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

RIPHEAN AND VENDIAN: TIMANIDES AND URALIDES

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 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

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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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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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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.

37

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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

39

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

40

D. GRAZHDANKIN

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

42

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

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D. GRAZHDANKIN

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 (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 39 462 100 |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.

129

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, ) 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

194

D. G. GEE & A. M. TEBEN'KOV

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

SVALBARD: LAURENTIAN CALEDONIAN MARGIN

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

SVALBARD: LAURENTIAN CALEDONIAN MARGIN

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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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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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

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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