The Neoproterozoic of Baltica—stratigraphy, palaeobiology and general geological evolution

Premmbrinn Reseurth ELSEVIER

Precambrian Research 73 (1995) 197-216

The Neoproterozoic of Baltica stratigraphy, palaeobiology and general geological evolution Gonzalo Vidal, Matgorzata Moczydtowska Institute of Earth Sciences, Uppsala University, Norbyvi~gen 22, S-752 36 Uppsala, Sweden

Received 10 May 1993; revised version accepted 15 December 1993

Abstract

Western Baltica contains numerous Neoproterozoic rock successions recording deposition in a variety of settings, including basinal, deep and shallow-shelf passive margin and intracratonic continental, fluvial and shallow-marine environments. Neoproterozoic pre-Varangerian episodic emplacement of dolerite swarms ( 1020-870 Ma; 720 + 260--665 + 10 Ma) and post-Varangerian (late Vendian or Ediacaran ~ 551 + 4.0 Ma) basaltic volcanism in western Baltica are related to episodes of rifting and paroxysmal stages of faulting that were followed by erosion and basin infilling. Tentative biostratigraphic correlation of successions in the western Baltoscandian basins and in the Volhyn Aulacogen in the East European Platform (EEP) suggests that initial rifting, formation of embryonic fault basins in western Baltica and infilling of rift basins within the core of Baltica might be largely contemporaneous events restricted to the late Riphean and bracketed in time to around 800-750 Ma. The biostratigraphic correlation and radiometric age control of pre-Varangerian rift deposition in the Lake V~ittern Graben in South Sweden with allochthonous rock succession in the Hedmark Basin in southern Norway (Vidal and Nystuen, 1990a) provide additional evidence to confirm the time frame of basin formation as a response to crustal fracturing of western Baltoscandia ~ 800-700 Ma ago. Sedimentation following late Riphean and Vendian rifting in Baltica was dominated by quartz arenites of considerable thickness on stable shallow shelves, thus indicating stabilization at the onset of the Early Cambrian transgressions. The earliest Cambrian successions on the western border of Baltica witness of moderate basin subsidence and crustal instability extending well into early Holmia times.

1. I n t r o d u c t i o n

Significant palaeobiological, physical and chemical environmental changes recognized in Neoproterozoic successions are under evaluation with the purpose to define a Neoproterozoic System. The upper boundary of this system coincides with the newly defined base of the Cambrian System (and the Phanerozoic Eon). An important part of this pursuit is to identify continuous N e o p r o t e r o z o i c - C a m b r i a n sedimentary successions for further study aimed at selecting a potential global stratotype section of the Neoproterozoic System. 0301-9268/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSDI030t-9268(94)00078-6

Satisfactory stratigraphic successions must fulfil a number of basic requirements. Sections meeting with required criteria of stratigraphic completeness, suitable facies associations and structural simplicity are not common. The aim of this article is to review the status of possible candidate successions in the realms of Baltica. This includes the best known sedimentary successions in the Fennoscandian Shield (Lake Viittern Graben) and down-faulted successions in the Muhos Basin on the Baltic coast of western Finland (Fig. 1). The marginal areas of the Caledonian fold belt in southern (Hedmark Basin) and northern Norway (Varanger

198

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Peninsula) are part of this review, whereas autochthonous successions along the eastern border of the Scandinavian Caledonides (e.g. Laisvall and Tornetrask in Sweden; Fig. 1), although commented upon, are left outside the scope of this study. Much attention is devoted to the apparently continuous Neoproterozoic to Lower Palaeozoic succession in the Varangerfjorden and Tanafjorden areas in East Finnmark (Varanger

Peninsula; Fig. 1), whereas only peripheral attention is paid to the important and very thick Neoproterozoic (late Riphean-Vendian?) rock succession in the northeastern half of the Varanger Peninsula (Fig. 1). Also part of this account is the sub-surface Neoproterozoic succession in the East European Platform (EEP; Lublin Slope) in southern Poland (Fig. l).

G. Vidal, M. Moczydtowska /Precambrian Research 73 (1995) 197-216

Subsurface Neoproterozoic successions in aulacogens and troughs in the eastern and southern parts of Baltica are outside the scope of this paper. These include the Volhynian Aulacogen (also known as Orsha Depression) in Russia, Belarus, eastern Lithuania, and the Ukraine, the Lake Ladoga and White Sea Depressions (the latter also known as Zimniy and Letniy Coast Depression), the Timan Aulacogen and coastal areas of northern Kola Peninsula (Kildin Island and Sredniy and Ribachiy Peninsula in northwestern Russia; Fig. 1). The Neoproterozoic-early Palaeozoic successions in the selected areas are fossiliferous, little or moderately affected by tectonic processes and virtually unmetamorphosed (see for example Vidal, 1976, 1981, 1985; Fcyn, 1985; Kumpulainen and Nystuen, 1985; Moczydiowska, 1991). The areas in question are largely covered by recent geological mapping scale 1:50,000 and the Neoproterozoic successions have been the subject of detailed structural, sedimentological and palaeontological investigations (see reviews by Foyn, 1985; Kumpulainen and Nystuen, 1985; Siedlecka, 1985; Vidal, 1985; Moczydtowska, 1991; Vidal and Nystuen, 1990a).

2. The Fennoscandian Shield

Neoproterozoic successions within the mainland of the Fennoscandian Shield are few and limited to isolated rift troughs within the Palaeoproterozoic basement of Baltica, e.g. the Lake V~ittern Basin in South Sweden and the Muhos-Hailuoto Basin on the northern Baltic coast of Finland (Fig. 1 ). The Lake V~ittern Basin occupies the Lake V/ittern depression, a graben structure trending NNE-SSW and fault-bounded basement areas south-southeast and northwest of the lake (Vidal, 1974, 1976). The Lake V~ittern Basin contains the Visings6 Group (Fig. 2), a succession exceeding 1000 m in thickness that consists of three informal lithostratigraphic units. The lower formation consists of about 145 m of well-sorted fluvial and littoral quartz sandstones (Vidal, 1974). The disconformably overlying middle formation is a detrital wedge attaining a minimum thickness of 315 m. It consists in its lower part of reddish and greenish-grey arkoses, conglomeratic sandstones and boulder-bearing debris-flow deposits locally attaining a thickness of

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240 m that were deposited as alluvial fans (Vidal and Bylund, 1981 ). Succeeding fluvial conglomerates and deltaic sandstones interdigitate with fossiliferous marine siltstones and shales. In its upper part, the middle formation consists of fine-grained pyritic arkoses deposited in a shallow-marine environment that grade into a succession of mudstone, phosphatic shale and dolomitic limestones (locally stromatolitic; Vidal, 1972), belonging to the upper formation. This unit represents an oscillating shallow-intertidal to supratidal depositional environment (Vidal, 1972, 1976). The maximum age range of the Visings8 Group is bracketed by K/Ar dates on detrital micas yielding ages of 985-1060 Ma (Magnusson, 1960). Furthermore, numerous Rb/Sr dates of dolerites intruding the surrounding Palaeoproterozoic basement rocks in southeastern Sweden yielded ages indicating emplacement during 870-1020 Ma, although some ages above 1000 Ma were believed to be anomalously high (Patchet and Bylund, 1977). The dolerite swarms might be a response to tensional stress in the Palaeoproterozoic (Svecofennian) crust of southwestem Baltica caused by marginal up-warping related to post-orogenic uplift of the adjacent Sveconorwegian region (Patchet and Bylund, 1977). The dolerites also yield pole positions that coincide with palaeolatitudes of 15° and 40 ° obtained for sandstones and siltstones of the middle Visings6 Group interpreted to represent acquired remanent magnetism (Vidal and Bylund, 1981 ). The minimum age limit of the Visings6 Group was obtained from Rb/Sr dates of shales and clay fraction from the upper formation suggesting ages of 663-703 Ma (Bonhomme and Welin, 1983). Isotopic age-bracketing is in agreement with a~late Riphean (or Kudashian; Vidal and Siedlecka, 1983) pre-Varanger age indicated by acritarchs from the middle and upper formations (Vidal, 1976, 1985). The Visings6 Group represents deposition within the western Baltoscandian basins formed in pre-Varangerian times through rifting that preceded a paroxysmal stage of faulting which was followed by erosion and basin infilling (Solyom et al., 1983; Kumpulainen and Nystuen, 1985; Vidal, 1985). Tentative biostratigraphic correlation of the Visings8 Group with successions within aulacogens (e.g. the Volhyn Aulacogen) in the EEP suggests that initial rifting, formation of embryonic fault basins in western Baltica and infilling of rift basins within the core of Baltica (Vidal, 1979b)

G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

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might be largely contemporaneous events restricted to late Riphean times (Vidal, 1979b; Kumpulainen and Nystuen, 1985), being bracketed to around 800-750 Ma. As suggested below, the biostratigraphic correlation and radiometric age control of the central Scandinavian pre-Varangerian rift succession (Visings~ Group) with the allochthonous rock succession in the Hedmark Basin in southern Norway (Vidal and Nystuen, 1990a) provide an additional basis for the formerly suggested time frame of basin formation as a response to fracturing of the Palaeoproterozoic crust of western Baltoscandia in pre-Varangerian times, mainly 800-700 Ma ago (Kumpulainen and Nystuen, 1985). The Muhos-Hailuoto basin (central-western Finland; Fig. 1) contains detrital sediments attributed to

201

the Muhos Formation attaining an approximate recorded thickness of 900 m. On microfossil evidence the succession was considered as possibly middle Riphean and late Vendian in age (Tynni and Donner, 1980). However, age control on the largely unconsolidated rocks of the Muhos Formation is rather poor. This is because a certain taxonomic assignment is generally not possible. Despite this, the microfossils from the Muhos Formation in the Bothnian Bay of Finland form assemblages largely resembling late Vendian cyanobacterial and acritarch associations (Tynni and Siivola, 1966; Tynni, 1978; Tynni and Donner, 1980; Tynni and Uutela, 1984) that were formerly recorded from late Vendian successions in various parts of the EEP (Vidal, 1985). The age of the Muhos Formation cannot be confidently established, neither on radiometric nor on palaeontological criteria. Notwithstanding this, it appears significant that the present distribution of the Muhos Formation on the Fennoscandian Shield is in WNW-ESE-trending fault-bounded troughs within Archaean basement rocks (Tynni and Uutela, 1984) that seem to parallel the recognized trend of trough basins within eastern Baltica containing late Vendian deposits (e.g. the White Sea Depression; Fig. 1).

3. Neoproterozoic basins within the Scandinavian Caledonides In southern Norway Neoproterozoic sedimentary rocks occur as allochthonous sequences incorporated in stacked nappes and thrust sheets in the eastern thrust belt of the Scandinavian Caledonides (Fig. 1). The sedimentary successions are dominated by coarsegrained feldspathic sandstones which since the days of Esmark (1829) are known under the informal name "sparagmites", thus yielding the name "Sparagmite Region". The best preserved portion of this important Neoproterozoic succession consists of strata within the lowermost nappe unit of the Osen-R0a Nappe Complex (Kumpulainen and Nystuen, 1985). The nappe complex consists of thrust sheets involving Proterozoic basement rocks (Gorbatschev, 1985). The Neoproterozoic-Lower Cambrian Hedmark Group attains an approximate thickness of at least 30004000 m and is overlain by ca. 2000 m of epicontinental CambrianSilurian strata (Bockelie and Nystuen, 1985). The

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G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

Hedmark Group (defined by Bjcrlykke et al., 1967) is widespread in the Lake Mjosa area (Fig. 1). Locally, tectonic deformation of rocks of the Hedmark Group includes imbricated fans, duplex structures, broad open folds and small-scale folds and thrusts (Nystuen, 1983; Morley, 1986). Distortion near thrust zones is the only deformation feature in the southern part of the Lake Mjcsa-Rena area (Fig. 1 ). However, in the northern part of the Sparagmite Region, penetrative fracture cleavage and foliation have modified sedimentary structures and textures. In general, metamorphism is very low (Morad, 1988). This feature is also reflected by the colour of kerogen and microfossils that suggest moderate to considerable thermal alteration corresponding to burial temperatures of around 175 to ~ 200°C (or higher) and stages of late catagenesis to early metagenesis (Vidal and Nystuen, 1990a). Six to eight formations form the Hedmark Group (Fig. 2) and other units were distinguished in the eastern and northern parts of the Sparagmite Region, including tholeiitic basalts (Sa~ther and Nystuen, 1981; Nystuen and Siedlecka, 1988). The Brottum Formation of the Hedmark Group comprises at least 3000-4000 m of coarse- to fine-grained feldspathic sandstones with intercalated organic-rich mudstones and shales and was suggested to have formed in a system of prograding fan deltas (Bjorlykke et al., 1976; S~ether and Nystuen, 1981; Nystuen, 1982, 1987). Brottum sandstones in the MjCsa-Gudbrandsdalen area (western part of the outcrop area) were deposited from turbidity currents (Englund, 1966, 1972, 1973b; Bj~rlykke et ai., 1976; Nystuen, 1982, 1987). Shales of the BrCttum Formation interpreted as hemipelagic low-density turbidites contain simple microfossils comparable to associations described from turbidite successions elsewhere (Vidal and Nystuen, 1990a). The succeeding Biskop~sen Formation (Fig. 1), also termed B iskopSsen Conglomerate (BjCrlykke et al., 1967), forms several wedge-shaped conglomerate bodies interfingering with turbidite sandstones of the BrCttum Formation and shales of the Biri Formation (Bjcrlykke, 1966; Englund, 1966, 1972, 1973b; Lcberg, 1970; Bjorlykke et al., 1976). Generally attaining a thickness of 200--400 m, lateral equivalents (Imsdalen Conglomerate) are at least 700 m thick (S~ether and Nystuen, 1981). Conglomerate beds and darkgrey, massive medium- to coarse-grained sandstone (that yielded microfossils; Vidal and Nystuen, 1990a)

make the bulk of the formation. Locally, the Biskop~sen Formation overlies the Biri Formation with erosional unconformity. Clasts derived from the underlying member of the Biri Formation are common in the basal portion of the conglomerate, including infrequent fossiliferous phosphorite pebbles (Spjeldn~es, 1963, 1967; Manum, 1967). At an early stage, wedge-shaped bodies of the BiskopSsen Formation were interpreted as formed in fluvio-deltaic settings (Skjeseth, 1963; Bjorlykke, 1966; Spjeldn~es, 1967; Bj~rlykke et al., 1967), whereas Skjeseth (1963) suggested that the conglomeratic unit was deposited during a fault-induced phase in the subsidence of the Hedmark Basin. Bjcrlykke (1966), Spjeldn~es (1967) and Bj~rlykke et al. (1976) favoured an interpretation that involves regional regression, also accounting for the presence of clasts of limestone, phosphorite and calcareous sandstone that were most likely derived from eroded portions of the Biri Formation and also from the lower member of this unit underlying the Biskop~en Formation (Bjorlykke et al., 1967, 1976). This part of the Biri Formation formed during an early stage of carbonate deposition. The Biskophsen Conglomerate wedges were also interpreted as subaqueous fans formed in front of deltas by gravity flow processes (R. Otter in Nystuen, 1982). The marine origin of the conglomerate is demonstrated by the presence of microfossils in dark shales (Vidal and Nystuen, 1990a) and the subaqueous origin is indicated by the interfingering between the Biskop~en Formation and the fossiliferous shales and sandstones of the Biri and BrCttum Formations (Vidal and Nystuen, 1990a). Hence, transport and deposition from highdensity turbidites and subaqueous mass flows were proposed by Vidal and Nystuen (1990a), who rejected subaerial fluvial deltaic counterparts of the subaqueous facies association. The boulder-bearing basal part of the formation is a subaqueous channel-fill deposit that may have been cut down during a regressive event involving erosion of calcareous deposits of the lower member of the Biri Formation that formed in shallowmarine platforms along the basin margin (Vidal and Nystuen, 1990a). High numbers of extrabasinal clasts and the architectural style of the Biskop~sen Conglomerate wedges imply strong tectonic control, whereas the recorded association of basaltic volcanism indicates that the "Biskop~sen depositional event" reflects a rifting climax in the Hedmark Basin that on micropa-

G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

laeontological evidence seems contemporaneous with rift sedimentation in the Lake V~ittern Graben (Figs. 1, 2; Nystuen, 1987). Notwithstanding this, Vidal and Nystuen (1990a) indicated that an increased run-off in the inferred drainage area could explain the flux of gravel-sized debris to the basin, whereas hemipelagic deposition of organic-rich mud took place between surges of gravity-flows carrying coarse-clastic debris. The Biri Formation (Fig. 2) is a complex unit attaining 50-200 m in thickness and containing a variety of interfingering carbonate- and detrital lithofacies. The carbonate-dominated facies represent deposition in shallow and basinal settings (Tucker, 1983). Arenaceous rocks and black shales are basin-slope and basinfloor deposits (BjOdykke et al., 1976; Nystuen, 1982), but calcareous sandstones formed in high-energy coastal environments and intraformational limestone breccias may have originated along margins of a carbonate platform (Nystuen, 1982). A lower member of the Biri Formation consists of shales, siltstones and calcareous rocks beneath the Biskop~sen Formation (Fig. 2; Bjcrlykke et al., 1967, 1976), but is missing in parts of the Hedmark Basin where the Biskop~sen Formation rests on and interfingers with the turbidite sandstones of the Brcttum Formation. Basinal organicrich shales and calcareous mudstones ( B iri Formation) succeed the Br¢ttum sandstones (or laterally equivalent formations) wherever the Biskop~sen Formation is absent (Nystuen, 1982, 1987). The Biri Formation was deposited during a regional transgression possibly resulting from eustatic sea-level rise (Spjeldn~es, 1967; BjCrlykke et al., 1976). The transgressive phase is observable in the northern and eastern Hedmark Basin, where fluvial deposits (Rendalen Formation) are overlain by the Biri Formation (Nystuen, 1982, 1987). The organic-rich black shales in the Biri Formation (Bjcr~nes Shale Member) have TOC values as high as 3% in the central and northern Hedmark Basin and were probably deposited in a sub-basin formed by a sudden tectonic subsidence of the basin floor (S~ether and Nystuen, 1981; Nystuen, 1987). The Ring Formation consists of several coarse-clastic wedges along the basin margin in the western part of the Hedmark Basin. Shales, thin-bedded turbidite sandstones, coarse-grained, conglomeratic arkoses and massive to graded quartz-pebble conglomerates of channel-fill origin make the bulk of the Ring Formation. The conglomerates carry clasts of dark-grey

203

mudstones and the contact with the shale-dominated underlying Bid Formation is locally sharp and erosional (L¢berg, 1970; Englund, 1972, 1973b), whereas at the type locality at Lake MjOsa the boundary is transitional from organic-rich shales and mudstones to the lowermost conglomerate channel-fills of the Ring Formation. The Ring Formation was interpreted as fan delta deposits in which subaerial and subaqueous facies are present (Bjcrlykke et al., 1976). Later, Nystuen (1987) suggested that the Ring Formation (like the Biskop~isen Formation) represents mainly subaqueous deposition by gravity flow on fronts of prograding fan deltas, a view supported by the occurrence of microfossils in mudstones in the lowermost Ring Formation (Vidal and Nystuen, 1990a). The Moelv Tillite (Fig. 2) marks the Varangerian glaciation (dated by Rb-Sr whole rock on shale to ~ 653 _ 7 Ma; Sturt et al., 1975) in the Hedmark Basin and in autochthonous areas of the eastern margin of the Scandinavian Caledonian fold belt (Fig. 1). Deposits of the Moelv Formation have a wide distribution in the Hedmark Basin. Underlain by a regional erosional unconformity, the Moelv Tillite rests on the Ring, Biri or Rendalen Formations or on the autochthonous crystalline basement (Nystuen, 1976; Bjcrlykke et al., 1976; Bjcrlykke and Nystuen, 1981; Bockelie and Nystuen, 1985). It comprises subglacially deposited massive diamictite and glaciomarine laminated mudstone with dropstones (Bjcrlykke et al., 1976; Nystuen, 1976) and the glacial sequence accumulated during the advance and retreat of a continental ice sheet across most of the Hedmark Basin and other rift- and shelfbasins at the margin of the Baltoscandian craton (Nystuen, 1985). The ice-sheet was probably buoyant in the deepest part of the basin and the ice might have eroded underlying unconsolidated sediments without imposing any high burden, as indicated by reworked and wellpreserved age-diagnostic acritarchs (Vidal and Nystuen, 1990a). The Ekre Formation ( or Ekre Shale; Fig. 2) succeeds conformably the glaciomarine upper facies of the Moelv Tillite. The transitional boundary is marked by an upward decrease in ice-dropped clasts (Nystuen, 1976, 1982). The Ekre Formation is dominated by greyish-green or red, laminated siltstones. In the upper part it is interbedded with deltaic sandstone beds of the overlying Vangs~s Formation (Bjcrlykke et al., 1976; Nystuen, 1982; Dreyer, 1988). A marine origin for the

204

G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

lower part is suggested by the conformity with the underlying Moelv Tillite (Nystuen, 1982, 1985), although the general absence of acritarchs may perhaps suggest unsuitable environmental conditions at the time of deposition and/or post-burial conditions unsuitable to the preservation of acritarchs (Vidal and Nystuen, 1990a). Geochemical studies on the Ekre Formation indicate deposition in oxygen-rich conditions in a wellaerated water body, either marine or lacustrine (Englund, 1973b). In summary, complex stratigraphic relationships and considerable lateral facies variations in the Biri and adjacent formations of the Hedmark Group suggest that deposition was controlled both by sea-level fluctuations and tectonic control contemporaneous to sedimentation (Vidal and Nystuen, 1990a). Black, organic-rich shales in the Biri Formation were interpreted as having accumulated during transgressive events which led to enhanced carbon burial through increased organic productivity and subsequent development of oxygendepleted bottom waters (Tucker, 1983). Early sediments of the Biri Formation (lower member) may have formed in a shelf slope where cold oxygendepleted and nutrient-rich waters carried by ascensional currents may have developed plankton blooms (Vidal and Nystuen, 1990a). Deposition in a deep basin in which organic debris has been accumulated and concentrated in hemipelagic mud over long periods was inferred from sedimentological and microfossil evidence (Vidal and Nystuen, 1990a). Hence, the conditions shaping deposition and organic productivity in the Hedmark Basin were considered propitious to enhanced carbon burial and accumulation of sedimentary phosphates. Microfossils from the BrCttum, Biri and Biskop~sen Formations consist largely of biostratigraphically longranging taxa. However, two age-diagnostic taxa abundant in the Biskop~sen and Biri Formations ( Octoedryxium truncatum and Vandalosphaeridium varangeri) seem important, since previous occurrences are in Neoproterozoic rocks preceding the Varangerian glaciation (Vidal and Nystuen, 1990). As also suggested by the stratigraphic position under the Varangerian Moelv Formation, Biri and Biskop~en facies associations originated in late Riphean times and can be coherently accommodated in a simple model of turbidite sedimentation, formation of coarse-clastic subaqueous fans in a deep-marine basin with upwelling

onto shelf platforms and enhanced organic production and burial of organic carbon in hemipelagic muds (Vidal and Nystuen, 1990a). The microfossils allow correlation with rift-related deposition of the Visings~ Group (Vidal and Siedlecka, 1983). The uppermost unit of the Hedmark Group, the Vangsfis Formation, consists of the Vardal Sandstone and the Ringsaker Quartzite Members (Fig. 2). The former displays major lateral facies and thickness variations and includes coarse- to fine-grained turbidite sandstones, subaqueous debris-flow deposits and black organicrich shales overlain by deltaic and fluvial, braidedstream sandstones and by conglomerates (Bjorlykke et al., 1976; Nystuen, 1982, 1987; Dreyer, 1988). Feldspathic arenites of deltaic and braided-stream origin grade into the overlying shallow-marine quartzites of the Ringsaker Member through a zone of marine reworking (Bjcrlykke et al., 1976; Nystuen, 1982; Dreyer, 1988). The Ringsaker Member comprises beach facies, tidal channel deposits, shallow-marine storm-deposits and offshore bars (Nystuen, 1982; Dreyer, 1988) and is a transgressive unit onlapping the autochthonous Precambrian basement (Bockelie and Nystuen, 1985). It contains bioturbation structures, such as Scolithos, Monocraterion and Diplocraterion, that occur in the uppermost part of the member (Skjeseth, 1963). Black, pyrite-rich mudstones of the Vardal Sandstone Member of the Vangs~s Formation yielded specimens of microfossils considered as colonial clusters of bacteria and non-diagnostic spheromorphs attributed to Leiosphaeridia sp. The uppermost part of the Vardal Sandstone Member yielded the only agediagnostic acritarchs recovered, Fimbriaglomerella minuta, a confident indicator of an Early Cambrian age (Vidal and Nystuen, 1990a, b), suggesting that the base of the Cambrian is within the Vardal Sandstone Member. The Ringsaker Member is overlain by fossiliferous shales of the Lower Cambrian Holmia "stage" (Skjeseth, 1963; Ahlberg et al., 1986; Vidal and Nystuen, 1990b).

4. The northern rim of Baltica Varanger Peninsula; the Varangerfjorden and Tanafjorden regions The northern rim of Baltica in Varanger Peninsula (northern Norway; Fig. 1) is underlain by unmeta-

G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

morphosed Neoproterozoic rock successions. Siedlecka and Siedlecki (1967) distinguished two major geological provinces in Varanger Peninsula, each characterized by distinctive sequences of sedimentary rocks (Siedlecka, 1985; Fig. 1). In the southern part, a predominantly terrigenous succession rests with profound unconformity on basement rocks. The Vads¢ Group is disconformably overlain by detrital rocks with subordinate carbonates assigned to the Tanafjorden Group (Fig. 2). Both groups of strata represent fluvial and shallow-marine depositional environments (F~yn, 1985). West of Tanafjorden (Fig. 1) the Tanafjorden Group is included in the Caledonian Gaissa Nappe Complex. The Vads¢ and Tanafjorden Groups (Fig. 2) represent onlapping sedimentation episodes on Proterozoic basement rocks initiated with transport from the present south and west. At least two breaks exist within the Vads~ Group (R~e, 1975; Vidal, 1981 ). The Tanafjorden Group consists mainly of shallow-marine siliciclastic overlain by a complex carbonate and siliciclastic unit (Grasdalen Formation). The age of the Vads¢ Group is indicated by a whole-rock Rb/Sr isochron from shales of the Klubbnasen Formation in the lower portion of the Vads¢ Group yielding an age of 805 + 7 Ma (Sturt et al., 1975). Several formations of the Vads¢ Group yielded an acritarch association including age diagnostic taxa that appear in agreement with a late Riphean age (Kudashian; Vidal, 1981; Vidal and Siedlecka, 1983). The lower five units of the Vads¢ Group are late Riphean in age, but the top of the group, the EkkerCya Formation, yielded complex and diagnostic acritarchs indicating a terminal Riphean (Kudashian; Vidal and Siedlecka, 1983) age. They corroborate the proposed existence of a formerly postulated hiatus (R~e, 1975). Stromatolite (BertrandSarfati and Siedlecka, 1980) and acritarch data (Vidal, 1981 ) indicate that the overlying Tanafjorden Group is equivalent to Kudashian strata elsewhere in Baltoscandia and the EEP (Vidal, 1981; Vidal and Siedlecka, 1983; Vidal, 1985; F~yn, 1985) with radiometric ages around 700 Ma. Rocks of the overlying Vestertana Group rest on the Vads~ and Tanafjorden Groups with low-angle regional unconformity (Fcyn, 1985). In ascending stratigraphic order, the Vestertana Group is build up by the Smalfjorden, Nyborg, Mortensnes, Stappogiedde and Breivika Formations ( Fig. 2). All but the latter two units are attributed to the terminal Proterozoic, Var-

205

angerian, glacial event (Edwards, 1975, 1984). In a detailed sedimentological study of the Vestertana Group, Edwards (1984) concluded that the glacial formations rest on a glacially scoured regional angular unconformity. Glacial deposits include lodgement tillite with subordinate laminated mudstones, often containing dropstones and repeated successions of lodgement tillite overlain by laminated mudstone were interpreted to be submarine glacial retreat sequences (Edwards, 1975). As is the case for numerous North Atlantic Neoproterozoic sections, in East Finnmark carbonate rocks are associated with deposits of Vendian tillites. Carbonate rocks are comparatively rare in the Neoproterozoic sequences in East Finnmark, but dolomitic carbonates containing stromatolites were recorded in the middle part of the Grasdalen Formation in the upper part of the Tanafjorden Group on Varanger Peninsula and in the Porsanger Dolomite (in the Porsangerfjorden Group, a succession nearly time equivalent to the Tanafjorden Group; White, 1969; Tucker, 1977; Bertrand-Sarfati and Siedlecka, 1980). Additionally, buff-weathering dolostones occur near the middle of the Nyborg Formation (Edwards, 1984; Fig. 2). The dolostones display cryptalgal lamination. The Nyborg Formation is interpreted as a transgressive to shallowing-up sequence and its shallow-water deposits originated in a tide- and wave-dominated delta front (Banks et al., 1971; Edwards, 1975). The unit is interglacial and, as in the case of the postglacial Stappogiedde Formation at the top of the succession, it shows evidence of local isostatic uplift and eustatic rise (Banks et al., 1971 ). Age control for the Varangerian glaciation derives from a Rb/Sr whole-rock isochron for the Nyborg Formation that yielded a recalculated age of 650 ___7 Ma (Sturt et al., 1975), whereas micropalaeontological evidence from the Vestertana Group is limited to scattered finds of obviously redeposited acritarchs in the matrix of tillites of the Smalfjorden and Mortensnes Formation (Vidal, 1981). Taxonomically low-diversity associations of acritarchs from older units were found in the Smalfjorden Formation and erratics of chert (believed to derive from the top of the Tanafjorden Group; Grasdalen Formation; Fig. 1 ) in the Mortensnes Formation yield vase-shaped microfossils formerly reported from Neoproterozoic, pre-Varangerian, rocks at various locations (e.g. the Visings6 Group, Knoll and Vidal, 1980; Svalbard, Knoll and Calder,

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1983; eastern Greenland, Vidal, 1979a). Additionally, acritarchs were reported from the top of the Nyborg Formation and from the Stappogiedde Formation (Vidal, 1981; Fig. 2). The Stappogiedde Formation also yielded organic ribbon-shaped fossils (Vendotaenia sp.) which in co-occurrence with organic tubular fossils (Sabellidites cambriensis; Farmer et al., 1992) may represent the lower portion of the stratigraphic range of both taxa, thus indicating a late Vendian age (Vidal, 1981). A poorly diversified fauna of possible medusoids occurs in the middle of the Stappogiedde Formation, within the Innerelva Member (Farmer et al., 1992), conforming with the inferred late Vendian age. Probably agglutinated tubular fossils occur in the lower Innerelva Member, well below the medusoid fossils (Moczydlowska and Vidal, unpublished data). The occurrence of the ichnofossil Phycodespedum is within the lower 3 m of the Lower Breivika Formation and Platysolenites antiquissimus occurs 150 m above the ca. 250 m thick Lower Breivika Member (Foyn, 1985). These two sources of biostratigraphic evidence bracket the position of the Proterozoic-Cambrian boundary within the Lower Breivika Member and sparse acritarchs also support this interpretation (Moczydlowska and Vidal, unpublished data). The correlation of the succession in the Tanafjorden area (Fig. 1) with the thin autochthonous succession of the Dividal Group occurring along the rim of the Scandinavian Caledonides (Fig. 1) is hampered by lateral facies variation, and by the discrepant nature of the biostratigraphic elements to be compared. Hence, there are considerable problems in comparing the rock units and fossil and ichnofossil associations with the Kullingia Zone in northern Baltoscandia (Jensen and Grant, 1992). Towards the northwest part of Varanger Peninsula (Fig. 1) the sedimentary sequence is in tectonic contact with the Barents Sea and Lokvikfjellet Groups along the Trollfjorden-Komagelva Fault Zone (Siedlecka and Siedlecki, 1967; Johnson et al., 1978; Siedlecki and Levell, 1978). These groups of strata are not discussed in this review. This is because, although relatively well dated by microfossils and isotopic dating (Beckinsale et al., 1975; Rhheim, pers. commun, in Siedlecki, 1980), the age of the 9000 m thick preVarangerian Barents Sea Group is insufficiently bracketed in the absence of precise dates for the overlying Lokvikfjellet Group (Siedlecki and Levell, 1978; Vidal and Siedlecka, 1983; Vidal, 1985; Siedlecka, 1985).

However, as inferred from available age estimates (Beckinsale et al., 1975; R~heim, pers. commun, in Siedlecki, 1980), regression prior to tilting of the Barents Sea basin may have been connected with the Varangerian (Vendian) glaciation followed by a transgression marking de deposition of the shallow-marine and fluvial LCkvikfjellet Group (Siedlecka, 1985).

5. The southwestern rim of Baltica--the Lublin Slope in southeastern Poland

An unmetamorphosed sedimentary succession has been penetrated by numerous drillholes in the Lublin Slope in the western part of the EEP (Figs. 1, 2), (Moczydtowska, 1991). The age of probably continental conglomerates and reddish quartz sandstones of the Polesie Formation at the base of the succession (Fig. 2) has been inferred to be late Riphean, but decisive evidence is not available. The succeeding Slawatycze Formation at the base of the Vendian sequence consists of a terrigenous lower portion of variable lithologic composition that is overlain by basaltic lavas. In certain areas, volcanic rocks rest directly on Palaeoproterozoic crystalline basement complexes (Mansfeld et al., 1993). The detrital basal member includes conglomerates, conglomeratic sandstones and arkoses, whereas the succeeding volcanic portion consists of basalts, tufts, volcanic breccias and agglomerates. Juskowiakowa ( 1971 ) reported that the thickness of the basaltic volcanics varies from several decimetres to several tens of metres. The maximum thickness attained by the Stawatycze Formation is 372 m. Conglomerates belonging to the lower part of the Slawatycze Formation were formerly interpreted as possibly glacigenic and equivalent to putative tillites of the Vilchan Formation in Russia and the Gruska Formation in the Ukraine (Arefi, 1968, 1981; Chumakov, 1981). However, examined portions of the succession display no features that could indicate a glacial origin. Possibly long-transported well-rounded pebbles consisting of granitoids, gneisses and quartz are enclosed in an arkosic matrix and probably underwent several cycles of erosion and redeposition. In our opinion a fluvial origin can be inferred for the lower part of the Stawatycze Formation and a glacial origin can be confidently outruled. The lower Stawatycze Formation can be accommodated in a sedimentation regime of

G. Vidal, M. Moczydtowska / Precambrian Research 73 (1995) 197-216

debris flows and braided delta-plain deposition bordering the southwestern extension of the Volhynian Aulacogen (Fig. 1; Po~aryski, 1977). Recent single-grain SHRIMP U/Pb datings on zircons from volcanic tuff of the upper Stawatycze Formation yielded an age of ~ 551 + 4.0 Ma (Compston et al., 1995). A sharp contact was initially claimed to separate the Stawatycze Formation from the overlying Siemiatycze and Bialopole Formations, composed of arkoses and sandstones. The Siemiatycze or Biatopole Formation, laterally interfingering with each other (Fig. 2), wedge out in southerly direction in the Lublin Slope (Fig. 1 ) (Arefi et al., 1979; Fig. 2). However, at least locally, our observations revealed a gradual transition from reddish-coloured felsic (S. Claesson, pers. commun., 1993) tuffs into mudstones across the junction between the Slawatycze and Biatopole Formations. The contact between the Bialopole Formation and overlying Lublin Formation is also transitional (Fig. 2). The Lublin Formation consists of alternating silty and argillaceous rocks that yield large concentrations of Vendotaenia antiqua (Arefi and Lendzion, 1978; Urbanek and Rozanov, 1983) and simple acritarchs and cyanobacterial microfossils (Moczydtowska, 1991). The Lublin Formation is conformably overlain by the Wlodawa Formation, consisting of poorly sorted glauconitic sandstones with minor interbedded mudstone and argillite. The latter unit is assigned to the Sabellidites cambriensis Zone and was formerly considered as basal Cambrian (Arefi and Lendzion, 1978; Lendzion, 1983). More recently, an association of acritarchs indicative of a late Vendian age was reported from the Wtodawa Formation and the underlying Lublin and Biatopole Formations (Moczydtowska, 1991). Presumed cyanobacterial microfossils and possible cyanobacterial envelopes and clusters of colonial spherical vesicles occur in the Biatopole and Lublin Formations. The Wlodawa Formation yielded numerous sheaths of vendotaenids, abundant filamentous cyanobacteria and leiosphaerid acritarchs. The first specimens of ornamented acritarchs assigned to diagnostic Early Cambrian taxa appear in the uppermost part of the formation and consist of species part of the Asteridium tornatumComasphaeridium velveturn acritarch Zone ( Platysolenites antiquissimus faunal Zone) that also characterize the overlying Mazowsze Formation, where they are more frequent (Moczydtowska, 1991 ). The Mazowsze Formation (Arefi and Lendzion, 1978;

207

Fig. 2) consists of alternating siltstone and sandstone and contains abundant and diverse acritarchs. The lavishly ornamented acritarch taxa exhibit processes, spines, ridges, muri, membranes and pores. These morphologic features persist among acritarch taxa from younger Cambrian strata (Moczydiowska, 1991 ). The Precambrian-Cambrian boundary and the base of the Cambrian System and the Phanerozoic Eon (recently ratified by the IUGS Executive Committee) are recognized by comparison of ichnofossils, small shelly fossils and acritarchs (Moczydtowska, 1991). On such grounds, the upper part of the Wtodawa Formation is correlated with the boundary level in eastern Newfoundland (Moczydtowska, 1991; Fig. 2).

6. Late Neoproterozoic Varangerian and postVarangerian biostratigraphy in Baltica Neoproterozoic deposits associated with the Varangerian glaciation are characterized by extreme morphological simplicity and the virtual lack of diagnostic taxa are salient features (Vidal, 1985). These features are observed in planktonic assemblages of microfossils from the top of the interglacial Nyborg Formation (Fig. 2) in Varanger Peninsula and the late Vendian Biatopole, Lublin and W~odawa Formations (Fig. 2) in the Lublin Slope. They are also recognised in associations reported from the late Vendian in Estonia, Lithuania, Belarus, the Ukraine, Russia (Vidal and Knoll, 1982), Namibia (Germs et al., 1986) and Australia (Damasa and Knoll, 1986). Varangerian and Ediacaran successions (Fig. 2) generally lack the distinctive, complex and relatively large acanthomorphic acritarchs reported from terminal Neoproterozoic (Yudomian, Sinian and Ediacaran) rocks in Siberia and China (Yin Leiming, 1985; Awramik et al., 1985; Pjatiletov and Rudavskaya, 1985; Moczydlowska et al., 1993) and Australia (Jenkins et al., 1992; Zang and Walter, 1993). The assemblages include variable proportions of cyanobacterial or bacterial components which normally seem to dominate assemblages related to glacial (Knoll et al., 1981; Vidal and Knoll, 1982) and basinal deposits (Mansuy and Vidal, 1983; Palacios, 1989; Vidal and Nystuen, 1990a) elsewhere. In general, acritarch assemblages from rock units associated with the Varangerian glacial event(s) have notoriously low diversity (Vidal, 1979a; Knoll et al.,

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1981; Vidal and Knoll, 1982; Palacios, 1989), either in response to stressed habitats (Knoll et al., 1981 ) or, alternatively, due to plankton diversity decline favouring cyanobacterial populations as a consequence of global eutrophic conditions (Knoll et al., 1981; Mansuy and Vidal, 1983). No absolutely satisfactory explanation to the evident extinction of pre-Varangerian protists (Vidal and Knoll, 1982) has been so far provided and Varangerian strata are characterized by a remarkable and disturbing lack of micropalaeontological information.

7. Discussion 7.1. Late Neoproterozoic biostratigraphy

At the regional scale, the comparison of patterns of change in the Vendian biota relating to the Varangerian glacial event is severely hampered by the difficulty of detailed correlation of packages of late Neoproterozoic strata within the various areas of Baltica. Hence, despite detailed and confident biostratigraphic correlation of Early Cambrian biostratigraphic units across Baltica (Hagenfeldt, 1989; Vidal and Nystuen, 1990b; Moczydtowska, 1991), substantial problems remain concerning the general stratigraphy and understanding of depositional environments of Vendian strata in parts of the EEP. This stems in part from the considerably uncertain stratigraphic ranges of scant "index fossils", some of which, e.g. vendotaenids and sabelliditids, appear to posses wide stratigraphic ranges in Vendian and Early Cambrian lithostromes (Moczydlowska, 1991). Correlation of post-Varangerian, upper Vendian strata in northern Baltica (Stappogiedde Formation and Lower Breivika Member; Fig. 2) could be attempted with Valdai rock units (Redkino and Kotlin "stages" or "horizons"; see below) in the EEP on the basis of overall comparisons of faunal record and longdistance extrapolation with strata in the White Sea Depression (Fig. 1 ) and the western portion of the EEP. However, this demands caution as recent developments indicate that certain "Ediacaran" faunal components may span substantial stratigraphic intervals (Hofmann et al., 1990; Jensen and Grant, 1992; Conway Morris, 1993; T.P. Crimes, A.N. Insole and B.P.J. Williams, unpublished data; Moczyd~owska and Crimes, unpublished data). Regional lithostratigraphy demands addi-

tional caution as the Valdai "Series", including Redkino and Kotlin "horizons" (Sokolov and Fedonkin, 1990), is generally used in a iithologic and timestratigraphic context. The units are by no means recognizable in simple fossiliferous sections, but are rather the result of regional compilations over hundreds of kilometres across the huge area of the EEP, extending from the White Sea over the Moscow Syneclise to the Volhynian Aulacogen (Fig. 1). The Vendian of the EEP is stringently defined in subsurface sections in the Lublin Slope (Figs. 1, 2), but in adjacent areas to the east the Vendian is defined as a "synthetic regional stratotype"; a composite of fragmentary sections spread over an area extending from the Baltic Sea in the west, the Moscow Syneclise in the east, the Black Sea in the south and the White Sea in the north (Sokolov and Fedonkin, 1990; Fig. 1). The quality of biostratigraphic control within this gigantic lithostrome is extremely variable. Although much data are available, confident understanding of the terminal Neoproterozoic evolution of Baltica is in need of extensive compilation and analysis of the litho- and biostratigraphy of local stratotypes, understanding of facies relationships and the development of sequence stratigraphy. Meanwhile, the absence of an overall synthesis suggests that the conceptual usage of geochronologic names relating to poorly constrained local chronostratigraphic units, being part of the late Vendian lithostrome (such as Kotlin and Redkino), should be abandoned. This is even more so as these "lithostromic" units were also used in recent studies with geochronologic and "event" connotations. 7.2. Palaeogeography

Palinspastic restorations suggest that the successions occupying the Hedmark Basin formed at locations amounting minimum tectonic displacements of 130150 km (Oftedahl, 1943; Strand, 1960; Morley, 1986; Fig. 1). More recently, palaeogeographic position of the Hedmark Basin was restored to a position 200-250 km northwest of its present location, where it formed as one of several rift basins (Nystuen, 1981; Bockelie and Nystuen, 1985). Judging from available radiometric ages the basins appear to have opened in Neoproterozoic times ( ~ 800-650 Ma; Kumpulainen and Nystuen, 1985) in connection with the break-up of the Baltoscandian-Laurentian supercontinent (e.g. Gale

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and Roberts, 1974; Gee, 1975; Schwab et al., 1988). Although located in a continental setting, the rift basins were interconnected by a marine seaway in the west which subsequently, during Cambrian and Early Ordovician times, evolved into the Iapetus Ocean. The fossil microbiotas from the Hedmark Group indicate that in late Riphean and Vendian times the original Hedmark Basin was connected to a marine seaway for most of its history (Vidal and Nystuen, 1990a, b). Palinspastic restorations of the decollement sheet also suggest that the allochthonous Lower Cambrian of the Osen-R~a Nappe Complex at Lake Mjcsa was deposited 130-150 km to the north-northwest relative to the autochthonous Lower Cambrian beds at the eroded Caledonian nappe front (Oftedahl, 1943; Nystuen, 1981; Morley, 1986). Lower Cambrian arenaceous and shaly units in these two tectonostratigraphical positions were all deposited within the same epicontinental foreland sea east of the Iapetus Ocean and the Caledonian mobile belt (Kumpulainen and Nystuen, 1985). This shallow-marine Early Cambrian epicontinental seaway rimmed the northwestern periphery of the Baltoscandian craton and transgressed the craton from northwest to southeast (Vidal and Nystuen, 1990b). In Early Cambrian times an axial-oriented land ridge probably separated areas of epicontinental sea flooding northwestern Baltoscandia from the epicontinental seas of southeastern Baltoscandia (Thorslund 1960; Skjeseth 1963; Martinsson 1974; Bergstr6m 1980). This land barrier terminated towards the southwestern border zone of Baltoscandia. The two epicontinental seas merged into the Iapetus seaway system along the Teisseyre-Tornquist Lineament. During the Early Cambrian, the distribution of faunas displays a coherent pattern supported by good biostratigraphic control (Moczydtowska, 1991) that implies free exchange between the various areas of deposition rimming the western border of Baltica. A more complex picture emerges for the terminal Neoproterozoic (late Vendian) as various rift basins infilled with late Vendian deposits seem to have developed along the perimeter of Baltica and as intracratonic troughs (Fig. 1). Although witnessing of a comparable development, there are indications of substantially discrepant basin history. The most homogenous evolution seems to have occurred in the passive southwestern (Lublin Slope) and northern (Varanger Peninsula) margins of Baltica (Fig. 1). The central portions of Baltica are blanketed

209

by late detrital Vendian successions that originated in extensive aulacogens that opened during a late Riphean (800-700 Ma) event of rifting and intracratonic basin development (Kumpulainen and Nystuen, 1985; Fig. 1 ). This early stage is documented in the western Baltoscandian basins, where it has been correlated with deposition in rift basins (Visingsti Group and Hedmark Group; Kumpulainen and Nystuen, 1985) and contemporaneous emplacement of basic dikes (Patchett and Bylund, 1977; Vidal and Bylund, 1981). A second postulated episode of rift sedimentation and emplacement of tholeiitic dyke swarms (Ottfj~illet dolerites; Solyom et al., 1979) is believed to post-date the Varangerian glacial event in southwestern Baltoscandia, as the Ottfjiillet dolerites, intruding the Toss~sfjallet Group that contains a glacial unit (Kum pulainen, 1980), yielded Rb-Sr ages of 720 + 260 Ma (Claesson, 1976; 1977) and 4°Ar-39Ar ages of 665 + 10 Ma (Claesson and Rodick, 1983). Accepting an approximate age of 650-655 Ma for the Varanger Glaciation, Kumpulainen and Nystuen (1985) concluded that the minimum age of dyke emplacement suggests improbable rates of subsidence and sedimentation and the need for additional age determinations in Baltoscandia. Representing a cooling age, the 4°Ar39Ar 665 + 10 Ma age postdates glacial deposition in the TossAsfj~illet Group and may suggest that the generally accepted ~ 650 Ma age for the Varanger glaciation could be a minimum age irrespective of which of the two glacial events is represented in the Toss~sfj~illet succession. This is clearly at contrast with a recently proposed age range of 590--620 Ma for the Varanger event (Knoll and Walter, 1992). Late Vendian basaltic volcanism is extensive in areas underlying the western margin of the EEP (Poland, Belarus, Russia, the Ukraine and Moldava), covering an approximate area of ca. 140,000 km 2 (Juskowiakowa, 1971). The ~551 Ma age of tuffs in the upper S{awatycze Formation in the Lublin Slope (Compston et al., 1995; Fig. 2) suggests that this volcanic event is connected with the final rifting of Baltica prior to stabilization at the onset of the Early Cambrian transgressions. This event is recorded in the thermal history of organic matter preserved in the late Vendian-early Cambrian succession (Moczydiowska, 1988). Hence, late Neoproterozoic deposits in the region display thermal alteration indexes (TAI) overall higher than those of kerogens from immediately overlying Cambrian

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strata. This feature is consistent with higher rates of heat flow witnessing the opening of the Volhynian Aulacogen that is considered as an aborted rift (Guterch, 1977; Moczydlowska, 1988). On lithologic grounds the Stawatycze Formation was correlated with the volcanogenic Volhyn Group in the Ukraine (Juskowiakowa, 1971 ). The latter is overlain by a succession with an Ediacaran fauna (Redkino "horizon"; Sokolov and Fedonkin, 1990) preceding the latest Vendian-Cambrian successions in Poland and in the Ukraine. The isotopic age of the S~awatycze Formation implies that volcanic rocks in Poland may be younger than those in the Ukraine. It also suggests the existence of episodic rifting, related to the opening of the Iapetus Ocean, at the margin of the Ukrainian Shield. This is consistent with recent palaeogeographic reconstructions (Torsvik et al., 1992) assuming that the Lublin and Podolian Slopes of the Ukrainian Shield were oriented towards Laurentia before rotation of Baltica which commenced in Vendian times. The preliminary U/Pb age of the S{awatycze Formation appears in harmony with K-Ar datings on glauconite from the Redkino "Horizon" (late Vendian=Ediacaran) yielding ages in the range of 584-568 Ma (tuffites of the Kairovo subseries), 591-580 Ma (globular phylosilicates of the Redkino Formation) and 575-546 (globular phylosilicates of the Yaryshevo and Nagoryanka Formations in Podolia; Sokolov and Fedonkin, 1990, p. 186). These datings seem to conclusively suggest an older age for the underlying Volhyn Group, generally regarded to have an age range of 590-625 Ma (Sokolov and Fedonkin, 1990, p. 186). Units overlying the dated Stawatycze tufts are considered to span late Vendian-Early Cambrian ages, as based on palaeontological evidence (see Moczydlowska, 1991, for a recent review). In respect to taxonomic composition, the earliest lbssil records of eastern Newfoundland and Baltica have a great deal in common, and allow detailed biostratigraphic correlations (Moczydtowska, 1991). Recent palaeomagnetic reconstructions for the Vendian to early Palaeozoic have placed Baltica "upside down" (Torsvik et al., 1992), thus locating the Tornquist shelf along the present Teisseyre-Tornquist Lineament (qTL) facing the "eastern" shelf of Avalonia. No palaeomagnetic work has been undertaken in the Lublin Slope, but recent results from Varanger Peninsula suggest the late Riphean-early Vendian palaeo-

pole to have been located at lat. 24°N, long. 207°E and in agreement with contemporaneous poles from Kola Peninsula and the Lake V~ittern Graben (Bylund, 1994a). The late Vendian (Ediacaran) Stappogiedde Formation (Farmer et al., 1992) appears to have a pole position generally close to former Vendian-Early Cambrian poles (Bylund, 1994b). Age data and palaeontological evidence of deposits above and below the S[awatycze volcanics are interpreted to represent the age of final rifting on the margins of Baltica and the initiation of the opening of the Tornquist Sea way between the shelves of Baltica and Avalonia. How this preliminary isotopic age relates to U/ Pb datings on zircons from tufts in the late Neoproterozoic, Ediacaran, Mistaken Point Formation in eastern Newfoundland (565_+3 Ma; Benus, 1988) remains unclear. Together with the absence of glacial deposits below Stawatycze volcanics, the radiometric age of Stawatycze volcanism provides additional support to formerly suggested (Vidal, 1981) post-Varangerian uplift and deep erosion. If ever deposited in the EEP and the western torso of Baltica (Sweden, Finland and the Baltic Depression), Varangerian glacial deposits were probably denudated by deep erosion preceding the late Vendian (Ediacaran) transgression. The comprehensive late Riphean and Vendian rifting and associated sedimentation in Baltica were succeeded by deposition dominated by quartz arenites 50-200 m in thickness, blanketing stable shallow shelves and witnessing the stabilization at the onset of the Early Cambrian transgressions. The age of the Precambrian-Cambrian boundary in the EEP is within the range of 551 + 4 Ma for the late Vendian Stawatycze Formation, ~ 545 Ma for the lowermost Cambrian Rusophycus avalonensis Zone in Newfoundland (Landing, 1992; recently reported as 531 _+ 1 Ma; Bowring et al., 1993) and a 544 Ma U Pb date for zircons from basal Cambrian strata in eastern Siberia (Bowring et al., 1993). On biostratigraphic grounds, quartz arenite deposition was established to be nearly contemporaneous (Early Cambrian Schmidtiellus age) over most of Baltica (Vidal and Nystuen, 1990b; Moczydlowska, 1991 ). It is stratigraphically preceded by units consisting of coarse- to fine-grained turbidite sandstones, subaqueous debris-flow deposits and organic-rich shales (e.g. in the Hedmark Basin and the Holy Cross Moun-

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tains in southern Poland; Fig. 1), deltaic and fluvial, braided-stream sandstones and conglomerates (e.g. in southern Sweden and Denmark). There are, however, major differences in facies associations and sequence thickness between the discrete regions of Cambrian sedimentation. In the Lublin Slope (Fig. 2) the basal portion of the Cambrian System is developed largely as shallow shelf sandstones and fossiliferous mudstones referred to the Wtodawa and Mazowsze Formations. In the Hedmark Basin, the lower member of the Vangs~s Formation is 200-800 m thick and generally consists of arkoses and conglomerates which in the lower part are deltaic, coarsening upwards into braided-stream facies that laterally and vertically grade into littoral and shallowmarine sandstones (Dreyer 1988). However, locally, the lower member consists of turbidite sandstones that grade into black shales (Nystuen, 1982). The palaeogeographic position of the Holy Cross Mountains (southern Poland; Fig. 1) in Early Cambrian times is uncertain. This is largely because of poorly known subsurface relationships and structural complications. The area was probably detached from Baltica along the Teisseyre-Tornquist Lineament ( T I L ) . Despite evident faunal similarity with Baltoscandia during Cambrian times (Dadlez, 1983; Bergstr/3m, 1984), long-distance displacements prior to docking in late Palaeozoic times was inferred (Brochwicz et al., 1981, 1986; Po~aryski et al., 1982; Lewandowski, 1993). Early Cambrian (largely early H o l m i a age) turbidite sedimentation attains in the Holy Cross Mountains an estimated thickness of 5 km and contrasts in respect to the Hedmark and Lublin basins by the absence of Neoproterozoic rocks (Po~aryski et al., 1981). The areas bordering the TFL are subjected to intensive scientific debate. In one model the southern Baltoscandian marginal areas and the EEP are viewed as part of a large single basin of which the Holy Cross Mountains sequence is an integral part (Dadlez, 1983; Bergstr/Sm, 1984). A contending model explains the sedimentary sequences on either sides of this major crustal feature in radically differing terms, claiming that the southern Holy Cross Mountains Basin was brought into its present vicinity to passive marginal basins of southern Baltoscandia and the EEP through extensive strike-slip displacement mainly along the T I L and the Hamburg-Krakow Lineament (HKL;

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Brochwicz et al., 1981, 1986; Po~aryski et al., 1982; Lewandowski, 1993). These models propose strikeslip displacement along the T I L , but disagree as for the sense of the displacement, which is proposed as either dextral (Lewandowski, 1993) or sinistral (Brochwicz et al., 1981, 1986; Po~aryski et al., 1982). Without taking party for any of the two main contending models, it must be pointed out that "relative" close palaeogeographic proximity has to be sought in the late Proterozoic and Early Cambrian sedimentary cover blanketing the Proterozoic metamorphic basement in the involved areas. It is there that major possible discrepancies due to proposed comprehensive lateral crustal displacement along lineaments might be found. To this must be added that the Cambrian trilobite faunas of Baltoscandia, the EEP and the Holy Cross Mountains seem palaeogeographically closely related and that this feature could outrule the possibility of lateral displacement in the order of thousands of kilometres (Bergstr~m, 1984). This is even more so as Early Cambrian trilobite taxa recovered from drillcores penetrating the Lower Cambrian in the Upper Silesian Massif in southern Poland display remarkable similarity with contemporaneous faunas in Baltoscandia and the EEP (Ortowski, 1975; Kowalczewski et al., 1984). Despite this, net strike-slip displacement of a few hundred kilometres is plausible (Pegrum, 1984), All told, the earliest Cambrian successions on the western border of Baltica witness of moderate basin subsidence and minor crustal instability well extending into early H o l m i a times (Vidal, 1985). It seems that a parallel can be found in the Holy Cross Mountains, and supported by faunal similarity in Cambrian times, may indicate deposition in response to extended instability along the western limits of Baltica in Early Cambrian times.

8. Conclusions

(1) Neoproterozoic successions in the western, northern and southern periphery of Baltica in Varanger Peninsula (northern Norway), the Hedmark Basin (southern Norway) and the Lublin Slope (eastern Poland) seem to be particularly well documented. These successions contain slightly different palaeontological records and cannot be considered separately, but are rather complementary to each other.

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(2) The stratigraphic relationship between Varangerian and pre-Varangerian Neoproterozoic strata can be understood only considering the geologic record in various rift basins in western Baltica and passive margin basins in the northern (Varanger Peninsula; Fig. 2) and southern (Lublin Slope; Fig. 1 ) periphery of western Baltica. (3) The H e d m a r k Basin and Varanger Peninsula in southern and northern Norway, respectively, represent an excellent and well-studied record of the Varangerian glacial episode. (4) The Varanger Peninsula and the Lublin Slope in eastern Poland provide sedimentologically well-studied, fossiliferous and isotopically dated portions of post-Varangerian (Ediacaran) deposition. The successions display transitional junctions with fossiliferous Early Cambrian beds marking the Precambrian--Cambrian boundary. However, no single section possesses all the positive attributes that could be desirable for defining a Neoproterozoic stratotype section. (5) Together with field and biostratigraphy relationships of Neoproterozoic successions of the Varanger Peninsula and the Scandinavian Caledonides, newly acquired preliminary isotopic dates of volcanism in western Baltica give further support to the existence of post-Varangerian uplift and denudation affecting the EEP and the western torso of Baltica. The Varangerian glacial deposits appear to be absent in the subsurface of the Lublin Slope, and in view of the above this might either reflect non-deposition or deep erosion prior to the late Vendian (Ediacaran) transgression. ( 6 ) The Neoproterozoic sections of western Baltica are particularly significant in that they present a consistent picture of repeated rifting, rift-related volcanism, basin evolution and sedimentation during the period ~ 800 M a to the base of the Cambrian. Successions in western Baltica present evidence of subsidence and deposition in response to extended instability along the western limits of Baltica in earliest Cambrian Holmia times.

Acknowledgements To a considerable extent this study is a compilation of data produced by research projects supported over the years by grants from the Natural Science Research Council ( N F R ) to G. Vidal and M. Moczydtowska. A

substantial portion of the work reviewed in this paper was generated in collaboration with Drs. Anna Siedlecka (Trondheim) and Johan Petter Nystuen ( O s l o ) . Excellent companionship, generous input and discussions over several years of collaboration are acknowledged. W e like to thank constructive reviews by Drs. Anna Siedlecka and Ian Fairchild.

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