S. 1994; fig. 2 in Gee et al. 1999). However, the major gabbro massifs in Nordmarka and on Stor0ya leave no signature on the magnetic map. Magnetic susceptibility measurements in the Nordmarka area confirmed low values for the gabbro; indeed, the only rocks yielding a strong positive magnetic signature were boulders of syenite in the Quarternary glacial deposits. These boulders might be derived from the north-trending belt of magnetic

Fig. 2. Schematic E-W-profile across Nordaustlandet showing geological relations. ESZ, Eolussletta Shear Zone. Inset shows highly schematic block diagram across the Nordaustlandet Terrane from west to east, indicating the inferred westward tilt of the folded basement complex, as well as the southward plunge of the Caledonian anticlines and synclines, with the deepest crustal levels being exposed in the north and east. WNF, West Ny Friesland Terrane.

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bodies in the vicinity of Stor0ya. Like the Djupkilsodden pluton they are also of Caledonian age (see below). Local geology, rock types and samples Porphyritic granites and augen gneisses Coarse porphyritic grey to pink granites and augen gneisses have been observed on the Sju0yane islands, the Damflya area in Orvin Land, Nordmarka and S0rmarka (Hjelle 1978a, b\ Hjelle et al 1978; Sandelin 2001). The samples analysed here include a red and a grey augen gneiss from Parry0ya (one of the Sju0yane group of islands; samples 898:122 and 123), two samples of augen gneiss from Nordmarka (G95:049 and 050) and one sample of strongly foliated pink augen gneiss from Isispynten (898:130). These are composed of quartz, K-feldspar, plagioclase, muscovite and biotite, with accessory apatite and zircon and occasional garnet (sample G95:049). The size of the feldspar megacrysts may reach several centimetres. Petrographically, these rocks are comparable to the Grenvillian Laponiafjellet augen granite on Laponiahalv0ya (c. 960 Ma, Gee et al. 1995) and the Grenvillian Fonndalen and Ringgasvatnet augen gneisses of central Nordaustlandet (c. 950 Ma, Johansson et al. 2000). From Isispynten, Sandford (1954) described a 'pink granodiorite' crosscutting the other rocks, suggesting a late (Caledonian) age for this granite, and it may be that sample 898:130 represents a locally foliated variety of this younger granite, rather than a normal augen gneiss. Metasupracrustals, paragneisses, amphibolites and migmatites Much of the area is composed of a complex of migmatized paragneisses with inclusions of various metasupracrustals, as well as rare amphibolites and marble layers. Such rocks occur on Sju0yane islands, Damflya and other parts of Orvin Land, Nordmarka, Isispynten, and Andreeneset on Kvit0ya (Sandford 1950, 1954; Hjelle 19780, b\ Hjelle et al. 1978; this study). The field descriptions of the migmatites provide evidence of in-situ melting, segregation of neosome, and mobilization of melt to form anatectic granites. Pelitic palaeosomes contain metamorphic assemblages of high grade such as garnet-cordierite-sillimanite, with occasional orthopyroxene indicating granulite facies conditions (Hjelle et al 1978). The migmatites occur together with the above-mentioned augen gneisses as well as more finegrained massive grey granites, with which they have been grouped together as the 'Duvefjorden Complex' by Tebenkov (in Dallmann & M0rk 1991). The stratigraphic position of the metasupracrustal inclusions and the protoliths of the paragneisses is unclear: they may represent highly metamorphosed parts of the Neoproterozoic Murchisonfjorden Supergroup or the underlying, late Mesoproterozoic Brennevinsfjorden Group and Helvetesflya Formation, or some even lower stratigraphic units not exposed on the surface, as favoured by Hjelle (1978&). Our samples include a biotite schist from Parry0ya (898:124), containing quartz, K-feldspar, plagioclase, biotite, muscovite and garnet with accessory apatite, titanite, zircon and opaque minerals, and an amphibolite from Andreeneset, Kvit0ya (898:129), dominated by hornblende with some K-feldspar, plagioclase, biotite and minor quartz. Several samples of migmatized paragneiss from Andreeneset have also been investigated. Sample 898:128 is a relatively massive, dark biotite-rich rock with quartz, K-feldspar and plagioclase forming millimetre-sized aggregates, and minor hornblende, titanite and opaques, which appears to represent a migmatite palaeosome. Samples AJ94:004 and 94045 contain mixtures of dark palaeosome and lighter grey

or pink neosome sampled close to the Andree monument on Andreeneset (Fig. 3c, d), and are composed of quartz, K-feldspar, plagioclase, biotite and amphibole, with some opaques, apatite, titanite, and accessory zircon. A sample of migmatite from Damflya (31-5), collected by Russian colleagues, has also been investigated. Structural, stratigraphic and isotopic evidence at Innvika in central Nordaustlandet indicate the migmatization to be Caledonian, affecting the base of the Murchisonfjorden Supergroup and having an age of c. 420 Ma (Tebenkov et al. 2002). However, the presence of older episodes of migmatization in eastern Nordaustlandet and the islands beyond can not be excluded.

Migmatite neosomes, red and grey granites At the above-mentioned Innvika locality in central Nordaustlandet, the dated migmatite neosome (samples S98:049A and 050) in hand specimen appears as a massive, fine-grained grey granite, containing a mosaic of quartz, K-feldspar, plagioclase, muscovite and biotite. Medium-grained red and grey granite (samples 898:055 and 056) occur nearby, and all these rocks were included for geochemical and Sm-Nd isotopic investigation. Similar, massive to weakly foliated, fine- to medium-grained grey granites are common further to the north and east, and have been sampled at Parry0ya (898:121), Foyn0yane (898:125 and 126), Isispynten (898:131) and Andreeneset on Kvit0ya (aplitic granite 898:127). They are composed of quartz, K-feldspar, plagioclase and biotite, with muscovite sometimes occurring in more subordinate amounts. According to Sandford (1954), Hjelle (19780, b) and Hjelle et al. (1978), such granites are closely related to the migmatites and may thus be interpreted as mobilized migmatite neosome. At Parry0ya, the medium-grained grey granite cross-cuts the coarse augen gneiss (Sandelin 2001), whereas at Isispynten, the grey granite is cut by pink granodiorite of the Rijpfjorden granite type (Sandford 1954). The grey granites of northeasternmost Svalbard may tentatively be correlated with the Nordkapp granite in northwestern Nordaustlandet, dated to c. 440 Ma (U-Pb monazite; Johansson et al. 2002). Similar grey granites are also common on NW Spitsbergen, where they are cross-cut by the post-tectonic Caledonian Hornemantoppen batholith (c. 415 Ma, Rb-Sr; Hjelle 1979; Balasov et al. 1996). Aplitic dykes Cross-cutting pegmatitic and aplitic dykes are observed within the gneiss and migmatite complex at several localities: Sju0yane islands, Orvin Land, Nordmarka, Isispynten and Andreeneset on Kvit0ya (Hjelle 1978a, b\ Hjelle et al. 1978; this study). They are mostly highly leucocratic with grey-white or pink colour, although according to Hjelle et al. (1978), the grey-white dykes on Isispynten have a quartz-dioritic composition. Their undeformed and cross-cutting nature has led to the suggestion that they are related to the cross-cutting pink granodiorite or Rijpfjorden granite (Sandford 1954; Hjelle et al. 1978). In the present study, aplitic dykes have been sampled at Nordmarka (G95:051) and Andreeneset (AJ94:005). The Nordmarka dykes cut the augen gneiss, and are composed of mediumgrained mosaic of quartz, K-feldspar and plagioclase, with biotite as thin elongated laths, and muscovite as more scattered flakes. The dykes on Andreeneset are a few decimetres to about one metre wide, and cut the surrounding migmatitic paragneiss in different directions, with sharp contacts suggesting that they intruded after the migmatite had solidified completely (Fig. 3c). The sample is light pink, medium-grained and very leucocratic,

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Fig. 3. Field photographs of rock types from northeasternmost Svalbard. (A) Augen gneiss with cross-cutting aplite dyke, Nordmarka. (B) Grey granite with cross-cutting pegmatite dykes, Foyn0yane. (C) Aplitic dykes cross-cutting migmatitic paragneiss, Andree monument, Andreeneset, Kvit0ya. (D) Close-up of the migmatitic paragneiss at Andreeneset, Kvit0ya. (E) Layered gabbro, Norvargodden, Stor0ya. (F) Massive gabbro, Kraemerpynten, Kvit0ya.

consisting of quartz, K-feldspar and plagioclase, with a few scattered small flakes of biotite.

Syenite Several, usually well-rounded, boulders of red, massive, mediumand even-grained syenite were encountered in the Nordmarka area and one was sampled (G95:048). These rocks have the highest magnetic susceptibilities in the area. In thin section, the sampled rock consists of centimetre-sized hypidiomorphic K-feldspar crystals in a matrix of more fine-grained K-feldspar and subordinate plagioclase but no visible quartz, with relatively fresh to strongly chloritized biotite and almost totally altered amphibole. In another thin section, biotite occurs in subordinate amounts and is heavily altered, whereas ampibole is much more common and only

slightly altered, and forms large aggregates together with titanite and opaque minerals. Since the syenite has not been found in outcrop, nothing is known about its relation to other rock types in the area. However, its undeformed nature would suggest a young (Caledonian) age. As noted above these syenites may provide evidence for the character and age of magnetic bodies occurring offshore to the east of Nordaustlandet. Gabbroic rocks Unlike in the rest of the Nordaustlandet Terrane, gabbroic rocks form an important and distinctive element in the bedrock of northeastern Nordaustlandet and adjacent islands. Gabbroic bodies occur in eastern Orvin Land, and the northern half of Nordmarka consists of a massive, partly layered metagabbro (Hjelle 1978b;

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A. JOHANSSON ETAL

Hjelle et al. 1978). Diorite and amphibolite have been reported from Isispynten (Sandford 1954; Hjelle 1978&; Hjelle et al. 1978), as well as cross-cutting metaporphyrite dykes (Hjelle 1978Z?; Hjelle et al. 1978). The exposed part of Stor0ya is made up entirely of a large (at least 7 km across) layered gabbro-anorthositediorite complex, cross-cut by numerous east-west-trending dolerite dykes (Hjelle et al. 1978; Ohta 1978). Similar mafic rocks of noritic to dioritic composition are exposed on Kraemerpynten and Hornodden on Kvit0ya (Hjelle 1978&; Hjelle et al. 1978). A detailed petrographic and petrochemical investigation of the Stor0ya and Kvit0ya metagabbros has been made by Ohta (1978), who considered them to be late-tectonic Caledonian rocks due to their relatively undeformed nature, with wellpreserved magmatic layering ranging from olivine gabbro through norite to diorite on Stor0ya. According to Ohta (1978), they belong to the tholeiitic rock series, although they have been partly converted to calk-alkalic rocks by addition of sodium, and have possibly formed in a well-developed island-arc setting with thick continental crust. The Kvit0ya gabbros are cut by pink pegmatitic and aplitic dykes, suggesting that they formed prior to the latest phase of Caledonian granitic magmatism (Hjelle et al. 1978; Ohta 1978). In the present study, gabbroic rocks have been sampled at Isispynten (94062C), Norvargodden on Stor0ya (AJ94:003; Fig. 3e) and Kraemerpynten on Kvit0ya (AJ94:006; Fig. 3f). The Isispynten sample is a massive gabbro composed of pale green amphibole and brown biotite, with subordinate plagioclase and olivine and accessory opaques. The Norvargodden sample is a greenish, medium-grained, dioritic rock composed of plagioclase and K-feldspar as hypidiomorphic crystals showing preferred orientation, relatively altered amphibole, and minor opaques. The Kraemerpynten sample is a dark green, massive, mediumgrained gabbroic rock composed of plagioclase, K-feldspar, minor quartz, and two generations of amphibole: as large (1-2 mm), pale green, irregular and relatively altered grains similar to sample AJ94:003, and as smaller, partly idiomorphic, olive green fresh crystals in association with opaque minerals.

AFM diagram of Irvine & Baragar (1971), they mainly show a calc-alkaline trend. On the normalized trace and rare earth element spidergrams (Fig. 5a, b), there is understandable and considerable scatter. Some of the most deviating samples have been especially labelled: the more depleted gabbroo samples (AJ94:003 and AJ94:006) and the two aplite samples (AJ94:005oand G95:051), as well as the more enriched migmatite sample AJ94:004. Most of the granitic samples, however, show similar trace element patterns to the Grenvillian and Caledonian granitoids of western and central Nordaustlandet (Johansson et al. 2000, 2002), with negative spikes for Ba, Ni, Sr and Ti, as well as a more or less marked negative Eu anomaly on the REE diagram. TotaloREE contents vary widely between 14 ppm (gabbro sample AJ94:003) and 400 ppm (migmatite sample AJ94:004). Enrichment factors relative to chondritic compositions vary between 20 and 200 for La (c. 1.5 for AJ94:003), and 1 to 40 for Lu. On the tectonic discrimination diagrams by Pearce et al. (1984) and Pearce (1996), only the granitic samples (including migmatite neosomes, aplitic dykes, augen gneisses, and syenite) have been plotted (Fig. 6a-d). As with the Grenvillian and Caledonian granitoids further west (Johansson et al 2000, 2002), they plot in the volcanic arc or syn-collisional granite fields, that also overlap with the post-collisional granite field added by Pearce (1996) in Fig. 6c, making the results less conclusive. In this context one has to remember, however that samples representing two entirely separate orogenic cycles have been mixed on these plots. In summary, the granitic rocks from easternmost Nordaustlandet and adjacent islands show similar geochemical compositions to the Grenvillian and Caledonian granitoids further west, indicating a crustal anatectic origin from sedimentary precursors, with formation in a volcanic arc, syn- or post-collisional tectonic setting. Since the geochemistry of Grenvillian and Caledonian granites are so similar in the Nordaustlandet Terrane, it is difficult to use geochemical criteria to discriminate between them.

U-Pb and Pb-Pb dating Geochemistry The samples have been analysed for major and trace elements at SGAB, Lulea, Sweden, using ICP-AES and ICP-MS. The results are reported in Table 1 and illustrated in Figures 4-6. This is a very heterogeneous suite of samples, including some rocks (paragneisses) that are not even of igneous origin, but for comparative purposes all analysed samples have been plotted in Figure 4 (major elements) and Figure 5 (trace and rare earth elements). In the total alkali v. silica diagram (Fig. 4a; Le Maitre 1989), most of the augen gneisses, the red and grey granites, the migmatite neosomes and aplitic dykes plot as true granites, with a couple of samples falling in the syenite field. The syenite sample G95:048 falls on the border between syenite and monzonite, the three migmatite samples from Kvit0ya have a granodioritic composition and the gabbro and amphibolite samples plot in the gabbro field. In the P-Q diagram (Fig. 4b; Debon & Le Fort 1983), there is more scatter, from granite via adamellite and granodiorite to tonalite. Sample G95:048 plots as syenite, the Isispynten gabbro (94062C) plots as quartz diorite, whereas o the Andreeneset amphibolite (S98:129), the Stor0ya gabbro (AJ94:003) and the Kraemerpynten gabbro (AJ94:006) plot as diorite or gabbro. Most of the granitic rocks have peraluminous compositions (Fig. 4c, d), similar to the Grenvillian and Caledonian granitoids from western and central Nordaustlandet (Johansson et al 2000, 2002), suggesting a sedimentary source for the magmas. On the K2O v. Na2O diagram (Fig. 4e; Chappel & White 1974), however, they scatter across the I- to S-type boundary. On the

Analytical methods Because of the problems with inherited zircons, only a few conventional multi-grain U-Pb analyses have been made, using monazite or titanite. Methods have been described in Johansson et al (1995, 2000) and Larionov et al. (1995). For monazite, separate 233~235U and 206Pb tracers were used, for titanite a mixed 233-235u_208pb tracer. Uranium and lead were both analysed on a Finnigan MAT261 thermal ionization mass spectrometer in multicollector mode at the Laboratory for Isotope Geology in Stockholm. Results are reported in Table 2 and illustrated in Figure 7. Single zircon Pb-evaporation dating according to the method of Kober (1986,1987) were carried out on selected grains, being analysed on the Finnigan MAT261 mass spectrometer in Stockholm. Details of the analytical procedure are given in Larionov et al. (1998). Some single titanite and monazite grains were also analysed in the same way. Monazite, being a non-silicate, required addition of silica gel as an emitter to the mounted monazite grains, in order to achieve a stable Pb signal for analysis (Larionov 2000). The results, most of them plateau ages, are summarized in Table 3, and shown in histogram form in Figure 8. Weighted average 207Pb/206Pb ages for each sample were calculated from the age results of the indvidual grains using the ISOPLOT program of Ludwig (1991). U-Pb spot analyses on selected zircons were made using the NORDSIM Cameca 1270 ion microprobe in Stockholm, following methods described in Whitehouse et al (1997, 1999). U, Th and Pb contents and U/Pb ratios were calculated with reference to

213

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

tion has negligible effect. Cathodoluminescence images of selected analysed zircons with superposed ages are shown in Figure 9, to illustrate internal structure. For ages above 500 Ma, the 207Pb/206Pb age is shown, since it is independent of uncertainties in U-Pb calibration. However, for younger ages, the amount of 207Pb becomes very small and the measured uncertainty large,

the 91500 zircon standard, with an accepted age of 1065 Ma (Wiedenbeck et al 1995), that was analysed repeatedly during each session. The measured lead isotope ratios were corrected for common lead with an age of 420 Ma or 950 Ma (growth curve of Stacey & Kramers 1975), depending on the inferred geological age. However, in most cases the common lead correc-

Table 1. Chemical composition of rocks from central and eastern Nordaustlandet and adjacent islands, NE Svalbard Innvika

Locality Rock type Mig. neosome S98:49A Sample

Mig. neosome 898:50

Major elements (wt%) Si02 74.5 14.5 A12O3 CaO 0.688 1.38 Fe2O3-tot K2O 5.49 MgO 0.236 MnO 0.019 3.15 Na2O 0.257 P2O5 0.158 TiO2

74.2 14.6 0.831 1.34 5.53 0.277 0.026 3.46 0.154 0.180

Red granite 898:55

Grey granite 898:56

74.6 14.6 0.579 1.30 4.94 0.182 0.032 3.56 0.272 0.148

73.2 14.8 0.756 1.49 5.49 0.298 0.021 3.37 0.274 0.208

Parry0ya

Foyn0yane

Grey Augen Augen Biotite granite gneiss gneiss schist 898:121 898:122 898:123 898:124

Grey Grey granite granite 898:125 898:126

73.8 14.4 1.01 1.52 5.78 0.289 0.023 3.19 0.149 0.182

100.2 0.8

99.9 0.7

Trace elements (ppm) 210 Ba 4.60 Be <5.4 Co <5.4 Cu 21.4 Ga 2.25 Hf <2.1 Mo Nb 14.9 14.8 Ni 306 Rb 1.22 Sc 9.47 Sn Sr 53.8 1.38 Ta Th 7.71 U 4.89 V <2.1 3.25 W Y 12.9 45.8 Zn Zr 64.8

195 18.9 <5.2 <5.2 22.6 2.08 <2.1 13.2 13.5 321 1.48 19.3 53.5 3.03 4.77 14.1 3.85 4.29 7.81 91.2 58.9

201 7.18 <5.8 <5.8 25.7 3.37 <2.3 10.8 16.9 340 <1.2 12.7 60.3 1.55 11.2 3.44 4.05 2.18 6.76 517 96.2

551 4.05 <5.6 <5.6 19.8 3.38 <2.2 6.20 18.8 242 1.99 4.06 92.1 0.845 11.8 2.68 10.5 1.62 25.4 304 115

Rare La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

9.96 24.8 2.70 10.4 3.03 0.436 3.68 0.441 1.47 0.250 0.554 <0.10 0.558 0.068

18.2 47.3 5.27 21.4 4.44 0.356 3.92 0.499 1.59 0.238 0.552 <0.12 <0.23 0.036

22.4 51.5 5.63 22.6 4.41 0.826 3.74 0.636 3.35 0.723 1.91 0.234 1.58 0.218

Total LOI

100.4

0.7

100.6 0.5

336 5.96 <5.7 <5.7 20.4 3.68 2.54 7.35 25.9 237 2.29 5.43 82.5 0.883 16.7 6.15 8.79 1.09 7.61 36.4 115 earth elements (ppm) 11.0 29.8 68.1 26.6 3.01 7.66 25.5 11.6 3.16 5.54 0.407 0.593 4.43 3.27 0.612 0.457 3.08 1.80 0.463 0.331 0.753 0.749 0.135 <0.11 0.453 0.650 0.086 0.089

Total REE

65.8

144.5

58.4

103.8

100.3 0.7

119.8

72.1 15.3 2.73 2.44 1.82 0.596 0.040 4.56 0.130 0.320

Nordmarka Syenite (block) G95:048

Augen gneiss G95:049

70.1 15.8 2.81 2.64 2.89 0.840 0.042 4.21 0.124 0.449

59.0 16.5 3.98 5.23 6.66 2.41 0.122 3.71 0.436 0.609

71.8 14.0 0.984 3.08 4.94 0.598 0.049 2.89 0.219 0.322

73.6 13.4 0.816 2.83 5.02 0.468 0.043 2.62 0.218 0.301

75.7 13.4 0.788 0.845 4.09 0.086 0.016 4.06 0.070 0.046

99.9 0.4

98.7 1.0

98.9 0.6

99.3 0.7

99.1 0.3

432 3.13 <5.8 8.0 13.6 4.4 <0.69 9.03 27.0 250 3.56 6.43 56.9 1.17 13.5 2.76 15.9 2.59 27.4 62.0 138

42.9 6.66 <5.4 <5.4 <10.7 2.35 < 0.645 10.9 <10.8 271 1.27 4.37 35.2 1.56 18.0 8.48 2.38 2.33 3.55 29.5 43.7

73.2 14.1 1.12 2.25 5.38 0.394 0.030 2.92 0.163 0.297

72.0 14.6 0.999 2.26 5.38 0.454 0.029 3.06 0.189 0.269

60.3 17.3 1.45 8.72 4.25 2.55 0.159 2.95 0.074 1.09

99.9 0.9

99.2 0.7

98.8 1.1

590 3.14 <5.7 <5.7 20.7 4.92 <2.3 9.38 30.1 246 2.96 5.36 78.6 1.07 14.3 4.06 20.2 1.70 36.8 395 168

520 4.10 <5.8 <5.8 22.1 5.32 <2.3 10.1 27.0 264 3.93 5.33 80.8 1.12 14.4 3.09 19.8 2.52 35.1 39.6 185

605 3.94 15.9 18.2 25.9 8.15 <2.1 17.3 55.7 341 18.1 6.51 135 1.68 15.8 3.55 137 2.52 37.8 133 280

99.7 3.73 <5.9 <5.9 22.2 3.72 <2.4 13.4 15.8 120 3.77 8.64 182 1.24 9.70 2.64 23.2 0.711 15.2 54.0 121

532 2.94 <5.7 <5.7 19.4 4.50 <2.3 5.55 18.5 142 3.56 5.00 333 1.01 8.24 1.68 36.5 0.863 6.99 95.0 163

1710 6.44 11.7 8.55 21.5 2.15 <0.73 8.66 26.8 275 9.58 7.58 1010 0.840 3.70 1.68 96.3 0.735 14.5 86.2 56.4

473 2.39 <5.7 10.3 13.6 4.07 <0.69 9.07 <11.4 202 4.14 3.95 91.7 0.828 14.5 2.61 17.2 4.02 24.5 56.2 126

48.5 109 12.5 50.0 9.19 1.79 8.30 1.14 6.23 1.25 3.51 0.504 3.02 0.557

18.8 44.5 4.58 19.0 3.37 0.525 3.80 0.499 2.19 0.442 1.62 0.207 1.37 0.243

22.2 47.4 5.06 19.5 2.42 0.747 2.13 0.298 1.20 0.259 0.696 <0.11 <0.23 0.069

36.7 76.3 8.96 36.7 5.97 2.11 4.79 0.770 3.56 0.662 1.84 0.289 1.60 0.228

28.0 64.0 7.93 30.5 6.24 0.963 6.27 1.16 6.92 1.25 3.52 0.368 2.75 0.358

27.6 63.5 7.15 30.6 5.58 0.887 6.13 0.839 5.80 1.18 2.64 0.488 2.43 0.318 155.1

Analysed by ICP-AES or ICP-MS at Svensk Grundamnesanalys AB, Lulea Fe2O3-tot, Total Fe as Fe2O3; LOI, Loss on ignition

27.2 65.0 7.17 26.5 6.12 0.866 5.94 1.02 5.69 1.22 3.37 0.496 3.09 0.439 154.1

255.5

100.0 0.5

101.2

102.0

180.5

160.2

Augen gneiss G95:050

23.7 55.6 6.73 27.7 6.25 0.544 5.19 1.13 6.82 1.34 3.03 0.475 2.68 0.337 141.5

Aplite dyke G95:051

9.13 20.2 2.22 9.29 2.26 0.250 1.72 0.442 2.54 0.449 0.983 0.176 1.08 0.159 50.9

214

A. JOHANSSON ETAL

Table 1. Continued Locality

Rock type Sample

Isispynten

Augen gneiss 898:130

Gabbro

Gabbro

94062C

AJ94:003

Aplitic granite 898:127

66.2 15.8 3.09 4.24 3.14 1.20 0.066 4.39 0.249 0.622

49.9 5.74 11.4 9.74 1.59 17.5 0.215 0.837 0.159 0.517

49.4 16.7 12.8 7.27 0.246 9.19 0.152 2.44 0.079 0.796

75.6 13.0 0.659 1.57 6.41 0.286 0.023 2.76 0.059 0.164

98.2

99.0

97.6

99.1

1.0

0.6

1.1

1.7

Trace elements (ppm) 1090 Ba 3.05 Be 6.97 Co <5.8 Cu 30.1 Ga Hf 10.0 Mo <2.3 Nb 15.9 27.2 Ni 158 Rb 8.21 Sc Sn 1.45 378 Sr 1.50 Ta Th 28.1 3.67 U 62.8 V 1.67 W 34.4 Y Zn 84.6 343 Zr

909 2.11 <5.6 <5.6 25.0 9.14 <2.2 10.4 15.1

134

458 0.73 60.6 72.7 10.0 2.78 7.28 3.50

257

3.60 2.90

65.4 39.0 2.80

486 0.952

206 0.289

18.5 3.09 58.3 0.850 18.5 74.2

5.18 1.05

360

141 0.937 16.8

109 84.7

Rare earth elements (ppm) 68.5 65.8 La 155 133 Ce 13.4 Pr 17.0 65.4 Nd 47.9 Sm 9.95 6.78 1.42 Eu 1.97 9.32 5.04 Gd 1.32 0.670 Tb Dy 6.53 3.10 1.24 0.641 Ho Er 2.85 1.55 Tm 0.550 0.273 2.55 1.63 Yb 0.402 0.243 Lu

14.8 37.6 4.67 21.9 3.93 1.00 3.43 0.578 3.18 0.546 1.42 0.269 1.43 0.182

Total REE

94.9

342.6

281.5

10.8 <0.59 39.6 16.1 <11.9 1.13 <0.71 0.544 84.4 2.52 42.3 0.98 132 0.401 0.208 0.192 250 0.704 7.86 211 9.34

0.529 2.12 <1.2 1.70 0.678 0.555 0.929 0.596 2.16 0.712 1.28 0.419 1.08 0.369 14.3

Kraemerp., Kvit0ya

Andreeneset, Kvit0ya

Grey granite 898:131

Major elements (wt%) Si02 61.9 A12O3 17.3 CaO 1.87 5.62 Fe2O3-tot K2O 5.26 MgO 1.52 MnO 0.077 Na2O 3.78 P205 0.114 TiO2 0.798 Total LOI

Stor0ya

Migmatite

Amphibolite

Migmatite

Migmatite

898:128

898:129

AJ94:004

940045

67.3 14.0 2.65 5.79 2.30 2.01 0.079 3.64 0.118 0.839

46.4 13.9 11.0 13.7 1.09 6.54 0.291 2.29 0.282 2.59

65.6 14.6 4.03 6.72 2.02 1.63 0.120 3.83 0.248 0.950

100.5

98.7

98.1

99.7

100.6

99.4

100.7

0.6

0.6

0.7

0.7

0.3

0.3

0.6

540

735

0.916 <5.8 <5.8 16.2 3.15 <2.3 7.67 17.0

2.60 <1.2

1.21 8.28 <5.4 21.4 8.26 <2.2 9.46 20.0 90.9 7.75 2.12

145 0.607

341 0.470

37.0 2.45 3.47 0.883 12.3 33.1

12.1 1.35 85.7 0.816 18.9 83.4

166

101

321

39.0 91.4 10.2 38.6 7.34 0.831 6.16 0.731 2.80 0.475 1.07 <0.12 0.574 0.071 199.2

37.3 88.6 10.2 43.1 6.26 1.31 7.15 0.802 4.03 0.695 1.88 0.132 1.36 0.132

203.0

91.1 <0.55 33.1 <5.5 16.8 4.53 <2.2 7.84 41.9 23.7 39.7 1.85 223 0.725 1.59 0.501 384 0.684 41.7 114 154 11.2 30.3 4.09 18.9 4.86 1.79 6.21 0.986 6.70 1.48 4.17 0.611 4.09 0.522 95.9

737 2.37 8.61 27.6 13.4 17.5 <0.72 21.8 20.3 69.6 24.6 6.27

287 2.31 23.7 4.01 52.5 1.04 88.7

139 682 70.4

156 19.0 77.3 16.1 3.20 15.6 2.74 16.5 3.50 8.46 1.35 8.20 1.25

399.6

70.6 15.9 3.51 3.05 1.28 0.825 0.048 4.93 0.090 0.350

409 4.74 <5.2 60.3 <10.5 7.30 <0.63 6.75 15.6 48.4 4.42 1.84

314 1.17 8.91 4.66 31.1 0.365 13.5 70.0

279 21.8 50.8 6.05 24.1 4.81 0.978 3.97 0.677 3.83 0.736 1.78 0.353 1.31 0.281 121.5

Aplite dyke AJ94:005

AJ94:006

74.6 13.8 0.866 0.948 5.02 0.117 0.030 3.91 0.080 0.067

49.8 19.6 10.4 10.2 0.457 5.25 0.211 3.91 0.219 0.630

271 5.04 <5.6 29.2 <11.2 3.11 <0.67 6.78 <11.2

169

Gabbro

126 <0.54 32.5 33.2 22.0 0.718 <0.65 1.17 21.5 3.63 26.4 0.724

2.20 2.10 97.3 1.26 23.3 9.31 <2.2 0.561 11.7 66.2 69.2

0.153 0.404 0.156 249 0.379 8.78 131 12.6

16.2 35.0 4.07 14.5 3.57 0.288 3.23 0.770 3.22 0.785 1.99 0.408 2.42 0.414

7.43 16.8 2.12 10.1 1.59 0.399 2.07 0.389 1.95 0.453 0.939 0.191 0.940 0.239

86.9

45.6

833

Analysed by ICP-AES or ICP-MS at Svensk Grundamnesanalys AB, Lulea Fe2O3-tot, Total Fe as Fe2O3; LOI, Loss on ignition

making the 206Pb-238U age more reliable. Also in the age calculations in Figure 10, the emphasis is on the 207Pb-206Pb ages for the older samples, and the 206pb-238U ages for the younger ones, although sometimes both sets of ages are shown. The results are listed in Table 4 and illustrated in Figure 10. All uncertainties stated in the text, whether for conventional U-Pb ages, Pb-evaporation ages, or ion microprobe ages, are at the 2cr level.

Augen gneisses Augen gneisses from Nordmarka (samples G95:049 and G95:050) have been analysed both with the single zircon Pb-evaporation method and the NORDSIM ion microprobe. A prismatic zircon from sample G95:049 yields a weighted average 207pb/206 Pb-evaporation age of 947 ± 1 8 Ma, and a zircon from sample

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

Fig. 4. Major element geochemistry of investigated rocks from NE Svalbard: (A) Total alkali-silica diagram (Le Maitre 1989). (B) P-Q diagram (Debon & Le Fort 1983). (C) B-A diagram (Debon & Le Fort 1983). (D) Shand's index (Maniar & Piccoli 1989). (E) I/S-type classification (Chappel & White 1974). (F) Multicationic R1-R2 diagram (Batchelor & Bowden 1985). (G) AFM diagram with calc-alkaline/tholeiitic division line from Irvine & Baragar (1971). (Data from Table 1.)

215

216

A. JOHANSSON ET AL

augen gneiss protolith at c. 950 Ma. There are occasional older inherited zircons (898:122 zr 15a; c. 1300 Ma) as well as a few yielding seemingly younger ages, that may have experienced lead loss. Augen gneiss sample 898:130 from Isispynten does not give any consistent ion microprobe age, but a number of semiconcordant points spread out along the concordia (Fig. lOe): two points from a metamict zircon grain at 300-400 Ma, three zircons at 600-700 Ma, and one zircon having a rim age of c. 1000 Ma and a core age of c. 1400 Ma. It may be that more analyses would have produced a group of magmatic zircons with a reasonably well-defined crystallization age. At present, this augen gneiss may be interpreted either as a Grenvillian granite, with superposed metamorphism leading to extensive lead loss from some of the zircons, or as a foliated Caledonian granite (pink granodiorite of Sandford 1954) with older inheritance. The intermediate, 600-700 Ma ages are, however, puzzling, especially since at least two of them (zr 02, 12) are from well crystallized prismatic zircons with simple internal structure and are only a few percent discordant. Similar ages have been recorded on zircons from the Parry0ya augen gneisses above, albeit from outer parts of crystals that may have undergone lead loss, and could possibly be recording some cryptic, late Neoproterozoic event affecting eastern Svalbard.

Biotite schist Three grains of titanite from the biotite schist sample 898:124 from Parry0ya have been analysed by the Pb-evaporation method. They show considerable scatter, with a weighted average 207Pb/206Pb age of 930 ± 100 Ma (Table 3, Fig. 8b). If this rock represents metamorphosed late Mesoproterozoic Brennevinsfjorden Group sediments, the titanite age would indicate metamorphic crystallization of titanite in late Grenvillian time. Alternatively, if the biotite schist originated from Neoproterozoic Murchisonfjorden Supergroup sediments, the titanite would be of detrital origin. In either case, Caledonian metamorphism was not strong enough to cause resetting of the titanite U-Pb age. Fig. 5. (A) MORE normalized trace element variation diagram; (B) Chondrite normalized REE diagram, for all investigated rocks from NE Svalbard. (Data from Table 1.)

Migmatites

G95:050 yields a 207Pb/206Pb age of 959 ± 19 Ma (Table 3, Fig. 8k; recalculating the 2cr of the mean error in Table 3 to the 2a level by multiplying it with the square root of the number of scans). On the ion microprobe, prismatic zircons from both these samples plotting concordantly or almost concordantly between 900 and 1000 Ma, yield a combined weighted average 207 Pb/206Pb age of 939 ± 7 Ma (n = 12, MSWD = 0.90; Fig. 10k) and a weighted average 2b6Pb/238U age of 962 ± 10 Ma (n= 12, MSWD = 2.1; Fig. 10k). These data together are taken to indicate a magmatic crystallization age of c. 950 Ma for the protolith of the Nordmarka augen gneiss. A few zircons yield clearly older ion microprobe ages, indicating inheritance from older crustal rocks. The augen gneisses from Parry0ya (samples 898:122 and 898:123) and Isispynten (sample 898:130) have been analysed with the ion microprobe only. The red augen gneiss from Parry0ya (898:122) yields a weighted average 207Pb/206Pb age of 944 + 17 Ma (n = 10, MSWD = 1.9; Fig. lOb), and a weighted average 206Pb-238U age of 939 ± 26 Ma (n = 10, MSWD = 7.5; Fig. lOb). The grey augen gneiss sample from Parry0ya (898:123) yields a weighted average 207Pb/206Pb age of 936 + 12 Ma (n = 8, MSWD = 0.57; Fig. lOc), and a weighted average 206Pb_238u age of 961 ± 17 Ma (n = 8> MSWD = 2.3; Fig. lOc). These ages are taken to indicate magmatic crystallization of the

Five zircons from migmatite sample 31-5 from Damflya have been dated with the Pb-evaporation method and yielded a weighted average 207Pb/206Pb age of 432 ± 7 Ma (Table 3, Fig. 8i), interpreted to date migmatization. Several age datings using different methods have been performed on migmatites from Andreeneset on Kvit0ya. Single grain Pb-evaporation dating of zircons from samples 898:128, AJ94-.004 and 94045 (Table 3, Fig. 4e, f, h) has yielded weighted average 207Pb/206Pb ages of 438 ± 16 Ma (seven zircons, an additional zircon yields an age around 900 Ma), 457 + 9 Ma (five zircons), and 452 ± 1 9 Ma (two out of four zircons), respectively. These ages are interpreted as dating migmatization. Ion microprobe dating of zircons from sample S98:128 yielded a weighted average 206Pb-238U age of 430 ± 21 Ma (n = 8, MSWD = 12; Fig. lOg). There is substantial scatter, as evident from the high MSWD value, with some suggestion of a small but systematic difference between cores (447 ± 28 Ma, n = 4, MSWD =o5.7) and rims (415 ± 37 Ma, n = 4, MSWD = 11). Samples AJ94:004 and 94045 are very similar and from the same locality and have been grouped together, yielding a weighted average 206Pb-238U ion microprobe age of 465 ± 11 Ma (n = 19, MSWD = 5.8; Fig. lOh), again with some suggestion of a small difference between cores (472 ± 13 Ma, n = 14, MSWD = 5.5) and rims (447 ± 8 Ma, n = 5, MSWD = 0.59). A conventional U-Pb murtigram dating on four fractions of titanite from samples 94045 and AJ94:004 yields slightly discordant

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

217

Fig. 6. Tectonic discrimination diagrams, based on Pearce et al (1984) and Pearce (1996), for granitic rocks from NE Svalbard. ORG, Ocean ridge granite; VAG, Volcanic arc granite; syn-COLG, Syncollision granite; WPG, Within-plate granite; postCOLG, Post-collision granite. (Data from Table 1.)

ages. The 206Pb-238U ages of the individual fractions fall between 424 and 433 Ma (Table 2, Fig. 7c), slightly younger than the zircon ages above. Alternatively, a regression through the four fractions produces a lower intercept age of c. 408 Ma, which could be taken to indicate cooling after metamorphism, and an upper intercept age of c. 980 Ma, which may indicate a small inherited component. However, the intercept ages are not well defined and should be interpreted with caution. Taken together, the data favour a Caledonian migmatization event at c. 460 to 430 Ma (late Ordovician to early Silurian), somewhat earlier than the intrusion of the late Silurian to early Devonian Rijpfjorden granite at 410-420 Ma. Although the slightly higher age may have been caused by inheritance in the zircons; the ion microprobe rim data from samples AJ94:004 and 94045 indicate that the migmatization really occurred at c. 450 Ma, in accordance with the Pb-evaporation data. No substantially older inherited ages (>500 Ma) were recorded during the ion microprobe analyses, in spite of many zircons having complex internal structures with apparent detrital cores visible in CL (e.g. 94045 zr 01, 04, 05, 07, 09; AJ94:004 zr 02, 03, 04, 05; cf. Fig. 9). However, ion microprobe analysis of such core-like internal parts produced ages only some tens of million years older than the rims, in sharp contrast to the situation in many of the Caledonian granites that contain abundant inherited

Proterozoic zircons (Johansson et al. 2002). It would seem more likely with the opposite relation: more of inherited old zircons should be present in a migmatite largely consisting of palaeosome material than in a more homogeneous granite. Only the Pbevaporation analyses produced one clearly inherited age of c. 900 Ma. Two possible explanations are offered: (1) the migmatite underwent heating and partial melting at depth with high fluid activity during a sufficiently long time period to almost reset the U-Pb system in the older zircons, whereas older zircons in the granites were picked up from the wall rocks at a late stage of magma ascent and solidification, and did not have time to equilibrate their isotope system, or (2) the protoliths of the migmatitic paragneiss are young (Ordovician ?) supracrustals with little inherited material, suggesting an island-arc setting outboard of the continental margin for the protolith rocks. Grey granites The homogeneous granitic migmatite neosomes from Innvika, samples S98:49A and S98:50, have already been dated by ion microprobe to c. 420 Ma (Tebenkov et al. 2002). Similar fine- to medium-grained grey granites from Parry0ya (S98:121), Foyn0yane (S98:125 and 126), Kvit0ya (S98:127) and Isispynten

Table 2. Conventional U-Pb data for monazite and titanite from northeasternmost Svalbard No.1

Sample/Fraction2

Weight (mg)

U (ppm)

Pb-rad (ppm)

Sample Mzl Mz2 Mz3

G95:051: Aplite dyke, Nordmarka, eastern Nordaustlandet G95:051 90-200 0.6064 8381 2480 G95:051 90-200 0.4386 8836 2409 G95:051>200 1.0394 8249 1791

Sample Til Ti2 Ti3

G95:048: Syenite boulder, Nordmarka, eastern Nordaustlandet G95:048 150-200 dark 0.3263 202 16.5 4.31 G95:048 150-210 light 0.3478 38.6 5.96 G95:048 106-150 0.3389 55.0

Sample AJ94:004: Migmatite, Andreeneset, Kvit0ya Til AJ94:004 150-210 0.5434 158 Ti2 AJ94:00490-150 0.3791 164 Sample 94045: Migmatite, Andreeneset, Kvit0ya Til 94045 150-210 0.9434 Ti2 94045 106-150 0.5989 !

117 113

10.8 11.2 8.40 7.92

Pb-com (ppm)

Measured ratios3

Radiogenic isotope ratios4

206pb

206pb

206pb

207pb

204pb

207pb

208pb

235U

-\-1ffff ±-

206pb 238U

-\-Jrr -L-ff

207pb 206pb

Error5 corr. -4-^.fr -L~ff

Model ages (Ma) 207pb

206pb

207pb

235U

238U

206Fb

447 439 302

468 464 484

330 484 251

9.952 11.59 8.719

0.2768 0.2975 0.2489

Sample 0.5588 0.5474 0.3760

3.94 3.50 3.41

237

8.641 3.715 4.489

2.211 0.8970 0.9806

Sample 0.5246 0.5171 0.5604

428

431

4J3

423 452

446 435

302 540

1.71 1.78

407 401

5.093 4.979

Sample AJ94:004: Migmatite, Andreeneset, Kvit0ya (com. Pb420 Ma) 431 0.5283 38 0.06835 11 0.05606 37 0.458 0.5266 80 0.06794 14 0.05621 79 430 0.538

426 424

455 461

1.95 1.75

276 291

3.704 4.024

Sample 94045: Migmatite, Andreeneset, Kvit0ya (com. Pb420 Ma) 0.5468 119 0.06942 15 0.05712 115 0.679 0.5364 42 0.06906 11 0.05633 41 0.449

433 430

496 466

121 83.5

106

67 87

10.97 10.92 9.143 9.479

G95:051: 46 27 26

Aplite dyke, Nordmarka (com. Pb420 Ma) 0.07188 19 0.05638 41 0.573 0.07052 16 0.05630 24 0.525 0.04802 14 0.05680 34 0.526

451 443 324

G95:048: Syenite boulder, Nordmarka (com. Pb420 Ma) 138 0.06916 17 0.05502 135 0.721 0.572 239 0.07160 35 0.05237 229 252 0.614 0.06974 36 0.05827 244

443 436

Mz, Monazite; Ti, Titanite Size fractions in micrometres 3 Corrected for fractionation and blank (206Pb/204Pb only corrected for fractionation) 4 A11 ratios corrected for fractionation (U according to analyses with 233~235U-tracer, Pb with 0.11 ± 0.04% per AMU), blank (0.01 ng U and 0.013 (titanite) to 0.075 (monazite) ng Pb with 206Pb/204Pb = 18.7, 207 Pb/204Pb - 15.6, 208Pb/204Pb - 38.6) and common lead (206Pb/204Pb = 18.048,207Pb/204Pb - 15.593, 208Pb/204Pb = 37.857; Stacey and Kramers (1975) model for 420 Ma). Errors reported as 2 standard deviations in the last digits. 5 Error correlation 207Pb/235U-206Pb/238U.

219

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD Table 3. Pb-evaporation data for zircon, monazite and titanite from northeasternmost Svalbard Sample/ gram

Mineral

Measured 207T»K Pb ~ „ 206p b ±- "*

S98:121: Grey granite, Parry 0ya E zircon 0.07176 0.80 32124 F zircon 0.07115 0.34 97880 H zircon 0.07194 0.58 9147 L zircon 0.07129 0.31 17230 MA monazite 0.05544 0.71 17546 MB monazite 0.05574 0.63 20456 MC monazite 0.06555 0.17 1510 MD monazite 0.05625 0.94 10437 898:124: Biotite schist, Parry0ya C titanite 0.07600 16 3380 D titanite 0.07133 0.18 12020 E titanite 0.07560 4.2 2145 898:125: Grey granite, Foyn0yane A zircon 0.05558 0.12 130330 B zircon 0.05612 0.11 28010 C zircon 0.05675 0.81 19580 D monazite 0.05678 0.33 23680 F monazite 0.05729 0.12 198890 898:127: Aplitic granite, Andreeneset, Kvit0ya F titanite 0.05611 0.59 15780 898:128: Migmatite, Andreeneset, Kvit0ya A zircon 0.05689 1.7 8686 C zircon 0.07040 3.5 16459 F zircon 0.05890 4.6 3859 G zircon 0.05705 1.7 22299 K zircon 0.05665 0.47 20530 L zircon 0.05654 0.92 18337 M zircon 0.05658 1.2 14825 N zircon 0.05720 1.4 8552 AJ94:004: Migmatite, Andreeneset, Kvit0ya A zircon 0.05690 0.20 16010 B zircon 0.05734 0.42 14384 C zircon 0.05805 0.48 7690 M zircon 0.05749 0.24 12070 N zircon 0.05653 0.70 22100 AJ94:005: Aplite dyke, Andreeneset, Kvit0ya A zircon 0.05740 3.5 18876 B zircon 0.05684 0.18 10080 C zircon 0.05627 0.23 29533 L zircon 0.05647 1.4 17731 M zircon 0.05700 3.3 16495 N zircon 0.08454 0.32 543 V zircon 0.05608 0.56 53836 W zircon 0.05610 0.35 21600 Z zircon 0.05591 0.41 67170 94045: Migmatite, Andreeneset, Kvit0ya A zircon 0.05770 3.7 7576 B zircon 0.05660 4.3 85420 K zircon 0.06924 3.5 840 L zircon 0.05877 0.84 2850 31-5: Migmatite, Damflya A zircon 0.05723 0.61 9860 B zircon 0.05685 1.1 12180 C zircon 0.05740 4.8 10405 D zircon 0.05720 3.7 7020 F zircon 0.05549 0.38 159400 G95:048: Syenite boulder, Nordmarka B zircon 0.05616 1.9 43880 Fig. 7. Conventional U-Pb monazite and titanite results. (A) Aplite dyke, Nordmarka (monazite). (B) Syenite boulder, Nordmarka (titanite). (C) Migmatite, Andreeneset, Kvit0ya (titanite). Size of symbols reflects 2cr analytical error, numbering refers to Table 2.

Age (Ma)

±2crm

48 27 32 12 63 74 2 36

966 958 940 942 396 414 449 406

17 3 17 5 7 4 9 6

280 22 58

979 933 896

46 4 6

18 10 56 8 25

432 437 453 459 500

6 6 11 13 6

11

420

6

46 200 33 51 12 45 23 49

421 916 417 468 450 442 437 433

22 54 10 14 6 14 18 7

11 37 8 14 20

452 466 459 464 448

4 15 5 18 9

136 4 50 112 220 1 102 200 21

478 429 444 439 458 520 445 430 441

12 3 15 14 10 27 10 14 4

83 225 7 22

444 469 277 356

24 34 23 9

26 11 180 110 69

443 440 453 418 429

15 18 60 20 9

222

446

4

10 190

947 959

2 3

Measure d 206¥>K Pb 2
G95:049, G95:050: Augen gneiss, Nordmarka 049A zircon 0.07079 0.15 86700 050A zircon 0.07124 2.3 78730

220

A. JOHANSSON ETAL.

(898:131) have been dated with the Pb-evaporation or ion microprobe techniques during this study. From Parry0ya sample 898:121, four zircon crystals analysed with the Pb-evaporation technique have yielded a weighted average 207Pb/206Pb age of 951 ± 19 Ma, and four monazite crystals a weighted average 207Pb/206Pb age of 416 ± 37 Ma (Table 3, Fig. 8a). The limited ion microprobe data (five points in three zircon crystals) are not conclusive, with 207Pb/206Pb ages of individual points scattering between 840 and 1560 Ma. The Pb-evaporation data may either be taken to indicate that the Parry0ya grey granite is a late Grenvillian c. 950 Ma old granite, comparable to the Kontaktberget granite (Gee et al. 1995), in which monazite underwent recrystallization and isotopic resetting during Caledonian metamorphism at c. 420 Ma, or that it is a Caledonian granite with considerable inheritance of Grenvillian zircons, possibly from the nearby augen gneiss which it cross-cuts.

From the Foyn0yane grey granite (sample 898:125), both zircon and monazite yield Caledonian Pb-evaporation ages, the zircon age being 440 + 26 Ma (average of three crystals; Table 3, Fig. 8c). The two monazite crystals show considerable scatter, with a somewhat bimodal distribution, yielding an ill-defined average of 480 + 250 Ma. Ion microprobe analyses of zircons both from sample 898:125 (medium-grained) and sample 898:126 (fine-grained grey granite) from Foyn0yane yield a group of concordant to semi-concordant zircons with ages close to 400 Ma. They may be subdivided into two groups, one with a weighted average 206pb_238u age of 437+ 12 Ma (n = 4, MSWD = 1.1), taken to indicate the magmatic crystallization age, and one at 390 ± 27 Ma (n = 4, MSWD = 3.5), in which the zircons may have undergone slight lead loss. From the grey aplitic granite from Andreeneset at Kvit0ya (sample 898:127), only one titanite grain has been analysed by

Fig. 8. Histograms of Pb-evaporation results for zircon, monazite and titanite from NE Svalbard. For detailed analytical data, see Table 3. N, number of crystals.

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

221

Fig. 8. Continued

the Pb-evaporation method. The analysed titanite yields a 207 Pb/206Pb age of 420 ± 54 Ma (Table 3, Fig. 8b, error expanded according to number of scans), probably reflecting cooling shortly after the Caledonian migmatization at Kvit0ya and crystallization of this aplitic granite. The grey granite from Isispynten (sample 898:131) yields individual ion microprobe zircon ages ranging from c. 250 Ma (affected by lead loss) to c. 900 Ma (Fig. 10). Only two points in one huge (c. 700 |xm)-, clear, CL-grey zircon with simple internal structure (zr 01) yield Caledonian ages of c. 400 Ma (Fig. 9). Although hinting at a Caledonian age for the Isispynten grey granite, these two points are not enough to produce a welldefined date. In summary, the Foyn0yane grey granite has a Caledonian age at c. 440 Ma, and the aplitic granite from Kvit0ya as well as the Isispynten grey granite are probably Caledonian as well. For the

Parry0ya grey granite, two possible interpretions of the age data have been given above, of which the simplest would be that this is also a Caledonian granite (monazite age of c. 420 Ma), with inherited Grenvillian zircons.

Aplite dykes Aplite dykes have been sampled and dated at Nordmarka, where they cross-cut the augen gneiss, and Andreeneset on Kvit0ya, where they cross-cut the migmatitic paragneisses. The Nordmarka aplite sample (G95:051) contained little zircon, but three monazite fractions were analysed by conventional U-Pb analysis. Two fractions are slightly discordant and one fraction highly discordant, with 207Pb/206Pb ages between 464 and 484 Ma, and a

222

A. JOHANSSON ETAL.

three-point discordia upper intercept age of 463 ± 9 Ma (lower intercept -45 ± 3 9 Ma, MSWD = 0.19; Table 2, Fig. 7a). However, this age is uncertain, considering the few points involved and the high discordancy of one of them, and in the light of the cross-cutting nature of the aplites as well as Ar/Ar data on the same sample (Ar/Ar plateau age on muscovite of 417 ± 5 Ma; Johansson et al 2001), the c. 460 Ma U-Pb

monazite age may be too high (cf. discussion in Johansson et al. 2001). o From Andreeneset on Kvit0ya, zircons from aplite sample AJ94:005 have been dated using both Pb-evaporation and ion microprobe techniques. Eight zircon crystals analysed by the Pb-evaporation method yielded a weighted average 207Pb/206Pb age of 445 ± 14 Ma (Table 3, Fig. 8g). Seven points in six

Fig. 9. Cathodoluminescence images of selected zircons used for ion microprobe analysis, with analytical spots and obtained 206Pb-238U ages (below 500 Ma) or 207 Pb/ ° Pb ages (above 500 Ma) in million years indicated. Ages in brackets are highly discordant, and thus lack geological significance.

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

223

Fig. 9. Continued

zircon crystals analysed by the NORDSIM ion microprobe yielded a weighted average 206Pb-238U age of 431 ± 7 Ma (n = 7, MSWD=1.9) and a slightly lower 207Pb/206Pb age of 4 1 1 + 2 0 Ma (n = 7, MSWD = 1.6; Fig. lOi), with no signs of inheritance. An age of c. 430 Ma for the Kvit0ya aplite dykes, and an age of c. 450 Ma for the surrounding migmatite, would be in accordance with observed field relations, with the aplite dykes cutting the migmatite (cf. Fig. 3c).

Syenite Zircons from a syenite boulder in Nordmarka (sample G95:048) have been analysed by Pb-evaporation and ion microprobe. The zircons are short and stubby, with the two pyramids at the ends almost meeting without any intervening prism, which is typical for zircons crystallized from alkali-rich syenitic magmas (Pupin & Turco 1975). They have simple internal structures with

Table 4. NORDSIM ion microprobe U—Pb data on zircons from augen gneisses, migmatites, granites and aplites from northeasternmost Svalbard Sample/ spot number

Description1

Measured ratios2

%

Err.3 corr.

Disc.4 %

0.1527 0.1707 0.1362 0.2241 0.1329

1.69 1.70 1.69 1.70 1.69

0.93 0.85 0.43 0.96 0.91

-4.1 -5.5 -2.1 -17.9 -12.5

1.79 1.53 1.87 2.07 .77 .80 .99 .71 .97 :-.22 .83 .86 .93

0.1545 0.1664 0.1550 0.1509 0.1636 0.2237 0.0932 0.1683 0.1491 0.1547 0.1532 0.1535 0.1387

1.47 1.48 1 .41 1.55 1.47 1.47 1.51 1.48 1.65 1.47 1.47 1.55 1.50

0.82 0.97 0.79 0.75 0.83 0.82 0.75 0.87 0.84 0.66 0.80 0.83 0.78

1.624 0.677 1.546 1.525 1.581 1.583 1.276 1.529 1.314 1.558 1.501

.58 .86 .80 .91 .68 .66 .63 .64 .89 .72 .97

0.1673 0.0844 0.1619 0.1581 0.1633 0.1632 0.1337 0.1572 0.1365 0.1598 0.1566

1.48 1.47 1.48 1.52 1.47 1.48 1.48 1.47 1.69 1.47 1.50

898:125. Medium-grained grey granite, Foyn0yane (com. Pb 420 Ma) 0.60 898: 125-05a Centre dark/turbid CL-dark 0.05465 898: 125-06a Centre clear CL-striped/dark 0.05474 0.50 1.84 898: 125-09a Outer clear CL-dark 0.05349 2.06 S98:125-10a Centre semi-turbid CL-grey 0.05642 0.87 898: 125-lOb Tip zoned CL-zoned 0.05349 2.15 898: 125-20a Centre clear CL-bright 0.05462 2.77 S98:125-24a Centre zoned CL-dark 0.05531

0.542 0.544 0.454 0.454 0.512 0.465 0.517

2.93 2.90 3.45 3.53 2.99 3.58 3.99

0.0719 0.0721 0.0615 0.0583 0.0694 0.0617 0.0678

898:126. Fine-grained grey granite, Foyn0yane (com. Pb 420 Ma) S98:126-01a Intermediate clear CL-bright 0.05453 1.61 S98:126-04a Round core? CL-patchy 0.09578 0.67 898: 126-04b Inner rim CL-bright 0.04020 5.59

0.486 2.083 0.269

2.36 1.82 5.84

0.0647 0.1578 0.0485

Pb 235U

±lCT Vl

206pb

898:121. Medium-grained grey granite, Parry 0ya (com. Pb 950 Ma) 0.69 S98: 121-02a Intermediate clear CL-zoned 0.07081 1.03 S98: 121-05a Centre clear CL-zoned 0.07507 3.53 S98: 121-05b Outer clear CL-zoned 0.06704 0.52 S98: 121-06a Core? fract CL-patchy 0.09641 0.75 S98: 121-06b Rim fractured CL-dark/zoned 0.06943

1.490 1.767 1.259 2.979 1.272

1.82 1.99 3.92 1.77 1.85

898:122. Red augen gneiss, Parry0ya (com. Pb 950 Ma) 898: 122-Ola Centre clear CL-grey 0.07052 S98:122-03a Centre clear CL-grey 0.07103 898: 122-06a Centre clear CL-grey 0.07022 898: 122-07a Centre clear CL-grey 0.06965 898: 122-1 la Centre clear CL-grey /zoned 0.06901 S98:122-15a Centre clear CL-zoned 0.08250 898: 122-15b Tip turbid CL-zoned 0.06126 898: 122-16a Centre clear CL-grey 0.07069 898: 122-18a Long centre clear CL-grey 0.06918 S98:122-27a Long centre clear CL-bright 0.06819 S98:122-29a Centre clear CL-grey /zoned 0.07097 898: 122-30a Centre clear CL-zoned 0.07100 898: 122-30b Tip clear CL-zoned 0.07061

1.01 0.39 1.15 1.37 0.99 1.03 1.31 0.86 1.07 1.66 1.10 1.03 1.21

1.502 1.629 1.501 1.450 1.557 2.544 0.787 1.640 1.422 1.454 1.500 1.503 1.350

898:123. Grey augen gneiss, Parry0ya (com. Pb 950 Ma) 898: 123-Ola Centre clear CL-grey 0.07041 S98:123-01b Tip clear CL-zoned 0.05814 898: 123-06a Centre clear CL-grey 0.06925 898: 123-07a Long tip fractured CL-grey 0.06995 898: 123-09a Long centre clear CL-grey 0.07023 S98:123-12a Oval centre clear CL-zoned 0.07034 S98:123-13a Centre turbid/black CL-grey 0.06919 S98:123-15a Centre clear CL-grey 0.07057 898: 123-15b Tip turbid CL-dark 0.06984 898: 123-16a Centre clear CL-striped 0.07072 898: 123-22a Centre clear CL-striped 0.06952

0.55 1.15 1.02 1.15 0.82 0.75 0.70 0.73 0.86 0.88 1.28

207pb 206pb

+ la «fc

Calculated ages

Concentrations 207

238U

f1206 5

207pb

207pb

206pb

U ppm

Th ppm

Pb ppm

Th6 If

206pb

0.03 0.95 4.42 0.61 0.55

330 626 633 975 1686

92 173 328 761 138

58 121 99 290 243

0.28 0.28 0.52 0.78 0.08

952 1070 839 1556 911

14 21 72 10 15

927 1034 827 1402 833

13 22 14 11

916 1016 823 1304 804

14 16 13 20 13

-2.0 3.8 -0.7 -1.4 9.3 3.9 -11.9 6.1 -1.0 6.5 -4.2 -4.1 -12.3

0.20 0.01 0.13 0.10 0.08 0.08 1.29 0.08 0.20 0.26 0.07 0.11 0.19

425 1396 248 264 258 374 1250 314 313 151 284 383 596

48 43 55 60 28 138 32 91 50 78 37 56 85

73 251 44 45 46 103 125 61 52 28 48 66 92

0.11 0.03 0.22 0.23 0.11 0.37 0.03 0.29 0.16 0.52 0.13 0.15 0.14

944 958 935 918 899 1257 648 949 904 874 957 957 946

21 8 24 28 20 20 28 17 22 34 22 21 25

931 982 931 910 953 1285 590 986 898 912 930 932 868

11 10 11 13 11 13 9 11 12 13 11 11 11

926 992 929 906 977 1301 574 1002 896 927 919 921 837

13 14 13 13 13 17 8 14 14 13 13 13 12

0.94 0.79 0.82 0.80 0.88 0.89 0.90 0.90 0.89 0.86 0.76

6.5 -2.5 7.3 2.2 4.6 4.2 -11.2 -0.5 -11.4 0.7 2.8

0.04 0.71 0.16 0.38 0.11 0.10 0.30 0.24 0.55 0.08 0.41

662 1617 280 359 397 427 1164 546 1683 305 256

47 107 55 98 44 86 64 168 183 135 55

121 146 51 65 71 79 168 99 250 59 45

0.07 0.07 0.20 0.27 0.11 0.20 0.05 0.31 0.11 0.44 0.21

940 535 906 927 935 938 904 945 924 949 914

11 25 21 23 17 15 14 15 18 18 26

980 525 949 941 963 964 835 942 852 954 931

10 8 11 12 11 10 9 10 11 11 12

997 522 967 946 975 975 809 941 825 956 938

14 7 13 13 13 13 11 13 13 13 13

2.87 2.86 2.92 2.86 2.86 2.87 2.87

0.98 0.99 0.85 0.81 0.96 0.80 0.72

12.8 12.1 10.4 -22.7 24.4 -2.8 -0.4

0.06 0.09 0.41 1.45 0.47 0.24 1.53

1411 2456 1073 600 3534 176 4950

31 257 83 76 109 54 317

108 192 71 38 259 12 359

0.02 0.10 0.08 0.13 0.03 0.31 0.06

398 402 350 469 350 397 425

13 11 41 45 20 47 61

439 441 380 380 420 387 423

11 10 11 11 10 12 14

447 449 385 366 432 386 423

12 12 11 10 12 11 12

1.72 1.69 1.69

0.73 0.93 0.29

3.0 -41.7 194.9

0.16 0.80 5.19

378 1187 945

549 926 834

35 214 48

1.45 0.78 0.88

393 1543 -338

36 12 138

402 1143 242

8 13 13

404 944 305

7 15 5

±lCT

%

±1
±lCT

235JJ

±lo

238U

11

898:128. Bi-rich S98:128-03a S98:128-03b S98:128-05a S98:128-05b 898:128-1 la S98:128-llb S98:128-12a S98:128-12b

porphyroblastic migmatite, Andreeneset, Kvit0ya (com. Pb 420 Ma) 2.11 Centre clear CL-zoned 0.05600 1.24 0.529 3.25 Outer clear CL-bright 0.05724 2.77 0.516 2.26 Core? clear CL-zoned 0.05378 1.42 0.560 2.40 Rim clear CL-grey 0.06041 1.60 0.583 3.29 Centre clear CL-bright 0.05969 2.82 0.602 Outer clear CL-grey 0.06499 1.46 0.631 2.35 2.86 Centre clear CL-zoned 0.05513 2.30 0.539 Outer zoned CL-zoned/bright 0.05583 3.92 0.482 4.27

898:130. Foliated augen gneiss, Isispynten (com. Pb S98:130-02a Long centre clear CL-zoned S98:130-06a Core turbid CL-patchy S98:130-06b Rim clear CL-zoned S98:130-07a Rim zoned CL-zoned S98:130-12a Centre clear CL-striped S98:130-14a Core? clear CL-patchy S98:130-14b Tip clear CL-grey /bright

950 Ma) 0.06223 0.05417 0.05181 0.06286 0.06343 0.08896 0.07285

0.0685 0.0653 0.0755 0.0700 0.0732 0.0704 0.0710 0.0626

1.71 1.69 1.76 1.79 1.69 1.84 1.69 1.70

0.81 0.52 0.78 0.74 0.51 0.78 0.59 0.40

-5.7 -19.1 30.7 -30.5 -24.0 -44.8 6.1 -12.6

0.10 1 .44 0.06 4.06 0.16 2.38 0.42 3.42

371 85 284 349 146 514 242 519

174 1 209 25 89 13 68 61

31 5 27 26 13 39 20 35

0.47 0.01 0.74 0.07 0.61 0.02 0.28 0.12

452 501 362 618 592 774 417 446

27 60 32 34 60 30 51 85

431 422 451 466 479 497 438 399

7 11 8 9 13 9 10 14

427 408 469 436 455 439 442 391

7 7 8 8 7 8 7 6

1.95 6.72 7.26 0.68 0.85 0.86 1.12

0.936 0.342 0.420 0.878 0.984 2.826 1.582

2.87 7.04 7.64 2.21 2.27 2.34 2.38

0.1091 0.0458 0.0588 0.1012 0.1125 0.2304 0.1575

2.10 2.10 2.37 2.10 2.10 2.18 2.10

0.73 0.30 0.31 0.95 0.93 0.93 0.88

0.56 -2.2 -24.2 23.74 2.07 34.0 -12.2 0.43 ^2 0.05 0.41 -5.3 0.34 -7.1

278 4884 1049 2359 601 559 690

35 9829 1395 1313 90 447 84

33 328 84 283 74 165 119

0.12 2.01 1.33 0.56 0.15 0.80 0.12

682 378 277 704 723 1403 1010

41 145 158 14 18 16 23

671 299 356 640 695 1362 963

14 18 23 11 11 18 15

668 289 369 622 687 1337 943

13 6 8 12 14 26 18

898:131. Medium-grained grey granite, Isispynten (com. Pb 420 Ma) S98:131-01a Huge (700um) clear 0.05553 1.20 S98:131-01b zoned CL-grey zircon 0.05509 0.91 S98:131-04a Centre clear CL-zoned 0.06824 1.16 S98:131-04b Outer dark CL-zoned 0.06736 1.30 S98:131-05a Core? clear CL-bright 0.06256 2.67 S98:131-05b Core? clear CL-bright 0.06421 14.58 S98:131-05c Rim clear CL-grey 0.05152 5.11 S98:131-08a Clear CL-bright 0.07040 2.17 S98:131-08b Clear CL-grey 0.06934 1.06

0.499 0.484 1.074 1.211 1.202 0.489 0.276 1.066 1.385

2.43 2.29 2.43 2.47 3.39 14.73 5.57 3.04 2.38

0.0651 0.0637 0.1142 0.1304 0.1394 0.0552 0.0388 0.1098 0.1449

2.11 2.10 2.13 2.10 2.10 2.11 2.24 2.12 2.13

0.87 0.92 0.88 0.85 0.62 0.14 0.40 0.70 0.90

-6.4 -4.4 -21.5 -7.3 22.7 -55.2 -7.1 -30.0 -4.3

0.00 0.05 0.10 0.26 0.39 9.48 7.49 0.24 0.17

570 1138 370 396 104 165 3314 125 526

225 282 125 263 157 96 137 118 463

43 81 49 63 21 13 137 18 98

0.40 0.25 0.34 0.66 1.50 0.58 0.04 0.94 0.88

434 416 876 849 693 748 264 940 909

27 20 24 27 56 281 113 44 22

411 401 741 806 802 404 247 737 883

8 8 13 14 19 50 12 16 14

407 398 697 790 841 346 246 672 872

8 8 14 16 17 7 5 14 17

94045. Migmatite (palaeosome + neosome), Andreeneset, Kvit0ya (com. Pb 420 Ma) 94045-Ola Centre semi-turbid CL-patchy 0.05540 0.94 0.556 2.22 94045-02a Centre semi-turbid CL-dark 0.05623 1.31 0.618 2.40 94045-02b Tip semi-turbid CL-dark 0.05440 3.07 0.546 3.66 94045-04a Centre semi-turbid CL-patchy 0.05443 2.51 0.591 3.22 2.64 94045-05a Centre clear CL-grey 0.05516 1.72 0.605 94045-05b Tip clear CL-zoned 0.05412 1.80 0.526 2.70 94045-06a Centre semi-turbid CL-patchy 0.05257 21.87 0.191 21.96 94045-06b Rim semi-turbid CL-zoned/dark 0.06903 16.79 0.188 16.92 94045-07a Centre clear CL-bright 0.05373 1.88 0.555 2.86 94045-09a Centre semi-turbid CL-patchy 0.05686 1.56 0.583 2.59 2.41 94045-09b Rim semi-turbid CL-dark 0.05617 1.30 0.552

0.0728 0.0797 0.0728 0.0788 0.0796 0.0704 0.0264 0.0197 0.0749 0.0743 0.0713

2.01 2.01 2.01 2.03 2.01 2.01 2.02 2.09 2.15 2.07 2.03

0.90 0.84 0.55 0.63 0.76 0.74 0.09 0.12 0.75 0.80 0.84

5.8 7.5 17.5 26.7 18.7 17.2 -46.5 -86.8 30.5 -5.1 -3.4

0.03 0.08 4.30 1.44 0.11 0.40 35.07 69.99 0.18 1.54 0.41

404 424 695 310 166 269 994 5970 127 826 769

3 255 154 141 123 5 279 794 37 171 25

31 42 57 29 17 20 30 171 11 69 58

0.01 0.60 0.22 0.45 0.74 0.02 0.28 0.13 0.29 0.21 0.03

429 461 388 389 419 376 310 900 360 486 459

21 29 67 55 38 40 434 313 42 34 29

449 489 443 472 481 429 178 175 448 466 446

8 9 13 12 10 9 36 28 10 10 9

453 495 453 489 494 439 168 126 465 462 444

9 10 9 10 10 9 3 3 10 9 9

AJ94:004. Migmatite (palaeosome + neosome), Andreeneset, Kvit0ya (com. Pb 420 Ma) AJ94:004-02a Core turbid CL-grey 0.05570 5.59 0.617 5.95 3.33 AJ94:004-02b Inner rim clear CL-bright 0.05581 2.65 0.562 2.46 AJ94:004-02c Outer rim clear CL-zoned 0.05410 1.33 0.543 3.59 AJ94:004-03a Centre semi-turbid CL-bright 0.05513 2.91 0.592 2.33 AJ94:004-03b Tip clear CL-dark 0.05580 1.14 0.548 2.76 AJ94:004-04a Round centre clear CL-zoned 0.05481 1.75 0.599 2.11 AJ94:004-05a Core dark/turbid CL-dark 0.05581 0.66 0.575 10.86 AJ94:004-05b Inner rim fract./clear CL-bright 0.05678 10.62 0.332 2.61 AJ94:004-07a Centre semi-turbid CL-grey 0.05694 1.60 0.625 4.30 AJ94:004-09a Centre clear CL-dark 0.05690 3.80 0.589 4.42 AJ94:004-10a Centre turbid/zoned CL-zoned 0.05663 3.87 0.533

0.0804 0.0730 0.0729 0.0779 0.0712 0.0793 0.0748 0.0425 0.0796 0.0751 0.0683

2.04 2.01 2.08 2.11 2.03 2.13 2.01 2.30 2.06 2.01 2.13

0.34 0.60 0.84 0.59 0.87 0.77 0.95 0.21 0.79 0.47 0.48

13.7 2.2 21.5 16.5 -0.2 22.3 4.7 -45.4 1.0 -4.4 -11.1

2.37 0.00 0.17 0.20 0.36 0.10 0.35 5.78 0.08 0.20 0.00

176 59 399 99 610 130 1851 95 174 229 545

151 11 7 47 9 130 440 18 139 246 830

19 5 31 9 46 14 157 5 18 23 53

0.86 0.19 0.02 0.48 0.01 1.00 0.24 0.19 0.80 1.07 1.52

440 445 375 417 444 405 445 483 489 488 477

120 58 30 64 25 39 15 219 35 82 83

488 453 441 472 444 477 461 291 493 470 434

23 12 9 14 8 11 8 28 10 16 16

498 454 453 484 444 492 465 268 494 467 426

10 9 9 10 9 10 9 6 10 9 9

(continued)

Table 4. Continued Sample/ spot number

Description1

Measured ratios2

0.0696 0.0695 0.0714 0.0682 0.0691 0.0686 0.0678

1.32 1.22 1.22 1.30 1.22 1.21 1.22

0.71 0.77 0.92 0.91 0.85 0.80 0.73

18.0 10.8 1.6 6.0 6.0 10.2 0.7

.65 .60 .91 .98 1.92 1.56 1.69 1.41 1.54 1.76 1.80

0.0700 0.0687 0.0684 0.0666 0.0670 0.0709 0.0711 0.0707 0.0702 0.0719 0.0721

1.22 .22 .27 .21 .24 .23 .22 .21 .23 .30 1.22

0.73 0.76 0.67 0.61 0.65 0.79 0.72 0.86 0.79 0.74 0.68

1.546 1.579 1.602 1.590 1.550 1.502 1.574 8.840 4.620 0.780 1.630 1.511

1.50 1.28 1.31 1.33 1.41 1.44 1.30 1.24 1.23 1.92 3.72 1.26

0.1614 0.1627 0.1647 0.1626 0.1587 0.1554 0.1616 0.3803 0.3133 0.0852 0.1357 0.1566

1.22 1.21 1.21 1.22 1.21 1.21 1.21 1.21 1.21 1.22 1.33 1.21

1.596 1.887 1.565 1.567 1.572 1.723

1.49 1.42 1.60 1.57 1.53 1.36

0.1640 0.1848 0.1619 0.1603 0.1626 0.1721

1.21 1.23 1.22 1.22 1.22 1.23

AJ94:005. Aplite dyke, Andreeneset, Kvit0ya (com. Pb 420 Ma) AJ94:005-01a Centre clear CL-grey 0.05396 AJ94:005-01b Rim clear CL-zoned 0.05451 AJ94:005-02a Core semi-turbid CL-dark 0.05564 AJ94:005-03a Centre clear CL-grey 0.05474 AJ94:005-04a Rim clear CL-zoned 0.05487 AJ94:005-05a Centre clear CL-zoned 0.05444 AJ94:005-06b Rim zoned CL-zoned 0.05519

1.31 1.02 0.53 0.61 0.76 0.90 1.15

0.518 0.522 0.548 0.515 0.523 0.515 0.516

.85 .59 .33 .43 .43 .51 .67

G95:048. Syenite boulder, Nordmarka (com. Pb 420 Ma) G95:048-01a Centre clear CL-zoned 0.05551 G95:048-02a Centre clear CL-dark 0.05494 G95:048-02b Outer clear CL-grey 0.05540 G95:048-03a Centre clear CL-grey 0.05605 G95:048-03b Outer clear CL-grey 0.05550 G95:048-04a Centre clear CL-dark 0.05485 G95:048-04b Outer clear CL-zoned 0.05416 G95:048-05a Centre clear CL-dark 0.05523 G95:048-05b Outer clear/zoned CL-zoned 0.05436 G95:048-08a Centre clear CL-grey/sectorzon. 0.05556 G95:048-09a Centre clear CL-dark/ sectorzon. 0.05441

1.12 1.04 1.42 1.56 1.47 0.95 1.17 0.72 0.94 1.18 1.33

0.536 0.520 0.522 0.515 0.513 0.536 0.531 0.539 0.526 0.551 0.541

G95:049. Augen gneiss, Nordmarka (com. Pb 950 Ma) G95:049-01a Centre clear/fract. CL-bright 0.06951 G95:049-01b Tip clear CL-zoned/dark 0.07038 G95:049-02a Long centre clear CL-striped 0.07054 G95:049-03a Centre clear CL-zoned 0.07090 G95:049-04a Long centre clear CL-striped 0.07084 G95:049-06a Centre clear CL-striped 0.07009 G95:049-06b Tip clear CL-zoned 0.07062 G95:049-07a Short centre inclusions CL-dark 0.16861 G95:049-10a Centre semi-turbid CL-patchy 0.10695 G95:049-10b Tip tubid CL-dark/zoned 0.06639 G95:049-l la Core dark/turbid CL-bright 0.08715 G95:049-l Ib Overgrowth turbid CL-dark 0.06997

0.88 0.41 0.50 0.53 0.72 0.77 0.47 0.24 0.22 1.48 3.48 0.33

G95:050. Augen G95:050-03a G95:050-05a G95:050-09a G95:050-09b G95:050-10a G95:050-l la

0.87 0.72 1.04 1.00 0.93 0.57

gneiss, Nordmarka (com. Pb 950 Ma) Long centre clear CL-striped 0.07054 Long centre clear CL-striped 0.07408 Core? semi-turbid CL-zoned 0.07013 Tip clear CL-zoned 0.07091 Long centre clear CL-grey 0.07016 Short centre s-turbid CL-zoned 0.07263

f 2 o6 5 %

%

207pb

206pb

Disc.4 %

Err.3 corr.

±lcr %

207pb

Calculated ages

Concentrations

235U

±lCT

%

206pb 238U

±lCT

u

Th ppm

Pb ppm

Th6 U

207pb

ppm

0.00 0.06 0.07 0.06 0.02 0.04 0.92

171 285 1008 644 507 281 672

88 247 20 283 284 152 252

14 26 76 52 43 24 53

0.51 0.87 0.02 0.44 0.56 0.54 0.38

369 392 438 402 407 389 420

29 23 12 14 17 20 25

424 427 443 422 427 422 422

6 6 5 5 5 5 6

434 433 444 425 431 428 423

6 5 5 5 5 5 5

0.8 4.7 -0.5 -8.8 -3.5 9.0 17.9 4.7 13.7 3.1 16.1

0.04 0.00 0.05 0.00 0.00 0.04 0.11 0.01 0.01 0.00 0.04

360 422 309 237 266 369 253 436 305 301 361

318 417 249 147 164 336 152 391 208 230 331

33 39 27 20 22 35 22 41 27 28 35

0.88 0.99 0.80 0.62 0.62 0.91 0.60 0.90 0.68 0.77 0.92

433 410 429 454 432 406 378 422 386 435 388

25 23 31 34 32 21 26 16 21 26 30

436 425 427 422 420 436 432 437 429 446 439

6 6 7 7 7 6 6 5 5 6 6

436 428 426 416 418 441 443 441 437 448 449

5 5 5 5 5 5 5 5 5 6 5

0.81 0.95 0.92 0.92 0.86 0.84 0.93 0.98 0.98 0.63 0.36 0.97

6.0 3.7 4.4 1.9 -0.3 0.0 2.2 -21.4 0.6 -37.0 -42.4 1.2

0.27 0.03 0.03 0.01 0.03 0.02 0.02 0.01 0.23 2.92 8.47 0.01

201 644 562 372 240 271 708 477 943 4033 166 903

42 75 191 77 43 37 72 386 211 1146 90 57

36 116 109 68 43 47 126 253 344 382 28 154

0.21 0.12 0.34 0.21 0.18 0.14 0.10 0.81 0.22 0.28 0.54 0.06

914 940 944 954 953 931 946 2544 1748 819 1364 927

18 8 10 11 15 16 10 4 4 31 66 7

949 962 971 966 951 931 960 2322 1753 586 982 935

9 8 8 8 9 9 8 11 10 9 24 8

964 972 983 971 950 931 966 2077 1757 527 820 938

11 11 11 11 11 11 11 22 19 6 10 11

0.81 0.86 0.76 0.77 0.80 0.91

4.0 5.1 4.0 0.4 4.4 2.1

0.03 0.03 0.04 0.03 0.03 0.02

181 324 214 241 304 444

92 90 30 38 34 108

36 70 38 43 54 87

0.51 0.28 0.14 0.16 0.11 0.24

944 1044 932 955 933 1004

18 15 21 20 19 11

968 1077 957 957 959 1017

9 9 10 10 10 9

979 1093 967 958 971 1023

11 12 11 11 11 12

±lCT

207pb

'Description: position of spot, appearance of spot area in transmitted light, appearance of spot area in cathodoluminescence. Corrected for common lead according to Stacey & Kramers (1975) model at 420 Ma (206Pb/204Pb - 18.048, 207Pb/204Pb - 15.593, 208Pb/204Pb - 37.857) or 950 Ma (206Pb/204Pb - 17.157, 208pb/204pb = 35 85g) 3 207

Error correlation Pb/235U-206Pb/238U. Percent discordance; negative value = normal discordant; positive value = reverse discordant. 5 Fraction of 206Pb derived from common lead. 6 Th/U ratio directly from measured Th and U concentrations.

4

±lCT

206pb

±lo-

238|j

235U

206pb

207

Pb/204Pb - 15.519,

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

oscillatory zoning, and sometimes also display sector zoning in CL (e.g. zr 01, 02, 09; Fig. 9), but have no signs of cores. One zircon analysed by the Pb-evaporation method yielded a 207Pb-206Pb age of 446 ± 40 Ma (Table 3, Fig. 8j, error expanded according to number of scans). Eleven points from seven zircons analysed by the NORDSIM ion microprobe yielded a weighted average

227

2o6pb_238u age of 434 + 8 Ma (n = 11, MSWD = 4.8) and a somewhat lower 207Pb/206Pb age of 413 ±14 Ma ( n = l l , MSWD = 0.80; Fig. lOj), with no signs of inheritance. An age of c. 440 Ma for the Nordmarka syenite boulder is 20-30 Ma older than similar highly magnetic syenite and associated granite at Djupkilsodden, central Nordaustlandet (Gee et al 1999).

Fig. 10. NORDSIM U-Pb ion microprobe results on zircons from NE Svalbard. Size of symbols reflects la analytical error, numbering of zircons refers to Table 4.

228

A. JOHANSSON ETAL

Fig. 10. Continued

Sm-Nd isotope geochemistry Whole-rock powders of the investigated rocks have been analysed for Sm and Nd isotope composition. The results are reported in Table 5, and illustrated in Figure 11. The preparation and analyses took place in two different ways, designated A and B in Table 5. In both sets, 100-200 mg of rock powder was mixed with an appropriate amount of mixed 147Sm-150Nd tracer, and dissolved in HF

and HNO3 (concentrated 10:1 mixture) in teflon capsules inside steel bombs in an oven at 205 °C for a few days. REE as a group were then separated using standard cation exchange procedures with HC1 and HNOa as media. For set A, the second ion exchange step, separating Sm from Nd, was carried out with a-hydroisobutyric acid in overpressurized columns. Sm and Nd were then both loaded on Re double filaments and analysed in static mode on a Finnigan MAT261 mass spectrometer. For set

229

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

Table 5. Sm-Nd isotope data for rocks from central and eastern Nordaustlandet and adjacent islands, northeasternmost Svalbard Sample

Anal .

Rock type/locality

Sm1 (ppm)

Nd1 (ppm)

^Sm/144™1

143

Nd/144Nd ± 2am2

sNd3 (present)

S98:49A 898:50 898:55 898:56 898:121

B B B B B

Migmatite neosome, Innvika Migmatite neosome, Innvika Red granite, Innvika Grey granite, Innvika Grey granite, Parry0ya

4.43 7.30 3.31 5.84 5.08

14.3 36.4 12.3 24.5 24.6

0.1872 0.1211 0.1633 0.1444 0.1250

0.512167 0.512038 0.512104 0.512054 0.512057

± 13 ±9 ± 11 ± 13 ± 11

-9.2 -11.7 -10.4 -11.4 -11.3

898:122 898:123 898:124 898:125 898:126 G95:048 G95:049 G95:050 G95:051 898:130

B B B B B A A A A B

Red augen gneiss, Parry0ya Grey augen gneiss, Parry0ya Biotite schist, Parry0ya Grey granite, Foyn0yane Grey granite, Foyn0yane Syenite (boulder), Nordmarka Augen gneiss, Nordmarka Augen gneiss, Nordmarka Aplite dyke, Nordmarka Augen gneiss, Isispynten

6.96 6.74 11.6 4.25 3.93 8.11 7.96 6.72 3.06 15.1

32.7 30.8 59.9 21.2 23.5 43.2 33.5 27.5 10.4 79.3

0.1287 0.1322 0.1166 0.1214 0.1011 0.1135 0.1435 0.1479 0.1783 0.1152

0.512070 0.512086 0.511929 0.512157 0.512147 0.512134 0.512131 0.512148 0.512260 0.512082

± 12 ± 15 ± 12

±5 ±4 ±5 ±5 ±9

-11.1 -10.8 -13.8 -9.4 -9.6 -9.8 -9.9 -9.6 -7.4 -10.8

898:131

B

Grey granite, Isispynten

8.34

55.3

0.0912

0.512325 ± 7

-6.1

94062C AJ94:003 AJ94:003 898:127 898:128 898:129

B A B B B B

Gabbro, Isispynten Gabbro, Norvargodden, Stor0ya Gabbro, Norvargodden, Stor0ya Aplitic granite, Andreeneset, Kvit0ya Migmatite, Andreeneset, Kvit0ya Amphibolite, Andreeneset, Kvit0ya

5.27 1.03 1.02 9.69 10.6 6.05

23.8 2.21 2.20 46.2 53.3 21.6

0.1341 0.2804 0.2807 0.1268 0.1198 0.1696

0.512213 0.513355 0.513350 0.512166 0.512155 0.512645

± 11 ± 14 ± 19 ±8 + 13 ± 13

-8.3 + 14.0 + 13.9 -9.2 -9.4 +0.1

94045 AJ94:004 AJ94:005 AJ94:006 AJ94:006 BCR-1 BCR-1

A A A A B A B

Migmatite, Andreeneset, Kvit0ya Migmatite, Andreeneset, Kvit0ya Aplite dyke, Andreeneset, Kvit0ya Gabbro, Kraemerpynten, Kvit0ya Gabbro, Kraemerpynten, Kvit0ya Basalt standard Basalt standard

6.09 20.9 3.57 2.61 2.63 6.65 6.61

28.4 94.8 15.2 11.4 11.5 29.0 28.8

0.1297 0.1330 0.1424 0.1382 0.1385 0.1386 0.1386

0.512273 0.512198 0.512261 0.512689 0.512705 0.512622 0.512626

±5 ±4 ±5 ±5 ±8 ±6 ± 12

-7.1 -8.6 -7.4 + 1.0 + 1.3 -0.3 -0.2

± 10 ± 10

^Nd

3

(initial)

T

4

ICHUR (Ga)

7.41 -8.6 (450 Ma) -7.4 (450 Ma) 1.21 -8.5 (450 Ma) 2.43 -8.4 (450 Ma) 1.70 -7.2 (450 Ma) 1.23 -2.6 (950 Ma) -2.8 (950 Ma) 1.27 -2.9 (950 Ma) 1.30 -4.1 (950 Ma) 1.35 -5.1 (450 Ma) 0.97 -4.1 (450 Ma) 0.78 -5. 0(450 Ma) 0.92 -3.4 (950 Ma) 1.45 1.53 -3. 6 (950 Ma) -6.4 (430 Ma) 3.12 -6.4 (430 Ma) 1.04 -0.9 (950 Ma) -0.0 (450 Ma) 0.45 +6.7 (950 Ma) -2.7 (700 Ma) 1.03 +6.5 (700 Ma) 1.31 +6.4 (700 Ma) 1.29 -5.2 (450 Ma) 1.03 -5. 0(450 Ma) 0.96 + 1.7 (450 Ma) -0.04 +3.4 (950 Ma) -3.3 (450 Ma) 0.83 -4.9 (450 Ma) 1.05 -4.4 (430 Ma) 1.06 +6.2 (700 Ma) -0.13 +6.5 (700 Ma) -0.18 0.04 — 0.03

T

5

ADM (Ga)

8.01 1.65 2.90 2.18 1.69 1.74 1.79 1.75 1.46 1.21 1.38 1.98 2.07 3.57 1.48 0.90 1.58 0.70 0.68 1.53 1.44 1.37 1.39 1.59 1.67 0.73 0.70 0.86 0.86

'Sm and Nd contents and l47Sm/144Nd ratio from isotope dilution analysis with combined 147Sm-150Nd tracer. Estimated analytical uncertainty of 147Sm/144Nd ratio 0.5%. 2 143 Nd/ l44 Nd ratios calculated from ID run, corrected for Sm interference and normalized to 146Nd/144Nd = 0.7219. Sixteen runs of the La Jolla Nd-standard during measurement period A gave an average 143Nd/144Nd ratio of 0.511837 + 20 (2a), eight runs during period B gave 0.511845 + 8 (2a). Error given as 2 standard devi ations of the mean from the mass spectrometer run in the last digits. 3 Present-day and initial s Nd values at the indicated age, according to Jacobsen and Wasserburg (1984): present-day chondritic 147Sm/144Nd ratio 0.1967, present-day chondritic 143Nd/144Nd ratio 0.512638. 4 Model age calculated relative to the chondritic uniform reservoir (CHUR) of Jacobsen and Wasserburg (1984). 5 Model age calculated relative to the depleted mantle curve (DM) of DePaolo (1981).

B, the second ion exchange step was instead carried out with HC1 in open columns using the Ln-spec method (Pin & Zalduegui 1997), and Nd was analysed in multi-dynamic mode on the MAT261 mass spectrometer. Results of repeated runs of the La Jolla Nd-standard, using both methods, are reported in the footnote to Table 5. Duplicate analyses of three samples (gabbro samples AJ94:003 and AJ94:006 and basalt standard BCR-1) suggest excellent reproducibility between the two methods and sets of data (Table 5). In Table 5, TCHUR (chondritic uniform reservoir) model ages have been calculated according to Jacobsen & Wasserburg (1984), and TDM model ages according to the depleted mantle curve of De Paolo (1981). However, these are mostly of little geological significance. Of more interest are the initial eNd-values. These have been calculated at 950 Ma for the augen gneisses, inferred to be of Grenvillian age, and at 430-450 Ma for the migmatites, grey granites, syenite and aplite dykes, inferred to be of Caledonian origin. For a few rocks, where the age is uncertain, sNd-values are reported both for 950 and 450 Ma, and for the gabbros for an intermediate age of 700 Ma (Table 5). Since Grenvillian and Caledonian granitoids have similar Nd isotope characteristics, with the Caledonian granites largely

representing remelted Grenvillian crust (cf. Johansson et al. 2000, 2002), the Nd isotope composition cannot be used to discriminate between them. The augen gneisses in this study overlap with the range of eNd-values recorded for Grenvillian magmatic rocks in western and central Nordaustlandet ( — 2 to — 4.5; Johansson et al. 2000), indicating a similar crustal anatectic origin. One sample, S98:130 from Isispynten, has a somewhat less negative value of -0.9 at 950 Ma. However, as discussed earlier, the age of this rock is uncertain, and it may be a Caledonian granite. Both sample 898:130 and the other augen gneiss samples would have 8Nd-values overlapping with the field of Caledonian granites at 420 Ma from Johansson et al (2002). For the biotite schist sample 898:124 from Parry0ya, an 8Nd-value of -4.1 at 950 Ma was calculated, following its Grenvillian U-Pb titanite age. If it is a Mesoproterozoic metasediment, it represents a potential source rock for the magmas forming the protoliths of the Grenvillian augen gneisses. The migmatites, grey granites, aplitic dykes as well as the G95:048 syenite yield eNd-values in the range — 3 to —8.5 at 430-450 Ma, similar or slightly less negative than the —5 to —10 range recorded for Caledonian granites from western and central Nordaustlandet at 420 Ma by Johansson et al (2002).

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Fig. 11. eNd against time diagram for Grenvillian and Caledonian rocks from NE Svalbard. Initial 8Nd values are marked at 950 Ma (Grenvillian rocks), 430-450 Ma (Caledonian rocks), or 700 Ma (gabbros of uncertain age), dotted lines show extrapolation of the Nd evolution lines back to the depleted mantle curve of De Paolo (1981), shaded fields show initial Nd isotope compositions of Grenvillian and Caledonian magmatic rocks from western and central Nordaustlandet and their evolution trends to the present for comparison (from Johansson et al 2000, 2002).

The three migmatite samples from Kvit0ya fall in the upper part of the recorded range in this study with values at —3.3, — 4.9 and —5.0. However, these values still indicate substantial input from older continental crust, and would argue against them having formed from juvenile Ordovician sediments. Also syenite sample G95:048 has a strongly negative initial 8Nd-value of — 5.0 at 450 Ma. Only one granitic sample of presumed Caledonian age, Isispynten grey granite 898:131, shows a markedly more juvenile initial 8Nd-value of 0 at 450 Ma, thereby deviating from all other granites. For the mafic rocks, the age is highly uncertain. Ampibolite sample 898:129 from Kvit0ya would have had an sNd-value of +3.4 Ma at 950 Ma and +1.7 Ma at 450 Ma, when the surrounding migmatites were formed. The gabbro samples from Norvargodden on Stor0ya (AJ94:003) and Kraemerpynten on Kvit0ya (AJ94:006) show highly deviating Nd isotope characteristics (present-day eNd-values of +14 and +1, respectively), but their Nd evolution lines intersect each other and the depleted mantle curve of De Paolo (1981) at c. 700 Ma (Table 5, Fig. 11), hinting at the possibility that these gabbros formed from magmas of depleted mantle origin at around 700 Ma, perhaps in connection with the opening of the lapetus Ocean. The third gabbro sample, 94062C from Isispynten, however, has a totally deviating pattern, and would have an 8Nd-value of —2.7 at 700 Ma.

Discussion and conclusions This study shows that the Duvefjorden Complex of eastern Nordaustlandet consists of a mixture of Grenvillian and Caledonian rocks. The northeasternmost part of Svalbard and the Barentsia microcontinent is neither a stable Precambrian craton, unaffected

by Caledonian tectono-magmatic events, nor a juvenile Caledonian mobile belt, but a complex mixture of Grenvillian and Caledonian elements. The augen gneisses from Parry0ya, north of Nordaustlandet, and Nordmarka in easternmost Nordaustlandet have similar ages of around 950 Ma, and similar geochemical and Nd isotopic characteristics, as the Grenvillian granites and augen gneisses from western and central Nordaustlandet. This indicates that they formed during the same, late Grenvillian tectono-magmatic event, and that the Grenvillian magmatism was widespread throughout NE Svalbard. The Barentsia microcontinent, of which the Nordaustlandet Terrane is the main exposed part, may be composed largely of Grenvillian crust, overlain by Neoproterozoic and Early Palaeozoic shallow marine successions, as seen in central and western Nordaustlandet. Migmatization within the so-called Duvefjorden Complex in central Nordaustlandet (Innvika; Tebenkov et al. 2002) as well as on Damflya and Kvit0ya (this study) is clearly of Caledonian age. No signs of an earlier, e.g. Grenvillian, migmatization event have been recorded, although it cannot be excluded that earlier migmatization has occurred. The age of the protoliths of the migmatized paragneisses on Damflya and Kvit0ya is uncertain; these rocks may be metamorphosed Mesoproterozoic equivalents of the Brennevinsfjorden Group, or involve younger, Neoproterozoic metasediments. Migmatization on Kvit0ya appears to have occurred at a relatively early stage of the Caledonian Orogeny, at around 450 Ma, significantly earlier than the intrusion of the Rijpfjorden granite (c. 410 Ma) in central Nordaustlandet (Johansson et al 2002). The migmatization was followed by intrusion of anatectic grey granites, which largely originated from melting within the migmatite complex, and cross-cutting aplitic dykes, at c. 440-430 Ma. Also syenite formed at this time, as indicated by the c. 440 Ma age of sample G95:048. Grey granites, as well as aplitic dykes,

TECTONO-MAGMATIC ACTIVITY IN NE SVALBARD

are widespread throughout the area, from Innvika and Parry0ya in the west, via Foyn0yane, Nordmarka and Isispynten, to Kvit0ya in the east. This shows the regional extent of the Caledonian magmatism and migmatization within the Nordaustlandet Terrane. The apparent increase in Caledonian metamorphic grade eastwards, from greenschist facies in western Nordaustlandet, to migmatitic conditions in central Nordaustlandet and further east, may partly reflect the deeper level of erosion in the east. However, since the migmatization reaches higher structural and stratigraphic levels in the east (base of the Murchisonfjorden Supergroup at Innvika in central Nordaustlandet) than in the west (greenschist facies preserved in underlying Kapp Hansteen and Brennevinsfjorden Groups on Botniahalv0ya), there appears to be a true horizontal gradient of increasing metamorphic grade towards the east. The age of the gabbroic intrusions remains highly uncertain as no U-Pb data are available from these rocks. A Caledonian age, following Ohta (1978), is still possible. Sm-Nd data on the Stor0ya and Kraemerpynten gabbros hint at a Neoproterozoic age of c. 700 Ma, but this cannot be considered a proper radiometric age date, as it is only based on two sample points. However, from a geotectonic point of view, it would seem attractive to relate these mafic rocks to crustal rifting connected to the formation of the lapetus Ocean in late Neoproterozoic time. Possibly related are the enigmatic 600-700 Ma zircon ages discovered in prismatic zircons that the Isispynten augen gneiss, zircons that may have been inherited from late Neoproterozoic rift-related rocks at depth. Similar ages of c. 650 Ma were recorded by Peucat et al (1989) and Gromet & Gee (1998) from zircons in felsic and mafic rocks from the Richarddalen Complex in NW Spitsbergen, and interpreted by Gromet & Gee (1998) to reflect a late Proterozoic rifting event. It may be that this rifting event can be correlated from NW Spitsbergen to easternmost Svalbard, although evidence in the east is still meagre. Alternatively, these zircons may be of detrital origin, having been picked up by the granitic magma from some early Palaeozoic sedimentary rocks, their ultimate origin being distant late Neoproterozoic (Timanian) magmatism, e.g. in the Timan range of northern Russia. In summary, northeastern Svalbard consists of late Grenvillian (c. 950 Ma) continental crust, formed by melting of earlier Mesoproterozoic metasedimentary rocks. The Grenvillian crust may have undergone a cryptic late Neoproterozoic (c. 600-700 Ma) rifting event with intrusion of gabbroic massifs, prior to massive late Ordovician (c. 450 Ma) migmatization and subsequent granitic to syenitic magmatism. Sedimentary evidence indicates that eastern Svalbard was a stable platform area with more or less continuous sedimentation during the late Neoproterozoic and early Palaeozoic, lacking signs of any Timanian orogenic activity. Thus, there is no obvious connection between eastern Svalbard and the Timanide Orogen in northern Russia, with its late Neoproterozoic magmatism. As eastern Svalbard (Barentsia) was probably part of the Laurentian margin in pre-Caledonian time, its position relative to the Timanides in the late Neoproterozoic is unknown; these areas may have been located far apart, and are separated, today, by Caledonian sutures trending NE through the central Barents Sea. Sample 31-5 from Damflya was obtained by our Russian colleagues during fieldwork for the Russian Polar Marine Geological Research Expedition, with zircons and monazite separated at the Geological Institute, Kola Science Center, Apatity; sample 94045 from Kvit0ya was collected by photographer L. Aby; and sample 94062C from Isispynten by Captain P. Engwall of M/S Origo. The remaining samples were collected during SWEDARCTIC fieldwork financed by the Swedish Polar Research Secretariat. We want to specially thank the crews of M/S Origo and M/S Lance, as well as the Russian helicopter pilots, for making it possible for us to reach these remote places in the high Arctic. M. Fischerstrom (Stockholm) skilfully carried out the Sm-Nd chemical preparation work. The NORDSIM ion microprobe analyses were performed under the guidance of T. Sunde, M. J. Whitehouse, K. Linden and K. Hogdahl of the NORDSIM laboratory in Stockholm, with the epoxy mounts originally prepared by J. Vestin, and SEM/CL observations carried out at Stockholm University

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under the guidance of M. Ahlbom. This paper has benefited from careful reviews by L. P. Gromet (Providence) and A. Andresen (Oslo). AJ has received funding from the former Swedish Natural Science Research Council (NFR). ANLs stay in Stockholm was initially financed by a guest researcher grant from the Swedish Institute, and later by a post-doctoral grant from NFR. The NORDSIM ion microprobe laboratory is financed under contract with the Nordic Natural Science Research Councils and the Swedish Museum of Natural History. This is NORDSIM publication no. 86, and a contribution to the SWEDARCTIC programme on Svalbard, the TIMPEBAR (Timan-Pechora Barents Sea) project of the Europrobe programme and INTAS-HALE project, and to IGCP-project 440: 'Assembly and Break-up of Rodinia'.

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Neoproterozoic Orogeny along the margins of Siberia 1

V. A. VERNIKOVSKY1, A. E. VERNIKOVSKAYA1, V. L. PEASE2 & D. G. GEE3 United Institute of Geology, Geophysics and Mineralogy, Siberian Branch, Russian Academy of Sciences, Koptyug Prosp., 3, 630090, Novosibirsk, Russia (e-mail: [email protected]) 2 Department for Geology and Geochemistry, Stockholm University, SE-10691 Stockholm, Sweden 3 Department of Earth Sciences, Uppsala University, Villavagen 16, SE-75236 Uppsala

Abstract: The Siberian Craton is bounded by fold-and-thrust belts involving Neoproterozoic (locally Mesoproterozoic) complexes on the southern (Baikal-Vitim), western (Yenisey Ridge and Turukhansk-Igarka), northern (Taimyr), and eastern (Verkhoyansk) sides. This paper focuses on the geological structure and evolution of these formations. Previous and new geochronological data show that passive continental margins existed around most of the Siberian Craton during the early Neoproterozoic and possibly the late Mesoproterozoic. Between about 850-760 Ma, the southern, western and northern passive margins of the Siberian Craton were transformed into active margins. Middle-late Neoproterozoic island arcs and ophiolites were formed between c. 750-650 Ma along these margins; they are inferred to have been obducted onto the Siberian continental margin at c. 600 Ma, prior to late Vendian deposition. New ion microprobe U-Pb ages of ophiolitic rocks from Taimyr's Central domain are presented. The Neoproterozoic record in the Cretaceous Verkhoyansk fold-and-thrust belt indicates that the eastern part of the Siberian Craton remained a passive continental margin during the Neoproterozoic and Palaeozoic. Baltica-Siberia relationships are also discussed.

This paper reviews and integrates recent geological and geochronological data of Neoproterozoic complexes around the Siberian Craton; these Precambrian rocks occur in the Yenisey Ridge, Igarka-Turukhansk, Taimyr, B ay kal-Vitim and Verkhoyansk fold-and-thrust belts. The new data allow us to recognize important stages in the formation of the Neoproterozoic Orogens and to correlate the successions of the different regions and the tectonic events of the Neoproterozoic magmatic complexes better. The results of these studies are useful for global palaeogeographic reconstructions relevant to the formation and breakup of the Rodinia supercontinent and relationships between Siberia and Baltica. The new geological and geochronological data are very important in this respect as palaeomagnetic arguments about the relative positions of Baltica, Laurentia and Siberia, and classic Rodinia reconstructions are contentious (Piper 1976; Sears & Price 1978; Bond et al 1984; Hoffman 1991; Dalziel 1991; 1997; Powell et al 1993; Condie & Rosen 1994; Torsvik et al 1996; Weil et al 1998; Pisarevsky et al 2000; Meert & Powell 2001; Hartz & Torsvik 2002; Wingate et al 2002). These Neoproterozoic events are often referred to in Russian literature as 'Baikalian events' or 'Baikalian Orogeny' and the orogen as the 'Baikalides' (Shatsky 1963; Khomentovsky et al 1969; Klitin et al 1970; Bulgatov 1983; Postel'nikov 1993; Khain & Rudakov 1995). The term 'Baikalian Orogeny' was introduced by Edelshtein and Shatsky (Edelshtein 1923; Shatsky 1963), who had in mind a tectonic era that started in the Late Precambrian and finished in the Early Palaeozoic, i.e. in a period after the Grenville epoch and lasting to the beginning of the Caledonian epoch (Khain & Rudakov 1995). In the type area of the Baikalides, superimposed Caledonian age events are much more widespread than the preceding Baikalian events. Some authors prefer to restrict 'Baikalian' events to those that took place in the narrower time interval of 850-650 Ma (Khomentovsky 1996, 2002). The terms 'Baikalian' and 'Baikalides' have also been used to designate a Late Precambrian stratigraphic system for Siberia, corresponding to the Cryogenian of the International Stratigraphic Chart (Plumb 1991). Thus, the Baikalian system would include passive margin, back-arc basin and island-arc units and ophiolites around the Siberian craton, that comprise the fold belts formed at 850-650 Ma (Khomentovsky 2002). This usage is controversial. Recent investigations have shown that Neoproterozoic accretionary processes along the margins of the Siberian Craton, with the formation of different fold-and-thrust belts, did not take place simultaneously (Dobretsov et al 1992, 2003; Vernikovsky

et al 1996, 19990, 2002; Kuzmichev et al 2001; Pease et al 2001; Khain et al 2002; Sklyarov et al 2003); in southern regions they continued into Palaeozoic time. We provide evidence that Neoproterozoic Orogeny along the southern, western and northern margins of the Siberian continent occurred over a wider time interval than 850-650 Ma (cf. Khomentovsky 1996, 2002). Although not entirely agreed between ourselves, we have decided not to use the term Baikalian in this paper, and instead refer only to Neoproterozoic Orogeny. Neoproterozoic complexes provide evidence of the existence of both passive and active margins around the Siberian Craton. Neoproterozoic ophiolites are present in all areas except those in the east and, even in these areas, the Neoproterozoic sedimentary successions indicate the probable existence of an outboard ocean (Fig. 1). The fold-and-thrust belts containing the Neoproterozoic complexes are described below.

The Baikal-Vitim fold-and-thrust belt The Palaeozoic Baikal-Vitim fold-and-thrust belt resulted from the collision of a microcontinent (Barguzin) with the Siberian Craton and involved granitoid magmatism, metamorphism and extensional deformation of both Palaeozoic and Proterozoic age (Konnikov et al 1994). The orogen is situated between Lake Baikal and the Vitim River; it comprises a major arcuate structure separating the Angara and Aldan-Stanovoy blocks of the Siberian Craton (Fig. 1). This re-entrant is bounded by the peri-Baikal thrust in the NW and by the Zhuin thrust in the east (Fig. 2). These faults are complicated by oblique movement, sinistral in the west and dextral in the east (Alexandrov 1990; Dobretsov & Bulgatov 1991; Gusev & Khain 1995). The minimum amount of displacement is 30 km along the peri-Baikal thrust and 25 km along the Zhuin thrust (Gusev & Khain 1995). The footwall of these faults is well exposed along their northern margins and consists of Neoproterozoic siliciclastic and carbonate units with abundant oncolitic and stromatolitic limestones and dolostones; these are inferred to have formed as part of the palaeoshelf of the Siberian Craton. These units are overthrust by strongly deformed, high-grade granite-intruded metamorphic terranes of Neoproterozoic and Palaeoproterozoic (?) age (Klitin et al 1970; Bulgatov 1983; Parfenov et al 1996; Konnikov et al 1999). The thickness of Neoproterozoic passive margin units

From: GEE, D. G. & PEASE, V. (eds) 2004. The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 233-247. 0435-4052/047$ 15 © The Geological Society of London 2004.

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Fig. 1. Sketch map of the Siberian Craton (approximate location of craton edge marked by broken line; tagged lines mark main thrusts).

increases southwards and they change their composition to mixed siliciclastic, calcareous and organic-rich formations typical of the deeper water environments of a continental slope (Khain 1979, 2001). Within the Baikal-Vitim fold belt, fragmented ophiolites and island-arc units (Baikal-Muya Complex) are thrust over the Neoproterozoic strata of the passive continental margin. Although these ophiolites have been inferred to be Proterozoic in age (Dobretsov 1982; Konnikov 1991; Gusev et al 1992), their Neoproterozoic age has been defined only recently. U-Pb zircon and Nd isotope studies of plagiogranites and island-arc volcanic rocks indicate that they were formed within the interval of 900 to 812 Ma (Neimark et al 1995; Izokh et al 1998; Rytsk et al 2001). Futher evidence for the pre-Vendian age of these epidote-amphibolite facies, island-arc units and ophiolites is, that they are unconformably overlain by late Neoproterozoic (including Vendian) and Cambrian successions of little metamorphosed bimodal subalkaline volcanic rocks, gritstones and sandstones (Gusev & Khain 1995) and also, that they are intruded by 556 ±16 Ma post-tectonic granitoids of the Lesnoy suite (Sryvtsev et al 1992). The early Neoproterozoic Baikal-Muya ophiolites and island arc rocks were formed outboard of the Siberian margin; they were accreted onto the margin of the Siberian continent later in the Neoproterozoic, in pre-Vendian or Vendian time (Khain et al 1997). However, the strongest deformation stage—folding, thrusting and strike-slip faulting—resulted from the Early-Middle Palaeozoic collision of the Barguzin microcontinent with Siberia. This collision resulted in the folding of the Cambro-Silurian sedimentary cover of the Siberian Craton; only the Middle Devonian

deposits of the Siberian platform remained undeformed (Khain 1979).

Yenisey Ridge The Yenisey Ridge is a complex Proterozoic fold-and-thrust belt with a dominantant NW-SE strike; it provides the western frame of the Siberian Craton and can be traced more than 700 km along the Yenisey River (Figs 1 & 3). The NW-trending structures of the belt are divided into two segments, separated by the ENE-trending, sinistral strike-slip Angara Fault. South of the Angara Fault, two allochthonous units have been recognized, the Palaeoproterozoic granulite-amphibolite facies Angara-Kan Terrane and the Neoproterozoic, mainly island arc, Predivinsk Terrane. North of the Angara Fault, the Yenisey Ridge is composed of thrust sheets predominantly with Neoproterozoic rocks, comprising the East Angara, Central Angara and Isakov terranes (Fig. 3 inset). The Angara-Kan Terrane consists of volcanic and sedimentary rocks that were generally metamorphosed to amphibolite facies, but also include two-pyroxene-plagioclase gneisses and highly aluminous garnet-biotite-sillimanite-cordierite gneisses (Nozhkin 1999; Popov 2001). These rocks were intruded by 1837 ± 3 Ma Taraka granites (Bibikova et al 1993). The Predivinsk Terrane was thrust eastwards onto the Angara-Kan terrane. It consists mainly of Neoproterozoic greenschist-facies

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Fig. 2. Schematic geological map of the Baikal-Vitim fold-and-thrust belt, after Bulgatov (1983), Parfenov etal. (1996), Khain (2001). MP, Mesoproterozoic; NP, Neoproterozoic.

island-arc assemblages, including basalt, andesite, dacite, and rhyolite. Gabbro-diorite-plagiogranite massifs are associated with the volcanic rocks. This oceanic allochthon includes serpentinized harzburgites (also serpentinite melange), gabbroids, and tholeiite basalts, that occur in the eastern and central zones of the terrane. Zircon U-Pb ages from a metarhyolite in the differentiated calc-alkaline series and a plagiogranite from this terrane yielded ages of 660-637 Ma (Vernikovsky et al 19990, 2003).

The East Angara Terrane, north of the Angara Fault, consists of Neoproterozoic siliciclastic and calcareous sediments (sandstone, siltstone, mudstone with sparse beds of dolostone and limestone), metamorphosed at greenschist, or lower, grade. These sediments are deformed into monoclines, box folds, and flexures. Khabarov et al. (1998) described the environments of deposition to be basin plain, slope, reef-bearing shelf margin, and peritidal and sufetidal inner shelves. In contrast to the other thrust sheets of the Yenisey Ridge, this allochthon is characterized by the absence

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Fig. 3. Geological sketch map of Yenisey Ridge, inset map shows locations of the five terranes discussed in the text. MP, Mesoproterozoic; NP, Neoproterozoic; PP, Palaeoproterozoic.

of Neoproterozoic magmatism. We consider this unit to be an essential part of the Neoproterozoic passive continental margin of Siberia. These greenschist-facies metasediments are unconformably overlain by late Neoproterozoic (including Vendian) molasse and by Vendian and Cambrian dolostone and limestone (Khomentovsky et al 1972; Postel'nikov 1980). The Central Angara Terrane was emplaced onto the East Angara terrane along the Ishimba thrust. At the base of the allochthon in the easternmost part of the Central Angara Terrane, ophiolite slices occur in the so-called Ribnaya-Panimba belt (Figs 3 & 4). However, most of the Central Angara terrane comprises Meso (?)-Neoproterozoic turbidites and carbonate sediments that were metamorphosed under greenschist and amphibolite facies (Kozlov & Lepezin 1995). These rocks were originally considered to be Archean or Palaeoproterozoic in age (Volobuev et al 1973; Kachevsky et al 1998). However, recent U-Pb ages of 880-865 Ma for Teya and Eruda granites (Nozhkin et al 1999;

Fig. 4. Schematic WSW-ENE cross-section of the Yenisey Ridge fold-and-thrust belt (see Fig. 3, A-A', for location) showing the divergent thrust character of the belt and the locations of the East Angara, Central Angara and Isakov terranes.

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NEOPROTEROZOIC OROGENY OF SIBERIA Table 1. Isotopic-geochronological data for the Neoproterozoic rocks ofYenisey Ridge

Age (Ma)

Interpretation

References

Central Angara terrane, Yeruda Granite whole rock Sm-NdTDM-2 zircon U-Pb (conventional)

2000-2100 878 ± 2

Model age of crustal source Age of crystallization

Vernikovskaya et al. 2002a

Central Angara terrane, Teya Granite zircon U-Pb (conventional)

866 ±16

Age of crystallization

Nozhkin^tf/. 1999

Central Angara terrane, Cherimba Quartz syenite whole rock Sm-NdTDM-2 zircon U-Pb (conventional) biotite Ar/Ar

2000 761+8 721 ± 1.6

Model age of crustal source Age of crystallization Cooling age following metamorphism

Vernikovskaya et al 20020

Central Angara terrane, Ayakhta Granite whole rock Sm-NdTDM_2 zircon U-Pb (conventional)

2130 750 ± 2

Model age of crustal source Age of crystallization

Vernikovsky et al. 2003

Central Angara terrane, Garevka Leucogranite whole rock Sm-NdTDM-2 zircon U-Pb (conventional)

1600 752 ± 3

Model age of crustal source Age of crystallization

Vernikovsky et al. 2003

Central Angara terrane, Glushikha Leucogranite whole rock Sm-NdTDM-2 zircon U-Pb (conventional)

1900 731+5

Model age of crustal source Age of crystallization

Vernikovskaya et al. 2003

Central Angara terrane, Strelka Leucogranite whole rock Sm-NdTDM-2 zircon U-Pb (conventional)

2200 718 + 9

Model age of crustal source Age of crystallization

Vernikovskaya et al. 2003

Ribnaya-Panimba ophiolite belt, Gabbro-amphibolite amphibole, plagioclase Ar/Ar

1050-900

Dated material

Method

Age of metamorphism

Vernikovsky et al. 2000

Yenisey Ophiolite belt, Isakov terrane, Porozhnaya Plagiogranite whole rock Sm-NdTDM_2 1272 zircon U-Pb (conventional) 697 + 4

Model age of crustal source Age of crystallization

Vernikovsky et al. 200la

Yenisey Ophiolite belt, Predivinsk terrane, Rhyolite zircon U-Pb (conventional)

Age of crystallization

Vernikovsky et al. I999a

Model age of crustal source Age of crystallization

Vernikovsky et al. 2003

638 + 6

Yenisey Ophiolite belt, Predivinsk terrane, Yagunov Plagiogranite whole rock Sm-NdTDM-2 1008 zircon U-Pb (conventional) 628 + 3

Vernikovskaya et al 20020) indicate that they are part of the Neoproterozoic evolution (Table 1). These granite-gneisses, gneisses and schists are unconformably overlain by siliciclastic and carbonate sediments (known as the Sukhopit and Tungusik formations) which are metamorphosed to greenschist and locally epidote-amphibolite facies. The Sukhopit and Tungusik units comprise metasandstone, metasiltstone, quartz-chlorite-sericite and carbonaceous schist, overlain by rhythmically laminated greywackes, limestones and dolostones. The Ribnaya-Panimba ophiolite-bearing belt consists essentially of a mafic-ultramafic complex, overlain tectonically by carbonaceous phyllite, arkose and quartz sandstone. Locally quartz-feldspar-actinolite schist, derived from mafic tuffs, occurs along the thrust contact. Elsewhere pyroxenite and gabbro, comprising differentiated sills, are present and a sheeted dyke complex has been recognized (Kheraskova 1999). In some places, cherts, conglomerates with ophiolitic clasts, and gritstone occur in the lower part of the section. In the middle and upper parts of the sections, altered pillow-basalts, metatuffs and hyaloclastites with pillow basalts and pillow breccias are present. The composition of the basalts is close to that of tholeiites formed in island arcs and marginal seas, as determined from high field strength (HFS) element ratios. Ar/Ar ages on hornblende and plagioclase from a gabbro-amphibolite in the Ribnaya-Panimba ophiolite belt are 1050-900 Ma (Vernikovsky et al. 2000) which has been inferred to reflect the time of their accretion to the Siberian craton. However, more isotopic data are required to determine the crystallization age of these, apparently Late Mesoproterozoic, ophiolites.

All these rocks of the Central Angara Terrane, including the Ribnaya-Panimba ophiolite, were intruded by the calc-alkaline Ayakhta-Chirimba granites at c. 760 Ma (U-Pb zircon data, Vernikovskaya et al. 20020) which have well-developed contact aureoles (Datsenko 1984) indicating their high level of emplacement and post-tectonic character, and by smaller bodies of Glushikha leucogranites at c. 750-720 Ma (Vernikovsky et al. 2002; Vernikovskaya et al 2003). These metamorphic and igneous rocks are overlain unconformably by little deformed Upper Neoproterozoic to Lower Cambrian sedimentary rocks of the Vorogovka, Chingasan and Chapa groups (Khomentovsky et al 1972; Kornev et al 1974; Postel'nikov 1980; Kachevsky et al 1998). The Isakov Terrane was thrust eastwards over the Central Angara terrane, the boundary being a complex transpressive fault zone with displacement of at least several tens of kilometres (Postel'nikov 1980). The tectonically imbricated terrane consists of Neoproterozoic volcanogenic and sedimentary units and ophiolite (Vernikovsky et al 19940) and is similar in composition to a modern island arc. The central part of the terrane comprises metamorphosed fragments of ophiolite (peridotite, gabbro and tholeiitic basalt), and also phyllite and carbonate-bearing quartz-mica schist. In some places, composite dykes and sills of diabase and diabase-porphyry occur among the metabasalts (Mironov & Nozhkin 1978; Postel'nikov 1980; Kuzmichev 1987; Volobuev 1993). A volcanic suite of calc-alkaline rhyolites, andesites and basalts, associated with metamorphosed tuff, sandstone, phyllite and limestone both overlie and underlie the ophiolitic units.

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Schists adjacent to the tectonic contacts are strongly refolded and injected with quartz veins, and a tectonic melange with carbonate and serpentine clasts has been mapped. Metamorphism in the southern and eastern parts of terrane reaches epidote-amphibolite facies with garnet in both the schists and metabasalts. Towards the north and west, higher in the Isakov allochthon, the metamorphic grade drops to greenschist facies. U-Pb zircon data from island arc Porozhnaya plagiogranite yielded a concordant age of 697 ± 4 Ma (Vernikovsky et al 20010). This is considered to be the upper age limit for the formation of island arc and ophiolitic rocks. Rb-Sr and K-Ar studies of metamorphic minerals from basites and pelites suggest that 620-600 Ma was the time of terrane obduction onto the passive margin of the Siberian Craton (Vernikovsky et al 19940). The Isakov Terrane is unconformably overlain by Late Neoproterozoic-Early Cambrian siliciclastic formation deposits (Postel'nikov 1980; Kuzmichev 1987), including Vendian molasse (Sovetov et al. 2000).

The Turukhansk-Igarka region Isolated inliers of Precambrian rocks, located between the Yenisey Ridge and Taimyr fold belts, occur within the Turukhansk and Igarka areas (Fig. 1). Within the Turukhansk area (often referred to as the Turukhansk Uplift), up to 4.5 km of Mesoproterozoic (?)-Neoproterozoic strata occur in three north-trending, faultbounded and east-transported blocks where the formations either dip gently westwards or occur in asymmetrical synclines with a gentle eastern limb and subvertical, partly truncated, western limbs (Petrov & Semikhatov 2001). These sandstones, shales, calcareous mudstones, dolostones, and limestones, including reef-bearing carbonate units were deposited on a passive continental margin (Petrov & Veis 1995; Petrov & Semikhatov 1997,2001; Sergeev et al 1997; Bogdanov et al 1998; Pisarevsky & Natapov 2003). The Turukhansk sediments include Neoproterozoic silicified microfossils and stromatolite assemblages (Petrov & Veis 1995; Sergeev et al 1997; Knoll & Semikhatov 1998). Limited isotopic-geochemical data constrain the age of these deposits to between about 1035 Ma (Pb/Pb carbonate; Ovchinnikova et al 1995) and 800 Ma (K-Ar glauconite; Semikhatov & Sebiryakov 1983). Vendian to Lower Cambrian strata were deposited unconformably across the different horizons of the Turukhansk succession (Petrov & Semikhatov 2001). The Igarka Terrane, composed mainly of volcanogenic rocks, was thrust from the west onto the Mesoproterozoic (?)-Neoproterozoic carbonate and siliciclastic deposits and overlying Vendian-Cambrian strata of the Turukhansk unit. This volcanogenic complex includes pillow-basalts, andesitic and basalt porphyrites, tuffs and spilites with interlayers of green and black schists, and also tuff-breccia. Diabase sills and dykes, quartz porphyries, and granodiorites are also present. These units are unconformably overlain by Neoproterozoic (?) and VendianCambrian carbonate and siliciclastic deposits (Postel'nikov 1989; Nekrasov & Berendeev 1991). Unfortunately, detailed petrological and geochemical and isotopic age data are lacking. However, the lithologies of this unit and the correlation of the overlying strata with those on the Turukhansk block suggest that the Igarka igneous and associated rocks were thrust onto the Siberian Craton margin in the Neoproterozoic. Comparison of the Igarka and Isakov island-arc complexes (Yenisey Ridge) suggest that the former are of Neoproterozoic age.

Taimyr The Taimyr, late Palaeozoic-Early Mesozoic, NE-striking foldand-thrust belt (Bezzubtsev et al 1986; Uflyand et al 1991; Vernikovsky 1996) can be traced almost 1000km along the

Kara Sea coast; it is situated between the Ural-Novaya Zemlya Orogen in the west and the Verkhoyansk fold belt in the east (Fig. 1). Its southern boundary is hidden under the Mesozoic and Cenozoic sediments of the Yenisey-Khatanga trough, and its northern part is exposed in the Severnaya Zemlya Archipelago. The Taimyr fold belt (Fig. 5) is divided into three NE-trending tectono-stratigraphic domains (southern, central, and northern), with the central and northern domains containing different Neoproterozoic terranes. The former was accreted to the Siberian margin during the Neoproterozoic and the latter (included in the Kara Terrane of Fig. 5) in the Late Palaeozoic. The three domains are separated by the major faults: the Pyasina-Faddey, Main Taimyr, and Diabasovy thrusts (Urvantsev 1949; Bezzubtsev et al 1986; Uflyand et al 1991; Vernikovsky 1996). The southern domain is composed of an unmetamorphosed Ordovician-Permian shallow marine platform (carbonatedominated) succession covered by Late Permian to Early Triassic sandstones, hosting flood basalts and dolerite dikes and sills (Bezzubtsev et al 1986; Uflyand et al 1991; Inger et al 1999). All rocks of the southern domain are folded and thrust southward, with a significant decrease in the intensity of both folding and faulting to the south. The domain was intruded by 249-245 Ma A-type unorogenic granites and syenites (U-Pb zircon data, Vernikovsky et al 2001Z?). Taimyr's northern domain is composed of a Neoproterozoic to Cambrian succession dominated by turbidites; these rhythmically alternating sandstones, siltstones and shales may be interpreted as continental slope sediments. The internal structure of this zone is complicated by Late Palaeozoic regional metamorphism and granite intrusion (300-265 Ma, U-Pb zircon; Vernikovsky et al 1995; Pease 2001), thrust faults of various scales, and SE-vergent folding. The Neoproterozoic-Cambrian deposits continue northwards to Severnaya Zemlya, where they are unconformably overlain by an Ordovician shallow-water, lagoonal, and coastal-marine facies with some rift volcanics and high level intrusions, thick Silurian limestones and Devonian sandstones (Kaban'kov & Sobolevskaya 1981; Makar'ev et al 1981; Gee & Pease 1999; Bogolepova et al 2001). The northern and southern Taimyr domains are separated by the central domain, also known as the Central Taimyr accretionary belt (Vernikovsky & Vernikovskaya 2001). It is composed mainly of Neoproterozoic sedimentary and volcanic rocks, including ophiolites and island arc magmatic suites and continental crust (Uflyand et al 1991; Vernikovsky et al 1996, 1998; Pease et al 2001). The tectonic contacts along thrust faults between individual terranes are often marked by cataclastic, mylonitic and autoclastic melange zones. These Precambrian rocks were folded and thrust together in the late Neoproterozoic and unconformably overlain by latest Neoproterozoic (Vendian)-Palaeozoic successions (Bezzubtsev et al 1986; Zabiyaka et al 1986; Uflyand et al 1991; Vernikovsky 1996). The older complexes of Taimyr's central belt occur within the Mamont-Shrenk and Faddey terranes (Fig. 5). They comprise metamorphosed (epidote-amphibolite and amphibolite facies) siliciclastic rocks, plagiogneisses and crystalline schists. Sheets and lenses of metabasic rocks transformed into biotite-amphibole crystalline schists and amphibolites, as well as quartzites and marbles, are present to a lesser extent (Bezzubtsev et al 1986; Zabiyaka et al 1986; Makhlaev 1988; Vernikovsky 1996). All are intruded by early Neoproterozoic granites, granite-gneisses and migmatites. Metapelites from both terranes show similar chemical compositions and high sediment maturity, corresponding to deposits of a passive continental margin (Vernikovsky 1995). The granites of these terranes form lenses and massifs that have been drawn out concordantly within the foliation; they occupy several hundreds of square kilometres. They are intensely foliated, locally cataclastic and mylonitic, and have an augen gneiss texture. They are typically two-mica, muscovite-bearing

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Fig. 5. Tectonic map of the Taimyr- Severnaya Zemlya fold-and-thrust belt.

porphyraceous granites of calc-alkaline and subalkaline affinities (Makhlaev 1988; Vernikovsky et al 1998; Pease et al 2001). U-Th-Pb zircon analyses of Shrenk granites were performed using a Cameca IMS 1270 ion-microprobe at the Swedish Museum of Natural History, Stockholm. Six granitic samples gave crystallization ages between 880 and 940 Ma (Pease et al. 2001). Xenocrystic zircons range in age from 1.2 Ga to 2.6 Ga, but are almost entirely Mesoproterozoic. A sillimanite gneiss, representing a metasedimentary raft entrained within granite, contains detrital zircons as young as c. 1000 Ma. Consequently, granite intrusion occurred shortly after the deposition of these metasediments and amphibolite facies metamorphism of the Mamont-Shrenk terrane is early Neoproterozoic in age (Pease et al. 2001). Sm-Nd isotope geochemistry shows that the formation of these granites included Mesoproterozoic to late Palaeoproterozoic material, the most ancient depleted mantle Nd model age for this continental crust being 1.6-2.1 Ga (Vernikovskaya et al. 20026). Late Grenvillian-age granites and metasedimentary rocks of the Mamont-Shrenk terrane are overlain unconformably by a Neoproterozoic succession of conglomerates and sandstones

(Krasnorechenskaya Formation), dolomites and stromatolitic limestones (Kolosovskaya Formation), volcanic and volcaniclastic rocks (Svetlinskay Formation) and sandstone and conglomerate (Posadochnaya Formation). Relationships between this terrane and the ophiolites have not been defined, but both the continental and oceanic crust were accreted to the margin of the Siberian Craton prior to the deposition of late Vendian and Palaeozoic cover successions. Isotope geochronology was carried out on the Faddey terrane granites using U-Pb (zircon, conventional) and whole-rock Sm-Nd methods with a Finnigan MAT 261 mass spectrometer in static mode at the Institute of Precambrian Geology and Geochronology, St Petersburg. U-Th-Pb zircon analyses of Faddey granites were also performed using Stockholm's Cameca IMS 1270 ion-microprobe and the Pb-evaporation technique was applied to single zircons. The Faddey (and related Zhdanov) granites yielded crystallization ages of c. 850 Ma (Vernikovsky et al. 1998; Pease & Vernikovsky 2000). According to depleted mantle Nd model age calculations, crust as old as Late Palaeoproterozoic (1.8-1.9 Ga; Vernikovsky et al. 1999&) may have been involved in the genesis of the Zhdanov granites.

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Taimyr's Neoproterozoic ophiolites occur in two separate NE-trending belts, the Chelyuskin and Stanovoy belts (Fig. 5). The recognition of these ultrabasic and basic rocks as ophiolite associations on the northern Taimyr Peninsula was established by Uflyand et al (1991) and described by Vernikovsky (1992, 1996) and Khain et al. (1997). The geochemical compositions of these ophiolites are similar to those from many other ophiolite complexes (Vernikovsky et al. 1996). U-Pb zircon (conventional) ages of c. 740 Ma and Rb-Sr (isochron) and depleted mantle Nd model ages of 850-785 Ma for Kunar plagiogranite in the Chelyuskin region, provided evidence of a Neoproterozoic age. Isotope data from the Stanovoy ophiolite belt has, to date, been unable to constrain its age.

New geochronological results for the Chelyuskin and Stanovoy ophiolite belts During international expeditions to Taimyr in recent years the Chelyuskin and Stanovoy ophiolite belts have been studied in detail (Gee 1998; Gee & Pease 1999; Pease 2002; Vernikovsky et al. 2003&). We present new ages for the Chelyuskin (Kunar River plagiogranite) and Stanovoya (Zimovochnaya gabbro) complexes. U-Th-Pb ion microprobe analyses were carried out using Cameca IMS 1270 at the Swedish Museum of Natural History's NORDSIM facility, Stockholm, and on SHRIMP-RG at the Stanford University, USGS ion microprobe laboratory in Menlo Park, Cali-

fornia. Methods for sample preparation, analytical procedures and data reduction at the NORDSIM facility follow those described in Whitehouse et al. (1997, 1999). Analytical methods and data reduction at the USGS/SU joint facility follow those described in Earth et al. (2001), with the exception that concentrations and isotopic ratios were calibrated against the 91500 standard at both facilities. Standard analytical errors were propagated to unknowns (Pb-U) and the 2°7Pb/206Pb error is derived from counting statistics. The common Pb estimate was based on the measured 204Pb and Stacey and Kramers' (1975) model isotopic compositions. Spot selection and spot documentation were carried out using cathodoluminescence (CL) and SEM images before and after analysis. The ion-beam spot size was typically about 25 (xm. Analytical results are presented in Table 2 and plotted on 207Pb/206Pb v. 238U-206Pb concordia diagrams (Fig. 6). The data are uncorrected for common lead. The Kunar plagiogranite (VP98-060) was collected from near Mod Bay (Fig. 5). It is strongly foliated and concordant to the surrounding metavolcanic rocks, which are present as xenoliths within the plagiogranite. In thin-section, evidence of dynamic deformation (sigmoidal porphyroclasts) is overprinted by static recrystallization of retrograde greenschist facies metamorphism. Twelve zircon grains were analysed from the plagiogranite. These grains were euhedral, had oscillatory-zoned interiors with bright rims in CL, and contained many inclusions. Inclusions were avoided during analyses with the small ion-probe beam size. Only cores were analysed. All analyses were concordant, generating a well-defined concordia age (Ludwig 1998) of

Table 2. U-Th-Pb ion-microprobe analytical data Sample U (grain no.) (ppm)

Th (ppm)

Pb (ppm)

Th/U

/206 %

207pb/206pb

± lcr%

206

Pb/238U ± lor%

Age estimates (Ma) Pb/206pb

207

VP98-060 (Kunar plagiogranite) 77°31.39 N. Lat., 104°31.48 E. Long. 1 55.4 (0.43) 0.729 8.6 37.7 (0.35) 0.593 9.7 2 62.4 43.1 (0.06) 0.448 7.3 23.2 3 47.9 (0.61) 0.676 7.5 29.1 4 47.9 0.374 (0.59) 16.5 6.8 5 45.7 (0.14) 7.3 25.2 0.519 6 47.8 7.7 0.683 (0.46) 30.6 7 49.2 (0.77) 0.369 5.0 8 34.0 13.0 (0.82) 6.3 0.558 20.6 9 41.5 0.666 (0.44) 7.2 10 46.7 32.6 0.564 7.4 (0.56) 11 47.4 30.3 0.796 (0.16) 10.3 51.8 12 65.1

0.06269 0.06674 0.06498 0.06230 0.06469 0.06279 0.06255 0.06467 0.06205 0.06416 0.06488 0.06287

2.42 1.89 2.39 1.86 1.75 2.18 2.55 2.36 2.01 2.49 2.31 1.98

0.12203 0.12285 0.12723 0.12489 0.12578 0.12664 0.12524 0.12445 0.12419 0.12197 0.12597 0.12296

1.70 1.37 1.29 2.30 2.30 1.32 2.49 1.33 2.30 1.28 1.36 1.40

697.8 829.8 773.7 684.4 764.3 701.2 693.0 763.7 675.8 746.8 770.3 703.9

50.7 38.9 49.4 39.2 36.4 45.7 53.3 49.0 42.4 51.7 47.8 41.6

VP98-093 (Zimovochnaya gabbro) 76°53.36 N. Lat, 107°19.42 E. Long. (0.04) 0.11848 0.368 1 954 350 351 0.787 (0.02) 0.06580 18 2 132 104 1.418 62 (0.37) 0.06613 3 403 572 0.923 (0.17) 0.06616 27 4 197 182 0.09466 0.346 26 (0.09) 5 105 36.1 (0.12) 0.06155 0.801 30 6 222 178 1.436 0.06273 112 (0.03) 7 723 1039 (0.02) 0.06398 0.831 25 8 182 151 1.697 (0.02) 0.05911 47 9 294 499 0.884 0.06237 (0.02) 27 181 10 205 (0.02) 0.05933 0.976 38 267 11 273 1.051 (0.02) 0.06298 17 121 12 115 0.05994 1.307 (0.02) 87 781 13 598 (0.02) 1.461 0.06208 38 14 246 359 1.926 (0.02) 0.06149 86 15 507 977

1.46 4.33 3.93 2.69 4.56 3.43 1.43 2.92 3.77 5.92 2.49 4.16 1.65 2.74 2.11

0.34959 0.11810 0.11809 0.11819 0.24034 0.12133 0.12027 0.11944 0.11831 0.11763 0.12135 0.12403 0.11823 0.12014 0.12056

1.93 1.83 1.39 1.86 3.21 1.40 1.20 1.68 2.54 1.33 1.69 1.81 1.27 1.61 1.35

1933.4 799.9 810.4 811.4 1521.2 658.6 699.0 740.9 571.2 687.0 579.3 707.7 601.5 677.0 656.5

26.4 93.4 84.5 57.2 88.6 75.2 30.8 63.2 84.2 131.7 55.1 90.9 36.2 59.6 46.0

206

Pb/238U

742.2 746.9 772.0 758.6 763.7 768.7 760.7 756.1 754.7 741.9 764.9 747.6

11.9 9.7 9.4 16.5 16.6 9.6 17.9 9.5 16.4 9.0 9.8 9.9

1932.6 32.2 719.6 12.5 719.5 9.5 720.1 12.7 1388.5 40.2 738.2 9.8 732.1 8.3 727.4 11.6 720.9 17.4 716.9 9.0 738.4 11.8 753.7 12.9 720.4 8.7 731.4 11.1 733.8 9.4

Notes: Analyses of VP98-060 were performed on the high-mass resolution, high-sensitivity Cameca IMS 1270 ion-microprobe at the NORDSIM facility in Stockholm, Sweden. Analyses of VP98-093 were performed on the high-mass resolution, high-sensitivity SHRIMP RG USGS Stanford University ion-microprobe facility at Stanford University. Data are reported at la. Errors in age estimates are quoted at la. All ages are calculated using the decay constants of Steiger and Jaeger (1977). Th/U ratios calculated from measured Th and U concentrations. /206(%X fraction of common Pb calculated from 204Pb; parentheses indicate no common Pb correction made.

NEOPROTEROZOIC OROGENY OF SIBERIA

241

The Chelyuskin and Stanovoy ophiolite belts apparently formed at about the same time, between 755 and 730 Ma, with an oceanic crust model age of 785 to 850 Ma (Table 3). These two belts may represent a single ophiolite/island-arc terrane fragmented by later thrusting and strike-slip faulting during collision with Siberia. Previous geochronological studies in the Taimyr Peninsula show that 630-620 Ma was the beginning of riftrelated magmatism (the Klyaz'ma River rhyolite, Pease & Vernikovsky 2000). Garnet-amphibolite metamorphism, connected with the obduction of the ophiolites and the island arcs onto the Siberian craton margin (Vernikovsky et al 1997) occurred thereafter at c. 626-575 Ma.

Neoproterozoic complexes along the eastern margin of the Siberian Craton

Fig. 6. U-Th-Pb results from plagiogranite and gabbro samples of NE Taimyr, (a) Kunar plagiogranite (VP98-060) concordia age (dark grey ellipse) with associated errors, (b) Zimovochnaya gabbro (VP98-093). Regression intercept derived from Stacey & Kramers (1975) model lead isotope composition anchored to 0 Ma (modern day composition).

755.1 ± 7 Ma (95% confidence limit). This age is interpreted to represent the crystallization age of the rock. The Zimovochnaya gabbro (VP98-093) within the Stanovoy belt was collected from the northern region of Faddey Bay (Fig. 5). It occurs in metavolcanic rocks as NNE-trending sheets 30-40 m thick, now stretched into discrete lenses. Fifteen grains were analysed from the gabbro. These grains were generally euhedral. They had oscillatory zoning from core to rim in CL, indicating a magmatic genesis. Two Meso- to Palaeoproterozoic grains are probably xenocrysts and were excluded from the final age determination. The results from the remaining thirteen analyses are well-clustered, but slightly discordant both above and below concordia, suggesting the need for lead correction. Consequently, these thirteen analyses were regressed using the model lead composition of Stacey and Kramers (1975) anchored to 0 Ma (modern day lead composition). The well-defined results provide an intercept age of 729.8 + 7.1 Ma (2a). This age is interpreted to be the crystallization age of the gabbro. These new geochronological results are important for understanding the formation of the Central Taimyr accretionary belt.

The Cretaceous, NW-trending, Verkhoyansk fold-and-thrust belt is located along the eastern margin of the Siberian Craton (Fig. 1). This orogen is composed of Carboniferous and Permian deposits, whereas Triassic and Jurassic rocks are exposed further to the east. In the southern part of this orogen, thrust sheets with Neoproterozoic and Early and Middle Palaeozoic deposits occur, made up of thick, carbonate shelf lithologies interpreted as part of Siberia's passive margin (Parfenov 1984; Parfenov & Prokopiev 1986; Parfenov et al 1995; Khudoley et al 2001; Khudoley & Guriev 2003; Pisarevsky & Natapov 2003). Within this southern part of the Verkhoyansk fold-and-thrust belt, the Palaeozoic and older rocks occur in three north-trending zones—Kyllakh, Sette-Daban and Allakh-Yun. They are composed of rocks of different age and characterized by different deformation styles (Fig. 7). The western Kyllakh zone consists of Meso-Neoproterozoic (including Vendian) and Cambrian rocks occurring in west-vergent listric thrusts, which are steep near the Earth's surface and become gently east-dipping detachments at depth (Parfenov & Prokopiev 1986). The central SetteDaban zone is composed of intensely deformed Vendian and Lower and Middle Palaeozoic formations occurring in a fan-like fold and thrust structure (Parfenov et al 1995). The eastern Allakh-Yun zone is dominated by a synclinorium containing Upper Palaeozoic and Triassic siliciclastic units of the same type as those outcropping further north within the NW-trending Verkhoyansk belt. These zones provide evidence for the eastward migration of the depocentres of the sedimentary basin during the Neoproterozoic to Palaeozoic. The Precambrian succession contains carbonate, sandstone, and shale units, which from oldest to youngest comprise the Uchur, Aimchan, Kerpyl, Lakhanda, Uy and Yudoma groups (Semikhatov & Serebryakov 1983; Khudoley et al 2001; Petrov & Semikhatov 2001). Most of these groups are bounded by regional unconformities. Their stratigraphic subdivision is based mainly on stromatolite and microphytolite assemblages and the transgressive clastic-carbonate cycles. Recent U-Pb baddeleyite analyses from mafic sills that cut the Lakhanda Group are dated at 974 ± 18 and 1004 ± 5 Ma (Rainbird etal 1998). A Sm-Nd isochron of the same sills yielded an age of 948 ± 1 8 Ma (Pavlov et al 1992). A Pb/Pb isochron on carbonates from the lower Lakhanda Group yielded an age of 1025 ± 40 Ma (Semikhatov et al 2000). Thus the Lakhanda Group and underlying strata are Mesoproterozoic in age, an interpretation supported by limited isotopic-geochemical data on the carbonates of the Lakhanda Group (Hartley et al 2001). The overlying Neoproterozoic Uy Group is distinguished from underlying units by its terrigenous composition and abrupt facies and thickness variations (the Uy Group increases in thickness eastwards from 400 m to 4500 m). The middle part of the Uy Group is the only unit in the Precambrian succession that contains relatively deep-water sediments and turbidites (Khudoley et al 2001). The upper part of the Uy

V. A. VERNIKOVSKY ET AL

242

Table 3. Isotopic-geochronological data for rocks of the central-Taimyr accretionary belt Age (Ma)

Interpretation

References

Mamont-Shrenk terrane, Mamont Granite zircon Th-U-Pb (SIMS) wr Sm-NdTDM muscovite Ar/Ar

880-940 1500-2100 816 ± 2.4

Age of crystallization Age of crust Age of metamorphism

Pease et al 2001 Vernikovskaya et al 2002Z? Vernikovskaya et al 2002&

Faddey terrane, Zhdanov granite zircon U-Pb (conventional) wr Sm-NdTDM

846 ± 11 1796-1902

Age of crystallization Age of crust

Vernikovsky et al 1998 Vernikovsky et al 1998

Dated material

Method

Faddey terrane, Faddey granite zircon Th-U-Pb (SIMS)

840

Age of crystallization

Pease & Vernikovsky 2000

Cheliuskin ophiolites belt, Kunar Plagiogranite zircon U-Pb (conventional) zircon Th-U-Pb (SIMS) wr Sm-NdTDM

740 ± 38 755 ±7 785-850

Age of crystallization Age of crystallization Age of crust

Vernikovsky et al I994b This paper Vernikovsky et al I994b

Cheliuskin ophiolites belt, Klyaz'ma metarhyolite zircon Th-U-Pb (SIMS)

627 + 7

Age of crystallization

Pease & Vernikovsky 2000

Stanovoy ophiolite belt, Zimovochnaya gabbro zircon Th-U-Pb (SIMS)

730 + 7

Age of crystallization

This paper

Stanovoy ophiolite belt, Garnet amphibolite gar, bt, pi, amf, wr Sm-Nd, isochron gar, bt, pi, amf, wr Rb-Sr, isochron bt Ar/Ar

573 ± 78 606 + 44 624 + 16

Age of metamorphysm Age of metamorphysm Age of metamorphysm

Vernikovsky et al 1997 Vernikovsky et al 1997 Vernikovsky et al 1997

Group is unconformably overlain by the Yudoma (Vendian) Group and this angular discontinuity cuts down progressively towards the west across older Meso-Neoproterozoic units (Semikhatov & Serebryakov 1983; Khudoley et al. 2001).

Summary of Neoproterozoic relationships around the Siberian Craton The Neoproterozoic complexes contained within the Palaeozoic and Neoproterozoic fold-and-thrust belts framing the Siberian Craton possess both remarkable similarities and significant differences. During the early Neoproterozoic (and probably also the late Mesoproterozoic), passive continental margins may have existed around most of the Siberian continent. This conclusion can be made confidently for the western (Turukhansk and Yenisey belts) and eastern (Verkhoyansk belt) margins. Structural, lithological, biostratigraphic and isotopic studies from the western region allow a reliable reconstruction of continental shelf and slope environments, with Neoproterozoic and late Mesoproterozoic age sediments and correlation of two key successions based on C and Sr isotopic data (Petrov & Semikhatov 1997, 2001; Hartley et al. 2001; Khudoley et al. 2001). In the Verkhoyansk belt, the overlying Vendian-Palaeozoic successions are also interpreted as passive continental margin deposits (Parfenov & Prokopiev 1986; Zonenshain et al 1990; Khudoley et al 2001); they are comparable with successions of similar age along the northern and western sides of the craton. Continental shelf depositional environments also existed along the southern margin of the Siberian continent in Neoproterozoic and perhaps late Mesoproterozic time. Major transgressive cycles of sedimentation, with coarse-grained sediments at the bottom and carbonate rocks in the upper part of each cycle have been recorded in the Baikal-Vitim-Patom region (Khomentovsky et al 1972; Shenfil 1991). The evidence for the age of these rocks mainly comes from stromatolite and microphytolite assemblages and several Pb/Pb dates of limestones (c. 860 Ma; Fefelov et al 2000). The age and time of accretion of palaeo-island-arcs suggest that the southern passive margin of Siberia may have developed as

an active margin in the early to middle Neoproterozoic (900-800 Ma; Gusev & Khain 1995; Rytsk et al. 2001). In the Yenisey Ridge area, the passive margin gave way to an active margin at c. 750 Ma, during the collision between the Central Angara terrane and the Siberian continent; this was followed by the development of island arcs and their accretion in the late Neoproterozoic (Vernikovsky et al 19990, 20010, 2003). It is not known whether an early Neoproterozoic (perhaps Mesoproterozoic) passive continental margin existed along the northern margin (Taimyr) of Siberia. In the southern domain of Taimyr only Palaeozoic passive margin rocks of Siberia, similar to those along the northeastern margin of Siberia, are exposed. These passive margin rocks are overthrust by allochthons of the Central Taimyr Neoproterozoic accretionary belt, that was formed between c. 750-650 Ma and inferred to have been obducted onto the Siberian continental margin prior to late Vendian sediment deposition at c. 600 Ma (Vernikovsky et al 1994Z?, 1997; Pease & Vernikovsky 2000; Vernikovsky & Vernikovskaya 2001). Thereafter, the northern margin of the Siberian continent developed as a passive margin up to Late Carboniferous time, before the collision of the Kara terrane with the Siberian margin (Vernikovsky et al 1995). The Neoproterozoic record in the Verkhoyansk region indicates that the eastern margin of the Siberian Craton remained a passive continental margin during the Neoproterozoic and Palaeozoic. Nevertheless, an oceanic domain may have existed further east (present day coordinates) of Siberia during the Neoproterozoic. The Neoproterozoic and Mesoproterozoic outboard terranes, both (micro)continental and oceanic, now occurring as allochthons emplaced onto the Siberian margin, are exotic and the former may have affinities with continents other than Siberia. Baltica- Siberia relationships As is clear from the descriptions of the fold-and-thrust belts of Neoproterozoic and Early Palaeozoic age around the Siberian Craton, Neoproterozoic Orogeny in most areas culminated prior to the late Neoproterozoic (early Vendian). Neoproterozoic Orogeny is characterized by a major unconformity in, or at the base of, the Vendian. These strata, together with the overlying

NEOPROTEROZOIC OROGENY OF SIBERIA

243

Fig. 7. Tectonic scheme of the south Verkhoyansk fold-and-thrust belt (A) and structural cross section A-A' (see Fig. 1 for location); (B) from Parfenov et al 1995.

Palaeozoic successions, were in some cases relatively little deformed, and in other cases were involved in Middle to Late Palaeozoic thrusting. It can be concluded that the Siberian Craton was surrounded, at least on the northern, western and southern sides, by oceanic domains in the Neoproterozoic and that their accretion to Siberia occurred shortly before the beginning of the Palaeozoic. The Early Palaeozoic faunal record is unambiguous (Khomentovsky et al 1972; Kaban'kov & Sobolevskaya 1981; Bogolepova et al. 2001): Strata, unconformably overlying middle to late Neoproterozoic terranes are all characterized by Siberian faunal assemblages. The isotopic evidence for the timing of island-arc development and the age of ophiolites indicates that, although most are clearly Neoproterozoic, their evolution was not synchronous. The magmatic and deformational histories vary around the margin; there appears to be significant diachronicity, which may partly reflect the paucity of the present isotope-age record. In the same way that Siberia was isolated from other continents in the Neoproterozoic and Early Palaeozoic, so Baltica apparently developed independently and was an isolated continent at least during the Early Palaeozoic. Along the margins of Baltica there is a variety of evidence of magmatic activity, some in oceanic domains, that was of similar age to the oceanic complexes

around Siberia. Thus the Engane-Pe ophiolites of the Timanide margin in the Polar Urals (Dushin 1997; Scarrow et al. 2001; Khain et al. 2003) are similar in age (c. 660 Ma) to those along the southwestern margin of Siberia (Fig. 8). Faunal evidence suggests oceanic separations between Siberia and Baltica by the late Cambrian, although the distance between the continents is much disputed (Cocks & Torsvik 2002; Gubanov 2002). Baltica's northeastern (Timanian) margin experienced orogeny in the late Neoproterozoic (Vendian), culminating with deposition of foreland basin molasse in the late Vendian. Thus, orogeny in parts of Siberia and Baltica may have been synchronous and certainly overlapped in time. The question naturally arises as to whether Siberia and Baltica were closer in the Neoproterozoic, perhaps connected via an active margin (e.g. the late Neoproterozoic Kipchak Arc of Sengor et al 1993; Sengor & Natalin 1996), the continents separating only after Vendian Orogeny to allow independent faunal developments in the Cambrian. A plausible alternative is that Siberia and Baltica were independent plates (cratons surrounded by oceanic domains) throughout the Neoproterozoic, the apparent synchronicity of orogeny along their margins controlled by changes in global stress regimes. The evidence is inconclusive. The outboard terranes around Siberia, occurring within Neoproterozoic complexes either as microcontinents or fragments of

244

V. A. VERNIKOVSKY ETAL.

Fig. 8. Geological sketch map of the western margin of the Siberian Craton and the easternmost margin of Baltica.

larger Proterozoic continents (such as the Mamont-Shrenk and Faddey complexes in Taimyr), may be related to other late Grenville-age orogens elsewhere in the Arctic (e.g. Svalbard, Greenland, Wrangle Island, etc.). More evidence, however, is necessary to define the relevance of these temporal correlations. Continued analysis of terranes marginal to Baltica and Siberia is needed to constrain their Neoproterozoic and Mesoproterozoic evolution further. This work was supported by grants 00-05-65398, 01-05-64732, 00-15-98562 from the Russian Foundation of Basic Research, by grants IR-97-1139, NEMLOR-765 from INTAS, by the EUROPROBE project (TIMPEBAR), and by the University of Western Australia through the Tectonic Special Research Centre and a Gledden Senior Visiting Fellowship for V. Vernikovsky. Financial support for fieldwork was received from the Swedish Polar Research Secretariat. Svetlana Bogdanova is thanked for helpful comments on the manuscript. This paper is a contribution to IGCP 440: Assembly and Breakup of Rodinia. We thank the SU-USGS ion microprobe laboratory, especially Joe Wooden, and the NORSIM facility staff for their assistance. The NORSIM facility is financed by the research councils of Denmark, Norway, Sweden, the Geological Survey

of Finland and the Swedish Museum of Natural History. This is NORDSIN contribution no. 94.

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Index Note: page numbers in italic, e.g. 164, refer to figures. Page numbers in bold, e.g. 72, signify entries in tables. acritarchs, Late Cambrian age constraints on Kolguev Island 159 Agma 39 Ai Formation 19-20 Aldan, River 234 Aldan shield 234 Amderma 2, 20 Anabar shield 234 Andreeland 792 Andreeland-Dicksonland Graben 792, 195-196 Andreeneset 208 Angara, River 234 Angara Block 234 Aral Sea 146 Arkhangelsk 2, 39, 146 palaeogeographic/palaeotectonic sketch map 49 Arsha Formation 23 Atomfjella Antiform 194-195,194, 209 Austfonna 209 Avzyan Formation 22 Bad'yashorskaya Formation 148 Baikal, Lake 234 Baikalides 1, 233 Baikibash Formation 24 Bakal Formation 20-21 Bakeevo Formation 23 Balto-Timanian triple junction (BTTJ) 185 Bangenhuken Complex 794 Barents Sea 2, 6, 59 Timanian and Caledonian trends 53-54 regional palaeotectonic setting in Late Permian 53 Barentsian Caledonides 799 Barents0ya 208 Barmin, Cape 8, 60, 70 Barminskaya Group 8-9, 9 Bashkirian Anticlinorium 19, 151, 752 Lower Riphean 19-21 Middle Riphean 21-22 Upper Riphean 22-23 Lower Vendian 23 Upper Vendian 23 stratigraphic correlation chart 27 Basinskaya Formation 151 Batsfjord Formation 777 Belaya 20 Belaya, River 30 Bellsund 792 Beloretsk 752 Beloretsk Terrane 38 Berezov 146 Berlevag 52 Billefjorden 792 Billefjorden Fault Zone 792, 194 Biri transgression 181 Biscayerhalvoya-Holtedahlfonna Horst 792, 196 Bj0rn0ya 197 Bolshoy Kameshek massif 70, 71 gabbro and granite isotope dating 71-73, 72 Bolshoy Rumyanichny, Cape 8, 60, 70, 71 Bolvan Creek Formation 29 Bolvansky Formation 7, 8

Bothnia, Gulf of 7 70 Botniahalvoya 208, 209 Breibogen-Bockfjorden Fault 792, 208 Brennevinsfjorden 793 Brennevinsfjorden Group 192 Bugrino 160 Burzyan 30 Burzyan Group 19-21 Buton Formation 26 Bystrino Group 29 Bystrynskaya Group 9-12, 72 Caledonian Suture 59 Caledonides 2 comparison with Laurentian margin 197-200, 198 relationship with Timanides 201 western Baltoscandian basins 178,178 type I basins 178-181, 779 type II basins 181-182 type III basins 182-183 Celsiusberget Group 193 Central Timan Fault (CTF) 6, 48, 48, 60, 170 palaeogeographic/palaeotectonic sketch map 49 Chaichiy Island 62 Chelyabinsk 2 Chelyuskin ophiolite belt 240-241 Chernaya, River 70 Chernyi Kamen Formation 26 Cheshskaya Bay 6, 8, 60, 70 Chetlas Group 28 Chidvia Formation 38 Churochnaya Formation 27 Damflya 193, 208 Demino Formation 26 Dimtemyol, River 77 Dorogor Formation 28 Dronning Louise Land 798 Duvefjorden 792, 793, 208 Dvina, River 2 Dvoretsk suite 26 Dzela complex 107, 121 geochemistry 112 blueschist- and greenschist-facies rocks 113-115, 773 mafic rocks 113, 773 ultramafic rocks 112-113, 773 geological setting 107-109, 108 magmatic rocks mafic rocks 112 ultramafic rocks 109-112 whole rock geochemical analyses 110-111 metamorphic rocks amphibolite facies 112 blueschist facies 112 greenschist facies 112 petrogenesis 118-119 blueschists and greenschists 119 tectonic setting 120-121 timing 119-120 U-Pb dating 115-118, 775, 776, 777, 118 microprobe data 114

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250

East European Craton 19, 20 Precambrian geology 75-76, 76 Riphean-Vendian tectonic and sedimentary events 31-32 lateral and vertical architecture map 31 Vendian succession 38 East Timan Fault (ETF) 6, 48,51, 60, 170 Edgeoya 792, 208 Ekaterinburg 2, 20, 30, 31, 146 Eletskaya Facies 150, 153 Engane-Pe Anticline 27, 88-89, 148, 148, 149 Enganepeiskaya Group 148 Engerdalen Basin 181 Eolussletta Shear Zone 194 Ernasalya, Cape 62 Eskolabreen Complex 194 EUROPROBE 1 Fedotovo Formation 25 Flatan-Instrumentberget Complex 194 Forlandsundet 792 Foyn0yane 208 Franklinsundet 793 Franklinsundet Group 793 Franz Josef Land 47, 59 Garevka Formation 26 Glitnefonna 793 Gnilsk Formation 29 Grubeinskaya Formation 149, 150 Hamberg Gletcher Foreland 798 Hedmark770 Hedmark Basin 178-181,180 Heidal770 Heidal Group 181 Helvetesflya Formation 192 Hinlopenstreten 208 Hinlopenstreten Supergroup 192, 793 Hinlopenstretet 792, 194 Hinlopenstretet Syncline 209 Hornsund 792 H0yvik 770 Igarka 234, 238 Il'yavozh Formation 27 Innvika 208 Innvikh0gda Syncline 209 Inostrantseva Bay 141 Inta 37, 760 Inzer 30, 37 Inzer Formation 22-23 Isfjorden 792 Ishimbai 30 Isispynten 793, 208 Ivdel 146 Izhma 37 Jameson Land 798 Kalak Nappe Complex 183 Kaltasa Formation 24 Kama 20 Kama, River 37 Kamen Rassolny Formation 26 Kamennye Sopki massif 70 Kamsk-Belsk depression 24 Kanin Kamen Ridge 6 Kanin Kamen Supergroup 7, 8, 15-16 Kanin Nos Cape 6

INDEX

Kanin Peninsula 2, 6-8, 6, 8, 48 Baltica margins 175-178, 775, 777 compared with northern Timan and Vymskaya Ridge 12-13 Early Palaeozoic unconformity 145-147, 147 Neoproterozoic high-grade metamorphism 59, 59, 67-68 conditions of metamorphism 65-67 geochronology, previous work 62-63 lithologies 63 metamorphism, previous work 61-62 regional geology 59-61, 60 sample descriptions 63 structure 63 timing of metamorphism 67 palaeogeographic/palaeotectonic sketch map 49 thermobarometry analytical data and interpretation 65 biotite compositions 64 garnet compositions 64 methods 63-64 plagioclase compositions 65 results 67 Kapitansky, Cape 8 Kapp Hansteen Group 192 Kara Sea 2, 47, 146, 234 Karatau Group 22-23 Karlin Formation 25 Karpinsky Peninsula 154 Katav Formation 22 Kazan' 2 Kernos Formation 26 Kharbei Complex 148 Khatalambinsky Suite 150 Khoidyshor Formation 27 Khoreyver terrane palaeogeographic/palaeotectonic sketch map 49 Kikvozhkaya Formation 11, 72 Kildin Island 52, 169-171, 770 Kirkenes 52 Klyktan Formation 25 Kocheshor Formation 27 Koiva Formation 26 Kola Peninsula 6, 47, 170 Baltica margins 174-175 palaeogeographic/palaeotectonic sketch map 49 Kolguev Island 2, 48 geological background 159-160, 160 stratigraphic levels 161 Late Cambrian age constraints by acritarchs 159 biostratigraphic implications 167 Cambrian rocks in Pechora Basin and adjacent areas 160-162 Cambrian-Ordovician successions 162-163 characterization of acritarchs 163-166, 164, 165 previous fossil records 163 study materials and methods 163 palaeogeographic/palaeotectonic sketch map 49 shelly fossils 166-161,166 Kolpakovskaya Formation 151 Kong Karls Land 208 Kongsfjord 52 Kongsfjord Formation 777 Kongsfjord Turbidite System (KTS) 13, 14, 15 Kotlas 37 Kotlin stratohorizon 44 Kozinskaya Formation 151 Kraemerpynten 208 Krasnaya Bay 62 Krasnokamsk Formation 25 Krayny Kameshek massif 70, 71 syenite isotope dating 71-73, 72

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INDEX

Krivaya Luka Formation 23 Kronprins Christian Land 198 Kurgashlya Formation 23 Kvaenangsfjorden 170 Kvarkush metamorphic basement 125, 132 ages basaltic dyke 131 blueschist-facies metamorphism 129-131, 130 post-Timanian processes 131 regional implications 131-132 tectonic setting of blueschist-facies metamorphism 131 analytical procedures 127-128 geoloigical setting of Kvarkush Anticline 125-127, 726, 727 results 128-129,129 sampling and sample characteristics 127 Kvarkush-Kamennogorsk Anticlinorium 25, 151 Upper Riphean 25-26 Lower Vendian 26 Upper Vendian 26 Kvitoya 208 Kvitvola Nappe Complex 181 Kyiv 38 Kykva Formation 25 Kykvozh Formation 29 Kyrmin Formation 25 Kyrpin Group 24 Kyzgey Formation 27 Lady Franklinfjorden 193 Lagoya 793 Laksefjord Nappe Complex 182 Laksefjorden 770 Lambert Land 198 Lamtsa 39 Lamtsa-Verkovka depositional sequence 41-43 Laponiahalvoya 793, 208, 209 Laptev Sea 59, 234 Laptopay Formation 27 Las'va Group 25 Laurentian margin 191, 203 comparison with East Greenland Caledonides 197-200, 798 comparison with North Greenland fold belt and Pearya 199, 200-201, 202 Svalbard 191 assembly of terranes 201 Bj0rn0ya 197 eastern terranes 191-195 interrelationships between the terranes 197 northern terranes 195-196 relationship between Caledonides and Timanides 201 southwestern terranes 196-197 Leaonidovo Formation 24 Lemvinskaya Facies 150 Lena, River 234 Lomfjorden Fault 792, 794 Lomfjorden Supergroup 192 Lomonosov Ridge 203 Lomonosova 59 Lomonosovskaya Series 141 Loven Syncline 209 Loz'va, River 31 Ludovatovskaya Formation 12 Ludovatye, Cape 60 Lunvozh Formation 29 Lunvozhskaya Formation 11, 72 Magnitogorsk 2, 146 Main Uralian Fault (MUF) 87, 89 Maka Bay 747

Makovskaya Series 141 Malochernoretskaya Formation 8-9, 9 Maly Kameshek massif 70 Man'inskaya Formation 148, 149 Manitanyrdskaya Group 148, 150 Man'ya Formation 27 Marun-Keu Complex 87, 101-102 analytical procedures 90 elemental chemistry 91 ion microprobe zircon analyses 90-91 single zircon Pb evaporation 90, 91 Sm-Nd and Rb-Sr analyses 91 analytical results comparative geochronology 99-100 ion microprobe U-Th-Pb 92- 93 major and trace element data 95, 700 Rb-Sr whole rock data 94, 99 Sm-Nd data and Nd model ages 94, 98-99 zircon geochronology 91-97, 98, 99 zircon images 96, 97 zircon stability in eclogite-facies metamorphism 97 geological setting 87-90, 88, 89 lithologies and sampling 90 regional correlations and geodynamic interpretation 100 Baltica passive margin rifting 101 Early Timanian evolution 100 Uralian convergence and arc—continent collision 101 sample characterization 102 Mashak Formation 21 -22 Matochkin Strait 754 Matusalya, Cape 62 Mezen 31 Mezen Basin 37, 38, 43, 44, 47 Mezen Formation 28 Mikulkin Antiform 61, 62 thermobarometry 66 Mikulkin, Cape 6, 62 Mikulkino Group 29 Mikulkinskaya Group 6, 7, 61, 66 Minsk 38 Minyar Formation 23 Mitushikha Bay 740 Mityushev Kamen' Massif 153 MoelvTillite 180-181 Mongol-Okhotsk belt 234 Morozovskaya Formation 153 Moscow 38 Moscow Basin 38, 43-44, 44 Motalafjellet 792 Murchisonfjorden Supergroup 192, 793 Murmansk 2, 52, 746 Nadezhdino Formation 24 Nafta Formation 28 Nar'yan-Mar 2, 37 Nenoxa Formation 38 Neoproterozoic northeastern and northwestern margins 169, 186 Baltoscandian margin western basins 178-183, 778 basinal, stratigraphic and palaeoclimatic features 183 -184 plate-tectonic aspects of developing margins 185-186, 785 Timanian margin coastal Kola Peninsula and White Sea 174-175 Kanin Peninsula 175-178, 775, 777 northern Norway and NW Russia 169-171, 770 westward extension 172-174, 773 Neoproterozoic passive-margin turbidites 5 central Timan and the Vymskaya Ridge 9-12

252

comparison between Kanin Peninsula, northern Timan and Vymskaya Ridge 12-13 comparison between Rybachi Peninsula and Varanger Peninsula 13-16 definition of submarine turbidite systems (STS) 5 Kanin Peninsula 6-8, 8 northern Timan 8-9, 8,10 regional setting 5 Neoproterozoic timescale 3 Nerpichiy Island 62 Nikol'skaya Series 135, 137, 142 Nizhnii Novgorod 2 Nizva Formation 26-27 Nordaustlandet Terrane 191, 192-194, 792 geology 193 structure 207-210, 208, 209 Nordmarka 208, 209 North Sosva, palaeogeographic/palaeotectonic sketch map 49 Northern Timan 170 compared with Kanin Peninsula and Vymskaya Ridge 12-13 Northern Timan Igneous Suite (NTIS) 69, 73-74 geological setting 69, 70 isotope dating 71-73, 72 study samples 69-71, 70 Novaya Zemlya 2, 47, 135, 142 Central Domain 139-141, 139 chemical composition of intrusive rocks 138 Early Palaeozoic unconformity 153,154 geological map 136 Northern Domain 141, 141 palaeogeographic/palaeotectonic sketch map 49 Southern Domain 135-139 Novobobrovsk Formation 28 Novozemelskaya terrane, palaeogeographic/palaeotectonic sketch map 49 Obdyr Group 28 Obeizskaya Formation 150 Olkhovo Formation 24 Omen Formation 28 Onega Graben 37, 38 Onega Peninsula 39 Orenburg 2, 30, 31, 146 Orsk 30, 31 Orvin Land 208, 209 Oslyanka Formation 25 Padun Formation 28 Pai-Khoi 48, 153 Pai-Khoi foldbelt 88 Parnuksky Suite 150 Parry0ya 208 Paun Formation 29 Paunskaya Formation 9-10, 12 Pav'yug Formation 29 Pearya 799, 200-201, 202 Pechora 20, 48,160 Pechora Basin 6, 47, 59 Early Palaeozoic unconformity 147 geochemistry 78 alteration 80 analytical methods 79-80 calc-alkaline affinity 81 classification of granitoid rocks 80-82, 80 isotopes 80, 82, 82 large ion lithophile abundances 81 major element oxides 78, 80-82, 81 trace and rare earth elements 79, 80-82, 81, 82, 82 Late Neoproterozoic granitoid magmatism 75, 84

INDEX

Palaeotectonic setting 83 tectonic discrimination diagram 83 Precambrian geology of Timan-Pechora region 75 Bolshezemel zone 76, 77 Devonian cross-section 76 Izhma zone 76,11 Pechora zone 76,11 pericratonic region 75-76, 76 Timan Range 76 sampling, isotopic ages and petrography 77, 77 Bolshezemel zone 78 Izhma zone 77-78 Pechora zone 78 sources, mantle and crust 82-83 Pechora, River 2, 31 Pechora-Kozhva Fault (PKF) 6, 48, 51 palaeogeographic/palaeotectonic sketch map 49 Perevalok Formation 26 Perm' 2, 20 Pez Formation 28 Piritovy Peninsula 737 Planetfjella Group 794 Platenhalv0ya 793, 208, 209 Pogureiskaya Formation 149 Pokju, River 77 Pokjuskaya Formation 9-10, 70, 72 Pok'yus Formation 29 Polhem Formation 794 Poludov Range 26 Upper Riphean 26-27 Lower Vendian 27 Upper Vendian 27 Pomorsk Strait 760 Porsangerfjorden 770 Prijutovo Formation 24 Prikamsk Formation 24 Prins Karis Foreland 792 Rakhovsky Suite 138-139 Raudfjorden Fault 792, 208 Raudfjorden Graben 792 Redkino stratohorizon 44 Richarddalen 792 Rijpdalen Anticline 209 Rijpfjorden 792, 793, 208 Rittervatnet Formation 194 Roaldtoppen Group 793 Rochug Formation 28-29 Rumyanichnaya Formation 8-9, 9, 70, 50 Rumyanichnaya, River 70, 71 gabbro isotope dating 71-73, 72 Rusanova Peninsula 737 Rusanovskaya Series 135, 137, 142 Rusanovsky Suite 137-138, 138-139 Rybachi Peninsula 6, 15, 52, 170 compared with Varanger Peninsula 13-16 palaeogeographic/palaeotectonic sketch map 49 Tsypnavolokskaya Formation 50 Rybachi Turbidite System (RTS) 13, 15 Rybachi-Sredni Peninsula 169-171, 774 Sablya Gora Formation 27 Safonovo Group 28, 37 Saledskaya Formation 150 Salekhard 88 Salikhovo Formation 25 Samara 2 Sandivey Formation 29 Sarv 770

INDEX

Satka Formation 20 Schegrovit Formation 25 Schugorskaya Formation 148 Scoresby Sund 198 Sed'iol'skaya Formation 147 Seduaykha Formation 29 Seiland Igneous Province (SIP) 183 Serafimovo Group 24 Seve Nappe Complex 183 Severnaya Zemlya 59 Severozemel 'sky Anticlinorium 141, 142 Shikhan Formation 24 Shokal'sky Glacier 141 Siberian margins 233 Baikal-Vitim fold-and-thrust belt 233-234, 234, 235 relationships with Baltica 242-244 Siberian Craton 234 Neoproterozoic complexes 241-242, 244 Taimyr 238-240, 239 geochronological data 240-241, 242 U-Th-Pb ion-microprobe data 240, 241 Turukhansk-Igarka region 238 Yenisey Ridge 234-238, 236 isotopic-geochronological data 237 Sim, River 31 Sinegorsk Formation 25 Sjuoyane 208 Smalfjord Formation 772 Smutsbreen Formation 194 Sokol'nikskaya Formation 153 Sokol'nino Formation 29 Solza 39 Sopka Bolvanskaya massif 70 S0rbreen Formation 194 Sorkapp 792 S0r0y-Seiland Nappe 183 sparagmite basin 178 Spitsbergen 792 Sredni Peninsula 52, 170 Sredni-Rybachi Fault Zone (SRFZ) 6, 13, 48, 52 Stanovoy 234 Stanovoy ophiolite belt 240-241 Staropetrovo Formation 24-25 Starye Pechi Formation 26 Stonsteinhalvoya 793 Storoya 208 Storsteinhalvoya 209 subduction westward beneath Baltica 75, 84 geochemistry 78 alteration 80 analytical methods 79-80 calc-alkaline affinity 81 classification of Pechora Basin granitoid rocks 80-82, 80 isotopes 80, 82, 82 large ion lithophile abundances 81 major element oxides 78, 80-82, 81 trace and rare earth elements 79, 80-82, 81, 82, 82 Palaeotectonic setting 83 tectonic discrimination diagram 83 Precambrian geology of Timan-Pechora region 75 Bolshezemel zone 76, 77 Izhma zone 76, 77 Pechora zone 76, 77 pericratonic region 75-76, 76 Timan Range 76 sampling, isotopic ages and petrography 77, 77 Bolshezemel zone 78 Izhma zone 77-78 Pechora zone 78

253

sources, mantle and crust 82-83 Subinebukta 209 submarine turbidite systems (STS) 5, 13 Suirovo Formation 23 Sulmenev Bay 739 Sulmenevskaya Series 139, 142 Suran Formation 21 Suzma 39 Svalbard 47, 59, 191,203 assembly of terranes 201 Bj0rn0ya 197 Caledonian terranes 792 comparison with East Greenland Caledonides 197-200, 798 comparison with North Greenland fold belt and Pearya 799, 200-201, 202 eastern terranes 191 -192 Nordaustlandet Terrane 192-194, 792, 793 West Ny Freisland Terrane 194-195, 194 geochemistry 212 composition of rocks 213- 214 major elements 275 tectonic discrimination diagrams 277 trace elements 216 geology, rock types and study samples 277 aplitic dykes 210-211 gabbroic rocks 211-212 metasupracrustals, paragneisses, amphibolites and migmatites 210 migmatite neosomes, red and grey granites 210 porphyritic granites and augen gneisses 210 syenite 211 interrelationships between the terranes 197 northern terranes 195 -196 relationship between Caledonides and Timanides 201 Sm-Nd geochemistry 228-230, 229, 230 southwestern terranes 196-197 tectono-magmatic activity in northeast 207, 230-231 structure of Nordaustlandet Terrane 207-210, 208, 209 U-Pb and Pb/Pb dating analytical methods 212-214 aplite dykes 221-223 augen gneisses 214-216 biotite schist 216 grey granites 217-221 migmatites 216-217 monazite and titanite, conventional data 218, 279 syenite 223-227 zircon, cathodoluminescence images 222- 223 zircon, ion microprobe data 224- 226, 227- 228 zircon, monazite and titanite, Pb-evaporation data 219, 220- 227

Svartrabbane Formation 192 Svetlino Formation 28 Syktyvkar 2, 48, 146 palaeogeographic/palaeotectonic sketch map 49 Syum-Keu ophiolite complex 89-90 Tabuevo Group 29 Tabuyevskaya Group 6, 7, 7, 60 Taimyr 59, 238-240, 239 geochronological data 240-241, 242 U-Th-Pb ion-microprobe data 240, 247 Tanin Formation 26 Tarkhanovo Group 29 Tarkhanovskaya Group 6, 7, 7, 60 Tchelyabinsk 30, 37 Tel'posskaya Formation 150 Tereklinskaya Formation 152 Timan foreland, Late Neoproterozoic sedimentation 37, 43-45

254

INDEX

depositional sequences 41-43 Pre-Vendian setting 37-38 sedimentology of the Vendian succession 38 alternating shale/siltstone 38-39, 40 channelized sandstone 40, 41, 42 interstratified sandstone and shale 39-41, 40 laminated shale 38, 40 stratigraphy 39 trough-cross-bedded sandstone 41, 42 Timan Range 20, 47 Precambrian geology 76, 76 Timan, central 9-12 Timan, northern 8-9, 8 Timan-Varanger Belt (TVB) 5, 6 definition of submarine turbidite systems (STS) 5 regional setting 5 Timanide Orogen 1 Baltica margins coastal Kola Peninsula and White Sea 174-175 Kanin Peninsula 175-178, 775, 777 northern Norway and NW Russia 169-171,170 westward extension 172-174, 773 Dzela complex 121 geological map 2 Kvarkush metamorphic basement 132 orogenesis age of deformation 51 -52 extent and character 49-51, 50 palaeotectonic setting 54 structural and tectonic development 47 palaeogeographic/palaeotectonic sketch map 49 pre-orogenic basinal evolution 47-48, 48 Timanides, Early Palaeozoic unconformity 145, 153-155 description 145,146 North of Polar Urals and Pechora Basin 152-153 Novaya Zemlya 153,154 Pai-Koi 153 Vaigach Island 153 Pechora Basin 147 Timan Range and Kanin Peninsula 145-147, 747 Urals mountains 147 Middle Urals 151 Northern Urals 150,757 Polar Urals 147-149,748 Southern Urals 151-152, 752 Sub-Polar Urals 149-150, 750 Timanides, relationship with Caledonides 201 Timanides, Riphean and Vendian sedimentary sequences 19, 20, 32 stratigraphic correlation of sedimentary sequences 29-31 lateral and vertical architecture maps 30 Timan-Pechora region 27-29 Izhma-Pechora depression 29 Kanin-Pechora zone 29 Khoreyver depression 29 Mezen-Vychagodsk zone 28 Obdyr-Chetlas zone 28 stratigraphic correlation charts 28 Tsil'ma-Ropchino zone 28-29 Vymsk-Volsk zone 29 Tolparovo Formation 23 Tooriksalya, Cape 62 Torozhma 39 Tossasfjallet 770 Tossasfjallet Basin 181 Troitsk Massif 26 Trollfjorden-Komagelva Fault Zone (TKFZ) 6, 13, 48, 52, 53, 169, 770 Tsypnavolokskaya Formation 50 Tuchkino Formation 38

Tukaevo Formation 24 Turukhansk 234, 238 Tury Cape 7 70 Ufa 2, 30, 31, 152 Uftuga Formation 37 Uk Formation 23 Ukhta 2, 48, 88, 160 palaeogeographic/palaeotectonic sketch map 49 Ural, River 2, 30, 31 Ural Basin 38 Uralian foredeep molasse 88 Uralides, Riphean and Vendian sedimentary sequences 19, 20, 32 Middle Urals 25 Upper Riphean 25-26 Lower Vendian 26 Upper Vendian 26 Poludov Range 26 Upper Riphean 26-27 Lower Vendian 27 Upper Vendian 27 southern Urals 19 Lower Riphean 19-21 Middle Riphean 21-22 Upper Riphean 22-23 Lower Vendian 23 Upper Vendian 23 stratigraphic correlation chart 27 stratigraphic correlation of sedimentary sequences 29-31 lateral and vertical architecture maps 30 Subarctic and Polar Urals 27 Lower Vendian 27 Late Vendian 27 stratigraphic correlation chart 27 Volga-Urals region 23 Lower Riphean 24 Middle Riphean 24 Upper Riphean 24 Upper Vendian 24-25 Urals 6, 20, 147 Middle Urals 151 Northern Urals 125, 132, 150, 757 ages for Kvarkush Anticline 129-132, 730 analytical procedures for Kvarkush Anticline 127-128 Kvarkush Anticline 125-127, 726, 727 results for Kvarkush Anticline 128-129, 129 sampling of Kvarkush Anticline 127 Polar Urals 87, 101-102, 107, 121, 147-149, 148 analytical procedures for Marun-Keu complex 90-91 analytical results for Marun-Keu complex 91-100 dating of Dzela complex 115-118 geochemistry of Dzela complex 112-115 geological setting of Dzela complex 107-109, 708 geological setting of Marun-Keu complex 87-90, 88, 89 lithologies and sampling of Marun-Keu complex 90 magmatic rocks of Dzela complex 109-112 metamorphic rocks of Dzela complex 112 petrogenesis in Dzela complex 118-119 regional correlations and geodynamic interpretation 100-101 sample characterization 102 tectonic setting of Dzela complex 120-121 timing in Dzela complex 119-120 Southern Urals 151-152,752 Sub-Polar Urals 149-150, 750 Usa, River 148 Usa Formation 24 Ust-Nafta Group 28, 37 Ust-Pinega Formation 28 Us'va Formation 25

255

INDEX

Vads0 52, 52 Vaigach Island 2, 153 Valdres 170 Valdres Group 179 Varanger Peninsula 6, 14, 169-171, 170, 171, 172 compared with Rybachi Peninsula 13-16 palaeogeographic/palaeotectonic sketch map 49 Timanian-Caledonian interrelationships 52-53, 52 Vard0 52, 52 Varysalya, Cape 62 Vassfaret Formation 194 Veinesbotn Formation 172 Velikaya, River 70 Velva Formation 25 Verchne-Kamsk depression 25 Vereschagino Formation 25 Verkhovka Formation 41-43 Vestfonna Anticline 209 Viluy, River 234 Vilva Formation 26 Vishera, River 31 Vizinga Formation 28 Voigatch, palaeogeographic/palaeotectonic sketch map 49 Volga, River 2 Volgograd 2 Vorkuta 2, 31, 88, 146, 148 palaeogeographic/palaeotectonic sketch map 49 Vozey Formation 29 Vymsk Group 29 Vymskaya Group 9-12, 12 Vymskaya Ridge 6, 9-12, 170, 178 compared with northern Timan and Kanin Peninsula 12-13 geological map 11

Wahlenbergfjorden 793, 208 Wandel Sea 798 Wedel Jarisberg Land 792 WestNy Freisland Terrane 191-192, 792, 194-195 geology 794 West Siberian Basin 234 West Timan Fault (WTF) 6, 48, 170 White Sea 6, 39, 170 Baltica margins 174-175 Winter Mountains 39 Yakorninsky Suite 139 Yambozerskaya Formation 8-9, 9 Yanei Formation 29 Yaneyskaya Formation 7, 8 Yenisey Ridge 234-238, 234, 236 isotopic-geochronological data 237 Yenisey-Khatanga Basin 234 Yorga depositional sequence 43 Yurmatau Group 21-22 Yusha Formation 21 Yuzhnonovozemel'sky Anticlinorium 135, 737 Zigala Formation 22 Ziganskaya Formation 151 Zigazino-Komarovo Formation 22 Zilmerdak Formation 22 Zimnegory depositional sequence 43 Zimnegory Graben 38 Zolotitsa 39