Njg volume 93 2013 no. 2

B-Economique B-Blad 2013 Number 2, Volume 93

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Norwegian Journal of Geology

Dierk Blomeier, Anna M. Dustira, Holger Forke & Christian Scheibner

www.geologi.no/njg

P.P. Norge/Norvège

No. 2, Vol. 93

NORWEGIAN JOURNAL OF GEOLOGY

Porto betalt ved innleveringen

Facies analysis and depositional environments of a storm-dominated, temperate to cold, mixed siliceous–carbonate ramp: the Permian Kapp Starostin Formation in NE Svalbard .................................................... 75

Anette E.S. Högström, Sören Jensen, Teodoro Palacios & Jan Ove R. Ebbestad New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark, northern Norway, from trace fossils and organic-walled microfossils ........................................................................... 95

Victor A. Melezhik, David Roberts, Svein Gjelle, Arne Solli, Anthony E. Fallick, Anton B. Kuznetsov & Igor M. Gorokhov Isotope chemostratigraphy of high-grade marbles in the Rognan area, North-Central Norwegian Caledonides: a new geological map, and tectonostratigraphic and palaeogeographic implications ............. 107

ISBN 978-82-92-39485-4

ISSN 0029-196X ISBN 978-82-92-39485-4

9 788292 394854

2013 Number 2 Volume 93

NORWEGIAN JOURNAL OF GEOLOGY Norsk Geologisk Tidsskrift

Norwegian Journal of Geology Norsk Geologisk Tidsskrift The main journal for Norwegian geological research, distributed to all members of the Norwegian Geological Society. First published in 1905, this geological journal is issued quarterly by the Norwegian Geological Society (Norsk Geologisk Forening). The journal publishes research articles, review articles, papers, notes and discussions relevant to the varied and complex geology of Norway, the Arctic regions, and adjacent offshore areas.

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Front cover A natural exposure of rhythmically layered, tectonically modified, variegated calcite marble and thin intercalations of calcareous schist known in Nordland as the Leivset marble.

A unique isotopic composition of this marble (δ13Ccarb = -11 to -7‰) that has been recorded only once in the Earth’s history links it to the Shuram-Wonoka negative excursion of carbonate carbon that occurred worldwide in the Late Ediacaran (580–550 Ma). The prominent, bright orangepink colour, unique δ13Ccarb ratio and wide lateral distribution in the Uppermost Allochthon of the Scandinavian Caledonides make this particular marble a very distinctive chemostratigraphic and chronostratigraphic marker horizon.

The width of the exposure is c. 1.3 m.

Photo by Victor Melezhik.

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NORWEGIAN JOURNAL OF GEOLOGY Facies analysis and depositional environments of the Permian Kapp Starostin Formation in NE Svalbard

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Facies analysis and depositional environments of a storm-dominated, temperate to cold, mixed siliceous–carbonate ramp: the Permian Kapp Starostin Formation in NE Svalbard Dierk Blomeier, Anna M. Dustira, Holger Forke & Christian Scheibner Blomeier, D., Dustira, A.M., Forke, H. & Scheibner, C.: Facies analysis and depositional environments of a storm-dominated, temperate to cold, mixed siliceous–carbonate ramp: the Permian Kapp Starostin Formation in NE Svalbard. Norwegian Journal of Geology, Vol 93, pp. 75–93.

Trondheim 2013, ISSN 029-196X. A facies model of the Permian Kapp Starostin Formation assigns various facies to specific depositional environments and thus shows the detailed spatial and temporal development of a temperate, mixed siliceous–carbonate ramp from the upper Cisuralian (Artinskian) into the Lopingian (Changhsingian). Calcareous, partly glauconitic and generally well-sorted sandstones are interpreted to represent shallow-marine sand flats within the most proximal, foreshore to shoreface areas of the inner ramp. These sediments indicate the uplift of a terrestrial, siliciclastic source area probably to the north or northeast of the study area. Highly diverse, commonly strongly silicified, skeletal limestones contain a typical heterozoan biotic assemblage, marked by a varying abundance of brachiopods, bryozoans and crinoids, as well as siliceous sponge spiculae. The carbonateproducing biota shows a specific distribution within the open-marine areas of the inner to mid ramp, where the bioclastic debris was reworked, redistributed and washed together by waves, tides and periodically occurring storms. While sandy brachiopod shell banks (coquinas) were mainly present within the inner ramp, bryozoan and crinoidal detritus accumulated within more distal areas, originating from scattered build-ups at the outer edge of the mid ramp. Spiculitic cherts, the most prominent facies of the Kapp Starostin Formation, are formed by the accumulation of abundant siliceous sponge spiculae, representing the major silica factory of the shelf. These deposits have the widest distribution, ranging from the inner ramp (light-coloured, massive to nodular cherts) around the fair-weather wave base to deeper-marine, outer ramp areas below the stormweather wave base (dark-coloured, bedded to massive cherts). Finely laminated to massive black shales generally indicate the most distal, deepmarine, toe-of-slope and basinal areas of the outer ramp, below the storm-weather wave base. The sediments originate from the accumulation of fine-grained, terrigenous matter under quiet-water conditions. The local preservation of primary lamination points to oxygen-depleted conditions, while bioturbation at other levels indicates the presence of bottom-feeding organisms under well-oxygenated conditions. The various facies were deposited on a stable, shallow submarine ramp marked by a subdued relief, gently sloping towards the south. The strata are arranged into four stacked parasequences (shallowing-upward cycles), which are interpreted as the result of short-term, possibly glacio-eustatic sea-level fluctuations superimposed on a long-term sea-level fall.

Dierk Blomeier, Norwegian Polar Institute, Geo Department, Hjalmar Johansens gate 14, 9296 Tromsø, Norway. Anna M. Dustira, Department of Geology, University of Tromsø, Dramsveien 201, 9037 Tromsø, Norway. Holger Forke, Millennia stratigraphic consultants, Lychenerstrasse 54, 10437 Berlin, Germany. Christian Scheibner, University of Bremen, Geosciences, Klagenfurter Straße 2, D-28334 Bremen, Germany.

E-mail corresponding author (Dierk Blomeier): blomeier@npolar.no

Introduction Sedimentary bedrock of Permian age records an exceptional time period in Earth history, characterised by fundamental climatic, oceanographic and environmental changes. Svalbard, comprising a part of an epicontinental shelf sea at the northern margin of Pangaea, drifted from approximately 25°N in the Late Carboniferous to around 45° in the Late Permian (Scotese & Langford, 1995; Ziegler et al., 1997; Golonka, 2002). The shelf was arranged into a series of platforms and basins, which are exposed today in various circum-Arctic regions, e.g., eastern North Greenland (Wandel Sea Basin), the Barents

Sea (Finnmark Platform, Stappen High), Arctic Canada (Sverdrup Basin) and Russia (Timan–Pechora Basin). All of these areas record a prolonged time period, termed the Permian Chert Event (Beauchamp & Baud, 2002), during which massive chert successions accumulated while temperate- to cold-water conditions prevailed within the shelf seas of the northern hemisphere. Marine warmwater carbonates disappeared completely during this ca.

30 Myr-long period (Artinskian–Changh­ singian), and were replaced by temperate- to cold-water depositional systems. The latter are marked by the predominance of siliceous sponges, whereas carbonate-producing organisms played only a minor role within the shelf

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ecosystems. The profound change from warm-water to temperate- and possibly cold-water environments has been associated with a number of palaeogeographic (e.g., closure of the Ural Ocean during the northward drift and final consolidation of Pangaea) and palaeoceanographic changes (e.g., upwelling of cold, nutrient-rich deep waters, salinity changes, anoxia, ocean acidification; Beauchamp, 1994; Stemmerik & Worsley, 1995; Beauchamp & Baud, 2002; Reid et al., 2007; Beauchamp & Grasby, 2012).

On Svalbard, Permian strata, consisting of the upper part of the Gipsdalen Group (Early Carboniferous to Early Permian), the Bjarmeland Group (Early Permian, restricted to Bjørnøya) and the Tempelfjorden Group (Early to Late Permian), are well exposed within a number of areas. Deposits of the Gipsdalen Group (Serpukhovian– Artinskian) mainly consist of carbonates and evaporites, representing a restricted- to open-marine, warm-water setting, marked by broad sabkhas, evaporite basins and carbonate platforms. The strata comprise a fully photozoan (Wordiekammen Formation) to reduced photozoan/ heterozoan assemblage (Gipshuken Formation, Blomeier et al., 2011). The strata of the overlying Tempelfjorden Group (Artinskian–?Changhsingian), on the contrary, are dominated by spiculitic cherts, as well as black shales, partly glauconitic sandstones and strongly silicified limestones as minor lithologies. The deposits reflect a generally deeper, temperate- to cold-water setting, comprising open-marine, shallow nearshore, to deeper offshore areas of a storm-dominated, mixed siliceous– carbonate shelf. A fully heterozoan biotic assemblage consists of siliceous sponges as the main silica factory and bryozoans, brachiopods and echinoderms as minor carbonate producers. Both lithostratigraphic groups are separated by a major unconformity (disconformity) comprising an Early Permian (Artinskian) hiatus, which resulted from the subaerial exposure of extended shelf areas and the subsequent erosion of the uppermost Gipsdalen Group strata with a renewed transgression at the beginning of the Tempelfjorden Group sedimentation (Ehrenberg et al., 2001; Blomeier et al., 2011).

Since the Tempelfjorden Group was defined by Cutbill & Challinor (1965), a number of publications have discussed the varied lithology, faunal assemblages, local and regional geological development, depositional environments, biostratigraphy and the sequence stratigraphic architecture of the strata (Burov et al., 1965; Szaniawski & Malkowski, 1979; Hellem, 1980; Knag, 1980; Biernat & Birkenmajer, 1981; Cutler, 1981; Lauritzen, 1981; Malkowski, K., 1982; Nakamura et al., 1987; Fredriksen, 1988; Henriksen, 1988; Stemmerik, 1988; Nakrem, 1988, 1991, 1994; Mangerud & Konieczny, 1991, 1993; Nakrem et al., 1992; Bottolfsen, 1994; Mangerud, 1994; Saalmann, 1995; Dallmann et al., 1999; Buggisch et al., 2001; Ehrenberg et al., 2001; Hüneke et al., 2001; Carmohn, 2007; Chwieduk, 2007; Grundvåg, 2008; Groen, 2010; Collins, 2012). However, uncertainties still exist with respect to age constraints, the regional

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and local palaeogeography, and the lithostratigraphic arrangement of the group.

This article presents a comprehensive facies study of the Tempelfjorden Group (Kapp Starostin Formation) in NE Svalbard. Four vertical sections with a thickness from around 80 to 100 m were measured and outcrop observations combined with detailed microfacies studies.

The latter enables the characterisation of a number of depositional environments defined by specific facies associations and their interpretation in terms of shelf position (distal/proximal), water depth (intertidal/ deeper submarine) and energy level (high/low). By identifying and correlating depositional cycles and their stacking pattern at each section location, the overall cyclostratigraphic architecture of the strata and the spatial and temporal development of the shelf area are discussed.

Facies analysis proves a powerful tool to interpret both depositional environments and the palaeogeographic and cyclostratigraphic development of the Permian strata of Svalbard.

Regional geological setting and litho­ stratigraphy of the Tempelfjorden Group Late Palaeozoic bedrocks crop out within the entire Svalbard archipelago with superb exposures on Spits­ bergen and Nordaustlandet complemented by small but important exposures on Barentsøya, Edgeøya and Bjørnøya (Fig. 1). While the sedimentary strata along the west coast of Spitsbergen are strongly deformed, steeply dipping and commonly thrusted (West Spitsbergen Fold Belt), outcrops within central and NE Spitsbergen as well as Nordaustlandet show mostly undeformed strata, with only gently dipping or horizontal bedding.

The Permian sedimentary record within NE Svalbard (NE Spitsbergen, Nordaustlandet) comprises two major lithostratigraphic units, the Early Carboniferous to Early Permian Gipsdalen Group (Serpukhovian–Artinskian) and the Early to Late Permian Tempelfjorden Group (Artinskian–?Changhsingian; Fig. 2). Both the lower and upper boundaries of the latter are distinct and sharp.

The boundary to the underlying Gipsdalen Group is marked by a major hiatus and shows a sharp facies change from warm-water carbonates and evaporites to mixed siliceous–carbonate deposits (Blomeier et al., 2011).

The boundary to the overlying, Early to Middle Triassic Sassendalen Group is characterised by the sharp change to fine-grained, marine siliciclastics such as black shales, siltstones and minor sandstones, combined with the termination of biotic deposits such as spiculitic cherts and bioclastic limestones. Controversy surrounds the transition into the Triassic, which is either conformable and marked by a condensed succession, or might comprise a major unconformity and hiatus. In this connection, the stratigraphic location of the Permian/Triassic boundary,

NORWEGIAN JOURNAL OF GEOLOGY Facies analysis and depositional environments of the Permian Kapp Starostin Formation in NE Svalbard

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Figure 1. General geological map of the Svalbard archipelago (Bjørnøya excluded), indicating the section locations within NE Svalbard (Z1, S1, H1, E1).

Note that all sections are located on the Eastern Svalbard Platform, a tectonic element east of the Lomfjorden Fault Zone.

which currently is defined at the sharp lithological boundary of the Tempelfjorden Group to the overlying Sassendalen Group, is also a topic of ongoing discussion (Steel & Worsley, 1984; Mørk et al., 1989; Stemmerik & Worsley, 1989; Mangerud & Konieczny, 1993; Mangerud, 1994; Stemmerik, 1997; Wignall et al., 1998; Ehrenberg et al., 2001; Nakrem et al., 2008; Dustira et al., 2013).

In most of Svalbard, the Tempelfjorden Group consists solely of the Kapp Starostin Formation. Only within the southernmost areas of Spitsbergen and on Bjørnøya

the formation is replaced by the contemporaneous Tokrossøya Formation (Hornsund), and Miseryfjellet Formation (Bjørnøya), which both formed in discrete depositional basins marked by specific sedimentological and palaeogeographical developments (Dallmann et al., 1999; Fig. 2).

The Kapp Starostin Formation shows strongly variable thickness and lateral facies changes, implying a stable marine shelf characterised by a pronounced palaeobathymetry from deeper, basinal depocentres to

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Figure 2. Permian time-scale (ICS, 2012), global sea-level curves and climate (yellow = warm; blue = temperate climate; Haq & Schutter, 2008) and modified lithostratigraphic system of Svalbard (Dallmann et al., 1999). Currently, considerable uncertainties exist with respect to age datings of the Tempelfjorden Group and its lithostratigraphic sub-units. SG = Sassendalen Group; Mb = Member.

shallow platform areas, controlling local sedimentation.

As no major tectonic activity has been reported from the Permian on Svalbard, the local variations in thickness and facies might be related to the existence of a number of pre-existing, structural elements, inherited from the reactivation of major tectonic lineaments during the Early Carboniferous, and the subsidence of a series of elongated, narrow, rift basins (St. Jonsfjorden Trough, Inner Hornsund Trough, Billefjorden Trough and Lomfjorden Trough) bounded by adjacent highs (Sørkapp–Hornsund High, Nordfjorden High, Ny Friesland High; Steel & Worsley, 1984; Dallmann et al., 1999). Accordingly, the local thickness of the Kapp Starostin Formation varies substantially, from a maximum of 460 m in central Spitsbergen (with 380 m at its type section in Festningen, outer Isfjorden) to only a few metres in the Hornsund area.

Farther south (Sørkapp Land), the strata completely wedge out against the margins of the Sørkapp–Hornsund High, a positive structural element, which remained emergent and acted as a terrestrial source area throughout the entire Permian.

The varied sedimentology and palaeogeography of Svalbard’s Permian depositional basin is expressed in the internal lithostratigraphic arrangement of the formation (Fig. 2). In most areas on Spitsbergen (apart from Brøggerhalvøya in the northwest of the island) and on Nordaustlandet, the basal part of the Kapp Starostin Formation is formed by the Vøringen Member, which

comprises a prominent, up to 40 m-thick (Dallmann et al., 1999) succession. The member consists mainly of partly silicified, locally sandy, fossiliferous or lithoclastic limestones and allochemical sandstones, featuring a diverse, fully heterozoan biotic assemblage mainly consisting of brachiopods, with lesser amounts of bryozoans, echinoderms and siliceous sponges as well as various trace fossils. The sediments record a major regional transgression, which resulted in the flooding of extended, previously subaerially emerged sabkha and carbonate platform areas across the whole of Svalbard (Hellem, 1980; Steel & Worsley, 1984; Blomeier et al., 2011). Above the Vøringen Member, the strata of the Kapp Starostin Formation are arranged into a number of informal, local members (Fig. 2). Within the type area of the formation in western central Spitsbergen, the Svenskeegga and Hovtinden members were proposed by Cutbill & Callinor (1965). The lower Svenskeegga member, which comprises the strata above the Vøringen Member, consists mainly of spiculitic cherts, silty shales and bioclastic limestone beds at the top. The overlying Hovtinden member is composed of shales, siltstones, sandstones and sandy limestones, laterally grading into the Revtanna member in southern Spitsbergen (Bellsund) and into the Stensiöfjellet member in central Spitsbergen (inner Isfjorden). The latter two members were proposed due to substantial facies changes in the upper part of the Kapp Starostin strata compared with the type area at Festningen (outer Isfjorden).

Accordingly, the Revtanna member, proposed by Knag

NORWEGIAN JOURNAL OF GEOLOGY Facies analysis and depositional environments of the Permian Kapp Starostin Formation in NE Svalbard

(1980) and defined by Dallmann et al. (1999), comprises siliciclastic-dominated strata, probably originating from the erosion of the neighbouring Sørkapp–Hornsund High. The member consists of three coarsening-upward sequences, each beginning with limestones and grading up via dark shales or siltstones into clayey, glauconitic sandstones. The Stensiöfjellet member, proposed and defined by Dallmann et al. (1999), is also marked by the predominance of siliciclastic material, characterised by distinct green, glauconitic sandstones (with up to 30% glauconite), sandy spiculitic cherts and sporadic brachiopodal limestone horizons.

On Nordaustlandet, parts of the Kapp Starostin strata are arranged into the Palanderbukta and Selanderneset members, which have only a local importance and are not correlated with the local members on Spitsbergen.

Whereas the Palanderbukta member, defined by Lauritzen (1981), is dominated by fossiliferous, sandy limestones with minor sandstones and cherts, the overlying Selanderneset member, originally introduced as the Selander suite by Burov et al. (1965), consists mainly of sandy fossiliferous limestones and glauconitic sandstones, which form three separate successions divided by minor chert intercalations.

However, it is rather hard to comprehend the internal lithological arrangement and boundary definitions of the various members in the field. In addition, neither the stratigraphic constraints of the informal members nor the overall lithostratigraphic framework of the Kapp Starostin Formation are yet fully understood.

Age constraints of the Tempelfjorden Group have been proposed by Forbes et al. (1958), Gobbett (1963), Cutbill & Callinor (1965), Szaniawski & Malkowski (1979), Biernat & Birkenmajer (1981), Nakamura et al. (1987), Stemmerik (1988), Nakrem (1988, 1991, 1994), Nakrem et al. (1992), Mangerud & Konieczny (1993), Buggisch et al. (2001) and Chwieduk (2007) based on small foraminifers, bryozoans, brachiopods, corals, conodonts and palynomorphs.

Accordingly, sedimentation of the Vøringen Member started at sometime in the late Artinskian/early Kungurian and continued until the late Kungurian, with the main part of the Kapp Starostin Formation probably deposited in the Guadalupian (Fig. 2).

Methods Four vertical sections have been established in key localities in northeastern Svalbard during geological summer fieldwork 2005 to 2007 (Fig. 1). The section sites comprise the stratotype locations for the Palanderbukta member at Zeipelfjella (section Z1) originally described by Lauritzen (1981) and for the Selanderneset member (section E1) at Eremitten (Hellem & Worsley, unpublished). At the section sites, the thickness, colours, lithologies, textures, sedimentary structures and macrofossils of each bed, as well as the overall large-scale stacking pattern of the entire strata, have been documented.

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In addition to the outcrop investigations, detailed microfacies studies have been carried out and 205 thinsections (Z1: 42, S1: 41, H1: 98, E1: 24) were prepared in order to investigate the microfacies and compositional variations of the strata. The deposits are highly diverse, showing strongly varying proportions of detrital, terrigenous, siliciclastic material (mainly quartz grains) and organic silica (sponge spiculae), as well as skeletal and non-skeletal carbonate components, embedded in carbonate or quartz matrixes and cements. Generally, carbonates with a content of less than 10% detrital quartz grains of the components are described according to the classification schemes of Dunham (1962) and Folk (1959).

The description of the mixed siliciclastic–carbonate deposits follows the classification of Mount (1985), based on the ratio of siliciclastic to carbonate material.

Siliciclastic sediments, consisting almost entirely of detrital quartz grains (>90%), are classified after Wentworth (1922). The frequency of the various components and matrix is estimated using visual comparison charts and is described after the following key: abundant (>60% of whole rock), frequent (60–40%), common (40–20%), occasional (20–5%), rare (<5%).

Facies Analysis Based on the field data and the findings of the microfacies studies, the strata of the Kapp Starostin Formation are arranged into four main facies associations. In the following, all facies are described, their spatial and regional occurrence indicated and their environmental significance discussed.

1 Sandstones (Figs. 3A, 3B) Description: Massive, thin- to medium-bedded sandstone beds and lenses predominantly show clear green colours, in addition to whitish and ochre colorations. Locally, the generally highly porous sediments are characterised by abundant Zoophycos or Skolithos burrows. Abundant, sand-sized, subangular to well-rounded, single-crystal quartz clasts are typically densely packed, forming a clastsupported fabric. Occasionally to frequently occurring, ruditic skeletal fragments of mainly thick-shelled brachiopods are locally embedded in the commonly well-sorted sediments. Rare to occasionally occurring glauconite minerals, peloids and larger trepostomebryozoans and crinoid fragments form minor constituents.

Non-skeletal grains and skeletal fragments are commonly cemented by one generation of blocky sparite, which is locally replaced by microcrystalline quartz. Within poorly washed areas and burrows, a greyish, micritic to microsparitic matrix is present.

Spatial and regional occurrence: Sandstones are common in NE Svalbard, where they form a considerable part of the basal Vøringen Member and characteristic greenish marker horizons in the overlying strata of the Kapp

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Figure 3A. Section S1/unit 28: Interbedding of distinct green, glauconitic sandstones (a, facies 1) with brachiopodal lime­stones (b, facies 2a).

Hammer for scale. 3B. Section H1/sheet 65: In addition to abundantly occurring, well-sorted, subangular to rounded, single-crystal quartz clasts, sandstones commonly to occasionally show coarser skeletal fragments of predominantly brachiopods (a: shell fragments; b: spines). 3C. Section S1/ unit 28: In the field, coarse-grained, brachiopodal limestones (facies 2a) are marked by the accumulation of more or less parallel-oriented, ruditic shell fragments, mostly from spiriferids and productids. Due to the local presence of glauconite, the commonly extensively silicified limestones in places show greenish colorations. 3D. Section H1/sheet 63: This brachiopodal rud- to floatstone (facies 2a) consists mainly of the coarse-grained debris of brachiopods (a: shells fragments; b: spines) enclosed by one generation of blocky spar. Here, the original fibrous microstructure of the bioclasts is still preserved. 3E. Section H1/sheet 25: Locally, in situ preserved colonies of trepostome bryozoans occur at the base of the Kapp Starostin Formation (Vøringen Member) in NE Svalbard. The boundary to the underlying Gipshuken Formation is formed by a bored hardground (arrows), which the colonies used as a growth substrate. Hammer for scale. 3F. Section H1/sheet 15: This bryozoan floatstone (facies 2b) is marked by coarse-grained skeletal fragments of timanodictyid (a: Gilmouropora timanodictyid) and fenestrate bryozoans (b), which are embedded in a spiculitic matrix formed by abundant megaspiculae and grey micrite. 3G. Eremitten location: This coarse-grained, crinoidal limestone (facies 2c) commonly shows whitish-weathering stem fragments (a), partly in natural articulation, filigree bryozoan bioclasts (fenestrate bryozoans, arrows) and occasional whole brachiopods (b). The diameter of the coin is about 2 cm. 3H. Section H1/sheet 12: The crinoidal bioclasts (a) of this rudstone (facies 2c) commonly show a syntaxial overgrowth (arrows). The original matrix is entirely replaced by microquartz. Note the trepostome bryozoan fragment (b).

Starostin Formation. The sediments are associated with coarse-grained, brachiopodal limestones (facies 2a) and fine-grained limestones (facies 2e) as well as more rarely with whitish, massive cherts (facies 3a), forming successions up to a few metres in thickness.

Environmental interpretation: The calcareous, commonly glauconitic and bioclastic sandstones are interpreted as a typical facies of proximal, shallow-marine foreshore to shoreface areas of the inner ramp, marked by a high siliciclastic input (Fig. 4). Within these nearshore areas, quartz clasts imported from a terrestrial source were constantly reworked and winnowed by tidal currents

and/or wave action above the fair-weather wave base (FWWB), resulting in generally well-sorted and selectively concentrated, detrital quartz grains. The local abundance of benthos, causing increased bioturbation and burrowing of the siliciclastic rocks, has been interpreted as a stormrelated feature (pipe rocks; Droser, 1991), which typically forms in the aftermath of storm events by the activity of opportunistic bottom feeders. Characteristic greenish colours of the sandstones are the result of a minor but constant amount of glauconite, the formation of which is usually associated with an open-marine environment, low sedimentation rates, low temperatures and deepermarine conditions (Odin & Letolle, 1980; Huggett & Gale,

quartz clasts peloids lithoclasts sponges foraminifers ostracodes corals bryozoans crinoids chetitides molluscs brachiopods Figs. 6, 7: sponge spiculae trepostome bryozoans fenestrate bryozoans bioturbation Zoophycos Skolithos siliceous concretions carbonate concretions

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Figure 4. Schematic shelf model for the Kapp Starostin Formation in NE Svalbard, showing the general distribution of the four main facies (sandstones, limestones, cherts, black shales) across the inner, mid and outer ramp sections. In addition, the distribution of the allochthonous, non-skeletal and skeletal components is indicated. Thick line = rock forming, thin line = significant component, dashed line = accessory. MLW = Mean low water level. FWWB/SWWB = Fair/storm-weather wave base.

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1997). This obvious contradiction to the herein proposed sedimentary environment is explained by the postdepositional, early diagenetic formation of glauconite from pore waters seeping through the uppermost horizon just beneath the sea floor. Accordingly, the glauconite minerals formed after the sandstone sedimentation, when deeper, quiet-water conditions marked by low sedimentation rates prevailed (Blomeier et al., 2011).

2 Limestones Limestones in NE Svalbard generally contain a substantial proportion of detrital quartz clasts and thus constitute sandy, bioclastic limestones to bioclastic sandstones (after Mount, 1985). The commonly strongly silicified sediments are generally thick- to thin-bedded and form laterally persistent successions up to a few metres thick with glauconitic sandstones (facies 1) and light, massive to nodular cherts (facies 3a). Slightly wavy bedding planes are locally replaced by low-relief stylolites. Colours vary from light- to dark-grey to ochre, with brownish or rare reddish discolorations. The limestones contain a diverse, fully heterozoan biotic assemblage consisting mainly of bryozoans, brachiopods and echinoderms, as well as molluscs (bivalves and gastropods), solitary rugose corals and siliceous sponge spiculae as minor elements.

Non-skeletal grains mainly consist of quartz sand and peloids, and minor proportions of larger lithoclasts (extraclasts). The depositional area of the limestones generally comprises intertidal to shallow submarine areas of the inner and mid ramp (Fig. 4). Based on the different primary habitat of the various marine benthos, the composition of the limestones is quite variable and indicates deposition within specific zones within these ramp areas. Based on fabrics, grain size and composition of the bioclastic and non-skeletal components, the limestones are arranged into five sub-facies (2aâ&#x20AC;&#x201C;e).

Coarse-grained, bioclastic limestones These moderately to poorly sorted, highly fossiliferous limestones are generally characterised by the predominance of coarse-grained skeletal components.

Strongly varying amounts of the main faunal elements allow the determination of brachiopod- (facies 2a), echinoderm- (facies 2b) or bryozoan-dominated (facies 2c), as well as mixed-bioclastic (facies 2d) limestones.

2a Brachiopod-dominated limestones (Figs. 3C, 3D) Description: This sub-facies is characterised by the abundant to frequent occurrence of thick-shelled brachiopods, commonly comprising spiriferids and productids, as well as more rare terebratulids, athyrids, rhynchonellids and strophomenids. While the coarsegrained, typically parallel-oriented shell debris is commonly accumulated within massive, laterally extensive brachiopod coquinas, whole specimens are enriched in more rarely occurring brachiopod pavements, commonly a few centimetres to several decimetres

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thick. Coarse-grained bioclasts of bryozoans, bivalves, gastropods and chaetetids are present in minor amounts, as are commonly strongly altered and Fe-stained lithoclasts (reworked extraclasts from the underlying Gipshuken Formation; Blomeier et al., 2011). The ruditic components are embedded in a micritic, microsparitic, spiculitic or sandy matrix consisting mainly of carbonate mud, sand-sized, edge-rounded to rounded, detrital quartz grains, smaller fragments of brachiopods (filaments), bryozoans and ostracods, as well as sponge spiculae (mostly megaspiculae) and peloids. Within well-washed areas, the components are cemented by blocky sparite.

Echinoderm fragments often show a syntaxial overgrowth.

The irregularly distributed, arenitic to ruditic components constitute mostly rudstones (minor packstones and grainstones) with a component-supported fabric, but also more loosely-packed floatstones (minor wackestones).

Glauconite minerals and mineral separations are rarely present. Fragments and matrix are commonly strongly silicified and the original material is, to varying degrees, replaced by microquartz or multigenerational chalcedony.

Spatial and regional occurrence: Brachiopodal coquinas are the typical facies of the Vøringen Member, commonly interbedded with fine-grained, bioclastic, peloidal limestones (facies 2e) or sandstones (facies 1). Higher up in the strata, brachiopod-dominated deposits (mostly coquinas, minor pavements) continue to be important facies elements, commonly occurring within the upper part of individual depositional successions.

Environmental interpretation: Within shoreface to transitional offshore areas, sandy brachiopod shell banks, marked by the accumulation of coarse-grained debris, were a prominent feature of the inner ramp (Fig. 4). The shells, probably washed together during storms, were constantly reworked due to wave action and tidal currents. Grainstone fabrics, the parallel orientation of elongated shells, their commonly high degree of abrasion (roundness) and relatively high sand proportions imply proximal, agitated areas above the FWWB. In contrast, brachiopod pavements, marked by the accumulation of mainly whole specimens, were probably present on more distal and deeper mid-ramp offshore plains between the FWWB and storm-weather wave base (SWWB). During storms, the shells were redistributed across the inner and mid ramp (proximal tempestites), to form the abovementioned brachiopod coquinas within nearshore areas.

A minor part of the bioclastic material was apparently also exported onto adjacent outer ramp areas (distal tempestites), where the brachiopods form a minor faunistic element.

2b Bryozoan-dominated limestones (Figs. 3E, 3F) Description: The main constituents of these poorly sorted and commonly strongly silicified skeletal limestones are frequently to commonly occurring fragments of bryozoans (mainly trepostome and fenestrate, minor cystoporate and timanodictyid). These main components

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are commonly supplemented by echinoderms (crinoids) and megascleres of siliceous sponges (bryonoderm biotic assemblage; Hﾃｼneke et al., 2001). In addition, fragments of brachiopods and solitary rugose corals as well as various small foraminifers and ostracodes are occasionally to rarely present. The irregularly distributed, arenitic to ruditic bioclasts are commonly embedded in a micritic to microsparitic and, more rarely, spiculitic, partly Fe-stained matrix. Due to a rather loose component package, a matrix-supported fabric (floatstones to wackestones) commonly dominates, besides occasional rudstone to packstone fabrics. Silt- to sand-sized quartz grains are present in varying amounts (rare to common), although generally fewer and finer as in the brachiopodal limestones (facies 2a). Components and matrix are commonly strongly silicified, with microcrystalline quartz replacing the original skeletal substance and/or embedding material.

skeletal fragments are poorly sorted and irregularly distributed, mainly forming matrix-supported floatstones (wackestones), in addition to rarer pack- and grainstones (rudstones). The matrix consists mainly of partly silty and Fe-stained micrite or microsparite. In well-washed areas, calcite cements are present, formed mainly by the syntaxial overgrowth of crinoid fragments or blocky sparite.

Cements and embedding material are locally replaced by microquartz or chalcedony.

Spatial and regional occurrence: In NE Svalbard, this facies plays only a subordinate role. Bryozoan-dominated, allochthonous limestones are present within basal sheets of the Vﾃｸringen Member, associated with in situ bryozoan patch reefs (Blomeier et al., 2011). Further up in the strata, this facies occurs only sporadically within single limestone units, associated with crinoidal (facies 2c) and mixedbioclastic (2d) limestones.

Environmental interpretation: The relatively high content of bryozoan fragments within this sub-facies as well as the association with bryozoan-dominated limestones point to the mid ramp as main habitat of the crinoids, including the bryozoan build-ups at the outer edge. In addition, crinoids occur within all limestone sub-facies as minor elements, implying that echinoderms and their bioclasts were widely distributed across the entire shelf. While rare grainstone fabrics reflect proximal, higher-energy areas of the inner ramp, more commonly occurring micritic matrixes imply deposition within distal, low- to moderate-energy areas below the FWWB of the mid ramp.

Environmental interpretation: Allochthonous, bryozoandominated limestones originate from the erosion of mostly trepostome and fenestrate bryozoan colonies. While the primary habitat of robust, bush-like trepostome bryozoans seems to include mainly the outer reaches of the mid ramp around the SWWB, smaller and more fragile fenestrate bryozoans were mostly distributed across adjacent deepermarine areas of the outer ramp, generally below the SWWB (Nakrem, 1994). From the primary habitat areas, coarse-grained bryozoan debris accumulated within deeper offshore banks and was further transported and distributed by storms, currents and waves. In comparison to the brachiopodal sub-facies (2a), the main distribution area of bryozoan-dominated limestones is located in a more distal shelf position, probably mainly comprising mid-ramp to proximal outer-ramp areas around and below the SWWB. This assumption is supported by the predominantly micritic or spiculitic matrix as well as by a generally finer and lesser proportion of terrigenous, detrital quartz, reflecting more distant, deeper-marine, low-energy conditions during deposition.

2c Echinoderm-dominated limestones (Figs. 3G, 3H) Description: This sub-facies is characterised by the predominance of frequently occurring echinoderm fragments (commonly from crinoids), showing arenitic to ruditic sizes of up to a few centimetres. Together with frequently to commonly occurring bryozoan fragments, crinoids form the main component category of this subfacies. Brachiopods, ostracods, sponge spiculae as well as rare, small foraminifers and glauconite minerals are minor constituents. The commonly partly silicified

Spatial and regional occurrence: Although crinoid fragments are present within all limestone facies as a minor constituent, echinoderm-dominated limestones are quite rare. Associated with bryozoan-dominated (facies 2b) and mixed bioclastic limestones (facies 2d), they are present within the Vﾃｸringen Member as well as in individual limestone successions higher up in the strata.

2d Mixed-bioclastic limestones Description: This sub-facies commonly comprises deposits that are marked by a highly diverse, heterozoan component composition. The main constituents of these moderately to poorly sorted, skeletal rud- to floatstones consist of fragments of brachiopods, bryozoans, echinoderms and sponge spicules without any dominant bioclast type. Fragments of chaetetids, gastropods and bivalves are locally present in minor amounts as well as brownish carbonate lithoclasts (extraclasts) and smaller peloids. In poorly washed areas, the matrix is composed of partly Fe-stained micrite/microsparite, sand-sized, edge-rounded detrital quartz grains or accumulations of unidentifiable, fine-grained skeletal fragments (filaments).

However, most of the inter- and intragranular pore spaces are cemented by blocky sparite or the syntaxial overgrowth of echinoderm fragments. Brachiopod shells, in particular, are commonly silicified.

Spatial and regional occurrence: This sub-facies, commonly occurring within the limestone successions of the Kapp Starostin Formation, is marked by a mixture of the most common skeletal fragments and is thus associated with all the other coarse-grained facies types (2a窶田).

Environmental Interpretation: The highly fossiliferous deposits comprise a fully heterozoan (bryonoderm) biotic assemblage, represented by brachiopods, bryozoans,

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echinoderms and/or siliceous sponge spicules. Due to the association with the other facies types, the coarsegrained character of the various bioclasts and the common presence of blocky sparite, sedimentation under highenergy conditions within open-marine and shallowmarine areas of the inner and mid ramp is assumed.

2e Fine-grained bioclastic, peloidal limestones (Figs. 5A, 5B) Description: This facies type comprises generally wellsorted, sandy limestones to allochemical sandstones with arenitic brachiopod shell fragments, peloids and sand-sized quartz grains as the main component categories. Ruditic bioclasts of brachiopods, bryozoans and echinoderms as well as glauconite minerals are occasionally to rarely present. The rather loosely packed and equally distributed components show no preferred orientation and are commonly cemented by one generation of blocky sparite, thus representing grainstones, in addition to more rarely occurring packstones and wackestones, which have a microsparitic matrix. Locally, the sediments show distinct cross-bedding or are intensely bioturbated, marked by abundant Skolithos or Zoophycos burrows with a microsparitic or more rarely spiculitic filling. Locally, the original matrix of these sediments is replaced by microquartz.

Spatial and regional occurrence: In NE Svalbard, this subfacies is prevalent within the Vøringen Member, where it is intercalated with coarse-grained brachiopodal limestones (facies 2a) and laterally grades into calcareous sandstones (facies 1), forming the upper unit of storm-related, stacked sediment couplets (Blomeier et al., 2011). Within limestone successions above the Vøringen Member, fine-grained bioclastic, peloidal limestones play only a subordinate role.

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the FWWB. As the upper unit of proximal tempestites (lower member: coarse-grained brachiopodal limestones), they represent sedimentation in periods without storm activity. Considering the relatively high proportion of detrital quartz in connection with the local abundance of various burrows, the sediments are comparable to pipe rocks (Droser, 1991), reflecting the post-storm activity of opportunistic sea-bottom grazers in near-shore environments.

3 Spiculitic cherts Within the Kapp Starostin Formation strata, spiculitic cherts display the most prominent facies in Svalbard, repeatedly forming successions up to several tens of metres in thickness. The sediments are marked by the accumulation of siliceous sponge needles (megaand microspiculae) and thus display the most obvious products of an immense, biogenic silica production characterising the sedimentation of the formation. During diagenesis, the biogenic silica was remobilised and skeletal and matrix substances were replaced to different extents, leading to a commonly strong silicification not only within the cherts themselves but also within all other facies of the strata. Due to their colour, sedimentary structures and association with other facies, two sub-facies of cherts are distinguished.

3a Light, massive to nodular cherts (Figs. 5C, 5D) Description: This sub-facies forms medium- to very thick-bedded chert units marked by generally lightochre, whitish and more rarely light-greenish colours.

The sediments are massive to nodular with discontinuous, strongly curved bedding planes, probably due to intense bioturbation and subsequent diagenetic pressuredissolution processes. The main component category Environmental interpretation: The fine-grained, well-­ is formed by abundantly occurring, densely packed, monazon megaspiculae, in addition to rarer microspiculae. sorted sediments are interpreted to represent sedi­ Other biota consist of arenitic to ruditic fragments of mentation within open-marine nearshore areas of the brachiopods (also commonly whole biogens), bryozoans, inner ramp. Here, the components were constantly washed crinoids and more rarely solitary corals, which are and winnowed due to wave action or tidal currents. irregularly distributed and commonly enriched in laterally Grainstone fabrics reflect agitated water conditions around

Figure 5A. Section Z1/sheet 23: This medium-bedded, well-sorted, sandy limestone (facies 2e) shows distinct cross bedding. The greenish colour is due to the presence of glauconite. Magnifying glass for scale. 5B. Section S1/sheet 43: This sandy, bioclastic grainstone (facies 2e) shows strongly reworked, elongated brachiopod shell fragments (a) and subangular to well-rounded, sand-sized quartz clasts (b) as the main component categories. 5C. Section Z1/unit 38: Interbedding of thick-bedded, whitish, nodular to massive cherts (facies 3a, a) and medium-bedded, grey limestones above (b). The irregular, wavy bedding planes are probably the result of diagenetic pressure/dissolution processes. 5D. Section H1/sheet 42: Abundant monazon megaspiculae commonly form light-coloured, massive to nodular cherts (facies 3a). The spicules, shown in various longitudinal and cross sections are marked by a central canal, which in places is filled by authigenic glauconite (arrow). The original matrix is replaced by brownish chalcedony. 5E. Section S1/sheet 35: Monotonous succession of dark, thin-bedded cherts (facies 3b) with black-shale partings on the slightly wavy bedding planes. Hammer for scale. 5F. Section S1/sheet 24: Dark cherts (facies 3b) are marked by the accumulation of sponge spiculae, which consist mainly of microspiculae in addition to macrospiculae. The darker micritic areas (a) of this spiculitic packstone probably represent flattened, micrite-filled burrows. 5G. Section H1/units 69/70: Light-coloured, massive cherts and greenish sandstones (facies 3a, 1; a) are overlain by dark, wavy-bedded cherts with shale partings (facies 3b, 4; b). Rusty discolorations due to the weathering of pyrite reflect condensation and low sedimentation rates at the boundary of the two units, which is interpreted as a sequence boundary between two sequences. 5H. Section S1/Vøringen Member: The upper part of the member (a), comprising glauconitic sandstones (facies 1), coarse-grained, sandy limestones (facies 2a, d) and light cherts (facies 3a), is capped by a distinct black-shale horizon, forming the basal sheet of the overlying sequence (b).

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restricted horizons or lenses within the cherts. Detrital quartz grains form a minor constituent and glauconite minerals are rare. The sediments are generally marked by an intense bioturbation mostly by variously oriented, single tubular burrows, backfilled with dense micrite or spiculite. A component-supported fabric (packstone) dominates due to the accumulation of sponge spiculae, in addition to matrix-supported areas (wackestones, also in burrows). Bioclasts as well as pore spaces are generally strongly silicified, often consisting of multigenerational chalcedony or microcrystalline quartz.

Spatial and regional occurrence: Light-coloured, massive to nodular cherts are the most prominent facies of the Kapp Starostin Formation in NE Svalbard. The sediments repeatedly form thick units within individual depositional cycles and are mainly associated with strongly silicified, skeletal limestones (facies 2aâ&#x20AC;&#x201C;e) and glauconitic sandstones (facies 1).

Environmental interpretation: Light-coloured cherts most probably formed on vast offshore plains above the SWWB, comprising distal inner-ramp areas (offshore transition) and the entire mid ramp (Fig. 4). Due to the predominance of siliceous sponges, the sediments within these areas consist mainly of their spiculae, which accumulated after the disaggregation of the skeletons. Bioclasts as well as whole biogens of minor biota, such as brachiopods, trepostome bryozoans, echinoderms and solitary corals are embedded within the cherts and partly form laterally restricted accumulations, washed together by waves, tidal currents or storms. Oxygenated sea-bottom conditions combined with agitated water conditions might also be the reason for the generally lighter colours of this chert sub-type, causing a lack of fine-grained, suspended and organic matter.

An intense bioturbation probably emphasised by post-­ sedimentary pressure/solution processes led to development of the massive to nodular fabric of these cherts.

3b Dark, bedded to massive cherts (Figs. 5E, 5F) Description: This chert sub-facies is characterised by generally dark-grey and, to a minor extent, dark-blue colours, commonly forming thick, internally thin- to medium-bedded and more rarely massive successions.

The chert units show thinly laminated to thinly bedded black-shale partings on the slightly wavy, discontinuous to continuous bedding planes. Locally, the sediments are marked by an intense bioturbation comprising Zoophycos spreiten, but also various other trace fossils, such as different grazing traces and single tubular burrows (dwelling structures?) of various sizes and orientations.

Whole siliceous sponges showing variable, compact to elongated growth forms with diameters of up to several decimetres are locally embedded in situ. Minor components such as arenitic to ruditic skeletal fragments of brachiopods, echinoderms, bryozoans or solitary corals are occasionally to very rarely present, often enriched within individual, massive beds or horizons, constituting spiculitic floatstones marked by brownish or ochre

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weathering colours. Generally, the cherts are formed by the accumulation of siliceous sponge spiculae (both megaand microspiculae) and fine-grained lime mud resulting in low-diversity, well-sorted, spiculitic packstones and wackestones. Glauconite grains are rarely to occasionally present. The deposits are affected by a strong silicification, including the complete or partial replacement of the dense, partly silty, micritic matrix, bioclasts and sponge spicules by microquartz.

Spatial and regional occurrence: While dark cherts form a substantial part of the Kapp Starostin Formation strata within central Spitsbergen, this sub-facies is generally far less common in NE Svalbard, where light-coloured cherts dominate. Dark cherts are commonly associated with black shales, which form thin partings on the bedding planes.

Environmental interpretation: This sub-facies reflects distal, deep, outer-ramp areas, characterised by quietwater, low-energy conditions far below the SWWB (Fig. 4). A low-diversity biotic association is represented by the widespread occurrence of siliceous sponges and the burrows of various bottom feeders, reflecting mostly oxygenated sea-bottom conditions. The original, micritic matrix probably formed due to the constant accumulation of fine-grained, suspended matter, filling the interspaces between the sponge spicules. During storm events, skeletal fragments (mainly brachiopods, bryozoans and echinoderms) as well as fine-grained lime mud were imported from more proximal, inner- and mid-ramp areas via distal tempestites.

4 Black shales Description: Dark-grey to black shales occur either as very thin partings or medium-bedded, massive or internally laminated horizons. A lamination results from the repeated lateral accumulation of silt-sized quartz grains within even finer-grained terrigenous material.

This primary structure is locally obliterated by an intense bioturbation, predominantly displayed by Zoophycos spreiten. In addition to occasionally to rarely occurring sponge spicules and glauconite minerals, arenitic to ruditic skeletal fragments or whole brachiopods are rarely present.

Spatial and regional occurrence: In NE Svalbard, black shales appear only locally at the base of the stacked depositional cycles. Here, the very fine-grained sediments are commonly associated with dark, bedded cherts (subfacies 3b), forming either discrete horizons or thin linings on the bedding planes of the cherts.

Environmental interpretation: Black shales are interpreted as the most distal facies, restricted to deep-marine areas of the outer ramp (Fig. 4). This depositional environment is characterised by quiet-water conditions far below the SWWB and a continuous background sedimentation consisting of very fine-grained terrigenous material (suspended matter). At the sea-bottom, variable oxygen

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levels prevailed through time. The preservation of a horizontal lamination points to oxygen-depleted conditions, whereas abundant burrows and bioturbation resulted in a massive texture, and indicate the presence of bottom-feeding organisms under oxidised sea-bottom conditions.

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The model presented herein is a further development of the one introduced by Blomeier et al. (2011), based solely on the strata of the Vøringen Member. Changes comprise the location of the bryozoan build-ups, which have been moved from the inner-/mid-ramp boundary to the outer mid-ramp margin, and the introduction of light cherts (facies 3a) representing the offshore plains of the mid ramp.

areas (foreshore to proximal shoreface), generally wellsorted and -washed sandstones (facies 1) were constantly reworked and winnowed within shoals and flats under agitated water conditions, due to tidal currents, wave action and periodic storm events. The occurrence of pure sandstones and the substantial sand proportion within the other nearshore lithologies point to a terrestrial source area in the vicinity, and might reflect the uplift and erosion of basement rocks (eastern basement province) in northern Nordaustlandet (Blomeier et al., 2011). Storm events are reflected by the local abundance of Skolithos and Zoophycos trace fossils in the sandstones and finegrained, sandy limestones (facies 2e), interpreted as an increased activity of opportunistic sediment feeders in the aftermath of tempests (pipe rocks; Droser, 1991). These fine-grained sediments commonly form the upper unit of stacked sediment couplets, which comprise coarsegrained brachiopod coquinas (facies 2a) as the lower member (Blomeier et al., 2011). The latter are interpreted as proximal tempestites and reflect the storm-related redistribution of skeletal debris (mainly from thick-shelled brachiopods in addition to minor molluscs, crinoids, sponges and robust, encrusting bryozoans) and nonskeletal components (quartz clasts, peloids, lithoclasts) within sandy shell banks across the shallow-, open-marine flats of the inner and mid ramp. With the transition into adjacent offshore plains, light-coloured cherts (sub-facies 3a) originating from the accumulation of siliceous sponges spiculae, become increasingly prominent.

Inner ramp The inner ramp comprises intertidal to shallow subtidal areas above and around the FWWB, including foreshore, shoreface and the transition into the more distal offshore plains of the mid ramp (Fig. 6). Within the most proximal

Mid ramp With increasing water depth in a seaward direction, the inner ramp gradually passes into the open, shallowmarine offshore areas of the mid ramp, roughly between the FWWB and the SWWB (Fig. 6). This shelf zone is

Ramp model Palaeoenvironmental reconstruction The overall depositional setting of the Kapp Starostin Formation corresponds to an open-water, intermediate- to high-energy, temperate to cold, mixed siliceous-carbonate shelf (Malkowski, 1982; Ehrenberg et al., 2001; HĂźneke et al., 2001; Blomeier et al., 2011). In NE Svalbard, the shelf can be separated into inner-, mid- and outer-ramp areas, each characterised by specific facies associations, depositional environments, sedimentary processes and biotic assemblages (Figs. 4, 6).

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Figure 6. Schematic 3D shelf model for the Kapp Starostin Formation in NE Svalbard, comprising a reconstruction of the ramp geometry. The diagram shows the distribution of facies and biota across the ramp. Fossil symbols in Fig. 4

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characterised by widespread, locally sandy to silty plains, covered by siliceous sponges that constituted the dominant silica factory of the shelf during the Permian Chert Event.

After disintegration of the skeletons, countless siliceous mega- and microscleres accumulated on the sea floor and eventually led to the formation of whitish, massive to nodular, spiculitic cherts (facies 3a). Minor carbonateproducing assemblages gradually changed from mainly thick-shelled into thin-shelled and smaller brachiopods, associated with crinoids, bryozoans and solitary rugose corals. Bryozoans (mainly trepostome) and crinoids are present in probably slightly more elevated areas at the outer mid-ramp margin, forming local build-ups around the SWWB. Their skeletal debris was distributed via storms, waves and tides across the mid-, and innerramp sections and exported onto the outer ramp (distal tempestites). In the process, the allochthonous bioclasts accumulated to form bioclastic limestones (coarse-grained bryozoan and crinoidal limestones, sub-facies 2b, c).

Outer ramp The outer ramp comprises the deepest areas from around to far below the SWWB (Fig. 6). A low-diversity biotic association consists mainly of siliceous sponges, fragile fenestrate bryozoans and various bottom feeders. Under generally low-energy, quiet-water conditions, dark, bedded spiculites (sub-facies 3b) and rare bryozoan limestones (facies 2b) formed due to the mixing of sponge and bryozoan bioclasts with fine-grained matter on the sea floor. Pure black shales (facies 4), displaying a constant background sedimentation and the accumulation of suspended matter, mark areas devoid of any biota.

A massive texture of the shales and cherts points to oxygenated sea-bottom conditions and the intense bioturbation of the muddy sediments by sediment feeders, whereas the preservation of a primary, slightly wavy to horizontal lamination might reflect oxygen-depleted conditions during sedimentation. Locally intercalated, thin- to medium-bedded, strongly silicified limestone beds (facies 2bâ&#x20AC;&#x201C;d), containing coarse-grained skeletal debris (mainly brachiopods, trepostome bryozoans and crinoids) from the mid and inner ramp are interpreted as distal tempestites. They mark the storm-related import of shallow-marine material (components and mud) into the most distal and deepest ramp zones.

Spatial and temporal ramp development All sections are located on the Eastern Svalbard Platform, a tectonic element comprising mainly Late Palaeozoic and Mesozoic bedrock east of the Lomfjorden Fault Zone (Fig. 1). The overall, similar, cyclostratigraphic development and sedimentation pattern in all sections implies a coherent, stable shelf area with a rather subdued relief and a ramp-like morphology. The shallow-marine ramp was marked by a siliciclastic input from a terrestrial source area, probably to the north, reflected by the highest sand contents and biggest grain sizes of quartz grains

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in section Z1. Generally thicker parasequences and a higher proportion of facies representing deeper midramp areas within the H1 and E1 sections imply more accommodation space and thus an inclination of the ramp towards the south.

Facies arrangement and cycle stacking pattern Within all sections, the Kapp Starostin Formation shows a pronounced cyclicity formed by stacked depositional sequences, which consist of conformable successions of genetically related bed sets. Each succession shows a specific order of the herein-defined facies, generally reflecting sedimentation from the deepest, most distal ramp areas at the base to the shallowest, most proximal ramp zones at the top (Fig. 7). Accordingly, the boundaries mark a fundamental and abrupt shift from the shallowest marine environments of the lower succession to the deepest depositional environments of the overlying sequence and thus are regarded as marine flooding surfaces, separating stacked depositional sequences (shallowing-upward sequences). Although no subaerial exposure is recorded at the sequence boundaries, a break in sedimentation during transgression, causing sediment starvation and condensation is commonly indicated by an intensive mineralisation (Fe staining) and glauconitisation of the sediments at the boundaries (Fig. 5G).

The investigated strata are arranged into four parasequences, each marked by an individual facies-set succession, varying substantially in thickness and in the facies and sub-facies which they contain. All sections start with an unconformity at the lower formation boundary to the underlying Gipshuken Formation (Gipsdalen Group), reflecting a fundamental change from warm-water to temperate- and possibly cold-water conditions connected with a major sea-level rise (Beauchamp, 1994; Reid et al., 2007; Blomeier et al., 2011). The lowermost, transgressive sequence, comprising the Vøringen Member (Fig. 5H), is marked by basal hardgrounds locally overgrown by in-situ bryozoan reef knobs, which in turn are embedded and overlain by sandy, skeletal limestones (facies 2aâ&#x20AC;&#x201C;e), skeletal, glauconitic sandstones (facies 1) and light, massive cherts (facies 3a). The ca. 8 m (section Z1) to 21 m (section H1) thick succession reflects the initial transgression after a prolonged subaerial exposure and hiatus during the Artinskian. The variable sediments reflect mid- to innerramp sedimentation under progressively shallower water depths (Blomeier et al., 2011).

This lowermost sequence is sharply overlain by a distinct black-shale horizon in all section sites (facies 4; Fig. 5H), indicating the termination of shallow-marine shelf sedimentation and the subsequent accumulation of suspended matter under quiet-water conditions below the SWWB. The basal black-shale horizon is followed by a ca.

30 to 40 m-thick succession of spiculitic cherts, which are locally dark and interbedded with shale horizons (facies 3b), or light and massive to nodular (facies 3a), reflecting

Facies analysis and depositional environments of a storm-dominated, temperate to cold, mixed siliceous–carbonate ramp: the Permian Kapp Starostin Formation in NE Svalbard

SECTION H1

sequences

SWWB

S3

100 T

T

90

90

T

100 Selanderneset member ?

100

S4

sandstone lithology shelf position limestone & facies chert 1 shale inner mid- outer g r,f shelf shelf shelf chert 2 m w p FWWB

scale (m)

S4

sandstone lithology shelf position limestone & facies chert 1 shale inner mid- outer chert 2 m w p g r,f shelf shelf shelf SWWB

scale (m)

sequences

SWWB

FWWB

scale (m)

sandstone lithology shelf position limestone & facies chert 1 shale inner mid- outer g r,f shelf shelf shelf chert 2 m w p

sequences

SECTION E1

SECTION S1

FWWB

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proximal tempestites

S3

90

T

distal tempestite ?

S3

T

S. mb ?

T

80

S3

70

Selanderneset member ?

T

80

S4

70

proximal tempestites

distal tempestite ?

80

Selanderneset mb ?

SWWB

sequences

80

sandstone lithology shelf position limestone & facies chert 1 shale inner mid- outer g r,f shelf shelf shelf chert 2 m w p FWWB

scale (m)

SECTION Z1

T

70

70

60 T

60

proximal tempestites

not exposed proximal tempestites

T

S2

T

distal tempestite ?

50

T

40

50

50 proximal tempestites

F

S2

T

Pb. mb ?

S2

Palanderb. mb ?

F

50

60

Palanderb. mb ?

60

Palanderbukta member ?

S2

40

40

40

T

proximal tempestites

distal tempestites ?

T

30

30

30

30

not exposed T

FT

20

20

distal tempestites ?

20

20

S1

S1

proximal tempestites T T

proximal tempestites

T

proximal tempestites

10

proximal tempestites T

0

T

0

T

Vøringen Member

T

0

10

S1

Vøringen Member

F

10 S1

T

Vøringen Member

10

0

distal tempestites ?

T

Vøringen Member

Figure 7. Simplified section correlation (Z1, S1, H1, E1; Fig. 1) and cyclic arrangement of the strata of the Kapp Starostin Formation. While the Vøringen Member is clearly recognisable, the stratigraphic locations and range of the Palanderbukta and Selanderneset members remain uncertain after current definition (Dallmann et al., 1999). Colours represent bedrock colours. Lithologies correspond to the main facies, which are assigned to specific ramp areas (sandstones = facies 1 = inner ramp, limestones = facies 2a–e = inner to mid ramp, chert 1 = facies 3a = mid ramp, chert 2 = facies 3b = outer ramp, shales = facies 4 = outer ramp). Symbols show the main components of the sediments (legend in Fig. 4). m = mudstone; w = wackestone; p = packstone; g = grainstone; r, f = rud-, floatstone.

not exposed

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deposition on outer to distal mid-ramp areas. Local limestone beds (rud- and floatstones), rich in fenestrate or trepostome bryozoans, minor crinoids or mostly whole brachiopods, are intercalated in the cherts and shales. Upwards, progressively more bioclastic limestones (facies 2a窶電) are present, grading into a several-metrethick succession of sandy, brachiopodal limestones and glauconitic sandstones at the top of the sequence, which shows an overall thickness between ca. 36 m (section S1) and 49 m (section H1). The limestone-dominated, upper part of this sequence, also constituting the informal Palanderbukta member (Lauritzen, 1981; Fig. 8), reflects inner-ramp sedimentation. The occurrence of light cherts (section H1) and an increased proportion of trepostome bryozoans (section E1) at the top of this succession might record a renewed deepening of the depositional environment at the transition into the overlying, third depositional sequence.

Separated by another marine flooding surface, the third sequence (ca. 22 m (section Z1) to 35 m (section S1) thickness) shows a basal black-shale horizon and dark, bedded cherts in all section sites, reflecting outerramp sedimentation due to a renewed, relative sealevel rise. The sediments grade into light, massive cherts, upwards increasingly intercalated by sandy limestone beds (brachiopod coquinas) and glauconitic sandstones, probably representing the informal Selanderneset member (Lauritzen, 1981).

The overlying, fourth sequence, from which only the lower part is exposed in sections Z1, S1 and E1, shows a similar facies arrangement, from black shales and dark, bedded cherts in the lower part to lighter, massive or nodular

cherts, followed by bioclastic limestones and glauconitic sandstones in the upper part.

Within the lower, chert-dominated parts of the individual sequences, a temporary shallowing is commonly indicated by minor facies variations, such as the deposition of single limestone beds, sandstones or the momentary transition from dark, bedded into lighter, massive cherts. These variations could have been caused by the deposition of distal tempestites, importing shallow-water material from the inner and mid ramp onto the outer ramp (Fig. 7). Another reason might be lower-order, short-term, sealevel fluctuations, superimposed on the higher-order sealevel fluctuations, which are reflected in the parasequence stacking pattern.

As no major tectonic activity has been reported during sedimentation of the Kapp Starostin Formation, the pronounced cyclicity is most likely the result of eustatic sea-level fluctuations, controlling accommodation space and sediment supply, as well as the lateral migration of adjacent ramp (facies) zones. This opinion is in accordance with Beauchamp & Baud (2002), who argued that the cyclicity of the Kapp Starostin Formation and coeval formations in Arctic Canada was driven by substantial ice-volume changes in southern Gondwana, probably accompanied by the development of ice caps in northern high latitudes. Thus, the shallowing-upward facies successions of the individual cycles and the stacking pattern of the latter were probably caused by glacio(?)eustatic sea-level fluctuations of probably the third order, superimposed on a long-term sea-level fall, which lasted from the Mid Permian (Roadian) until the latest Permian (Changhsingian; Haq & Schutter, 2008).

Palanderbukta mb ?

Selanderneset mb ?

S4

S3

S2

Palanderbukta mb ?

S1

Vﾃｸringen Mb

Figure 8. Outcrop at Selanderneset (section S1). The strata, showing a thickness of around 100 m, are dominated by light cherts with intercalated shale, limestone and sandstone beds or bed-sets. The Vﾃｸringen, Palanderbukta and Selanderneset members, as well as individual sequences (S1 to S4; Fig. 7), are indicated.

NORWEGIAN JOURNAL OF GEOLOGY Facies analysis and depositional environments of the Permian Kapp Starostin Formation in NE Svalbard

Conclusions During sedimentation of the Kapp Starostin Formation, a storm-dominated, temperate to cold, mixed siliceous– carbonate shelf prevailed over the depositional area of Svalbard. In the northeast, a shallow, distally deepening, homoclinal ramp developed on a tectonic element, the Eastern Svalbard Platform. With the ramp sloping gently towards the Lomfjorden Fault Zone, siliciclastic material was imported from a terrestrial source area probably to the north.

With increasing water depth and distance to the terrestrial mainland, the ramp is arranged into a nearshore, intertidal to shallow submarine inner ramp, a submarine mid ramp between the FWWB and the SWWB, and a deepermarine, outer ramp with areas below the SWWB, all marked by distinctive biota and sedimentary facies. A varied heterozoan biotic assemblage constituted a major silica and a minor carbonate factory. While the carbonate factory, consisting mainly of brachiopods, bryozoans and echinoderms (with subordinate bivalves, gastropods, chaetetids, foraminifers and ostracods), was restricted to inner- and mid-ramp areas, the silica factory comprised a prolific siliceous sponge fauna, whose primary habitat stretched from the distal inner ramp, across the entire mid ramp to the outer ramp.

On the inner ramp, glauconitic, skeletal sandstones and coarse, sandy brachiopodal coquinas were deposited.

These near-shore areas above and around the FWWB were characterised by sand flats, shoals and shell banks, intensively reworked by tides, waves and occasional storms. With the transition into the mid ramp, roughly between the FWWB and the SWWB, broad offshore plains gradually developed with increasing water depth.

These were populated by abundant siliceous sponges and more rarely by brachiopods, crinoids, bryozoans and solitary corals. Accordingly, sediments consist mainly of light, nodular to massive cherts besides minor coarsegrained, silicified limestones containing the debris of the carbonate producers. At the boundary to the outer ramp, allochthonous bryozoan and crinoidal limestones, originating from the erosion of robust trepostome bryozoan build-ups, are more common. The outer ramp contains the most distal and deep-marine areas below the SWWB. Dark, bedded to massive cherts and black shales formed there due to the accumulation of enormous quantities of sponge needles and fine-grained, suspended matter under quiet-water conditions. In places, interbedded limestones formed due to the accumulation of mainly fragile, fenestrate bryozoans, as well as bioclasts and lime mud imported via distal tempestites from the innerand mid-ramp areas. Changing sea-floor oxygen levels controlled bioturbation and the preservation of either a primary lamination or massive fabrics in the sediments.

The Kapp Starostin Formation strata display a cyclicity which was formed by stacked shallowing-upward

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successions bounded by marine flooding surfaces. The sequences are seen as the result of glacio(?)-eustatic sea-level fluctuations (of probably the third order) superimposed on a long-term sea-level curve, which shows an overall shallowing trend throughout the Mid and Late Permian.

Acknowledgements. We are grateful to Winfried Dallmann for the compilation of the geological maps of the study area. David Bond at the Norwegian Polar Institute provided important comments and discussion, substantially improving an earlier version of the manuscript.

We are thankful to the organisation and logistics department of the Norwegian Polar Institute for support in the field. Ralf Bätzel, technician at the University of Bremen, is thanked for the excellent preparation of numerous thin-sections forming the basis of the substantial microfacies investigations of this study. Alenka E. Crne and Hans Arne Nakrem are thanked for reviewing an earlier version of this article

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Isotope chemostratigraphy of high-grade marbles in the Rognan area, North-Central Norwegian Caledonides: a new geological map, and tectonostratigraphic and palaeogeographic implications Victor A. Melezhik, David Roberts, Svein Gjelle, Arne Solli, Anthony E. Fallick, Anton B. Kuznetsov, Igor M. Gorokhov

Figure 2. Detailed geological map (1:20,000 scale) of the Rognan area, with three geological profiles; A–A’, B–B’ and C–C’.

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New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark, northern Norway, from trace fossils and organic-walled microfossils Anette E.S. Högström, Sören Jensen, Teodoro Palacios & Jan Ove R. Ebbestad Högström, A.E.S., Jensen, S., Palacios, T. & Ebbestad, J.O.R.: New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark, northern Norway, from trace fossils and organic-walled microfossils. Norwegian Journal of Geology, Vol 93, pp. 95–106. Trondheim 2013, ISSN 029-196X. The Vestertana Group on the Digermul Peninsula, Finnmark, northern Norway, presents one of the few, potentially continuous Ediacaran– Cambrian sections in Scandinavia. Trace fossils provide the main age constraint, with the boundary traditionally placed at the base of the Breidvika Formation. Here, we provide trace-fossil evidence to show that this boundary is at least as low as the third cycle of the Manndraperelva Member, Stáhpogieddi Formation, where Treptichnus pedum is associated with trilobed trace fossils. Organic-walled microfossils from the same stratigraphic interval include Granomarginata prima and the first report from Scandinavia of Cochleatina. The second cycle of the Manndraperelva Member contains trace fossils, including treptichnids and ?Cochlichnus isp. tentatively interpreted as latest Ediacaran. Reports of palaeopascichnids suggest a late Ediacaran age for the first cycle. The age of lower parts of the Stáhpogieddi Formation is poorly constrained but discoidal Ediacara-type fossils, vendotaenids, and possible simple trace fossils, suggest that the middle part of the Innerelva Member is younger than c. 560 Ma. Anette E.S. Högström, Tromsø University Museum, Natural Sciences, 9037 Tromsø, Norway. Sören Jensen, Teodoro Palacios, Área de Paleontología, Facultad de Ciencias, Universidad de Extremadura, 06006 Badajoz. Jan Ove R. Ebbestad, Museum of Evolution, Uppsala University, Norbyvägen 16, 752 36 Uppsala, Sweden.

E-mail corresponding author (Anette Högström): anette.hogstrom@uit.no

Introduction The increase in size and complexity of trace fossils from about 555 Ma to 535 Ma is an important manifestation of the radiation of bilaterian metazoans that took place at this time (e.g., McIlroy & Logan, 1999; Budd & Jensen, 2000; Droser et al., 2002; Erwin et al., 2011).

This evolution of trace fossil morphotypes has been documented globally and is long known to have stratigraphical potential with the recognition of tracefossil-based zones (e.g., Crimes, 1987; Mángano et al., 2012), culminating in the decision to define the basal Cambrian global stratotype section and point (GSSP) on trace fossils (Narbonne et al., 1987; Brasier et al., 1994).

Although the broad patterns in trace-fossil evolution at this time are well established, sections where it can be observed without major breaks in sedimentation or facies changes are not common. The Ediacaran– Cambrian boundary type section on Fortune Head, Newfoundland, largely meets these criteria, with the GSSP defined at the then lowest known occurrence in the Chapel Island Formation of the three-dimensional burrow system Treptichnus pedum (as Phycodes pedum), and more generally at the base of the Treptichnus pedum Ichnozone, with the appearance of plug-shaped burrows

(Bergaueria) and vertical spiral burrows (Gyrolithes) (Narbonne et al., 1987). This contrasts with underlying beds where only morphologically simple trace fossils are found, and with the upper range of the problematic fossils Harlaniella and Palaeopascichnus a few decimetres below the GSSP. The ranges of Cambrian-type traces in the type section have subsequently been extended down-section, so that Treptichnus and Gyrolithes now overlap with Palaeopascichnus (Gehling et al., 2001). In addition, trace fossils similar (but not clearly identical) to Treptichnus pedum occur in strata traditionally considered Ediacaran in Namibia, and possibly elsewhere, where they overlap the stratigraphical range of some Ediacara-type fossils (Jensen et al., 2000). While this raises questions about stratigraphical resolution, it does not challenge the general validity of Ediacaran– Cambrian ichnostratigraphy, and may instead suggest a latest Ediacaran ichnozone with treptichnids (Jensen, 2003). Furthermore, recent revisions of radiometric ages from Namibia (Schmitz, 2012) suggest that some of these ‘Ediacaran’ occurrences of complex burrow systems (Streptichnus narbonnei; Jensen & Runnegar, 2005) are within the error of the currently accepted age of the base of the Cambrian (c. 541 Ma), and others (treptichnids; Jensen et al., 2000) may be only a few million years

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older. Trace fossils therefore remain a powerful tool in Ediacaran–Cambrian transition biostratigraphy in successions dominated by siliciclastic rocks (Buatois et al., 2013).

The Digermul Peninsula, northern Norway (Fig. 1), is probably the only fossiliferous site in Scandinavia with sedimentation across the Ediacaran–Cambrian transition without a significant hiatus. The transition occurs in the Vestertana Group (Reading, 1965; Banks et al., 1971), which is a thick, largely siliciclastic unit. It contains in its upper part glacial diamictites possibly related to the Gaskiers glaciation, followed by mudstone and sandstone with Ediacara-type fossils and, higher still, Cambrian-type trace fossils and skeletal fossils (e.g., Farmer et al., 1992; Crimes & McIlroy, 1999). In an early study on trace-fossil evolution, Banks (1970) documented the initial appearance of small, simple trace fossils and, successively up-section, larger, simple trace fossils followed by branching trace fossils including

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Treptichnus pedum, and higher still the appearance of trace fossils attributed to arthropods (Rusophycus).

Later studies on trace fossils in the Vestertana Group (Farmer et al., 1992; McIlroy, 1995; McIlroy & Logan, 1999) have resulted in greater ichnotaxonomic diversity, but the general trend observed by Banks (1970) has been maintained, and to this date trace fossils provide the main age control on the upper part of the Vestertana Group. The base of the Cambrian has generally been positioned near the base of the Breidvika Formation, at the appearance of trace fossils like Treptichnus pedum and Gyrolithes (Banks, 1970; Farmer et al., 1992), an assemblage that corresponds closely to that of the Treptichnus pedum Ichnozone on Newfoundland.

The appearance of Rusophycus further up-section in the Breidvika Formation indicates the succeeding Rusophycus avalonensis Ichnozone. Many workers have maintained the possibility of a slightly lower position of the Ediacaran–Cambrian boundary in the upper part of the Stáhpogieddi Formation (e.g., Farmer et al., 1992;

Figure 1. (A) Location of the Digermul Peninsula in northern Norway. Also indicated are selected locations in the East European Platform from which Lower Cambrian fossils comparable to those described from the Digermul Peninsula have been reported. 1: Kaplonosy drillcore in eastern Poland. 2: Stradech–17 drillcore in western Belarus. 3: Southern part of Podolia, Ukraine. (B) Map of the Digermul Peninsula, showing main geographical locations mentioned in the text. The stippled rectangle corresponds to an area detailed in Fig. 1C. (C) Map showing geology of the study area (based on Siedlecka et al., 2006) and fossil localities.

Location of measured section (Fig. 2B) south of the mouth of the Bárdelouvttjohka rivulet is indicated by the curved line with terminal bars. Locations of Manndraperelva Member cycle boundaries in coastal outcrop are indicated by vertical bars. Abbreviations: rv – river, rl – rivulet. References: 1 – Farmer et al. (1992), 2 – Crimes & McIlroy (1999).

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Crimes & McIlroy, 1999), based on the relatively large size of some of the trace fossils in that unit. Nielsen & Schovsbo (2011) proposed that an even greater portion of the Stáhpogieddi Formation is Cambrian on the basis of an exhaustive sequence stratigraphy-based analysis of the Cambrian of Scandinavia. In their scheme, two regressive sedimentary cycles recorded at the top of the Stáhpogieddi Formation represent eustatic signals that may correspond to sequences in the Rovno Regional Stage of the Baltic area.

Here, we report on new finds of trace fossils from the upper part of the Stáhpogieddi Formation, including an extended stratigraphical range of Treptichnus pedum. We also present the first data on organic-walled microfossils from the Manndraperelva Member. Organic-walled microfossils have been widely used to characterise the Ediacaran–Cambrian transition in Baltica, and elsewhere (e.g., Moczydłowska, 1991). The new fossil data shed new light on the location of the Ediacaran–Cambrian boundary in the Vestertana Group, and more generally invites additional studies on the age of the lower part of the Stáhpogieddi Formation.

Geological setting The Vestertana Group on the Digermul Peninsula is part of the basal Caledonian thrust sheet, the Gaissa Nappe Complex, formed during an early phase of the Caledonian Orogeny. It is found elsewhere in Finnmark, notably on Varanger Peninsula, but the upper part of the Vestertana Group is nowhere so completely developed as on the Digermul Peninsula. The Vestertana Group was deposited in a shallow basin close to the northern (present coordinates) margin of Baltica, interpreted by Gorokhov et al. (2001), Roberts & Siedlecka (2002) and Nielsen & Schovsbo (2011) as a foreland basin developed ahead of the deformation front of the Timanian Orogeny.

The lithostratigraphical terminology (Fig. 2) stems from Reading (1965) and Banks et al. (1971); the orthography used here follows Siedlecka et al. (2006), with common older spellings also indicated on the first mention below.

The basal two formations of the Vestertana Group, the Late Cryogenian, glaciogenic Smalfjord Formation and the succeeding Nyborg Formation (siliciclastics and minor carbonates), are not developed in the study area and will not be further considered here. The Mortensnes Formation is a glaciogenic diamictite, the base of which is a regional unconformity; this formation has recently been tentatively correlated with the c. 580 Ma Gaskiers glaciation of Newfoundland on the basis of carbon isotope signals in the Nyborg Formation (see Rice et al., 2011). There exists no biostratigraphical age constraint on the Mortensnes Formation as the only fossils are nonage diagnostic and reworked organic-walled microfossils (sphaeromorphic acritarchs and vase-shaped micro­ fossils) (Vidal, 1981; Vidal & Moczydłowska, 1995). The succeeding Stáhpogieddi (Stappogiedde) Formation is

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divided into three members. The Lillevannet (Lillevatn) Member, in abrupt but apparently continuous succession with the Mortensnes Formation, consists of a lower unit of laminated siltstone and sandstone and an upper sandstone-dominated unit. It has been interpreted as a fluvial to shallow-marine deposit. The Innerelva (Innerelv) Member is dominated by laminated mudstone with several coarsening-upward sequences terminating in sandstone. Banks (1973) interpreted this member to have formed on a quiet marine shelf with occasional storm influence. Discoidal Ediacara-type fossils were recorded from the approximate middle of this unit (Farmer et al., 1992) (Fig. 2A). The basal part of the Manndraperelva (Manndrapselva) Member consists of a thick package of reddish sandstone, followed by two coarsening-upward cycles, each starting with mudstone and fine sandstone (greywackes), and terminating in cross-bedded sandstone. This development has been referred to as three cycles within the Manndraperelva Member (Fig. 2A), with the fine-grained sediments interpreted as distal and proximal turbidites, and the sandstone-dominated portions as deposited in a shallowmarine environment (Banks et al., 1971). The Breidvika (Breivik) Formation is divided into two informal members: the Lower Breidvika member consists of alternations of mudstone, siltstone and sandstone interpreted as a shallow-water marine deposit, and the Upper Breidvika member, dominated by mudstone, has been interpreted as an offshore shelf deposit. Compared to the Manndraperelva Member, the Lower Breidvika member shows a more rapid alternation of sediment types, and there is a greater proportion of green-coloured sediments.

The sedimentary succession from the Innerelva Member to the Lower Breidvika member has been interpreted as recording several episodes of relative shallowing but to date without clear evidence for any substantial break in sedimentation (e.g., Banks et al., 1971).

Review of earlier reports of fossils from the Stáhpogieddi and Breidvika formations In the absence of reliable radiometric dates, fossils provide critical age constraints on the upper part of the Vestertana Group. In order to place in context the new fossil data reported here from the Manndraperelva Member, the principal published reports of fossils from the Stáhpogieddi and Breidvika formations will be briefly reviewed.

Innerelva Member. Farmer et al. (1992) reported discoidal Ediacara-type fossils as Nimbia sp.? in float from the lower part of the Innerelva Member, and Ediacara? sp., Beltanella sp., Hiemalora sp. and Nimbia? sp. from outcrops in the middle part of the member.

If a different approach to the naming of discoidal

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Figure 2. (A) Log showing a schematic representation of lithology and selected fossil occurrences in the Vestertana Group on the Digermul Peninsula. In general, only first occurrences of fossils are indicated. Fossil symbols in parenthesis refer to occurrences in the Vestertana Group outside of the Digermul Peninsula, placed in the approximate corresponding stratigraphical level. (B) Lithological log of upper part (part of third cycle) of the Manndraperelva Member in the coastal section south of the mouth of Bรกrdelouvttjohka rivulet.

Ediacara-type fossils is used, in which various discoidal taxa are considered preservational morphs (Gehling et al., 2000), the Finnmark material may, with the exception of Hiemalora, all belong to Aspidella (Narbonne in Vickers-Rich et al., 2007). Additional material illustrated here from the main fossil locality in the middle of the member includes a large disc-shaped Aspidella (Fig. 3A), and a twinned specimen (Fig. 3C) of a type that has been variously interpreted as asexual reproduction or as merged discs.

Banks (1970, 1973) recorded small simple trace fossils from the Innerelva Member. Banks (1973) reported, but did not illustrate, simple burrows to be relatively abundant at Kvalneset in the southeastern part of the Varanger Peninsula. Farmer et al. (1992) studied material from this locality and considered these structures to be better explained as pseudofossils formed by fluid escape within the sediment. This interpretation seems convincing but may not apply to material that Banks (1970, pl. 1a) illustrated from the Digermul Peninsula

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Figure 3. Fossils from the upper part of the Vestertana Group. (A–C) Discoidal Ediacara-type fossils from the middle part of the Innerelva Member in coastal outcrop north of the mouth of the Manndrapselva river. Scale bars are 20 mm. (A) Large Aspidella (TSGF 18239), about 8 cm wide, with two clearly developed discs and a faint central tubercle. (B) Discoidal form (TSGF 18240) with two well-developed concentric rings and several more faintly developed ones. (C) Two closely positioned (merged?) discoidal forms (TSGF 18241). (D) Small bow-shaped trace fossils (TSGF 18242) from the lower part of the second cycle, Manndrapereleva Member, from a section along a small rivulet north of the mouth of the Manndrapselva river. Scale bar is 10 mm. (E–F) Acritarchs (TSGF 18243 & 18244) from the approximate middle part of the Lower Breidvika member, along the Bárdelouvttjohka rivulet. Scale bar is 20 µm.

about 150 m above the base of the Innerelva Member.

These are preserved on bed bases as short sand-filled knobs 1–2 mm in diameter, not on bed tops and bed bases as is the material from Kvalneset described by Farmer et al. (1992, fig 6). Banks (1970) interpreted these knobs as portions of passively filled vertical burrows, and they have been compared with both Skolithos and Arenicolites (Crimes, 1987, p. 104). Short, sand-filled knobs were observed during this study on surfaces containing discoidal fossils (Fig. 3). Some of these appear to have been spherical (Fig. 3A) and may represent mineral replacements. Others (Fig. 3B, C) have at least a superficial similarity to short knobs that represent partially preserved trace fossils from the Mandraperelva Member (Fig. 3D). Finally, one such structure is located within an Aspidella specimen, and may represent a central tubercle (Fig. 3A). Future studies are necessary to shed new light on the presence of trace fossils in the Innerelva Member.

Vidal (1981) reported sphaeromorphic acritarchs and vendotaenids (as Vendotaenia sp. and Vendotaenia cf. antiqua) from the Innerelva Member on the Varanger Peninsula. Awaiting description are ‘probably agglutinated tubular fossils’ reported to occur well below the discoidal fossils (Vidal & Moczydłowska, 1995, p.

206).

Manndraperelva Member. Banks (1970) mentioned poorly preserved trace fossils from the basal part of the Manndraperelva Member that he compared with structures from the Late Ediacaran of South Australia now assigned to Palaeopascichnus. Antcliffe et al. (2011) provided an affirmative identification of Palaeopascichnus delicatus in the Manndraperelva

Member without additional detail. Palaeopascichnus is a geographically widespread form, now generally interpreted as a body fossil of uncertain and debated affinity (Jensen, 2003; Antcliffe et al., 2011). From the second cycle of the Manndraperelva Member, Banks (1970) identified regularly sinuous horizontal burrows (i.e., Cochlichnus), horizontal spirals as cf., Helicolithus, and simple vertical U-tubes (i.e., Arenicolites), and also remarked on tool marks that he suggested were formed by the exoskeleton of an arthropod. Banks (1970) recorded larger-sized trace fossils from the third cycle of the Manndraperelva Member, including forms with lateral grooves (i.e., ‘Curvolithus’). Trace fossils from the second and third cycles of the Manndraperelva Member are discussed in more detail below.

Lower Breidvika member. The Lower Breidvika member has yielded a greater diversity of trace fossils. Banks (1970) reported Treptichnus pedum from the basal three metres, and Rusophycus to first appear about 70 m above the base of the formation. Farmer et al. (1992) recorded the tubular fossil Sabellidites sp. (reported as Sabellidites cambriensis in Vidal & Moczydłowska, 1995 and Moczydłowska, 2002) from the Lower Breidvika member and also mentioned the presence of the vertical spiral trace fossil Gyrolithes. McIlroy (1995) listed 27 ichnotaxa from this member. The discoidal fossils Tirasiana disciformis and Nimbia occlusa, taxa otherwise known exclusively from the Late Ediacaran, have also been reported from the lower part of the Lower Breidvika member (Crimes & McIlroy, 1999). Trace fossils from the Upper Breidvika member include Teichichnus.

McIlroy et al. (2001) reported the probable foramini­ ferans Platysolenites antiquissimus and Platysolenites

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cooperi from near the base of the Upper Breidvika member on the Digermul Peninsula. The lowest occurrence of Platysolenites antiquissimus in the Vestertana Group may, however, be at a somewhat lower level. Føyn (1967, pp. 31, 32), Hamar (1967) and Føyn et al. (1983) recorded Platysolenites antiquissimus together with the spirally coiled Spirosolenites spiralis at Kunes in the Laksefjord area, Finnmark, from beds that they estimated to be some 100 to 150 m above the base of the Breidvika Formation. According to Føyn et al. (1983) only the Lower Breidvika member is present in this area. Correlation between the Laksefjord and Digermul successions is complicated by a significantly thinner development of the Stáhpogieddi Formation in the Laksefjord area, but a lower stratigraphical position of the Laksefjord Platysolenites is likely, and is indicated in Fig. 2A. The only other skeletal fossil reported from beds close to the transition between the lower and upper members are tubular fossils cited as Circotheca cf. annae (McIlroy, 1995) or Ladatheca cylindrica (McIlroy et al., 2001).

Vidal (1981) reported sphaeromorphic acritarchs and poorly preserved specimens of the acanthomorph Asteridium sp. (as Micrhystridum sp.) with vesicles 18–22 µm and processes 1.5–2 µm in length from the Lower Breidvika member in the Laksefjord area in samples containing Platysolenites antiquissimus.

Moczydłowska (2002, p. 204) listed a greater diversity of acritarchs from an unspecified level within the Lower Breidvika member at Manndraperelva, with Asteridium tornatum, Granomarginata squamacea, Lophosphaerdium tentativum, and Tasmanites tenellus, indicating the Asteridium–Comasphaeridium Zone.

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New fossil data from the Manndrape­ relva and Lower Breidvika members Sections studied and material This study is largely based on a coastal outcrop south of the mouth of the Bárdelouvttjohka rivulet (Fig. 1C), where the transition from the Manndraperelva Member to the Lower Breidvika member is well exposed (Fig. 2B). The lower part of the section consists of thick packages of fine-grained greywacke sandstone capped by mudstone (Fig. 4A) and thinner beds of coarser quartz arenite (Fig. 4B). There is an estimated additional 4–5 m of the third cycle below the measured section that could not be accessed. Reading (1965) defined the top of the Manndraperelva Member in the Manndrapselva river section in a 12–15 m-thick package of red quartzitic sandstone. A comparable thickness of red-coloured sandstone is present also in the Bárdelouvttjohka section, where it is followed by grey quartz arenite that Farmer et al. (1992) took to form the base of the Breidvika Formation. Locally, chaotic disruption of beds is observed in the upper part of the Manndraperelva Member in this section (Fig. 4C). Additional observations were made on the lower part of the Manndraperelva Member along coastal outcrops north of the mouth of the Manndrapselva river (Fig. 1).

Stratigraphic positions of trace fossils and samples collected for organic-walled microfossils are indicated in Fig. 2B. The designation Lower Cambrian is here used to encompass the Terreneuvian Series and the provisional Cambrian Series 2 of the recently proposed four-fold division of the Cambrian System. Illustrated material is stored with the palaeontological collection (TSGF) of Tromsø University Museum, Tromsø, Norway. Location details for sections and illustrated material is provided in Electronic Supplement 1.

Figure 4. Field images of the upper part (third cycle) of the Manndraperelva Member in coastal outcrop south of the mouth of the Bárdelouvttjohka rivulet. (A) View corresponding to c. 5–10 m in the log in Fig. 2B. The sample level 10 and in situ find of Treptichnus pedum are both located somewhat below shoulder height of the seated person. (B) Interval at 9 m in the log in Fig. 2B, with fine-grained sandstone grading into mudstone that is sharply overlain by medium-grained sandstone. Sample bag (c. 7 cm wide) is positioned within sample level 10. Treptichnus pedum, illustrated in Fig. 5D, originates from the thin sandstone ledge in the middle of the image. (C) View of interval with chaotic bedding at about 19 m in Fig. 2B.

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Trace fossils The lowest unquestionable trace fossils were observed in coastal outcrops (Fig. 1C) within the first two metres of the second cycle of the Manndraperelva Member in the form of a series of aligned sediment pods that are connected by a faintly visible horizontal ridge (Fig. 5A). A trace fossil from scree material west of Avžejohka rivulet, attributed to higher levels of the second cycle, consists of a series of circular knobs preserved on a bedding-plane (Fig. 5B) and represents a more distinctly developed form of the same type of trace fossil. This type of trace fossil is generally interpreted as sections through vertical shafts of a three-dimensional burrow system that consisted of a basal connecting tube, from which emerged vertical probes. For example, Geyer & Uchman (1995) identified trace fossils from the Lower Cambrian of Namibia with a morphology identical to that of the specimen in Fig. 5A as Treptichnus pollardi. This type of trace fossil has also been reported as Hormosiroidea, Saerichnites and Ctenopholeus (see Fürsich et al., 2006, for discussion).

McIlroy’s (1995) report of Hormosiroidea from the Stáhpogieddi Formation probably refers to a trace fossil of this type. Trace fossils of this type can also be included in a broadly defined Treptichnus pedum sensu Seilacher (2007). Here they are referred to as treptichnids.

A variety of small trace fossils were observed on a bed sole with prominent tool marks about 2.5 m into the second cycle. This bed can be traced laterally for many tens of metres and is probably the bed illustrated by Banks (1970, pl. 2b). On this surface, bow-shaped trace fossils are found (Fig. 5C), up to 4 mm wide and several centimetres long. These can be variously considered to be Arenicolites or short Palaeophycus/Planolites-type trace fossils. There are also a variety of smaller trace fossils, including horizontal spirally coiled burrows (cf., Helicolithus of Banks, 1970). Small forms of the same type of bow-shaped trace fossils were also observed in a section west of Avžejohka (Fig. 3D). Their interpretation is discussed below.

Unequivocal examples of Treptichnus pedum (Fig. 5D) were found in the third cycle, about 20 m below the top of the Manndraperelva Member in the coastal section south of the Bárdelouvttjohka rivulet. Probable additional examples of Treptichnus pedum were observed in vertical sections in the same succession but identification could not be confirmed because of lack of bed-sole exposure.

The same interval contains trace fossils with a trilobed basal surface (Fig. 5E, F) and poorly preserved, plugshaped trace fossils (cf., Bergaueria). The trilobed trace fossils present naming problems (Jensen & Grant 1998); Systra & Jensen (2006) proposed the term ‘Bure ichnocomplex’ for trace fossils that are united by a trilobed basal surface, but which may be straight and band-like, or form a succession of short segments that alternate in a regular manner. As such, they arguably span more than one ichnogenera and have been variously included in Curvolithus and Treptichnus. This informal

term was introduced to emphasise that these various forms likely had a common producer. Banks (1970, pl.

2d) illustrated examples of this type of trace fossil with a wide central lobe flanked by narrow marginal lobes from the third cycle, about 180 m above the base of the Manndraperelva Member. The material illustrated here all has proportionally wider lateral lobes, and compares more closely to better preserved material from the Lower Breidvika member (Fig. 5G, H).

The first Rusophycus (Fig. 5I) was observed in the Lower Breidvika member at a distance above the base of the Breidvika Formation closely comparable with that cited

by Banks (1970). These are large forms with coarse scratch marks, and, as such, typical of Rusophycus from the Terreneuvian.

Organic-walled microfossils Samples from the third cycle of the Manndraperelva Member in the section south of the mouth of the Bárdelouvttjohka rivulet yielded leiosphaerids, Granomarginata prima (Fig. 6C), horn-shaped Cerato­ phyton and various filaments, including forms identical to Eoschizotrix composita as recorded by Moczydłowska (2008) from the Włodawa Formation of the Lublin Slope, Poland. Spirally coiled filaments with serrated margins (Fig. 6A, B) are here interpreted as the first reports from northern Scandinavia of Cochleatina. Well-preserved material of Cochleatina from the East European Platform have been described as chitinous flat ribbons of unknown affinity, typically arranged in a logarithmic spiral in which the width of the ribbon increases outwards (Burzin, 1996). Vidal (1981, p. 39) described a spirally coiled, organic-walled microfossil from the more than c. 650 Ma Dakkovarre Formation at Skallneset on the southeastern coast of the Varanger Peninsula, where it was found with various sphaeromorphic microfossils.

The illustrated specimen is an opaque band about 2 µm wide forming a spiral that expands from about 20 µm to 40 µm. The Dakkovarre Formation specimen lacks the pronounced teeth seen in the Manndraperelva material and appears to be different. Species of Cochleatina are differentiated on the nature of vertical zonation within the ribbon (Burzin, 1996), not readily observed in the Manndraperelva Member material. However, Burzin (1996) observed that ribbon width decreased through time, so that basal Cambrian C. rudaminica, from the upper part of the Rovno Regional Stage and the lower part of the Lontova Regional Stage in Latvia, Lithuania and Belarus (Paškevičenė, 1980), is 3.5 times narrower than C. canilovica, a species reported throughout the Late Ediacaran Kanilovka Group of Podolia, and the basal part of the Rovno Regional Stage in Volyn (Burzin, 1996).

In the best preserved specimen from the Manndraperelva Member (Fig. 6A), the narrowest part of the ribbon is about 3 µm wide, expanding to slightly less than 10 µm at the fourth whorl, where the spiral has a diameter of about 40 µm. Paškevičenė (1980) reported ribbon widths

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Figure 5. Trace fossils from the second (A–C), and third (D–F) cycles of the Manndraperelva member, and the Lower Breidvika member (G–I). (A) Trace fossil consisting of aligned sediment pods on bed sole photographed in the field in a coastal outcrop. Scale bar is 10 mm. (B) Treptichnid (TSGF 18245) in scree material from the second cycle of the Manndraperelva Member, north of the mouth of the Manndrapselva river. Scale bar is 10 mm. (C) Base of bed c. 2.5 m above base of second cycle in coastal outcropc with several shallow, U-shaped trace fossils (TSGF 18246) and, near the lower part of image, a clear example of a spirally coiled trace fossil (?Cochlichnus isp). Scale bar is 10 mm. (D) Field photo of Treptichnus pedum (TSGF 18247). See Figs 2B and 4 for details. Scale bar is 10 mm. (E, F) Trilobed trace fossils in coastal section south of the mouth of the Bárdelouvttjohka rivulet. Scale bars are 10 mm. (G) Trilobed trace fossils in coastal section north of the mouth of the Bárdelouvttjohka rivulet. Scale bar is 10 mm. (H) Trilobed trace fossil (TSGF 18248) and Gyrolithes in coastal section north of the mouth of the Bárdelouvttjohka rivulet. Scale bar is 10 mm. (I) Rusophycus isp. in coastal section north of the mouth of the Bárdelouvttjohka rivulet. Scale bar is 30 mm.

NORWEGIAN JOURNAL OF GEOLOGY New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark 103

Figure 6. Organic-walled microfossils from the third cycle of the Manndraperelva Member in coastal section south of the mouth of the Bárdelouvttjohka rivulet. Scale bars represent 20 µm. For each slide, number and England finder coordinates are given. (A) Cochleatina sp. Sample 10, 10–2, E–45–3 (TSGF 18249). (B) Large fragment of Cochleatina sp. Sample 8, 8–1n, B–24–4 (TSGF 18250). (C) Granomarginata prima.

Sample 10, 10–2, O–41–2 (TSGF 18251). (D) Nano-scale filaments, probably representing a degraded sabelliditid. Sample 10, 10–1, P–25–3 (TSGF 18252).

in C. rudaminica and C. ignalinica to be 9–13 µm and 10–15 µm, respectively, whereas ribbons of C. canilovica are 15–20 µm wide at the inner end, expanding to 70 µm at the fourth coil (Burzin, 1996). Fragments of larger specimens were also found, with ribbons about 10 µm (Fig. 6B). In terms of size, the Manndraperelva Member specimens compare to the younger occurrences of Cochleatina from the East European Platform.

Samples from the basal part of the Lower Breidvika member, to the level of the first Rusophycus, in coastal outcrops north of the mouth of the Bárdelouvttjohka rivulet yielded poorly preserved Granomarginata prima, Ceratophyton sp., leiosphaerids and filaments. A sample from the middle portion of the Lower Breidvika member, from an outcrop along the Bárdelouvttjohka rivulet, yielded poorly preserved process-bearing acritarchs comparable to Polygonium (or Goniosphaeridium) varium (Fig. 3E, F), although with unusually short processes.

Sabelliditids Farmer et al. (1992) recorded the tubular fossil Sabellidites sp. from 9 and 16 m above the base of the Breidvika Formation in the Bárdelouvttjohka section.

In the present study, specimens of Sabellidites sp. were observed on additional bed surfaces within the same part of the section over an interval of 18 m (Fig. 2B). Samples prepared for organic-walled microfossils from the lower part of the sampled interval contain masses of very fine filaments (Fig. 6D). Similar masses were recovered from the Lower Cambrian Ratcliffe Brook Formation of New Brunswick, eastern Canada, where they were interpreted as possible degraded sabelliditids (Palacios et al., 2011).

Discussion Coastal exposures stretching from the mouth of the Manndrapselva river to just south of the Breidvika bay provide essentially uninterrupted exposure from the Ediacara fossil-bearing Innerelv Member to Cambrian fossils in the Breidvika Formation. Although there is sedimentological evidence for several episodes of relative shallowing, no indication of any substantial break in sedimentation has been identified. The present study extends the lowest occurrence of Treptichnus pedum in the Vestertana Group from the Lower Breidvik member to the upper part of the Manndraperelva Member, where it is found with trilobed traces and poorly preserved Bergaueria. This association of trace fossils is best attributed to the Treptichnus pedum Ichnozone, indicating that all or most of the third cycle of the Manndraperelva Member is Cambrian.

The trace-fossil assemblage from the basal part of the Breidvika Formation compares closely to those described from the basal Cambrian Khmelnitsky Formation of Ukraine, which includes Treptichnus pedum, Gyrolithes polonicus, and trace fossils of the Bure ichnocomplex (Treptichnus triplex, Curvolithus) (e.g., Palij et al., 1983).

The only trace fossils reported from the underlying Okunets Formation are Planolites and ‘Curvolithus’ (e.g., Gureev, 1988), although Kiryanov (2006) extended the Treptichnus pedum Ichnozone to the base of the Okunets Formation. On the basis of trace fossils, the third cycle of the Manndraperelva Member may correspond in time with either the Khmelnitsky or the Okunets formations.

The only body fossil recorded from the Khmelnitsky Formation is Sabellidites cambriensis, although it also contains problematic discoidal fossils as well as the scratch circle Kullingia, also found in the Lower Cambrian Dividalen Group of northern Sweden (Jensen & Grant, 1998; Jensen et al., 2002). The organic-walled microfossils from the third cycle of the Manndraperelva

104 A.E.S. Högström et al.

Member can be compared with associations reported from the East European Platform. In many areas of the East European Platform, the first appearance of Granomarginata prima is in beds attributed to the lower part of the Lontova Regional Stage, or the upper part of the Rovno Regional Stage. For example, in the Stradech–17 core of western Belarus, Granomarginata prima and Ceratophyton are found with Cochleatina rudaminica and C. ignalinica in the lower part of the Stradech Formation, in beds attributed to the lower part of Lontova Regional Stage, a short distance below the first Platysolenites antiquissimus (Paškevičenė, 1980).

Although the Manndraperelva Member Cochleatina compare in size to the younger occurrences of this genus from the East European Platform, a better understanding of its stratigraphical implications for the Vestertana Group must await a fuller sampling of the Stáhpogieddi Formation for organic-walled microfossils. In the Lublin slope, eastern Poland, Moczydłowska (1991) equated the Asteridium–Comasphaeridium Zone with the upper part of the Rovno Regional Stage and the Lontova Regional Stage, and proposed the Kaplonosy core as a regional reference section for the Ediacaran–Cambrian transition. Direct comparison of the Manndraperelva assemblage with the Asteridium–Comasphaeridium Zone is complicated by the absence of species of Asteridium and Comasphaeridium, and largely hinges on the first appearance of Granomarginata prima in the basal Cambrian Asteridium–Comasphaeridium Zone.

Cochleatina has not been reported from the Polish part of the East European Platform, but a fragmentary fossil from the Asteridium–Comasphaeridium Zone in the Kaplonosy core (Moczydłowska, 1991, p. 12, fig. 15D) could represent this genus. To date, a more diverse assemblage of taxa indicative of the Asteridium– Comasphaeridium Zone appears only in higher levels of the Lower Breidvika member.

Trace fossils comparable to the treptichnids reported here from the basal part of the second cycle of the Manndraperelva Member have a long stratigraphical range but specimens from Namibia are older than c. 543 ± 2.5 Ma and younger than c. 547 ± 0.65 Ma (Jensen et al., 2000; ages from Schmitz, 2012). Trace fossils comparable to Treptichnus (but not T. pedum) also occur in the GSSP section, Newfoundland, in the upper part of Member 1 of the Chapel Island Formation (Gehling et al., 2001). Small, horizontally coiled trace fossils from about the same level as these treptichnids on the Digermul Peninsula, previously compared to Helicolithus, are of interest as this would be the earliest record of this ichnogenus. The Late Ediacaran Harlaniella had been interpreted as a horizontal, spirally coiled trace fossil, but Jensen (2003) found that the morphology of Harlaniella does not match the expected geometry for a spiral and that it may be closer to Palaeopascichnus. Streptichnus narbonnei from beds of either Ediacaran or Cambrian age in Namibia (Jensen & Runnegar, 2005) appears to be, at least in part, spirally coiled but correspond to a more

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complex burrow system. A closer comparison of the Manndraperelva spirals can probably be made with the sinusoidal Cochlichnus, a form commonly reported from basal Cambrian and possibly also older beds. Seilacher (2007, p. 96) noted that “In a three-dimensional substrate, the undulation of Cochlichnus can easily switch to a corkscrew motion.” Palij et al. (1983, pl. 61:4) reported a “crawling trace close to Cochlichnus” from Late Ediacaran beds of the White Sea Region, Russia, that is larger but otherwise comparable to the Manndraperelva specimens.

There is, thus, no positive evidence from trace fossils for a Cambrian age for the second cycle of the Manndraperelva Member. Here it is suggested that the association of trace fossils corresponds to a latest Ediacaran trace-fossil zone (Jensen, 2003). Of great importance to further constrain the location of the Ediacaran–Cambrian boundary within the Vestertana Group will be the publication of detailed information on the exact occurrence of Palaeopascichnus within the Manndraperelva Member. In both Newfoundland and Ukraine the upper stratigraphic range of Palaeopascichnus and Harlaniella is close to that of Cambrian-type trace fossils. In Ukraine, Palaeopascichnus has a long stratigraphic range with the youngest occurrences in the Studenitska Formation (Fedonkin, 1983). In Newfoundland, Palaeopascichnus ranges from the upper part of the Fermeuse Formation (c. 560 Ma) to within a few decimetres of the basal Cambrian GSSP in the Chapel Island Formation (Narbonne et al., 1987; Gehling et al., 2001). As discussed above, the available information on the location of the Manndraperelva Palaeopascichnus supports an Ediacaran age for the first cycle.

Discoidal Ediacara-type fossils of the type found in the middle (and lower) portion of the Innerelva Member have long stratigraphical ranges. In Ukraine and in the White Sea Region, Russia, similar forms are known from beds about 555 Myr old (e.g., Grazhdankin et al., 2011), but both Aspidella- and Hiemalora-type fossils extend from c. 580 Ma. Narbonne (2005) included the Innerelva Member fossils in his category of Fermeuse-style preservation, with associations formed in slope or outer shelf environments where only bases of hold-fasts are preserved, reflecting a taphonomic bias rather than a true picture of the benthic assemblage (see also Laflamme et al., 2011). The absence of core Ediacara-type taxa such as Dickinsonia and Tribrachidium, that would indicate the younger portion of the stratigraphical range of Ediacaratype fossils, may therefore be taphonomic. Simple trace fossils (Planolites isp.) would indicate an age younger than about 560 Ma, but their presence in the Innerelv Member requires confirmation. In Ukraine, Vendotaenia antiqua ranges from the Nagoryany Formation (<553 Ma) into the Okunets Formation. Although tentative, the fossils from the middle part of the Innerelva Member suggest that this portion of the Stáhpogieddi Formation is younger than c. 560 Ma. If the Mortensnes Formation is related to the Gaskiers glaciation (c. 580 Ma), the c.

200 m comprising the Lillevannet Member and lower

NORWEGIAN JOURNAL OF GEOLOGY New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark 105

part of the Innerelva Member represent more than 20 million years, raising the possibility of some hitherto unrecognised break in sedimentation, as previously indicated by Nielsen & Schovsbo (2011, p. 288).

Conclusions New records of trace fossils, including Treptichnus pedum, and organic-walled microfossils from the upper part (third cycle) of the Manndraperelva Member, Stáhpogieddi Formation, show that this portion of the Vestertana Group is Cambrian. The middle part (second cycle) of the Manndraperelva Member contains trace fossils consistent with a latest Ediacaran age, which may in the future be corroborated when more detailed information on the stratigraphical level of earlier reports of palaeopascichnids in this section becomes available. The improved precision in the location of the Ediacaran–Cambrian boundary reported here, in addition to Ediacara-type fossils further down-section in the Stáhpogieddi Formation, and diverse Cambriantype trace fossils higher in the Breidvika Formation, demonstrate that coastal exposures of the Vestertana Group on the Digermul Peninsula provide a rare case in which metazoan evolution can be observed in a continuous outcrop.

Acknowledgements. The authors wish to thank Arne Thorshøj Nielsen for his most constructive review. AESH acknowledges support from Tromsø University Museum towards costs of fieldwork. SJ and TP acknowledge funding from the Spanish Ministry of Science and Innovation, through grant CGL–2008–04373 and CGL-2012-37237 (both co-financed by FEDER). The Museum of Evolution, Uppsala University, supported fieldwork for JORE. We are grateful to Trygve Larsen, Sjursjok, for logistic assistance.

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 107

Isotope chemostratigraphy of high-grade marbles in the Rognan area, North-Central Norwegian Caledonides: a new geological map, and tectonostratigraphic and palaeogeographic implications Victor A. Melezhik, David Roberts, Svein Gjelle, Arne Solli, Anthony E. Fallick, Anton B. Kuznetsov, Igor M. Gorokhov Melezhik, V.A., Roberts, D., Gjelle, S., Solli, A., Fallick, A.E., Kusnetzov, A.B, &. Gorokhov, I.M.: Isotope chemostratigraphy of high-grade marbles in the Rognan area, North-Central Norwegian Caledonides: a new geological map, and tectonostratigraphic and palaeogeographic implications. ­Norwegian Journal of Geology, Vol 93, pp. 107–139. Trondheim 2013, ISSN 029-196X. Carbon and strontium isotope chemostratigraphy (255 δ13Ccarb and δ18O, and 130 87Sr/86Sr analyses of carbonate components in whole-rock samples) in combination with detailed mapping at 1:20,000 scale in part of the Fauske Nappe in the Rognan area, Nordland, have led to a major revision of what had previously been considered to represent a continuous lithostratigraphy spanning over 100 million years of geological time: from Cambrian to Silurian. This new work has indicated that the high-grade, barren, marble and siliciclastic succession occurs in a series of thrust sheets which have diverse apparent depositional ages ranging from Early Cryogenian to Early Silurian. The new data support earlier interpretations that the rocks of the Fauske Nappe were deposited along the carbonate platform and adjacent continental slope of the eastern margin of Laurentia during the Neoproterozoic to Early Palaeozoic time interval. Most of the thrust sheets were generated in Early Ordovician time during the earliest stages of the main phase of the Taconian orogeny, and then immediately overlain unconformably by a carbonate breccia and conglomerate unit, the Øynes formation, also of Early Ordovician age. Just one formation (Rognan formation) in the Fauske Nappe is younger, of Early Silurian age, with Sr- and C-isotopic data that are comparable to those in fossiliferous, Early Silurian, metalimestones farther north in Troms. These particular carbonate rocks are considered to have accumulated in a post-Taconian successor basin, prior to their transport during the Scandian orogeny into the Uppermost Allochthon of the Scandinavian Caledonides.

Victor A. Melezhik, David Roberts, Svein Gjelle, Arne Solli, Geological Survey of Norway, Postbox 6315 Sluppen, 7491, Trondheim, Norway. Anthony E. Fallick, Scottish Universities Environmental Research Centre, Rankine Avenue, East Kilbride, Scotland. G75 0QF. Anton B. Kuznetsov, Igor M. Gorokhov, Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences, Makarova 2, 199034 St. Petersburg, Russia.

E-mail corresponding author (Victor A. Melezhik): victor.melezhik@ngu.no

Introduction The evolution of the Norwegian Caledonide orogen was preceded by a series of Archaean and Proterozoic crustforming events including creation of the palaeocontinent Baltica and fringing seas and oceans, later collision of Baltica with other plates, and a chain of contractional and extensional tectonic events. The history of Baltica, spanning c. 850–400 Ma, has been marked by a sequence of magmatic, biological and climatic events including several inferred Snowball episodes between c. 760 and 560 Ma (Hoffman & Schrag, 2002).

The timing of magmatic and metamorphic events has been reasonably well constrained. However, a large part of the detailed history of the Caledonides imprinted in metamorphosed sedimentary complexes has yet to be deciphered and fully understood. Several marble formations that were originally deposited in

seas apparently located between ancient continents represent missing links relative to shelf areas that once fringed Baltica and Laurentia. Assessing their age and provenance is crucial for a better understanding of the palaeogeographic evolution of Baltica and its interaction with Laurentia and Siberia.

Establishing the timing of various surface processes and basinal, biological and climatic events is seriously hampered by a scarcity of fossils and, until fairly recently, a general lack of methods which can be employed for the dating of sedimentary rocks. During the last decade, however, a concept of ‘carbon and strontium isotope stratigraphy’ has been employed for indirect dating of sedimentation of the carbonate protoliths of high-grade metamorphic terranes in the Scandinavian (Melezhik et al., 2001a, b, 2002a, b, 2005a; Slagstad et al., 2006), Scottish and Irish (Thomas et al., 2004; Prave et al., 2009a, b) Caledonides, in the Pan-African Mozambique

108 V.A. Melezhik et al

Belt (Melezhik et al., 2006, 2008a), and in East Antarctica (Satish-Kumar et al., 2008; Otsuji et al., 2013).

Previous chemostratigraphic studies in the Norwegian Caledonides have yielded results of fundamental significance including the recognition of: (i) new agegroup rocks (Melezhik et al., 2002a); (ii) a Laurentian ancestry for several marble formations (Melezhik et al., 2002b; Roberts et al., 2002); (iii) a much more complex tectonic development and an additional orogenic event in the history of the Norwegian Caledonides (Roberts et al., 2001, 2002; Melezhik et al., 2002a, b); and (iv) some links between the tectonostratigraphy and commercial marble deposits (Melezhik et al., 2002a, b). A robust test of the applicability of isotope stratigraphy for indirect dating, stratigraphic subdivision, correlation and reconstruction of the geological history of barren, high-grade marble complexes can only be achieved by a comparative study

NORWEGIAN JOURNAL OF GEOLOGY

of several areas that are known to show similar geological and tectonic developments. In 2003, for example, a chemostratigraphic investigation together with 1:20,000 scale geological mapping in the Ofotfjorden area proved to be successful in distinguishing different age groups of non-fossiliferous marble units and defining cryptic thrust-faults, and was considered to represent a novel approach for reconstructing the geological history and for exploration strategy in a high-grade marbledominated terrane (Melezhik et al., 2003).

The current contribution presents a new 1:20,000 map of the Rognan area (Figs. 1 & 2) and has several objectives: (i) to discuss the probable limitations of strontium isotope stratigraphy; (ii) to demonstrate the applicability of carbon and strontium isotope stratigraphy for constraining apparent depositional ages, and for subdividing and correlating barren, highgrade marble successions; (iii) to illustrate our further progress in the production of a new generation of maps of non-fossiliferous, high-grade metamorphic terranes by combining detailed mapping with carbon and strontium isotope chemostratigraphy; and (iv) to utilise chemostratigraphy in palaeotectonic and palaeogeographic reconstructions.

Regional geology, metamorphism, deformation and isotopic ages Geographically (Fig. 1), the studied and mapped area represents the western part of the 1:50,000 Rognan map-sheet 2129 III (Kollung & Gustavson, 1995), and the easternmost part of the 1:50,000 Misvær map-sheet 2029 II (Solli et al., 1992). The northern boundary of the mapped area is limited by Skjerstadfjorden and its eastern limit is defined by Saltdalsfjorden and the Saltelva valley (Fig. 2, fold-out map in back pocket).

Figure 1. Simplified outline map of the Rognan–Fauske district showing the extent of the Fauske Nappe and adjacent nappe complexes.

Areas in red represent granite and/or diorite plutons, in some places with minor gabbro and ultramafic rocks. The boxed area west of Rognan is the study area, shown in detail in Figure 2.

Figure 2 – enclosed on the inside of the back cover.

Detailed geological map (1:20,000 scale) of the Rognan area, with three geological profiles; A–A’, B–B’ and C–C’.

In terms of tectonostratigraphy, the rocks of the studied area constitute part of the Fauske Nappe. The term was introduced by Nicholson (1974, p. 184) for the mediumgrade, marble-rich successions lying structurally above the high-grade Gasak Nappe of the Upper Allochthon, and below rocks of the Rödingsfjället Nappe Complex (including the Beiarn Nappe) which are part of the Uppermost Allochthon (e.g., Roberts & Gee, 1985). The designation Fauske Nappe has been retained in most later map compilations (Solli et al., 1992; Gustavson et al., 1999; Kollung & Gustavson, 1995; Gustavson, 1996), and it is used in this contribution.

The Fauske Nappe includes several, formally-defined, lithostratigraphic units with the rank of formations.

Three such formations, namely the Rognan, Øynes and Kjerktinden, have been adopted in the current contribution, although all are used informally. There are two reasons for this informal usage. Firstly, their boundaries and lithological contents remain poorly

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defined, and secondly, the content of the originally defined Rognan Formation has been changed by subdividing it into three discrete units with different apparent ages (for an explanation, see section Apparent depositional ages of carbonate formations in the Rognan area), namely the Fjellengflåget and Leivset units, whilst for the third unit we will retain the name Rognan.

Hence, five informal formations are now recognised within the geographic limits of the study area - Rognan, Øynes, Fjellengflåget, Leivset and Kjerktinden (Fig. 2).

The Rognan, Leivset and Kjerktinden formations are composed mainly of calcite and dolomite marbles with minor mica, quartz and calcareous schists. In addition, graphite-bearing schists are present in the Rognan and Kjerktinden formations. In contrast, garnet-mica schists comprise the bulk of the Fjellengflåget formation where dolomite, calcite marbles and quartzites are minor lithologies. The Øynes formation is composed of polymict conglomerates and calcareous schists.

No fossils have been found in the carbonate formations, and no radiometric ages have been reported from the study area, hence the actual depositional ages of the studied units are unknown. Earlier workers (e.g., Vogt, 1927) considered that some of the Fauske Nappe marbles (the Fauske limestone of Vogt (1927), Strand (1972)); the Fauske marble of Rutland & Nicholson (1965); the Fauske conglomerate of Melezhik et al., 2000; part of the Øynes formation in the current contribution) could be extended southwards and then eastwards along the margin of the Nasafjellet tectonic window into the supposedly Middle Ordovician Pieske limestone of Sweden (Kulling, 1972, p.

263). Later mapping has not supported this interpretation.

Strand (1972) suggested another long-distance correlation of the Fauske limestone and its structurally overlying Øynes conglomerate with the Evenes limestone and Evenskjær conglomerate (the Evenestangen thrust sheet of Melezhik et al., 2003) of the Ofoten district of southern Troms. Chemostratigraphic studies, however, indicated a likely Mid Cambrian maximum depositional age for the Fauske conglomerate (Melezhik et al., 2000), whereas the Evenes limestone includes several imbricated marble units with apparent depositional ages ranging from Ediacaran to Silurian (Melezhik et al., 2002a, b, 2003, 2008b).

Detailed sedimentological, isotope-geochemical and structural investigations on the Fauske conglomerate have shown that the tectonic deformation sequence in the Fauske Nappe is both polyphase and polyorogenic (Melezhik et al., 2000; Roberts et al., 2001, 2002). These studies have provided the first direct indications that most of the lithostratigraphy now composing the Fauske Nappe originated on a shelf and continental slope on the Laurentian side of Iapetus (Roberts et al., 2001, 2002).

Although isotopic dating is lacking in the Fauske Nappe, the above-mentioned studies suggest that the rocks experienced both Taconian (Early to Late Ordovician) and Scandian (Late Silurian to Early/Mid Devonian) deformation and metamorphism (James et al., 1993;

Isotope chemostratigraphy of high-grade marbles in the Rognan area 109

Yoshinobu et al., 2002; Barnes et al., 2007; Augland et al., 2013; McArthur et al., 2013). These features are more noticeable within the Uppermost Allochthon over a wider, regional scale (Roberts et al., 2007).

Structural geology of the area The new geological map and profiles (Fig. 2) reveal a particularly complex structural development of the study area. Five thrust sheets are recorded, four of which are overlain unconformably by the conglomerates of the Øynes formation. Early, mesoscopic, tight to isoclinal folds are found throughout the tectonostratigraphic successions and are mostly coeval with the local, main schistosity, although some such folds are slightly later and deform the schistosity (e.g., Nicholson & Rutland, 1969).

Larger-scale sub-isoclinal folds are also represented, as shown on the profiles B–B’ and C–C’. As will be explained later, these early folds and foliation are likely to have been generated during an inferred Early Ordovician deformation episode on the Laurentian side of the Early Palaeozoic Iapetus Ocean. The only exception to this are the structures and main schistosity in the Rognan formation which, by virtue of the apparent age of the carbonate rocks in this formation, i.e., Early Silurian, were evidently produced during the Silurian, Scandian orogeny. Evidence for early NW–directed thrusting, as reported from the Øynes formation in the Løvgavlen quarry near Fauske (Roberts et al., 2001, 2002), has not been detected in the study area.

Structures that clearly deform the early schistosity and folds vary in scale from the microscopic to regionalscale antiforms and synforms (see profiles) with steep to vertical axial surfaces. These folds also deform the unconformity at the base of the Øynes formation as well as the internal early folds and metamorphic fabric in this same formation, and are therefore likely to be part of the Scandian cycle.

In addition to these diverse, polyorogenic, contractional structures, mesoscopic extensional shear bands and down-dip verging folds occur along or in the vicinity of several west-dipping surfaces, including original thrusts. Such extensional structures are now widely recognised throughout the western, central and northcentral Norwegian Caledonides as far north as Lofoten, and developed during a late-Scandian, extensional ‘collapse’ of the nappe pile (e.g., Hossack, 1984; Norton, 1986; Andersen, 1998; Fossen, 2000; Roberts, 2003), now isotopically dated to Devonian age (Terry et al., 2000; Nordgulen et al., 2002; Osmundsen et al., 2003, 2006; Kendrick et al., 2004; Steltenpohl et al., 2011). The nearest, documented, major extensional shear zone is the Breivika shear zone (Braathen et al., 2002), of inferred Devonian age, which forms the southwestermost limit of the study area near Ljøsnehammaren. This shear zone represents a late-Scandian reactivation of the basal

110 V.A. Melezhik et al

thrust of the Rödingsfjället Nappe Complex. Another comparable extensional shear zone has recently been reported from just northeast of Bodø, with preliminary isotopic indications of a Devonian age for these lateScandian movements (Steltenpohl et al., 2013).

Secular variations of δ13Ccarb and 87 Sr/86Sr in the Neoproterozoic and the Early Palaeozoic Secular variations of δ13Ccarb and 87Sr/86Sr in the Neoproterozoic and the Early Palaeozoic have been reconstructed by using diverse geological material (e.g., fossils, whole-rock samples) and by employing different methods for placing constraints on the age of deposition (e.g., radiometric dating, biostratigraphy, chemostratigraphy). For the Cambrian and younger time, the isotopic evolution of seawater is based on analyses of fossils composed of aragonite or high-Mg calcite (e.g., brachiopods, conodonts, foraminifera) and, hence, measured C- and Sr-isotope values represent close approximations to the seawater composition (e.g., Veizer et al., 1999), whereas radiometric and biostratigraphic ages provide robust time constraints on the deposition.

Neoproterozoic and older rocks commonly do not contain fossils suitable for C- and Sr-isotope analyses, and the reconstructed evolution of seawater at that time is based on whole-rock samples which are prone to diagenetic overprints of both C- and, in particular, Sr-isotope systems (e.g., Derry et al., 1992; Kaufman et al., 1993). However, the most important shortcoming for the Precambrian world is the dearth of reliable radiometric ages constraining the time of deposition of the measured successions (see discussion in Melezhik et al., 2001b).

Significant progress has been made in recent years in constructing the δ13Ccarb reference curve for seawater

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evolution through the Late Tonian, Cryogenian and Ediacaran (e.g., Halverson & Shields-Zhou, 2011). The currently available reference curve (Fig. 3A) shows frequent δ13Ccarb fluctuations with prominent magnitudes.

Superficially, although the curve looks robust it is not all based on radiometrically constrained data, as some parts involve a significant measure of intercontinental litho- and chemostratigraphic correlation (e.g., Akademikerbreen Group; Halverson et al., (2007)) and can thus be disputed, and may consequently experience future amendments.

Other details of the compilation of the Neoproterozoic δ13Ccarb record have been discussed in numerous publications (e.g., Kaufman et al., 1996; Hayes et al., 1999; Walter et al., 2000; Melezhik et al., 2001b; Halverson et al., 2005; Prave et al., 2009a, b). Despite its several shortcomings, the main structure of the Neoproterozoic δ13Ccarb record is now fairly clear. It is distinguished by generally high average δ13C values (~ +5‰) during the Late Tonian, and >+5‰ values throughout the Cryogenian punctuated by several, sharp, high-amplitude negative departures (e.g., Halverson et al., 2010) whose precise timing is not always constrained (for a detailed discussion, see Halverson & Shields-Zhou (2011)).

Obtaining reliable and well-constrained 87Sr/86Sr data through the Neoproterozoic represents an even more serious challenge. The Sr-isotope system is more prone to post-depositional resetting during diagenesis and metamorphism in comparison to the C-isotope system. Radiometrically-dated sections do not always contain high-Sr limestones, which are resistant to postdepositional alteration of 87Sr/86Sr, whereas high-Sr marine limestones are not always located within dated successions (e.g., Melezhik et al., 2009). The complexity of the issue is reflected in several of the reference curves that are available in the published literature. The earliest composite curve was presented in Melezhik et al.

(2001b). This has been later modified (Fig. 3A; Melezhik et al., 2008a) by involvement of new data obtained by

Figure 3. Temporal trends of 87Sr/86Sr and δ13Ccarb in seawater through the Neoproterozoic and Early Palaeozoic.

(A) 87Sr/86Sr and δ13Ccarb reference curves based on compilations made by Melezhik et al. (2001b) and Halverson & Shields-Zhou (2011).

(B) Two 87Sr/86Sr reference curves (red dots after Halverson & Shields-Zhou (2011); rectangles after Kuznetsov et al. (2013)) which differ between 680 and 660 Ma. The data, which are considered as the proxy to original Proterozoic seawater, are taken from: (1) Member I-6, Atar Group (Veizer et al., 1983); (2) Little Dal Group (Halverson et al., 2007); (3) Gillen Member, Bitter Springs Formation (Walter et al., 2000); (4) Inzer Formation, Karatau Group (Kuznetsov et al., 2003, 2006); (5) Minyar Formation, Karatau Group (Kuznetsov et al., 2003, 2006); (6) Shaler Group (Asmerom et al., 1991); (7) Akademikerbreen Group (Derry et al., 1989, 1992; Halverson et al., 2007); (8) Coates Lake Group (Halverson et al., 2007); (9) Rasthof Formation, Otavi Group (Yoshioka et al., 2003); (10) Uk Formation, Karatau Group (Kuznetsov et al., 2003, 2006); (11) Keele Formation, Windermere Supergroup (Narbonne et al., 1994; Halverson et al., 2007); (12) Ombaatjie Formation, Otavi Group (Halverson et al., 2007); (13) Hayhook Formation, Windermere Supergroup (James et al., 2001; Halverson et al., 2007); (14) Maieberg Formation, Otavi Group (Halverson et al., 2007); (15) Doushantuo Formation (Yang et al., 1999; Sawaki et al., 2010); (16) Blueflower Formation, Windermere Supergroup (Kaufman et al., 1993; Narbonne et al., 1994); (17) Witvlei Group (Kaufman et al., 1993); (18) Wonoka Formation (Calver, 2000; Walter et al., 2000); (19) Huqf Group (Burns et al., 1994); (20) Nama Group (Kaufman et al., 1993); (21) Tinnaya Formation (Gorokhov et al., 1995); (22) Ust-Yudoma Formation (Semikhatov et al., 2003); (23) Pestrotsvet Formation (Nicholas, 1996); (24) Tommotian, Atdabanian, Botomian and Toyonian type sections, Early Cambrian (Derry et al., 1994; Kaufman et al., 1996); (25) Macha and Tolbacha formations, Early Cambrian (Gorokhov et al., 1995).

(C) δ13Ccarb and combined 87Sr/86Sr reference curves based on compilations made by Kuznetsov et al. (2013) and Halverson & Shields-Zhou (2011). Pink rectangles designate ranges in 87Sr/86Sr, which allow a unique constraint of apparent age range for the Mid-Late Cambrian and Cryogenian-Ediacaran. Blue rectangles designate ranges in δ13Ccarb, which allow a unique constraint of apparent age range for the Mid and Late Ediacaran and, in combination 87Sr/86Sr data, for three ages in the Cryogenian.

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 111

112 V.A. Melezhik et al

Kuznetsov et al., (2003, 2005, 2006). Recently, Halverson & Shields-Zhou (2011), and Kuznetsov et al. (2013) published two new accounts on the evolution of 87Sr/86Sr in the Neoproterozoic (Fig. 3B). The 87Sr/86Sr curves of Halverson & Shields-Zhou (2011) and Kuznetsov et al.

(2013) show some similarities as well as clear differences.

The differences occur at around 830–810 and 670–660 Ma, and they are caused by far too low 87Sr/86Sr ratios measured from the Uk and Inzer formations of the Karatau Group in the Urals with respect to those obtained from supposedly correlative sections in the Shaler Supergroup in Canada and the Akademikerbreen Group in Svalbard (Fig. 3B). None of the groups and formations involved in these conflicting measurements has been precisely dated radiometrically. The depositional age of the Akademikerbreen Group remains radiometrically unconstrained. The depositional age of the Inzer Formation has been constrained by a Pb–Pb dating technique on carbonates (836 ± 25 Ma, Ovchinnikova et al., 1998) and the age of the Uk Formation has been constrained by a K–Ar (669 ± 16 Ma) and Rb-Sr (663 ± 9 Ma) dating of glauconite (Zaitseva et al., 2008). Moreover, no other measured sections which are correlative with the Uk Formation, are currently known. The overall disagreement between the two curves will apparently remain until the disputed and measured formations are radiometrically and precisely dated. Here, we take a conservative stance, by accepting the 87Sr/86Sr reference curve for Neoproterozoic seawater as a wide band (Fig. 3C) enveloping all reliable Sr-isotopic data reported in the recent compilations by Halverson & ShieldsZhou (2011), and Kuznetsov et al. (2013). For the Early Palaeozoic, the reference curve is more accurate especially where analytical data have been obtained from several fossiliferous limestone formations worldwide (e.g., Derry et al., 1994; Denison et al., 1994; Veizer et al., 1999).

Chemostratigraphic approach for constrain­ing apparent depositional ages ­of carbonate successions – general consider­ations Secular variations of δ13Ccarb and 87Sr/86Sr in the Neoproterozoic and Early Palaeozoic have already been widely used for the purpose of ‘isotope chemostratigraphy’ and indirect dating of carbonate sedimentation where no additional biostratigraphic or geochronological control could be provided (e.g., Melezhik et al., 2001b and references therein). With the task in hand, it is fair to state that the method has shown great promise for obtaining approximate, apparent depositional constraints when applied to highgrade metamorphic complexes. Even marbles that have experienced multiphase metamorphism and deformation may retain their depositional carbon-isotope values and preserve near-primary 87Sr/86Sr ratios (e.g., Melezhik et al., 2001a; Slagstad et al., 2006; Satish-Kumar et al., 2008;

NORWEGIAN JOURNAL OF GEOLOGY

Prave et al., 2009a,b; Otsuji et al., 2013). The reliability of C- and Sr-isotope chemostratigraphy clearly depends on (i) the precision of the reconstructed isotopic evolution of seawater and (ii) the potential of the tested rocks to preserve depositional isotopic values.

The compilation presented in Fig. 3C suggests that secular variation of 87Sr/86Sr in Tonian-Silurian seawater, even if used alone and with the conservative stance (wide 87Sr/86Sr band), can clearly discriminate carbonate formations deposited between c. 630 and 1000 Ma (87Sr/86Sr = < 0.7074) and in the Mid and Late Cambrian (87Sr/86Sr = >0.70895– 0.7094). The compilation presented in Fig. 3C also suggests that, in general, the carbonate formations deposited from the Tonian to the Silurian cannot currently be unambiguously dated indirectly by means of carbon isotope chemostratigraphy alone, with one exception. The exception is seen in the Mid Ediacaran when isotopic values suddenly plunge below -8.5‰. Such low values have been linked to the Wonoka (Halverson et al., 2005) or Shuram (Fike et al., 2006; Le Guerroué et al., 2006) or Shuram-Wonoka (Melezhik et al., 2008b) isotopic event. Since the event was first recognised as a global phenomenon (Halverson et al., 2005; Melezhik et al., 2005b) there have been continuing debates on the origin of such extraordinary values, with opinions ranging from depositional (Fike et al., 2006) to burial-diagenetic (Derry, 2010). Recently, however, Johnston et al. (2012) have reported paired, organic-carbonate, carbon isotopic data and shown that negative excursions of a similar magnitude (down to -8‰) in the Cryogenian are almost certainly depositional. Hence, there are growing lines of evidence suggesting that high-magnitude negative excursions of δ13Ccarb in the Neoproterozoic have been driven by syndepositional factors. The best estimates for dating the Shuram–Wonoka isotopic excursion, as reported in the literature, are either at around 551 Ma (Condon et al., 2005), between c. 600 and 550 Ma (Le Guerroué et al., 2006), or from c. 580 to c. 550 Ma (Fig. 4; Fike et al., 2006); hence, all point to the Mid–Late Ediacaran. Strontium isotope data from highly Sr-enriched marine limestones from southern Siberia apparently provide the most reliable constraint on 87 Sr/86Sr for Mid-Late Ediacaran seawater. At about the level of the nadir in the Shuram–Wonoka excursion, 87Sr/86Sr = 0.7082, and it gradually increases to 0.7086 up-section where it is still within the range of very low δ13Ccarb < -8‰ (Pokrovskii et al., 2006; Melezhik et al., 2009). The compiled seawater 87Sr/86Sr reference curve suggests a somewhat wider range of 0.7078–0.7089 (Fig. 3C). The combined application of δ13Ccarb < -8.5‰ and 87Sr/86Sr = 0.7082– 0.7089 for the Shuram-Wonoka seawater (Fig. 3C) can thus provide a reliable constraint on the apparent depositional age of sedimentary carbonates accumulated during the Mid-Late Ediacaran (c. 600–580 Ma). A similar range of the strontium isotopes combined with δ13Ccarb ranging between -8.5‰ and -5‰ now appears to be a characteristic feature for 580–570 Ma (Mid–Late Ediacaran) seawater.

Three other negative excursions of δ13Ccarb (down to -6 or -8.5‰), which have been imprecisely constrained at

Isotope chemostratigraphy of high-grade marbles in the Rognan area 113

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660, 730 and 750 Ma within the Cryogenian (Fig. 3C; for details see Halverson & Shields-Zhou (2011)), can be also applied for deciphering apparent depositional ages when δ13Ccarb values are used together with 87Sr/86Sr. However, no discrimination is possible between these three separate Cryogenian ages. The applications of the combined δ13Ccarb and 87Sr/86Sr data for reconstruction of their apparent depositional ages are presented in Table 1. In addition, Fig. 3C demonstrates that isotope chemostratigraphy does not allow us to discriminate confidently between Silurian and Ordovician, and some Cambrian and Ediacaran ages.

Analytical techniques and data Samples obtained from marbles from various formations have been analysed for major and trace element concentrations at the Geological Survey of Norway (NGU), Trondheim. Measurements of carbon and oxygen isotopes have been carried out at the Scottish Universities Environmental Research Centre (SUERC), Glasgow. Rubidium and strontium isotope analyses have been performed at the Institute of Precambrian Geology and Geochronology of the Russian Academy of Sciences, St. Petersburg. Analytical techniques for all measurements are specified in the Appendix. The obtained analytical data are presented in the Electronic Supplement.

Geochemical screening against postdepositional alteration of C, O and Rb–Sr isotope systems in high-grade marbles Figure 4. Detailed δ Ccarb profiles through the Ediacaran period (after Fike et al., 2006) showing the apparent age projection of the Leivset marbles based on the observed δ13Ccarb range. Formational names are from Oman.

13

Diagenetic and metamorphic alterations affect carbonate material in a similar way (Nabelek, 1991). These processes usually lower δ13C and δ18O, and introduce radiogenic strontium. In general, during post-depositional, opensystem recrystallisation, the δ13C of calcite and dolomite would be buffered by the dissolving precursor, while

Table 1. Range of δ13Ccarb and 87Sr/86Sr or combination of both ratios (shown in bold), which have a unique age resolution and can be confidently applied for indirect age constraints when projected on to the seawater reference curve through the Neoproterozoic-Cambrian. δ13Ccarb

Sr/86Sr

87

Apparent age, Ma

Comment Mostly Tonian–Cryogenian No discrimination is possible b ­ etween these three Mid–Late Cryogenian ages in 'blind' experiment Mid Ediacaran (Shuram–Wonoka excursion) Late Ediacaran (Shuram–Wonoka excursion) Mid to Late Cambrian

Any values between -8.5 and +9‰

0.7052–0.7074

1050–630

-5 to -8.5‰

0.7062–0.7074

c. 750, c. 730 or c. 660,

<-8.5‰

0.7078–0.7089

600–580

-5 to -8.5‰

0.7078–0.7089

580–570

0 ± 3‰

0.7091–0.7094

530–485

Isotope data shown in bold are specific for designated time intervals.

>20

>22

<0.0005

≤0.0001

As the rocks studied are polymetamorphosed, normal petrographic screening and cathodoluminescence are of very limited use as they can only detect the alteration associated with the last geochemical transformation.

Hence, the discrimination technique has been based essentially on geochemical criteria.

?

<50

>300

>1000

-

n.d.

<1.0

?

n.d.

<5

-

n.d.

n.d.

>0.5

Palaeozoic non-metamorphosed shelf limestones4

Cambrian non-metamorphosed ­limestones5

Vendian non-metamorphosed l­ imestones6

High-grade marbles7

5

Data are from Montañez et al. (1996) Mn/Sr <0.065 Data are from Derry et al. (1992) and Kaufman et al. (1993).

7 Data are from Melezhik et al. (2001a, 2002a, 2000b, 2003). ‘n.d.’ – not determined.

<0.02

5-81 154-966

<2 <10

Silurian brachiopods3

6

n.d.

<300 >900

-

Modern shells

Holocene brachiopods2

1

Data are from Kuznetsov et al. (2012).

Data are from Lowenstam (1961), Dittmar & Vogel (1968), Frank et al. (1982) and Grossman (1994).

3 Data are from Azmy et al. (1998).

4 Data are from Denison et al. (1994).

1

2

n.d. <0.065

<1.0

n.d.

n.d. ≤0.2

≤0.02

>24

n.d. n.d.

4-200 1000-2000

wt% -

ppm

n.d.

the δ18O, Mn and Sr contents would be partially shifted towards equilibrium with the ambient diagenetic fluids.

In order to alter oxygen, trace elements and carbon by diagenetic/metamorphic fluids, it would require a water/ sediment ratio of 101 : 102–103 : 104, respectively (Banner & Hanson, 1990; Land, 1992).

≤0.02

-

24-28

n.d. n.d.

<0.01

750-2000

n.d.

1-10

5-350

1020-2400

≤0.02

δ18O

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Rb/Sr Mn/Sr Mg/Ca Mn Sr Al2O3 SiO2 Sampling distance from contact with silica-aluminous rock m Grade of metamorphism/ material

Table 2. Empirically obtained selection criteria for the ’least-altered’ strontium isotopic values in high-grade marbles with an aragonitic precursor. Geochemical parameters used in low-grade or nonmetamorphosed carbonate rocks, non-metamorphosed brachiopods, and the marble formations previously studied are included for purposes of comparison. Based on data presented in Melezhik et al.

(2003).

114 V.A. Melezhik et al

In general, depending on metamorphic conditions and the chemical composition of the carbonate rocks, their premetamorphic carbon, oxygen and strontium isotopic values could either be overprinted (Nabelek, 1991; Romer, 1994; Bickle et al., 1995, 1997; Lewis et al., 1998) or preserved even under amphibolite-facies conditions (Ghent & O’Neil, 1985; Baker & Fallick, 1988, 1989a, b; Wickham & Peters, 1993; Boulvais et al., 1998; Melezhik et al., 2001a, 2003; Satish-Kumar et al., 2008; Otsuji et al., 2013).

The common geochemical assessment of postdepositional alteration of carbonate is largely based on the relative abundances of Mn, Fe, Rb and Sr (e.g., Brand & Veizer, 1980). Several elemental ratios (Mn/Sr, Fe/Sr, Ca/Sr and Rb/Sr), as well as carbon and oxygen isotopes, are widely used as geochemical criteria for detecting the least disturbed carbon, oxygen and Rb–Sr systems.

Different authors, however, have suggested and use not only dissimilar values of the same ratios, but also very variable combinations of these ratios (Asmerom et al., 1991; Derry et al., 1992; Kaufman et al., 1993; Semikhatov et al., 2002; Kuznetsov et al., 2003). In all cases the choice of the elemental ratios and their values is empirical and to some extent arbitrary (Table 2).

Conventional geochemical assessment of postdepositional alteration of carbonate has been specifically refined for high-grade rocks (Melezhik et al., 2001a, 2002a, 2003). Table 2 shows that the levels of empirically established screening criteria that have been used in our studies for the high-grade rocks is commonly significantly tighter than those used for nonmetamorphosed rocks. In addition, the Mg/Ca ratio was found to be a very sensitive parameter for identifying any disturbance of the Rb–Sr system in high-grade marbles, although the application of this parameter to non-metamorphosed limestones has been very rarely discussed in the published literature. Oxygen isotopes, being sensitive indicators of even minimal alteration in non-metamorphosed limestones, do not always appear to be useful parameters for screening disturbances of the more resistant strontium isotope system in both nonmetamorphosed limestones (Bickle & Chapman, 1990; Bickle, 1992; Jones et al., 1994; Montañez et al., 1996) and high-grade marbles (Melezhik et al., 2001a, 2002a, b).

0.5–5.8

wt%

<0.01

<0.01–0.23

Al2O3

<0.01–0.25

5.9

0.8

<0.01

<0.01

µg g-1

Mn 10-3

Mg/Ca 10-3

Mn/Sr

228

3540

260–484

662–1565

587–2686

5

-

10–140

9–34

19–82

9–50

3

6

7–50

10–40

20

-

50–820

10–50

30–80

0.09

0.02

<0.01

<0.01–0.2

3.2

<0.01

1.0

<0.01

1545

1925 22

13

34

18

2

3

630

5

20

7

70

9

2

high

7

1

2

3–14

2–5

1–4

10-4

Rb/Sr

+2.3 to +2.5

+7.3 to +7.7

-0.2 to +0.5

-2.5 to +1.2

-6.4 to -4.1

-13.7 to -6.1

-1.2 to +2.5

-0.1 to +2.2

+2.3 to +8.0

δ13C ‰

30

24

23

26

30

24

24

28

25

δ18O

0.70669

0.70675

0.70928

0.70881

0.70820

0.70798–(0.70889?)

0.70897–0.70940

0.70805–0.70814

0.70807–0.70826

Sr/86Sr

87

* This apparent depositional age is based on the chemostratigraphic correlation of 13C-rich dark-coloured marbles of the ’Tripartite Unit’ with a low-grade, fossiliferous limestone at Sagelvvatnet, which contains L ­ landovery-age brachiopods and corals (for details see Melezhik et al., 2008b). ** Number corresponds to that in the legend of the geological map (Fig. 2).

Numbers shown in bold exceed the empirically established elemental concentrations, elemental ratios and isotopic ratios that are indicative of the least-altered post-depositional Sr-isotope system in high-grade calcite marbles (see Table 2).

Subset 2

Cryogenian (800–660 Ma) grey and dark grey calcite marble (# 38**)

Subset 1

Kjerktinden formation Mid Cryogenian (775–750 Ma) grey and dark grey calcite marble (# 38**)

Subset 4

68

252

Mid Cambrian (530–500? Ma; dolomitisation age) white dolomite marbles (# 22, 23**)

Subset 3

Late Ediacaran (c. 560 Ma) or Early Cambrian through Ordovician (530–470 Ma) bedded, white, pale grey and grey calcite marbles (# 26, 27**)

Subset 2

Late Ediacaran (c. 550 Ma) white, grey and beige calcite marbles (# 24, 25**)

Subset 1

Leivset formation Mid Ediacaran (600–580 Ma) variegated and white calcite marbles (# 33, 32**)

0.08–5.4

Sr Rognan formation Early Silurian (c. 440 Ma*) grey and dark grey calcite marble (# 11**)

Øynes formation Early Ordovician, calcite matrix and calcite pebbles in conglomerates (# 13, 14**)

<0.01–9

Subset 2

SiO2

Subset 1

Formation

Table 3. The least-altered δ13C, δ18O and 87Sr/86Sr, and geochemical parameters employed for the selection of the best-preserved Sr-isotope ratios used for the reconstruction of apparent depositional ages.

NORWEGIAN JOURNAL OF GEOLOGY Isotope chemostratigraphy of high-grade marbles in the Rognan area 115

116 V.A. Melezhik et al

Table 3 summarises the geochemical characteristics of the studied high-grade marbles in the Rognan area and illustrates that most of them show a level of preservation of many of the geochemical parameters that is much higher than even in comparative non-metamorphosed limestones (Table 2). Quite surprisingly, Sr abundances in six out of nine, and Mn in all studied calcite marble formations correspond to those reported in nonmetamorphosed brachiopods. In all studied calcite marble formations, the Sr and Mn contents, and Mg/ Ca and δ18O ratios fall within the range of Cambrian limestones (Table 3). Tentatively, all marble formations having a Sr content above 1000 µg g-1 are assigned to an aragonitic protolith.

In high-grade, amphibolite-facies marbles, a limited number of analyses are considered to be insufficient for recognising alteration trends, even if all necessary geochemical, mineralogical and petrological precautions have been taken to identify alteration effects. Consequently, databases with a small number of analyses (<10) should not be used for the reconstruction of δ13C values, and particularly 87Sr/86Sr ratios, in both low- and high-grade carbonate rocks. In our study, 300 samples of calcite and dolomite marbles, representing all major marble units of the Fauske Nappe, were analysed

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for major and trace elements. Two hundred and fiftynine selected samples were measured for carbon and oxygen isotopes; and 87Sr/86Sr ratios were measured on 150 selected samples. The ‘least altered’ carbon, oxygen and strontium isotope ratios, selected by using multiple geochemical criteria (Table 2, high-grade marbles) with the additional aid of different cross-plots, are presented in Table 3 and discussed in some detail below.

Apparent depositional ages of ­carbonate formations in the Rognan area The analysed rocks represent mainly calcite and variably dolomitised calcite marbles (Fig. 5A). Dolomite and partially calcitised dolomite marbles are also present but are less abundant (Fig. 5A). Most of the analysed samples show Sr contents below 500 µg g-1 although high-Sr marbles are also common throughout the tectonostratigraphy (Fig. 5B, Electronic Supplement).

The overall carbon, oxygen and strontium isotope ratios show large variations. The δ13C values range between -13.7 and +8.0‰ (V-PDB), δ18O fluctuates between 15.6 and 30‰ (V-SMOW), whereas 87Sr/86Sr ratios vary between 0.70662 and 0.71087. When all the δ13C,

Figure 5. Various cross-plots illustrating the presence of several distinct C- and Sr-isotopic marble groups in the Fauske Nappe.

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δ18O and 87Sr/86Sr data are plotted together, they appear as several distinct, non-overlapping subsets (Fig. 5C, D). This might have been the combined result of postdepositional alteration and the presence of different age-group marbles. The discrimination between these two alternatives is considered below for each individual formation.

The Kjerktinden formation The Kjerktinden formation is composed of interbedded, pale-grey, garnet-mica schist (locally with staurolite), dark-grey, graphite-bearing mica schist, calcareous schist, dark-grey, graphite-bearing, calcite marble and quartzite (# 34–391). The formation occurs mainly in Kvanndalen–Kjerktinden and at Storvika, from where it continues along the western part of the mapped area.

Apparent Kjerktinden marbles also appear as a thin sliver in the Dverset area, extending farther south to western Nestbyfjellet (Fig. 2). In the study area, the structurally lower contact of the Kjerktinden formation is juxtaposed against marbles of the Leivset formation (# 24, 25, 26).

At Kvanndalen, a steeply dipping calcareous schist of the Kjerktinden formation is overlain unconformably by flatlying Øynes formation conglomerates. The Kjerktinden marbles (# 38) represent a geochemically homogeneous

Isotope chemostratigraphy of high-grade marbles in the Rognan area 117

group. All analysed samples (n = 41) are calcite marbles (Mg/Ca = 0.02–0.04) with low Al2O3 (<0.01–2.3 wt.%) and moderate SiO2 (0.15–15.3 wt.%) contents. The rocks have high Sr abundances (970–2540 µg g-1) and show a low Mn/Sr ratio (0.004–0.11). All analysed samples (except five) contain organic matter (0.11–0.28 wt.%) in the form of graphite. 87Sr/86Sr ratios range between 0.70669 and 0.70786 (Electronic Supplement).

Despite the geochemical homogeneity, the Kjerktinden marbles plot on a δ13C–δ18O diagram as two discrete subsets (Fig. 6A). The precise field relationship between Subset 1 (δ13C = -0.4–+2.5‰ and δ18O = 24.1–30.4‰) and Subset 2 (δ13C = +5.4–+7.7‰ and δ18O = 16.9– 24.4‰) marbles remains to be defined; however, in several outcrops throughout the mapped area, both geochemical types occur close to the contact with schists where they appear to be interbedded or tightly infolded. Various discrimination cross-plots do not suggest that the two subsets are the result of obvious post-depositional alteration processes. On the contrary, δ13C–δ18O, δ13C–Mn/Sr and δ18O–Mn/Sr cross-plots (Fig. 6A–C) show that each subset exhibits its own alteration trend. The cross-plots suggest that the least altered δ13C values for Subset 1 and 2 marbles are +2.3 to +2.5‰ and +7.3 to +7.7‰, respectively, whereas for δ18O the

Figure 6. Various crossplots illustrating apparent alteration trends in C-, O- and Sr-isotopic systems in the calcite marbles of the Kjerktinden formation. The apparent alteration trends are indicated by blue and pink arrows for Subset 1 and Subset 2, respectively.

1

Here and in the following text, these numbers refer to those in the legend of the geological map presented in Fig. 2.

118 V.A. Melezhik et al

values are 30‰ and 24‰ (Table 3). Similarly, 87Sr/86Sr– Mn/Sr and 87Sr/86Sr–δ13C cross-plots (Fig. 6D, E) identify alteration trends for Subset 2 marbles and indicate a least altered 87Sr/86Sr ratio of 0.70675. In contrast, in Subset 1, neither the used cross-plots (Fig. 6D, E) nor the empirically obtained discrimination criteria (Table 2) can assist in the reconstruction of the least altered Sr ratio.

Consequently, taking the conservative stance, the lowest value of 0.70669 has been accepted as the best proxy for the seawater composition.

When the least altered strontium isotope ratios of Subsets 1 and 2 are projected onto the 87Sr/86Sr reference curve for seawater, both suggest an apparent depositional age within the 800–660 Ma time interval, i.e., Cryogenian (Fig. 3C). The carbon isotope values provide further constraints for Subset 2 marbles because the δ13C > +7.7‰ combined with 87Sr/86Sr of 0.70675 is a characteristic feature of the Mid Cryogenian, 775– 750 Ma (Fig. 3C). The δ13C of +2.3 to 2.5‰ and 87Sr/86Sr of 0.70669 of Subset 1 marbles is a typical feature of the Early Cryogenian (800–785 Ma) although such a combination has also been registered in the Mid (745–735 Ma) and Late Cryogenian (c. 660 Ma; Fig. 3C). Further work is required to resolve the precise tectonostratigraphic relationship between Subset 1 and Subset 2 marbles.

The Leivset formation The Leivset formation is composed of interbedded white, pale-grey and variegated calcite marbles (Fig. 7A–F), grey and white dolomite marbles (Fig. 7G–J), calcareous mica schist, and a minor, thin, matrix-supported conglomerate unit resembling diamictite (# 22–33). The formation occurs in the northern, western and eastern parts of the mapped area, and appears in a series of tightly folded, dismembered slivers in its central part (Fig. 2).

Carbon, oxygen and strontium isotope systematics (Fig. 8) suggest that three distinct groups (hereafter Subsets 1, 2 and 3) of calcite marble (Fig. 8A) and one group of dolomite marble (hereafter Subset 4) can be identified in the Leivset formation (Fig. 8D-F). At Kvanndalen, all three subsets of calcite marble are erosively overlain by the Øynes formation conglomerate (Fig. 2).

Subset 1 includes banded, variegated and white calcite marble (# 33, Fig. 7A–F) passing into white and palegrey, coarse-grained, calcite marble (# 32), both of which have low δ13C ranging between -13.7 and -6.1‰ with one exception at -3.7‰ (n = 52), and δ18O =15–24‰ (with one exception at 27‰) (Fig. 9). Although the marbles crop out throughout the mapped area, they occur mainly at Nestbyfjellet. Subset 1 marbles have been sampled and studied from quarries at Leivset and Brenne, and from several natural outcrops in the Vensmoen, Nestbyfjellet and Dverset areas (Fig. 2).

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Subset 1 marbles are known from a previous study as a prominent chemostratigraphic marker horizon in the north-central Norwegian Caledonides. Melezhik et al.

(2008b) reported that its extremely 13C-depleted primary nature (-7.9 ± 1.2‰ on average, n = 93) together with a high Sr content (up to 8740 µg g-1) that buffered 87Sr/86Sr ratios between 0.70802 and 0.70872, strongly suggests correlation with the worldwide Shuram-Wonoka isotopic event (Halverson et al., 2005; Melezhik et al., 2008b) occurring within the 600–550 Ma time interval during the Ediacaran (Vendian) period (Fig. 4). Subset 1 marbles (δ13C = -13.7 to -6.1‰) apparently record the nadir of the Shuram-Wonoka isotopic excursion. These 13 C-depleted marbles occur over a great lateral extent in the Scandinavian Caledonides. A unique combination of variegated colour, 13C-depletion and Sr enrichment of the marker horizon combined with its apparent depositional ages and wide geographic distribution meet the requirements for a chronostratigraphic marker (Melezhik et al., 2008b). In several places, including the study area, isotope chemostratigraphy has identified a prominent, cryptic, stratigraphic discontinuity and suggests that the 13C-depleted Ediacaran marble was tectonically juxtaposed above Llandovery-age marbles during the Scandian orogeny. These Llandovery-age carbonate rocks occur in places as low-grade, fossiliferous limestones, as, for example, in the Sagelvvatnet area of Troms (Melezhik et al., 2008b). In the study area, the Llandovery-age fossiliferous marbles correspond to the Rognan formation (# 11; for details, see the Nestbyfjellet road section in Fig. 10). Figure 10 presents a C-isotopic cross-section through tectonically stacked, Llandovery and Ediacaran marble units.

The preservation potential of carbon and strontium isotope systems in the 13C-depleted marbles has been discussed in great detail in Melezhik et al. (2008b), including some data from the study area (the Leivset, Brenne and Ljøsenhammaren quarries and the Nestbyfjellet road section). The newly obtained and analysed material shows no major geochemical and isotopic differences from previously analysed material.

The marbles do not show a significant correlation of the Mn/Sr ratio with δ18O or δ13C, hence no obvious resetting of the δ13C and δ18O values can be readily identified (Fig. 9B–E). 87Sr/86Sr ratios tend to correlate positively with Mn abundances and Mg/Ca ratios, hence suggesting some alteration (Fig. 9G, H). Fifteen samples with Sr contents ranging between 1000 and c.

7000 µg g-1 have 87Sr/86Sr fluctuating between 0.70798 and 0.70889 (Fig. 9F). Such high Sr contents suggest an aragonitic precursor and should provide a robust buffer for post-depositional alteration of the Sr isotope system.

Moreover, these 15 samples with high Sr content meet all empirically derived selection criteria for the ‘least altered’ strontium isotopic values established for high-grade marbles with an aragonitic precursor (Table 2). The average contents of SiO2 (4.4 wt.%) and Al2O3 (0.5 wt.%) are low. Mn/Sr (0.03) and Mg/Ca (0.02) ratios, as well

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as the Mn content (67 µg g-1), are close to the required limits (Table 2). The average Sr content is >3000 µg g-1.

Consequently, the whole 87Sr/86Sr range of 0.70798– 0.70889 could be considered to represent the least altered ratios and represent the best proxy to seawater 87 Sr/86Sr. When these ratios are combined with the δ13C of -13.7 to -6.1‰ and projected onto the reference curve of seawater evolution, they agree within the 600 and 565 Ma time interval (Fig. 3C). However, the 87Sr/86Sr ratios correlate positively with the Mg/Ca ratios (r = +0.77, n = 18, >99.9%), strongly suggesting alteration of the Sr-isotope system by dolomitising fluids (Fig. 9G). Based on the 87Sr/86Sr–Mg/Ca cross-plot, the lowest 87Sr/86Sr ratio of 0.70798 is proposed to represent the best proxy to seawater composition. This ratio, combined with the δ13C value of -13.7 to -6.1‰, thus suggests a much narrower apparent depositional age range of 600–580 Ma (Fig. 3C).

Subset 2 includes grey calcite marble (# 24, 25) showing a higher though also negative δ13C (-6.3 to -4.1‰), but a generally higher δ18O (26–30‰) (Fig. 9). The marbles occur as a rather thick unit throughout the mapped area forming a large, complex, antiformal structure. They have been studied and sampled in a quarry at Ljøsenhammaren as well as in several natural exposures in the Gullurda, Dverset and Kvanndalen areas (Electronic Supplement, Fig. 2). Both Subset 1 and Subset 2 marbles may also occur in the Dverset area, but there they are not interbedded and separated by calcareous mica schist of the Øynes formation (Fig. 2). In Subset 2 marbles, neither δ13C nor δ18O correlates with Mn/Sr or Mg/Ca ratios, nor with any other geochemical parameter or measured elemental abundance. Thus, we have assumed any alteration to have been relatively insignificant. As in the case of the Kjerktinden marbles, the high δ18O values may represent a primary depositional feature; currently the local preservation of such high δ18O values remains enigmatic and unexplained.

A large range of the 87Sr/86Sr (0.70820–0.70988) ratios combined with a Sr content <1000 µg g-1 (Fig. 9F) is indicative of alteration of the Sr-isotope system. 87Sr/86Sr versus Mn and Mg/Ca cross-plots also suggest a certain measure of alteration. Moreover, 87Sr/86Sr ratios correlate positively with Mn abundances (r = +0.69, n = 16, >99%). This strongly indicates a severe alteration of the Sr-isotope system caused by Mn-bearing fluids (Fig. 9H).

Based on the 87Sr/86Sr–Mn cross-plot, the lowest 87Sr/86Sr ratio of 0.70820 has been accepted to represent the best proxy for the seawater composition. This ratio matches with the lowest Mn/Sr (0.02) and Mg/Ca (0.003) ratios, and with SiO2 and Al2O3 contents below 0.01 wt.% (Table 3). This Sr-isotope ratio and δ13C values ranging between -6.4 and -4.1‰ provide an apparent depositional age of c. 550 Ma when the data are projected onto the reference curve (Fig. 3C). If the inference that the Subset 1 marbles with δ13C = -13.7 to -6.1‰ record the nadir of the

Isotope chemostratigraphy of high-grade marbles in the Rognan area 119

Shuram-Wonoka excursion is correct, then the second subset (δ13C =-6.4 to -4.1‰) apparently documents the recovery period towards near-zero values (Fig. 4).

Subset 3 includes white, grey and dark-grey, massive to weakly banded, coarse-grained calcite marbles­ (# 26, 27). This group of marbles is characterised by very homogeneous δ13C values fluctuating around zero (ranging between -0.9 and +1.2‰ with one exception at -2.5‰; n = 23) and by δ18O values varying between 19 and 26‰ but mainly clustering around 21.5‰. Subset 3 rocks are rather pure (SiO2 = 0.12–4.4 wt.%; Al2O3 <0.01– 0.39 wt.%; n = 26) calcite (Mg/Ca = 0.005–0.03) marbles with low to moderate Sr abundances (255–910 µg g-1), and Mn/Sr ratios ranging between 0.02 and 0.07. Subset 3 marbles occur throughout the mapped area from north to south and appear as a tightly folded unit showing pinchand-swell structures and isoclinal folds. In general, a robust control on the stratigraphic base or top of this unit is not yet available, and the unit is in contact with various rocks including Leivset formation calcareous schist, isotopically light, white and variegated marbles (Subsets 1 and 2) of apparent Mid (600–580 Ma) and Late (c. 550 Ma) Ediacaran ages, and mica schist of the Fjellengflåget formation. At Kvanndalen, the marbles are erosively overlain by the Øynes formation conglomerate (Fig. 2).

Here, a thin unit of matrix-supported conglomerate with calcareous schist matrix, resembling diamictite, occurs at the structural and stratigraphic base of the Subset 3 marble. Where the conglomerate wedges out, the marble is in direct contact with the Kjerktinden formation of Neoproterozoic age (800–650 Ma). Interestingly, carbonate clasts in the diamictite at Kvanndalen are mainly calcite marble with low δ13C (-5.5 to -4.9‰) and high δ18O (>24‰, Electronic Supplement), hence their source has no connection with the Kjerktinden marbles (δ13C = +0.7 to +7.7‰; Electronic Supplement).

Various cross-plots do not suggest any immediately recognisable alteration affecting the carbon and oxygen isotope systems in the Subset 3 marbles. δ13C and δ18O ratios do not show any significant correlation (Fig. 11A).

Similarly, neither δ13C nor δ18O shows any correlation with Mg/Ca or Mn/Sr (Fig. 11B–C). Intriguingly, the most negative δ13C (-2.5‰) is coupled with the highest δ18O of 26.3‰ (Electronic Supplement, Fig. 11A).

Consequently, we assume that the entire δ13C range from -2.5 to +1.2‰ may represent a proxy for seawater composition.

In contrast, several cross-plots do suggest a fairly pronounced alteration of the 87Sr/86Sr ratios that fluctuate between 0.70881 and 0.70963 (Fig. 11D–H). 87Sr/86Sr shows a significant positive correlation with Mn (r = 0.68, n = 16, >99%) and Mg/Ca (r = 0.84, n = 16, >99.9%), and hence was altered to a certain extent by Mn-bearing, dolomitising fluids. Cross-plots of 87Sr/86Sr versus Mn abundance and Mg/Ca ratio offer a better guidance, and suggest that 0.70881 represents the least altered ratio

120 V.A. Melezhik et al

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 121

Figure 7. Photographs of selected rock types from the Leivset formation.

(A, B, C) Variegated, calcite marble with a distinctive metamorphic layering enhanced by a considerable tectonic thinning. The photographs were taken from near-vertical, smooth walls in quarries at Leivset, Ljøsnehammaren and Brenne, respectively; the scale bar here and in all subsequent photographs is 10 cm. (D) Intensely folded, variegated, calcite marble with pinch-and-swell structure; the photograph was taken from a nearvertical, smooth wall in a quarry at Storvika. (E) Detailed view of pinch-and-swell structure in variegated, calcite marble at Storvika. (F) Sheared, variegated, calcite marble exposed at Dverset; pen for scale is 14 cm. (G) Quarry wall with partially collapsed adit at Hammerfall illustrating variably banded, beige and pale grey dolomite marble; length of photograph is c. 65 m. (H) Quarry wall with adit at Løvgavlen illustrating massive, white dolomite marble. (I) Detailed view of thick-banded, white and pale grey, dolomite marble at the base, which gradually becomes thinner banded and platy at the top; quarry wall at Løvgavlen; the width of the photograph is c. 2 m. (J) Banded, white, beige and pale grey, dolomite marble forming the substrate to the pink, carbonate conglomerate of the Øynes formation. The irregular palaeorelief beneath the conglomerate has been affected by tectonic deformation involving folding and an associated spaced cleavage. The photograph was taken from a near-vertical, smooth wall in the Løvgavlen quarry at Fauske.

122 V.A. Melezhik et al

NORWEGIAN JOURNAL OF GEOLOGY

Figure 8. Various cross-plots illustrating the presence of several distinct C- and Srisotopic marble groups in the Leivset formation. A positive corre­lation between 87Sr/86Sr and δ13C (excluding the highest 87Sr/86Sr value of 0.70988) in calcite marbles is interpreted as a stratigraphic trend reflecting the isotopic evolution of Ediacaran seawater (see Fig. 3c).

Figure 9. Various cross-plots illustrating apparent alter­ ation trends in the Sr-isotopic system in Subset 1 (variegated and white) and Subset 2 (white and pale grey) calcite marbles of the Leivset formation. The apparent alteration trends are indicated by grey and pink arrows for Subset 1 and Subset 2, respectively. A positive correlation between 87Sr/86Sr and δ13C (excluding the highest 87 Sr/86Sr value of 0.70988) is interpreted as a stratigraphic trend reflecting the isotopic evolution of Ediacaran sea­ water (see Fig. 3c).

Isotope chemostratigraphy of high-grade marbles in the Rognan area 123

NORWEGIAN JOURNAL OF GEOLOGY

Figure 10. Simplified lithological sections and isotopic profiles through the calcite marble of the Llandovery Rognan formation (dark grey) and the variegated and white calcite marble of the Ediacaran Leivset formation showing the sharp offset of δ13Ccarb by 14‰ (from +6‰ to -8‰) at the contact between the two units; modified from Melezhik et al. (2008b).

24 22 20

0

24

-1

22

-2

20 18

18 -3

-2

0.7098

D

-1

0

1

13

C (‰, VPDB)

0.04

0.7098

E

0.06

0.08

0.1

Mn/Sr

0.12

0.02

87

87

-2

-1

0

C (‰, VPDB) 13

1

0.7092

G

0.7098

18

20

0.7098

H

0.7094

0.7094

87

r = +0.84 n = 16, >99.9%

0.7090 0.7088

18

22

24

26

O (‰, VSMOW)

28

0.7088

0

0.02

0.04

Mg/Ca

0.06

0

20

40

60

Mn, g g-1

Subset 3 13

C = -2.5 to +1.2‰ 18

O = up to 26‰ 87 Sr/86Sr = 0.70881

0.7092 0.7090

0.08

r = +0.68 n = 16, >99%

The least-altered values

86

86

Sr/ Sr

0.7096

Sr/ Sr

0.7096

0.7092

0.12

0.7092 0.7090

0.7088

2

0.1

Sr/86Sr

Sr/86Sr

Sr/86Sr

-3

0.08

Mn/Sr

0.7094

0.7090

0.7090

0.06

0.7096

0.7094

0.7092

0.04

F

0.7098

0.7096

0.7094

0.7088

-3 0.02

2

0.7096

87

1

26

13

C (‰, VPDB)

18

O (‰, VSMOW)

18

O (‰, VSMOW)

26

C

2

B

28

A

28

87

Figure 11. Various cross-plots illustrating apparent alteration trends in the Sr-isotopic system in Subset 3 calcite marbles of the Leivset formation. The apparent alteration trend is indicated by grey arrows.

0.7088 200

400

600

800

-1

Sr, gg

1000

80

124 V.A. Melezhik et al

(Fig. 11G), and, hence, a proxy for seawater composition.

The least altered 87Sr/86Sr ratios of 0.70881 with a limited fluctuation of δ13C values around zero would not appear to suggest a unique time resolution; such a combination is consistent with either Late Ediacaran (c. 560 Ma) or with several ages within Early Cambrian to Early Ordovician time (530–470 Ma, Fig. 3C). Since Subset 3 marbles were originally mapped jointly with Subset 1 and 2 marbles as the unit constituting the Leivset formation, and as they are lithologically indistinguishable from the adjacent Late Ediacaran (c. 550 Ma) Subset 2 marbles at Gullurda, we tentatively suggest that c. 560 Ma is a more likely apparent depositional age for the Subset 3 marbles.

However, considering that the marbles are overlain unconformably by the Cambrian–Early Ordovician (520, 510–505 or 475–470 Ma; see section Øynes formation) Øynes formation conglomerate, their apparent depositional age still remains uncertain; either c. 560 Ma or within the 530–470 Ma range.

Subset 4 marbles are the next isotopically and lithologically distinctive rock unit in the Leivset formation. They are represented by white and grey dolomite marbles (# 22, 23; Fig. 7G–J) with δ13C ranging between -0.2 and +0.5‰ and with δ18O fluctuating between 20.5 and 23.2‰ (Fig. 8D, E). The dolomite marbles occur in two different stratigraphic positions.

The oldest marble occurs as a thin continuous unit occurring in the central part of the mapped area where it rests on the structural and stratigraphic top of the isotopically light carbon, 600–580 Ma, Leivset marble (Subset 1). The younger dolomite marble occurs as a series of thick lenses in the western part of the mapped area where it appears to be associated with the stratigraphic top of the c. 550 Ma Leivset calcite marble (Subset 2). Here, these dolomite marbles also appear to be in tectonic contact with the Fjellengflåget formation, both units being overlain unconformably by the Øynes formation carbonate conglomerate. Thick lenses of Leivset dolomite marble and Øynes conglomerate also occur abundantly in the western part of the mapped area.

The dolomite marbles of Subset 4 are generally finegrained rocks with a rather massive appearance, although in places retaining remnants of bedding inherited from their limestone precursor (Fig. 7G–J). They represent a valuable raw material for the agricultural industry and they are currently extracted from the Hammerfall and Løvgavlen quarries near Fauske (Fig. 1). In the Løvgavlen quarry, the dolomite marble is erosively overlain by the Fauske carbonate conglomerate of the Øynes formation, hence indicating an exposure surface (Melezhik et al., 2000).

Nineteen samples obtained from the Hammerfall and Løvgavlen quarries show that the dolomite marbles are very pure rocks with an average Mg/Ca ratio of 0.63. They contain only small amounts of SiO2 (<0.01–1.3 wt.%) and Al2O3 (<0.01–0.8 wt.%). The rocks have low Mn and Sr contents averaging 34 and 68 µg g-1, respectively

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(Electronic Supplement). The dolomite marbles have very homogeneous δ13C (-0.2 to +0.5‰) and δ18O (20.5– 23.2‰) values, and radiogenic 87Sr/86Sr ratios ranging between 0.70928 and 0.70967 with one outlier at 0.71097 (Fig. 8D–F). Since in high-grade dolomite marbles, the strontium isotope system is commonly re-equilibrated by dolomitising fluids (e.g., Melezhik et al., 2005a), the only available constraint on the time of deposition of the dolomite marble precursor is the Early Cambrian to Early Ordovician Øynes conglomerate that erosively overlies the Leivset dolomite marbles. The homogeneous near-zero δ13C values and homogeneous and relatively high δ18O ratios are consistent with a high proportion of seawater in the dolomitising fluids. If such an inference is correct, then the least radiogenic 87Sr/86Sr ratios (0.70928–0.70939, n = 5) and the near-zero δ13C values suggest an Early to Mid Cambrian age (530–500 Ma) for the dolomitising fluids. This would not contradict the apparent depositional age of the unconformably overlying Øynes conglomerate which contains clasts of the underlying dolomite marbles, and whose other calcite marble clasts are not younger than 500 Ma (see section Øynes formation).

The Fjellengflåget formation The Fjellengflåget formation comprises diverse lithologies including grey and white, banded and massive calcite marbles, mica schists, minor dolomite marble and quartzite (# 15–21). Grey, massive, calcite marble (# 18), calcite-mica schist and garnet-mica schist (# 21) are the main lithologies of the formation. This rock assemblage occurs as an intensely folded, pinch-and-swell unit stretching from the northeastern part of the area through its central part to its southern limit. The calcite- and garnetiferous mica schist also occurs in the southeastern part of the mapped area. From the map it appears that at Nestbyfjellet, the Fjellengflåget marbles and schists may be structurally discordant to the Leivset formation marbles. At Ljøsnehammaren, the Fjellengflåget schist is overlain unconformably by the Øynes formation conglomerate.

The grey, massive, calcite marbles sampled at Nestbyfjellet are relatively pure rocks (SiO2 = 0.33– 2.7 wt.%; Al2O3 = 0.01–0.5 wt.%; n = 11; Electronic Supplement) with elevated Sr contents (580–1100 µg g-1, n = 11), moderate Mn/Sr ratios (0.05–0.38) and a low degree of dolomitisation (Mg/Ca = 0.01–0.06). δ13C values show a limited range around zero (-2.1 to +0.3‰, n = 7), whereas δ18O is high (24.6–29.2‰, with one outlier at 22‰) indicating a high degree of preservation. The latter inference is collaborated by the lack of significant correlation between δ13C and δ18O, and between these two parameters and the Mn/Sr and Mg/Ca ratios. 87Sr/86Sr ratios range between 0.70846 and 0.70907 (n = 5). The most radiogenic value has the highest Mn/ Sr and Mg/Ca ratios among samples analysed for Sr isotopes (Electronic Supplement), hence suggesting an

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alteration. If this single value is excluded, the remaining four samples are characterised by a rather limited 87 Sr/86Sr range of 0.70846–0.70856. Acknowledging the limitations of this database, this range is tentatively assumed to represent the best proxy for 87Sr/86Sr in seawater.

Projection of the least altered 87Sr/86Sr (0.70846– 0.70856) and δ13C (-2.1 to +0.3‰) ratios onto the seawater reference curves unfortunately provides no unique solution for the apparent depositional age. These two isotope systems best agree at two age intercepts, namely at Early Cambrian (c. 520 Ma) or at Early to Mid Ordovician (485–465 Ma) (Fig. 3). The carbonate rocks of the Fjellengflåget formation unfortunately remain understudied due to their limited occurrence and a general lack of well exposed, continuous sections. Consequently, our limited database is unable to discriminate between Early Cambrian and Early– Mid Ordovician ages. However, considering that the Fjellengflåget formation is overlain unconformably by the Øynes formation conglomerate, its depositional age cannot be younger than 470 Ma. The Øynes formation The Øynes formation is composed mainly of calcareous schist (# 12), which occupies a significant part of the mapped area and occurs as a wide, continuous unit from Dverset to Gullurda (Fig. 2). The next abundant lithology is a conglomerate occurring as a thick unit at Storvika and Kvanndalen (# 13, 14). From here, this thick conglomerate unit continues outside the mapped area farther to the north across Skjerstadfjorden to the type locality at Øynes (Fig. 1), and farther to the southwest just outside the mapped area. The conglomerate is less abundant in the central and southern parts of the study area where it occurs as numerous, thin, discontinuous beds and lenses at Ljøsnehammaren, and thin but continuous beds at Dverset and between the Gullurda and Nestbyfjellet areas (# 13, 14).

At Kvanndalen, the flat-lying polymict conglomerate lies erosively and unconformably (cf., Kollung & Gustavson, 1995) on steeply dipping calcareous mica schist of the Neoproterozoic (800–660 Ma) Kjerktinden formation (Fig. 12A; see Fig. 2 for the geological and geographic location of the photograph). Here, the conglomerate is also in primary stratigraphic contact with various marbles of the Ediacaran Leivset formation. In the Ljøsnehammaren, Dverset, Gullurda and Nestbyfjellet areas, the Øynes conglomerate is in contact with diverse rocks of different apparent depositional ages including Mid and Late Ediacaran, Leivset calcite marbles and Fjellengflåget garnet-mica schist. At Ljøsnehammaren, the Øynes conglomerate lies with a primary contact upon Fjellengflåget calcareous and amphibole-bearing mica schists, and Leivset formation dolomite and calcite marbles.

Isotope chemostratigraphy of high-grade marbles in the Rognan area 125

The Øynes conglomerate is both matrix- and clastsupported and in places is rhythmically interbedded with calcareous greywacke and mica schist (Fig. 12B–F).

The matrix is composed of either calcareous mica schist or calcite, whereas the clast composition can vary both laterally and vertically even within a single outcrop (e.g., Fig. 12D). At Kvanndalen, clasts include vein quartz, quartzite and subordinate dolomite and calcite marbles.

At Øynes, in the southern part of the peninsula, the conglomerate is mainly matrix-supported, and contains boulders (up to 80 cm in size) of quartzite and igneous rocks together with smaller pebbles of dolomite marble (Fig. 12B, E). In the northern part of the peninsula, the conglomerate is clast-supported and contains pebbles and cobbles of dolomite and pink calcite marble together with quartzite and amphibolite in a calcite matrix (Fig. 12D, G). In contrast, in the Ljøsnehammaren, Gullurda and Nestbyfjellet areas, the conglomerate contains clasts mainly of various dolomite and calcite marbles. This is also the case for a section exposed in the Løvgavlen quarry, just to the north of the mapped area. Here, the Øynes formation conglomerate (the Fauske carbonate conglomerate in Melezhik et al. (2000)) overlies the Leivset formation dolomite marble above an erosional contact (Fig. 12H). Although the Øynes conglomerate contains fragments of a dolomite marble derived from the eroded unit below, the bulk of the clasts (darkgrey or ‘blue’, white and pink calcite marbles; Fig. 12J, K) have been transported over a long distance from an unidentified source (Melezhik et al., 2000). Clastsupported conglomerates grade rapidly into gritstone and calcareous greywacke (Fig. 12L) both laterally and vertically.

Deposition of the carbonate conglomerate in the Løvgavlen quarry has been previously indirectly dated by means of isotope chemostratigraphy to a maximum age of c. 520 Ma (Melezhik et al., 2000). Sedimentological and structural studies have also shown that, at Løvgavlen, the dolomite marble and overlying carbonate conglomerate represent a rock assemblage which originated on the Laurentian carbonate shelf and continental slope and was later thrust upon nappes along the Baltoscandian margin of Baltica during the Baltica–Laurentia, Scandian collision in Siluro–Devonian time (Roberts et al., 2001, 2002).

In order to obtain a better insight into the provenance of carbonate clasts in the Øynes conglomerate, 28 new samples from marble pebbles and carbonate matrix were collected from different localities and analysed (Electronic Supplement). These analyses, together with 25 clast and matrix samples analysed previously from the Løvgavlen quarry (Melezhik et al., 2000), form the current database in our study of the provenance of the Øynes conglomerate. The analysed clasts are represented by approximately equal amounts of dolomite, variably calcitised dolomite and variably dolomitised calcite and calcite marbles (Fig. 13A); the calcite marbles include

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Figure 12. Photographs of selected rock types from the Øynes formation.

(A) Discordance between Neoproterozoic Kjerktinden calcareous schist in the foreground (dipping steeply to the left) and overlying polymict conglomerate of the Øynes formation, south of Kvanndalen. Mapping has shown this contact to be an unconformity (Fig. 2). (B) Natural exposure of clast-supported conglomerate; the clasts are of quartzitic sandstone (white) and dolomite marble (yellowish) in a calcareous schist matrix. (C) Natural exposure of matrix-supported conglomerate interbedded with gritty and silty greywacke; unsorted and unevenly distributed clasts are represented by arkosic and quartzitic sandstones (white) and dolomite marbles (yellow). (D) Natural exposure of clast-supported conglomerate; clasts are of quartzitic sandstones (pale grey), amphibolites (black), pink calcite marble and dolomite marble (yellow) in an

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 127

arkosic matrix. (E) Natural exposure of clast-supported conglomerate; unsorted clasts are foliated intermediate volcanic rocks (dark grey), quartzitic sandstones (bright) and dolomite marbles (yellow) in a greywacke matrix. (F) Natural exposure of gritty and silty greywacke interbedded with matrix-supported conglomerate containing unsorted and unevenly distributed clasts of arkosic and quartzitic sandstones (white) and dolomite marbles (yellow). (G) Natural exposure of clast-supported conglomerate; unsorted clasts are represented by foliated intermediate volcanic rocks (dark grey), white and beige calcite marbles, dolomite marbles (yellow) and minor quartzitic sandstones. (H) Pale yellow, banded dolomite marble of the Leivset formation below the base of the Ă&#x2DC;ynes formation carbonate breccias which pass upwards into poorly-sorted, carbonate conglomerate, and then into rhythmically bedded gritstone-calcareous greywacke; note that the irregular palaeorelief beneath the conglomerate (marked by red arrows) is affected by the tectonic deformation involving folding and a spaced cleavage; also note an apparent angular discordance between the basal contact (marked by red arrows) and the bedding plane of the uppermost bed of gritstone-greywacke rhythmites (marked by white arrows). (I) Chaotically deposited and unsorted blocks and angular fragments of dolomite (white) and calcite marbles in a calcareous schist matrix from a bed located c. 2 m above the contact with the underlying Leivset formation dolomite marble.

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Figure 12. Photographs of selected rock types from the Øynes formation. (Continued) (J) An inner corner of two subvertical walls of the quarry at almost 90˚ to each other demonstrating stretching of the pink calcite and white dolomite marble clasts. (K) Photograph illustrating that in some beds, fragments of pink calcite marbles are replaced by black and white calcite marbles suggesting a rapid switch in clast source. (L) Pale-grey and pink gritstone and small-pebble carbonate conglomerate interbedded with dark-grey greywacke. The calcareous greywacke rapidly replaces the carbonate conglomerate upwards in the succession, and becomes a major component in the deep basinal facies as documented in the Løvgavlen quarry; hammer handle for scale is 40 cm (arrowed).

Photographs B–G were taken along the northern shore of Skjerstadfjorden, east of Øynes. Photographs H–L were taken from near-vertical, smooth walls in the Løvgavlen quarry at Fauske. Photographs C–E, J and K are from Melezhik et al. (2000).

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 129

Figure 13. Various cross-plots illustrating the presence of different isotopic groups of marbles­in clasts of the Øynes formation conglomerate.

The pale-grey rectangle in (B) denotes the δ13C field that in Early Palaeozoic and Neo­ proterozoic time was unique for Ediacaran–Mid to Late Cryo­ genian seawater. The dark-grey rectangle in (E) denotes the 87Sr/86Sr field that in Early Palaeozoic and Neoproterozoic time was unique for Cambrian seawater. The pale-green rectangle in (E) denotes the 87Sr/86Sr field that in Early Palaeozoic and Neoproterozoic time was unique for Mid to Late Cryogenian seawater.

Sr contents in (A, C) obtained

by ICP–AES, and in (D) by isotope dilution.

high-Sr and moderate-Sr varieties. The carbonate matrix sampled in the Løvgavlen quarry and at Nestbyfjellet is a calcite marble with a Mg/Ca ratio of 0.02–0.16 (Electronic Supplement). In both localities, the carbonate matrix is characterised by a limited range of δ13C (-1.2 to +2.2‰, n = 12) and δ18O (19–23‰, n = 12) values, whereas the clasts show a larger range and a bimodal distribution in both δ13C and δ18O (Fig. 13B). The large mode includes both calcite and dolomite clasts and has δ13C values ranging between -1.9 and +5.7‰ and δ18O varying between 17.7 and 24‰. In contrast, the smaller mode has low δ13C (-5.4 to -4.8‰) and high δ18O (26.2– 27.6‰) (one sample from Kvanndalen and 3 samples from Ljøsnehammaren; Electronic Supplement). Three calcite marble clasts from Øynes and Nestbyfjellet with the most positive δ13C (+3.1 to +5.7‰) show the highest Sr content (Fig. 13C, D) and exhibit the least radiogenic 87 Sr/86Sr ratios (0.70741–0.70794), whereas the rest of the dolomite and calcite marble clasts have 87Sr/86Sr ratios varying between 0.70887 and 0.70980 (Fig. 13D, E). Two analysed samples of the calcite matrix show a narrower range (0.70910–0.70930; Electronic Supplement).

Assessing the degree of alteration of the C-, O-, and Sr-isotope systems in marble clasts is problematic

because they might have been derived from different sources with different depositional ages, and hence carried originally variable δ13C, δ18O and 87Sr/86Sr ratios which could, in turn, have been variably modified by diverse post-depositional processes. 87Sr/86Sr–δ13C and 87 Sr/86Sr–Sr cross-plots (Fig. 13D, E) indicate a negative correlation and thus are suggestive of alteration.

However, the correlation is driven by two different and unrelated subsets of samples and hence is unlikely to be significant. δ13C–Mn/Sr and Mg/Ca–Mn/Sr cross-plots corroborate this conclusion (Fig. 13F, G).

Despite all the obstacles involved in screening against post-depositional alterations, it becomes apparent that the two clasts with the lowest 87Sr/86Sr values (0.70741 and 0.70774) fall into the Mid to Late Cryogenian field (Figs. 3, 13E). If these two samples have been altered, then the apparent depositional age could be even older.

One sample with δ13C < -5‰, if primary, suggests that this clast has been derived from rocks deposited in Mid Cryogenian–Ediacaran time (Figs. 3, 13B). Interestingly, all three clasts with δ13C values ranging between -5.4 and -4.8‰ and δ18O = 26.2–27.6‰ are isotopically identical to the Late Ediacaran Leivset Subset 2 marble (Fig. 9 versus Fig. 13; Electronic Supplement).

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Figure 14. Various cross-plots illustrating apparent alteration trends in the Sr-isotopic systems in calcite marbles of the Rognan formation. The apparent alteration trends are indicated by pale-grey arrows.

The majority of the clasts and matrix material in the current study fall, however, within a unique 87Sr/86Sr Cambrian field (87Sr/86Sr = 0.7091–0.7094) within the 530–500 Ma time interval (FigS. 3C, 13E). Assuming that these measured values reflect the seawater signal, the obtained δ13C ratios also corroborate the presence of carbonate clasts originally deposited in Early to Mid Cambrian time (Fig. 13B).

The previous sedimentological and isotopic investigation did not suggest a unique solution for the apparent time of deposition of the Øynes formation conglomerate (475– 470, 510–505 or 520 Ma), although a slight preference was given to the maximum age of 520 Ma based on the least-altered, uniform δ13C values (Melezhik et al., 2000).

Even though the current work substantiates this result, both attempts have provided only the lower age limit for the deposition. Considering all the data, from both the current study and from the earlier work, and taking a conservative stance, the apparent depositional age of the carbonate rock now present as clasts range from the Mid to Late Cryogenian to Early Ordovician (Floian). Hence, the depositional age of the Øynes conglomerate cannot be older than Early Ordovician.

At this point it is pertinent to note that the Øynes carbonate conglomerate compares favourably with the redeposited limestone conglomerate and breccia debris sheets of the late Middle Cambrian to late Lower Ordovician Cow Head Group of the Appalachians of Newfoundland, Canada (Rodgers, 1968; James & Stevens, 1986; Knight et al., 1995; Roberts et al., 2001, 2002). As with the Øynes formation, the Cow Head conglomerates accumulated on the lower continental slope seaward of the edge of a wide carbonate platform, along the margin

of Laurentia. Several different facies of the Cow Head Group have been described (James & Stevens, 1986), many of which compare well with those documented for the Øynes formation by Melezhik et al. (2000).

The Rognan formation The Rognan formation is composed mainly of white, grey and dark-grey calcite marbles, with subordinate quartz schist, calcareous mica schist, graphite-bearing mica schist and dolomite marbles (# 5–11); all these lithologies occur in the eastern part of the map area (Fig. 2). In the southeastern area, the Rognan calcite marbles appear together with the less ambiguous, isotopicallylight, variegated and white marbles of the Leivset formation as a tectonically duplicated unit intruded by a large gabbro body (Fig. 2).

The currently available analytical database includes 38 samples obtained from grey and dark-grey marbles (# 11). The database includes 21 new samples and 17 samples published by Melezhik et al. (2008b). All analysed samples are calcite marbles (Mg/Ca = 0.006– 0.09) with low Al2O3 (<0.01–0.53 wt.%) and SiO2 (<0.01– 9.4 wt.%) contents (Electronic Supplement). The rocks have high Sr abundances (840 µg g-1 on average) and show a low Mn/Sr ratio (0.07 on average). Twenty-three samples contain organic matter (0.10–0.26 wt.% total organic matter) in the form of graphite.

Calcite marbles occurring at Brenne and Nestbyfjellet (hereafter Subset 1) have a highly positive δ13C (+4.2 to +8.0‰ with one outlier at +2.3‰; n = 24) and a moderate δ18O (17.2–25.1‰; Fig. 14A). In contrast, marbles occurring at Vensmoen (hereafter Subset 2)

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exhibit slightly positive δ13C (+0.7 to +2.2‰; n = 7) and high δ18O (25.7–28.0‰) values. However, among the Vensmoen marbles there is one sample that shows both δ13C (+7.1‰) and δ18O (19.3‰) values identical to those measured in marbles at Brenne and Nestbyfjellet (Fig. 14A), hence indicating that one marble group might have been infolded in another.

The strontium isotope ratios show no difference in the two subsets (Fig. 14B–E). If the most radiogenic outlier of 0.70861, which also has the greatest Mn/Sr of 0.15 (Fig. 14B), is excluded from further discussion, then the remaining eleven 87Sr/86Sr ratios range between 0.70805 and 0.70826 (Fig. 14B–E). No correlation can be seen between δ13C and δ18O and Mn/Sr and Mg/Ca ratios.

Consequently, we tentatively assume that all ranges observed in δ13C in both groups of marbles may represent a proxy for the seawater signal. Similarly, no significant correlation is observed between 87Sr/86Sr values and δ13C, Mn/Sr and Mg/Ca ratios, and Sr concentrations.

Weak alteration trends seen in Mn/Sr–, Mg/Ca– and Sr–87Sr/86Sr cross-plots (Fig. 14B, D, E) do not allow us to make any confident discrimination against postdepositional alterations. Hence, we tentatively assume that the entire range of 0.70805–0.70826 may reflect the seawater composition. If previously suggested geochemical criteria for the selection of the least altered Sr-isotope ratios in high-grade metamorphic marbles (Mg/Ca ≤0.02, Mn/Sr ≤0.02, Sr >1000; Melezhik et al.

(2003)) are routinely used, then a somewhat narrower range of 0.70805–0.70818 can be considered to represent the proxy for the seawater signal.

The grey and dark-grey calcite marble of the Rognan formation was examined in our previous investigation, and the Nestbyfjellet section was considered as one of the type sections in chemostratigraphic research in the Norwegian Caledonides (Melezhik et al., 2008b). It has been shown that this marble forms part of the ‘Tripartite Unit’ that occurs discontinuously over a distance of 450 km throughout the Uppermost Allochthon in NorthCentral Norway. The 13C-rich, grey calcite marble of the Rognan formation forms the structural base of the Tripartite Unit, whereas the 13C-depleted, variegated and white calcite marbles of the Leivset formation constitute its structural top (Fig. 10). Deposition of the 13 C-rich, grey calcite marbles has been constrained to the Llandovery (Melezhik et al., 2008), whereas the deposition of the structurally overlying, 13C-depleted, Leivset formation was linked to the Shuram–Wonoka isotopic event occurring between 580 and 550 Ma. The contact between the Rognan and the overlying Leivset marble is sharp. It represents a prominent, cryptic, stratigraphic discontinuity and suggests that the 13 C-depleted, Ediacaran Leivset marble was tectonically juxtaposed above the Llandovery-age, 13C-rich marbles during the Scandian orogeny (Melezhik et al., 2008b).

The new

Sr/86Sr data (0.70805–0.70826) obtained in

87

Isotope chemostratigraphy of high-grade marbles in the Rognan area 131

the mapped area are in agreement with the previously measured values (0.70807–0.70827) and fit with Early Silurian (Llandovery) time (Fig. 3). Comparison with the Phanerozoic δ13C reference curve indicates that the δ13C of both Rognan subsets (near-zero and 13C-rich) would agree with the C-isotopic composition of Early Silurian seawater that is characterised by a wide fluctuation in δ13C (Fig. 3). The Rognan formation represents the youngest rocks in the study area, and its Llandovery depositional age seems to be confidently constrained.

Apparent depositional ages of marble formations in the Rognan area and their implications Stratigraphic implications The mapped area covers mainly the Fauske Nappe of Nicholson (1974), the rocks in which were considered to be of Cambro–Silurian age. The carbon and strontium isotope data presented here strongly suggest that the Fauske Nappe consists of several marble formations, which were deposited over a wide period of time from c. 800 to 440 Ma. The youngest, Llandovery-age marbles occur close to the structural base of the nappe, whereas the oldest unit, the Kjerktinden marble, of Cryogenian (800–660 Ma) age, occurs at its structural top (Fig. 15).

The Fauske Nappe includes several formally established groups of which only the Kjerketinden, Rognan and Øynes, all originally assigned to the Cambro-Silurian, occur within the mapped and studied area (Fig. 15).

The chemostratigraphic study suggests that the marbles previously assigned to the Kjerketinden Group have an apparent Cryogenian depositional age, whereas the Øynes rocks were apparently deposited in Early Ordovician time.

It appears that the former, undivided, Cambro-Silurian Rognan Group of Gustavson (Gustavson, 1996) is composed of several units with different apparent depositional ages. The youngest marbles, which have been informally assigned in this study to the Rognan formation, have an apparent depositional age of c. 440 Ma and form the structural base of the original Rognan Group (Fig. 15). The intermediate structural level is occupied by the Leivset and Fjellengflåget formations, which contain the main volume of marbles in the study area. As suggested by the 87Sr/86Sr and δ13C data, some of these Leivset calcite marbles (variegated and overlying white marbles, Subset 1, # 33, 32) were originally deposited in the Mid Ediacaran (600–580 Ma). The time of deposition of the two remaining lithostratigraphic units (with distinctly different but homogeneous 87Sr/86Sr and δ13C values) within the Leivset formation can be constrained to c. 560–550 Ma. The calcite marbles of the Fjellengflåget formation have an apparent depositional age of c. 520 Ma (Fig. 15). The upper structural level of the Cambro-Silurian Rognan Group of Gustavson

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Figure 15. A simplified tectonostratigraphic column for the Fauske Nappe in the Rognan area.

The suggested depositional ages are based on C- and S-isotope chemostratigraphy, details of which are given in the text. Several thrusts are required in order to reconcile the observed disparities in age between several of the formations. The unconformity at the base of the Øynes formation cuts across almost all formations, the exception being the Llandovery Rognan formation.

(1996) includes Cryogenian (800–660 Ma), and Mid (600–580 Ma) and Late (c. 550 Ma) Ediacaran, calcite and major dolomite marbles (Fig. 2). These are unconformably overlain by the Early Ordovician, Øynes formation conglomerate with marble clasts of Mid–Late Cryogenian to Early Ordovician age.

Tectonic implications Accepting the age assignments suggested by isotope chemostratigraphy for the carbonate formations, it follows that at least the Rognan Group of Gustavson (1996) does not represent a coherent stratigraphic unit, and is instead composed of tectonically imbricated marble formations of diverse age groups that were

emplaced in a non-chronostratigraphic order (Fig. 15). The oldest marble unit (800–660 Ma), constituting part of the Kjerktinden formation, has its structural base above the Leivset formation marble (# 27) which has an apparent depositional age of 550 Ma. Hence, the contact between these two formations represents either a thrust or a hiatus in a structurally inverted succession.

Since the inferred non-depositional break should have an apparent duration of over 100 Myr, as based on the chemostratigraphic depositional ages, this is a very unlikely scenario. Consequently, a thrust contact is proposed to reconcile the observed age disparity.

A major thrust contact should also be invoked in order to reconcile the apparent disparity in age between

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the structurally lowest Rognan formation (c. 440 Ma, Llandovery) and the overlying Leivset marbles (600–580 Ma, Ediacaran). This prominent cryptic discontinuity, marked by a sharp break in the C-isotopic record (Fig. 10), has been identified earlier in several places in the North-Central Norwegian Caledonides by means of isotope chemostratigraphy. Such a discontinuity has been resolved by inferring tectonic juxtaposition of the Ediacaran Leivset marbles above Llandovery-age rocks during the Scandian orogeny (Melezhik et al., 2008b).

Additional tectonic contacts should also be considered in order to reconcile the apparent disparity in age between the Fjellengflåget formation (c. 520 Ma) and structurally subjacent (600–580 Ma) and overlying (c. 560 or 530– 520 Ma) marbles of the Leivset formation (Fig. 15).

Consequently, all distinctive chemostratigraphic units considered in this article as ‘formations’, except for the Øynes formation, could well be redefined as thrust sheets.

Geological relationships indicate that most of the proposed tectonic contacts were comparatively early phenomena in the structural history of the region, with the exception of the thrust at the top of the Rognan formation, i.e., thrusting occurred prior to the unconformable deposition of the Early Ordovician Øynes formation conglomerates, and that all these faults were affected by later episodes of deformation and associated metamorphic recrystallisation. Because marbles, in general, have an extremely high propensity for being recrystallised during the latest tectonometamorphic episodes, they rarely retain fabrics associated with earlier thrusting or metamorphic events. This could, perhaps, have been a reason why the layer-parallel, conjunctive fault contacts were not recognised during previous mapping and structural investigations. In the study area, and as discussed earlier for the Ofotfjorden area, the tectonic emplacement of Neoproterozoic formations upon rocks of Early Silurian age can readily be related to Scandian orogenic deformation. However, the tectonic processes that resulted in the juxtaposition of Cryogenian and Cambrian rocks may well relate to an early deformation event in the Late Cambrian to Late Ordovician, Taconian orogenic cycle (Roberts et al., 2001, 2002; Melezhik et al., 2002a; Yoshinobu et al., 2002; Barnes et al., 2007).

In the Løvgavlen quarry near Fauske, for example, NW-vergent folds and thrusts in the Øynes formation have been suggested to relate to Taconian orogenesis. Accepting an Early Ordovician age for the Øynes formation, and as the Early Cambrian Fjellengflåget formation is also delimited by thrusts, then most of the thrusts cut by the basal Øynes unconformity must have been generated at some stage between c. 520 and c. 475 Ma. In Newfoundland, the ‘Taconic-2’ orogenic event first affected the carbonate shelf and adjacent slope in earliest

Isotope chemostratigraphy of high-grade marbles in the Rognan area 133

Tremadocian time, with imbricate thrusting resulting from loading of the margin by NW-transported oceanic lithosphere (Waldron & van Staal, 2001; van Staal et al., 2009; van Staal & Barr, 2012). It is therefore highly likely that the pre-Øynes thrusting recorded in the Rognan area is also of earliest Ordovician, Taconic-2 age. The NW-vergent thrusts and folds documented from near Fauske clearly postdate deposition of the Øynes formation. They may date either to a late phase of Taconic-2 or to a Mid Ordovician component of Taconian accretion.

In summary, the detailed geological mapping in combination with the isotope study suggest that the Fauske Nappe preserves a complex, tectonically dissected package of rocks. The isotope data have helped to provide a crude chronostratigraphic subdivision of the marble units and thus allowed us to produce a chronologicallybased geological map (Fig. 2), acknowledging the apparent limitations of indirect dating by employing the chemostratigraphic approach.

Palaeogeographic implications Within the mapped area and farther to the north, near Fauske, there are a few rare localities where primary erosional contacts are preserved. In one such locality, at Fauske, rocks of the Øynes formation (the Fauske carbonate conglomerate in Melezhik et al. (2000a)) lie unconformably upon dolomite marble of the Leivset formation. Sedimentological research revealed channelling and cross-bedding in the Øynes formation which indicate southeast-directed palaeocurrents and transport of carbonate clasts, and accumulation of the debris on a basinal slope deepening to the southeast (Melezhik et al., 2000a), i.e., in an opposite direction with respect to the Baltoscandian margin of Baltica. This anomalous situation, foreign to Baltica, together with the evidence of NW-directed thrusting, has been interpreted to indicate a Laurentian palaeogeographic ancestry for the rocks of the Fauske Nappe (Roberts et al., 2001, 2002, 2007).

Another case of a preserved, primary erosional contact has been documented within the study area, at Kvanndalen (Fig. 12A). Here, the flat-lying Øynes polymict conglomerate lies erosively upon steeply dipping, calcareous mica schist of the Kjerktinden formation of Cryogenian age (Fig. 12A). At Kvanndalen and Ljøsnehammaren, this same conglomerate rests with an apparent unconformity upon various marbles of the Leivset formation of Ediacaran age (600–550 Ma).

In these areas, the Øynes conglomerate contains clasts showing identical C- and O-isotopic compositions to those of the c. 550 Ma Leivset calcite marbles. It appears that the Laurentia-derived Øynes formation shows erosive contacts with all marble units except the Early Silurian Rognan formation. Since the Øynes formation conglomerate at Fauske was identified as part of the

134 V.A. Melezhik et al

shelf break and continental slope of the Laurentian margin (Roberts et al., 2001, 2002), then the entire tectonically-stacked, marble-schist thrust complex with the unconformably overlying, Early Ordovician, Øynes conglomerates was detached from its Laurentian roots and transported into the higher levels of the Caledonian orogenic wedge during the Scandian orogeny. The provenance of the Llandovery-age marbles of the Rognan formation is less definitive, simply because they are nowhere in juxtaposition with the Øynes formation.

However, in view of the preponderance of marbles in the tectonostratigraphy of the Fauske Nappe, we surmise that the Rognan marbles, too, were formed as limestones and dolomites along the eastern margin of Laurentia. In this case, however, they probably accumulated on the shelf of a successor basin, postdating the Taconian orogeny.

Interpretation of the marble-dominated tectono­ stratigraphy of the Fauske Nappe in the study area, and in much of the Uppermost Allochthon in other parts of Nordland and Troms (e.g., Melezhik et al., 2000, 2003; Roberts et al., 2002, 2007), as having originated along the Laurentian margin in Cryogenian to Ordovician time invites a brief comparison with successions in other parts of the Caledonide–Appalachian orogen.

The Dalradian Supergroup of the Scottish and Irish Caledonides is one such macro-unit, ranging in age from Early Cryogenian (and possibly latest Tonian) to Early Ordovician, and deformed in the Taconian-equivalent, Grampian orogeny (Oliver, 2001). Several Cryogenian to Ediacaran limestone formations are recorded, with ages now supported by δ13C chemostratigraphy (Thomas et al., 2004; Prave et al., 2009a, b). Moreover, the c. 550 Ma Shuram-Wonoka, negative δ13C carbonate excursion has also been registered (Prave et al., 2009b), as in Norway (Melezhik et al., 2008b). Three glaciogenic units in the Dalradian succession are now regarded as equating with the worldwide Sturtian (750–690 Ma), Marinoan (c. 630 Ma) and Gaskiers (580 Ma) glacial events. In a palaeogeographic context, the Scottish–Irish segment of the Laurentian margin is inferred to have been located off Southeast Greenland (Leslie et al., 2008). In the East Greenland Caledonides, the Eleonore Bay Basin records deposition of the 14 km-thick, Cryogenian–Ediacaran, Eleonore Bay Supergroup which includes a thick succession of limestones and dolostones representative of an extensive carbonate platform (Sønderholm et al., 2008). The supergroup succession is overlain by diamictites of the ‘Tillite Group’, which are now considered as representatives of the Marinoan glaciations (Kristiansen, 2007). A return to carbonate platform conditions then characterised Cambrian time. Taking an overall view of the chrono- and tectonostratigraphy of the East Greenland and Scottish–Irish Caledonides and Canadian Appalachians, these regions clearly have much in common, and also show several similarities with features described here from the Fauske Nappe and from other parts of the Uppermost Allochthon. There are also differences, ascribed partly to segmentation of the

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Laurentian continental margin into several promontories and reentrants (e.g., Lavoie et al., 2002) resulting in diachroneity of depositional and deformation events, but these features are outside the subject of our present contribution.

Implications for mineral exploration There are five commercially exploited dolomite deposits in Norway, all located in the county of Nordland. Within the study area, the Fauske and Hammerfall deposits, hosted by the Leivset formation, are considered to be the most important (Karlsen, 1998). The major and thick dolomite strata and lenses are located immediately below the prominent erosional surface (Fauske) or very close to it (Ljøsenhammaren), and overlain unconformably by the Øynes formation conglomerates and schists. Such a stratigraphic position of the commercially exploited dolomite marble units suggests that the regionalscale exposure surface and long-lasting period of emergence might have played an important role in their formation; e.g., dolomitisation of limestones by Mg-rich fluids. The stratigraphic position of, and inferred palaeoenvironmental constraint for, the major dolomite units provide prospectors with exploration criteria, namely that the search for new dolomite deposits of the Fauske–Hammerfall type in the Rognan–Fauske area should be restricted to the c. 550 Ma Leivset marble unit located below the unconformity at the base of the Øynes conglomerate.

Conclusions Detailed geological mapping in the Fauske Nappe in the Rognan area of Nordland, combined with isotope chemostratigraphy of non-fossiliferous, high-grade, polydeformed marble formations, has demonstrated progress towards producing a new generation of geological maps in dissected metamorphic terranes.

Although depositional age constraints derived by Cand Sr-isotope stratigraphy are comparatively imprecise, the method has now been tested sufficiently widely in different orogenic belts worldwide, also on fossiliferous units of known age, to demonstrate its high potential for quantitatively-based chronostratigraphic subdivision and for geological correlation of high-grade marble successions.

Concerning the present study, the thick marble and siliciclastic succession of the Rognan area had previously been considered to represent a stratigraphically homogeneous assemblage of Cambro–Silurian age, spanning over 100 million years of geological time.

However, this interpretation is inconsistent with the isotopic evidence presented here for diverse Cryogenian to Early Silurian depositional ages for these thick marble units, which now occur as polydeformed thrust sheets stacked in a complex imbricate manner.

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The newly obtained Sr- and C-isotopic data together with previously performed sedimentological, structural and analytical work, in conjunction with an assessment of carbonate platform successions elsewhere in the Caledonide–Appalachian orogen, supports the earlier suggestion of a Laurentian ancestry for the tectonostratigraphy of the Fauske Nappe. Diverse marble units within the nappe range in age from Early Cryogenian to Cambrian and are dissected by thrusts in a non-chronostratigraphic order. The package of thrust sheets is overlain unconformably by shelf-edge and continental-slope breccias and conglomerates of the Øynes formation, of apparent Early Ordovician age.

Both the pre-Øynes thrusting and early NW-vergent folds in the Øynes formation are considered to date to the second and main phase of the Taconian orogeny, along the Laurentian margin. Just one formation - the Rognan formation - is younger, of Early Silurian age, with Sr- and C-isotopic data that are comparable to those in fossiliferous, Early Silurian, limestones farther north in Troms. These particular carbonate rocks are considered to have accumulated in a post-Taconian successor basin, prior to their ultimate incorporation during the Scandian orogeny into the Uppermost Allochthon of the Norwegian Caledonides.

Acknowledgements. The fieldwork for these investigations was supported by the Geological Survey of Norway (NGU) (Projects 304800 and 291200). The laboratory work and isotope study was financed by NGU (Project 291200), the Scottish Universities Environmental Research Centre, the Earth Sciences Branch of the Russian Academy of Sciences (Programme No. 4), and the Russian Foundation for Basic Research (Project 13–05–01059; ABK). The help of G.V. Konstantinova, E.P. Kutyavin and N.N. Melnikov in the Rb–Sr analytical work is greatly appreciated. T.L. Turchenko provided the results of the XRD analysis for the siliciclastic constituents of the carbonate rocks. The C- and O-isotope analyses were performed at the Scottish Universities Environmental Research Centre supported by the Consortium of Scottish Universities and the Natural Environment Research Council.

A special round of thanks go to Appalachian colleagues John Waldron, Ian Knight, Cees van Staal, Denis Lavoie, Michel Malo and Jon Kim for their patience in providing information and new references on the Laurentian, Early Palaeozoic carbonate platform and adjacent slope. We acknowledge a constructive review of the manuscript by Calvin Barnes, and the Editor of the journal, Trond Slagstad.

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Isotope chemostratigraphy of high-grade marbles in the Rognan area 139

Appendix Analytical techniques.

Major and trace elements were analysed by X-ray fluorescence spectrometry at the Geological Survey of Norway (NGU), Trondheim, using a PANalztical Axios at 4 kW X-ray spectrometer. The precision (1σ) is typically 2% of the major oxide present. Acid-soluble Fe, Ca, Mg and Mn were analysed by ICP-AES at NGU using a Thermo Jarrell Ash ICP 61 instrument. Detection limits for Fe, Mg, Ca and Mn are 5 ppm, 100 ppm, 200 ppm and 0.2 ppm, respectively. The total analytical uncertainty including element extraction (1σ) is ± 10% rel.

Oxygen and carbon isotope analyses of wholerock marble samples were carried out at the Scottish Universities Environmental Research Centre, Glasgow, using the phosphoric acid method of McCrea (1950) as modified by Rosenbaum & Sheppard (1986) for operation at 100°C. Carbon and oxygen isotope ratios in carbonate constituents of the whole-rock samples were measured on a VG SIRA 10 mass spectrometer. Analyses were calibrated against NBS 19, and precision (1σ) for both isotope ratios is better than ± 0.2‰. Oxygen isotope data for dolomites were corrected using the fractionation factor 1.00913 recommended by Rosenbaum & Sheppard (1986). The δ13C data are reported in per mil (‰) relative to V-PDB and the δ18O data in ‰ relative to V-SMOW.

Rb-Sr analyses were carried out at the Institute of Precambrian Geology and Geochronology of the Russian Academy of Sciences, St. Petersburg, as specified in Gorokhov et al. (1995). The Rb and Sr concentrations were determined by isotope dilution. Sr concentrations obtained by isotope dilution are systematically 15% higher compared with those determined by ICP-AES at NGU. Rb isotopic composition was measured on a single-collector MI 1320 mass spectrometer. Strontium isotope analyses were performed in static mode on a multi-collector Finnigan MAT-261 mass spectrometer.

All 87Sr/86Sr ratios were normalised to a 87Sr/86Sr of 0.1194 and measurements of the NIST SRM-987 run with every batch averaged 0.710255 ± 5 (2σmean, n=36). During the course of the study, the value obtained for the 86Sr/88Sr ratio of the U.S.G.S. EN-1 standard was measured at 0.709187 ± 6 (2σmean, n = 10).

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Facies analysis and depositional environments of a storm-dominated, temperate to cold, mixed siliceous–carbonate ramp: the Permian Kapp Starostin Formation in NE Svalbard .................................................... 75

Anette E.S. Högström, Sören Jensen, Teodoro Palacios & Jan Ove R. Ebbestad New information on the Ediacaran–Cambrian transition in the Vestertana Group, Finnmark, northern Norway, from trace fossils and organic-walled microfossils ........................................................................... 95

Victor A. Melezhik, David Roberts, Svein Gjelle, Arne Solli, Anthony E. Fallick, Anton B. Kuznetsov & Igor M. Gorokhov Isotope chemostratigraphy of high-grade marbles in the Rognan area, North-Central Norwegian Caledonides: a new geological map, and tectonostratigraphic and palaeogeographic implications ............. 107

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