BasinResearch The Marinoanglaciation …late Neoproterozoic Polarisbreen Group (Fig.1) sits atop 2km of conformable carbonate strata, the Akademikerb-reenGroup,andwithin its mixed

TheMarinoan glaciation (Neoproterozoic) innortheast SvalbardGalen P. Halverson, Adam C. Maloof and Paul F. Hoffman

Department of Earth andPlanetary Sciences, Harvard University, Cambridge, Massachusetts, USA

ABSTRACT

Two separate and distinct diamictite-rich units occur in the mixed carbonate-siliciclasticPolarisbreenGroup, which comprises the top kilometer of47 km of Neoproterozoic strata in thenortheast of the Svalbard archipelago.The platformal succession accumulated on the windward,tropical to subtropical margin of Laurentia.The older PetrovbreenMember is a thin glacimarinediamictite that lacks a cap carbonate. It contains locally derived clasts and overlies a regional karsticdisconformity that was directly preceded by a large (410%) negative d13C anomaly in the underlyingshallow-marine carbonates.This anomaly is homologous to anomalies in Australia, Canada andNamibia that precede theMarinoan glaciation.The younger and thickerWilsonbreen Formationcomprises terrestrial ice-contact deposits. It contains abundant extrabasinal clasts and is draped by atransgressive cap dolostone 3^18m thick.The cap dolostone is replete with sedimentary featuresstrongly associatedwith post-Marinoan caps globally, and its isotopic pro¢le is virtually identical tothat of otherMarinoan cap dolostones. From the inter-regional perspective, the two diamictite-richunits in thePolarisbreenGroup should represent the ¢rst and ¢nal phases of theMarinoan glaciation.Above the Petrovbreen diamictite are �200m of ¢nely laminated, dark olive-coloured rhythmites(MacDonaldryggenMember) interpreted here to represent suspension deposits beneath shorefast,multi-annual sea ice (sikussak).Above the suspension deposits and below theWilsonbreen diamictitesis ao30-m-thick regressive sequence (SlangenMember) composed of dolomite grainstone andevaporitic supratidal microbialaminite.We interpret this sabkha-like lagoonal sequence as an oasisdeposit that precipitatedwhen local marine ice melted away under greenhouse forcing, but while thetropical ocean remained covered due to in£ow of sea glaciers from higher latitudes. It appears that thePolarisbreenGroup presents an unusually complete record of theMarinoan snowball glaciation.

INTRODUCTION

In the northeastern part of the Svalbard archipelago, aCaledonian (Silurian) fold belt exposes a magni¢cent 7-km-thick succession of Neoproterozoic through MiddleOrdovician sedimentary strata (Fig. 1). The Hecla HoekSuccession (Harland &Wilson, 1956; Knoll & Swett, 1990;Harland etal.,1993) is similar to that in the EastGreenlandCaledonides (Hambrey & Spencer, 1987; S�nderholm &Tirsgaard, 1993; Fairchild & Hambrey, 1995), and theyjointly formed a platform contiguous with the windward,tropical to subtropical margin of Laurentia (Fig. 1). Thelate Neoproterozoic Polarisbreen Group (Fig. 1) sits atop2 km of conformable carbonate strata, the Akademikerb-reen Group, and within its mixed terrigenous and carbo-nate sequences are two separate and distinct glacialdiamictite-rich units (Harland et al., 1993).

It is nowwidely accepted that at least one largeNeopro-terozoic ice sheet reached tidewater close to the equator(Embleton & Williams, 1986; Schmidt & Williams, 1995;

Sohl et al., 1999; Evans, 2000) ^ a climatic scenario thatcannot be understood in terms of Quaternary analogues.Invoking various lines of geological and geochemical evi-dence, as well as simple energy-balance climate models(e.g. Budyko, 1969; Sellers, 1969; Caldeira & Kasting,1992), Kirschvink (1992), Klein & Beukes (1993) and Ho¡-man et al. (1998) proposed that the entire ocean froze overduring the Sturtian andMarinoan ice ages.

Much attention focused on marine sedimentary andisotopic proxy records of the conditions under whichlow-latitude glaciation was initiated (Halverson et al.,2002; Schrag et al., 2002) and later terminated (Fairchild,1993; Grotzinger & Knoll, 1995; Kennedy, 1996; Ho¡manet al., 1998; Kennedy et al., 1998; James et al., 2001; Ho¡man& Schrag, 2002; Higgins & Schrag, 2003). In three of thebest-studied regions, the Adelaide rift complex in SouthAustralia, the northern Canadian Cordillera, and the Con-go cratonic margin in Namibia, the deposits associatedwith the younger (Marinoan) glacials are directly under-lain by an unique carbon isotope anomaly (McKirdy et al.,2001; Halverson et al., 2002; Ho¡man & Schrag, 2002), andare directly overlain by a lithologically unique cap-carbo-nate sequence (Kennedy, 1996; So¡er, 1998; James et al.,2001;Ho¡man&Schrag, 2002).The older glacial sequence

Correspondence: Galen P. Halverson, Department of Earth andPlanetary Sciences, Harvard University, 20 Oxford St., Cam-bridge,MA 02138, USA. E-mail: [email protected]

BasinResearch (2004) doi: 10.1111/j.1365-2117.2004.00234.x

r 2004 Blackwell Publishing Ltd 1

has no similar premonitory isotope anomaly and itscap-carbonate sequence is isotopically and lithologicallyunlike the Marinoan cap (Kennedy et al., 1998). Accord-ingly, pre- and post-glacial strata in combination shouldo¡er a more reliable basis for correlation than the glacialdeposits alone, which are usually discontinuous and spa-tially variable.

Here we present detailed stratigraphic and geochemicaldata spanning what previously were regarded as two dis-crete glaciations in the Polarisbreen Group (Hambrey,1982; Knoll et al., 1986; Harland et al., 1993; Fairchild &Hambrey, 1984, 1995; Kaufman et al., 1997). The glacialunits are di¡erent lithologically and the younger one (Wil-sonbreen Formation) has a cap-carbonate sequence (basalDracoisen Formation) indistinguishable from other Mar-inoan cap carbonates. On the other hand, the sub-Mar-inoan isotope anomaly consistently appears beneath theolder glacial unit (Petrovbreen Member) and never be-neath the younger glacials.Two interpretations are possi-ble. Either the correlation criteria developed in Australia,Canada and Namibia are faulty, or both glacial units inSvalbard plus the intervening strata (including open-waterlagoonal deposits) belong to a single Marinoan glaciation.

In this paper, we explore the latter interpretation and at-tempt to reconcile it with evidence from elsewhere thattheMarinoan glaciation was a snowball cycle (Kirschvink,1992; Ho¡man et al., 1998; Ho¡man & Schrag, 2002). Webrie£y discuss the implications of our proposed correla-tions for the Sturtian glaciation in Svalbard.

GEOLOGICAL BACKGROUND

The Polarisbreen Group (Fig.1) comprises the mixed sili-ciclastic-carbonate Elbobreen, Wilsonbreen and Dracoi-sen Formations in the uppermost Neoproterozoic of theHecla Hoek Succession of northeastern Spitsbergen (Wil-son & Harland, 1964; Fairchild & Hambrey, 1984). It con-formably overlies a thick carbonate platform, theAkademikerbreenGroup, which grades downward into si-liciclastic strata of the underlyingVeteranenGroup (Fig.1).A cryptic unconformity separates the Polarisbreen Groupfrom the CambrianTokammane Formation (Harland etal.,1993), but biostratigraphy suggests a signi¢cant hiatus(Knoll & Swett, 1987).The Polarisbreen Group is exposedin a north-south trending belt that extends from southern

Fig.1. Generalized lithostratigraphy of the Proterozoic successions in northeast Svalbard (Harland et al., 1966; Fairchild &Hambrey,1984; Knoll & Swett, 1990; this paper) and East Greenland (Katz, 1960; Henriksen &Higgins, 1976; Hambrey & Spencer, 1987;Herrington & Fairchild, 1989; S�nderholm &Tirsgaard, 1993; Frederiksen et al., 1999; Smith & Robertson, 1999a). Correlations betweenEast Svalbard and East Greenland are based on original lithological (e.g. Harland et al., 1993; Fairchild &Hambrey, 1995) and carbon-isotopic (e.g. Knoll et al., 1986) comparisons from the literature. Black regions denote glacial diamictites.The maximum age of 950Mafor theVeteranenGroup is based onU-Pb dating of detrital zircons fromNE Spitsbergen (Larionov et al., 1998) and igneous zirconsfrom sub-VeteranenGroup granites in Nordaustlandet (Gee et al., 1995).The 780Ma age for the Svanbergfjellet-Draken formations is aloose approximation based on biostratigraphic correlations with Australia (Hill et al., 2000) and an unpublished Rb-Sr isochron(S. Jacobsen, pers. comm.) on the middle Grusdievbreen Formation.The 575Ma minimum age for the top of the succession is based onthe absence of Ediacaran fossils in Svalbard and East Greenland (Knoll & Swett, 1987) and an age of �575Ma for the oldest knownEdiacaran fossils, in Newfoundland (Bowring et al., 2003).

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G. P. Halversonet al.

Olav V Land in Spitsbergen to western Nordaustlandet(Fig. 2). A di¡erent stratigraphic nomenclature was de-¢ned in Nordaustlandet, but as correlation with Spitsber-gen is unambiguous (Harland & Gayer, 1972; Knoll et al.,1986; Fairchild & Hambrey, 1995), the Spitsbergennomenclature (Fig. 1) will be used throughout this paper.Three informal names are proposed for establishedmembers.

The older glacigenic unit in the Polarisbreen Group isthe Petrovbreen Member (Harland et al., 1993) of theElbobreen Formation (Fig. 3).Younger and typically muchthicker glacigenic diamictites make up the WilsonbreenFormation (Fig. 3).The glacial units have been studied se-dimentologically (Chumakov, 1968; Hambrey, 1982; Fair-child & Hambrey, 1984; Harland et al., 1993) andgeochemically (Fairchild & Spiro, 1987, 1990; Fairchildet al., 1989).They are often correlatedwith Neoproterozoicdiamictite pairs in other parts of theNorthAtlantic region(e.g. Hambrey, 1983; Nystuen, 1985; Fairchild & Hambrey,1995), but geochronological con¢rmation is lacking. Thecorrelation with theTillite Group in East Greenland (Fig.1) is the most widely accepted (Kulling, 1934; Harland &Gayer, 1972; Knoll et al., 1986; Hambrey & Spencer, 1987;Fairchild & Hambrey, 1995). The Neoproterozoic succes-sions in eastern Svalbard and East Greenland are similarin their entirety (Fig. 1) and there is consensus that theyformed in contiguous sedimentary basins. Knoll et al.(1986) produced an early record of the secular variation ofd13C in carbonate and organic matter carbon in the twosuccessions, which supported the correlation and drew at-tention to the connection between Neoproterozoic glacia-tions and large negative d13C anomalies. Fairchild & Spiro(1987) and Kaufman et al. (1997) have since provided addi-tional isotopic data for PolarisbreenGroup carbonates, in-cluding stromatolitic limestones within the Wilsonbreenglacials.The data presented here con¢rm and build uponthis previous work.

Tectonic setting

The Svalbard archipelago comprises three tectonic ter-ranes that were juxtaposed through sinistral transpressionduring theSilurian^DevonianNyFriesland orogeny (Har-land & Gayer, 1972; Harland et al., 1992; Gee & Page, 1994;Lyberis &Manby,1999;Dewey&Strachan, 2003).The He-claHoek Succession (Neoproterozoic-MiddleOrdovician)underlies the northern part of the eastern terrane, whichincludes parts of the Barents Sea and is called Barentsia.The correlation of the Hecla Hoek with the succession inthe southern East Greenland Caledonides (Fig. 1) under-pins the consensus view that Barentsia originated on theLaurentian margin adjacent to East Greenland (Harland& Gayer, 1972), forming part of a joint East GreenlandÊast Svalbard (EGES) platform.The origin of the EGESplatform is obscure, but presumably itwas associatedwiththe fragmentation of Rodinia, which in the region of theEGES platform involved the separation of four cratons

Fig. 2. Geological map of the north-south trendingNeopro-terozoic outcrop belt in northeastern Svalbard (OlavV Land, NyFriesland, and northwestern Nordaustlandet). Numbered boxeskey measured sections in Figs 4 and 8. See the Supplementarymaterial for location names and coordinates. ESZ5 EolusslettaShear Zone; LFZ5Lomfjorden Fault Zone. Location names:(1) Backlundtoppen, (2) Slangen-Ormen, (3) Klofjellet, (4)Ditlov-toppen, (5) Dracoisen, (6) Liljequisth�gda, (7) Sveanor, (8) S�reRuss�ya, (9) Kinnvikka, (10) Br�vika and (11) Skyttelodden.

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Marinoan glaciation in Svalbard

including Laurentia, Baltica, and Amazonia (Torsvik et al.,1996; Hartz & Torsvik, 2002).

The north^south trending outcrop belt in northeastSvalbard (Fig. 2) is believed roughly to parallel the EastGreenland cratonic margin (Knoll &Swett,1990; Fairchild& Hambrey, 1995). Following the siliciclastic VeteranenGroup (44 km), a broad carbonate platform (Akademi-kerbreen Group) developed (Fig.1).The conformable nat-ure of the 2-km-thick carbonate succession suggests broadthermal subsidence, and the paucity of terrigenous inputimplies either a subtropical location with little runo¡ or asemi-isolated platform like the Bahamas. The strata thinnorthward, particularly in the upper Svanbergfjellet andDraken Formations (Wilson, 1961), consistent with asoutherly opening basin (Hartz &Torsvik, 2002). Renewedcrustal stretching is inferred from deepening of the plat-form in East Greenland (Herrington & Fairchild, 1989)and an in£ux of ¢ne terrigenous sediment (KinnvikkaMember, Fig. 1) in the uppermost Backlundtoppen For-mation (Fairchild &Hambrey,1995), althoughwe postulatean additional paleoclimatic factor at this time.The Polar-isbreen Group is dominantly siliciclastic (Fig. 3), but

shallow-water carbonate banks were intermittently rees-tablished (Fairchild &Hambrey, 1995).

Paleopoles constrain the paleogeographic position ofLaurentia at 723Ma (Franklin Dikes; Buchan et al., 2000;Heaman et al., 1992) and 615Ma (Long Range Dikes;Murthy etal., 1992;Kamo&Gower,1994), but the apparentpolar wander path between those ages is unconstrained.Both poles place the EGES platform in low southernlatitudes, with the East Greenland margin facing thesoutheasterly trade winds and the depositional strike ofthe basin trending roughly southwest^northeast. Duringparts of this interval, the margin may have been under thein£uence of the dry downwelling arm of theHadley circu-lation, and onshore trade winds (from the present-daynortheast) may have dominated the margin throughout.

Geochronology

Precise ages from theNeoproterozoic of Svalbard are lack-ing. The only ¢rm age constraint is a maximum of�950Ma for the base of the Veteranen Group from detri-tal zircons in the upper Planetfjella Group (Fig. 1) in NyFriesland (Larianov et al., 1998) and primary zircons inthe underlying granites in northern Nordaustlandet (Geeet al., 1995; Johannson et al., 2000). The lack of Ediacaranfossils in the Polarisbreen (Knoll & Swett, 1987) suggestsa minimum age of �575Ma (Bowring et al., 2003). Derryet al. (1992) applied a time line to the Neoproterozoic stra-tigraphy based on a single, unpublished Rb^Sr isochron( �780Ma) in the lower Akademikerbreen Group (Grus-dievbreen Formation; S. Jacobsen, pers. comm., 2002)and correlationwith the ShalerGroup in northernCanada(Asmeron etal., 1991).Derry etal. (1992) estimated an age of630Ma for the base of the Polarisbreen, but this age hingeson correlation with other poorly dated Neoproterozoicglaciations. Hill & Walter (2000) suggested an age of�780Ma for the Svanbergfjellet Formation (Fig.1) basedon the occurrence of the acritarch Cerebrosphaera buickii inboth the Svanbergfjellet and Draken formations in Sval-bard (Knoll etal., 1991; Butter¢eld etal., 1994) and the mid-dle Burra Group in South Australia (Cotter, 1999; Hill etal., 2000), the base ofwhich is dated at 777 � 7Ma (Preiss,2000).

METHODS

Sample selection and preparation

All sampleswere collected in the course of measuring stra-tigraphic sections. Fresh dolomite and limestone sampleswith little visible veining, cleavage or siliciclastic compo-nent were targeted. Samples were prepared by cuttingslabs perpendicular to lamination and polishing one faceto clear saw marks and clarify internal texture. Between 5and 20mg of powder were microdrilled from the indivi-dual laminations (where visible) on the polished slab, col-lected and stored. All subsequent isotopic and elementalanalyses were performed on aliquots of this powder.

Fig. 3. Composite stratigraphic and d13C pro¢le through thePolarisbreenGroup. See text for discussion of how the compositewas constructed and Fig. 4 for the legend.



Isotopic analysis

d13C and d18O isotope data (see the Supplementary mate-rial) were acquired simultaneously on a VG Optima dualinlet mass spectrometer attached to aVG Isocarb prepara-tion device (Micromass, Milford, MA, USA) in theHarvard University Laboratory for Geochemical Oceano-graphy. Approximately, 1-mg microdrilled samples werereacted in a common, puri¢ed H3PO4 bath at 90 1C.Evolved CO2 was collected cryogenically and analysedusing an in-house reference gas. External error (1s) fromstandards was better than � 0.1% for both d13C and d18O.Samples were calibrated toVPDB using the Cararra Mar-ble standard. Potential memory e¡ect resulting from thecommon acid-bath system was minimized by tripling thereaction time for dolomite samples. Memory e¡ect is esti-mated ato0.1%based onvariability of standards run afterdolomite samples.

Elemental analyses

Elemental analyses (Ca, Mg, Sr, Mn, Fe, Na) were per-formed on a Jovin Yvon 46P ICP-AES mass spectrometerin the Harvard University Laboratory for GeochemicalOceanography. All samples were prepared by dissolving�4mg of carbonate powder in �4ml of 2% nitric acid.SCP multielement standards were used for element-spe-ci¢c calibration at the beginning of each run. External er-ror (1s), determined by repeat analyses, waso7% for allelements other than Na, which had an error of �28%and is regarded as only an approximate ¢gure.

RESULTS AND INTERPRETATIONS

Here we summarize the stratigraphic, sedimentological,and d13C (carbonate) records in northeastern Svalbard, be-ginningwith the ¢rst in£ux of terrigenous sediments in theupper Backlundtoppen Formation (Fig.1) and ending withthe cap-carbonate sequence (Dracoisen Formation) abovetheWilsonbreen glacials.We introduce informal names forthree previously unnamedmembers, as shown inFig. 3: theKinnvikka Member (Member B5 of Harland, 1997) of theBacklundtoppen Formation, the Russ�ya (E1 of Harlandet al., 1993) Member of the Elbobreen Formation, and theBr�vikaMember of theWilsonbreen Formation.

Figure 3 presents a composite stratigraphic sectionconstructed so as to provide the most complete isotopiccoverage available, using splices of the upper Backlund-toppen Formation at Br�vika (measured section 10); thelowerRuss�yaMember atBacklundtoppen (measured sec-tion 1) and Sveanor (measured section 7); the upper Rus-s�ya Member at Sveanor (measured section 7); thePetrovbreen Member throughWilsonbreen Formation atDracoisen (measured section 5, including our only com-plete section of the MacDonaldryggen Member); and theDracoisen Formation at Ditlovtoppen (measured section4). Signi¢cant erosional relief is present at the base of thePetrovbreen diamictite, but the underlying Russ�yaMem-ber appears to be nearly complete in most Nordaustlandet

sections. No other prominent erosional truncation(excluding that beneath the contact with the Cambrian) isapparent.Thus, the main gaps in the isotopic record corre-spond to the two glacial intervals and siliciclastics in theupper Kinnvikka Member, the MacDonaldryggen Mem-ber, and in the Dracoisen Formation.

Kinnvikka Member

The upper Backlundtoppen Formation is an �200-m-thick stack of uninterrupted stromatolites (Wilson, 1961;Knoll et al., 1989; Knoll & Swett, 1990) that is strongly en-riched isotopically (d13C �8%; Fig. 4a).This unit is locallycapped by a bed of irregular and discontinuous micro-bially laminated dolomite (microbialaminites) with tepeestructures, indicating emergence and the end of the long-lived Akademikerbreen carbonate platform. A £oodingsurface above this microbialaminite bed marks the base ofwhat is de¢ned here as the Kinnvikka Member.The lowerpart of the KinnvikkaMember is a 5^20-m-thick intervalof thin parasequences of shales, ribbon carbonates (ribbo-nites), and microbialaminites. The carbonates at thebase of the Kinnvikka Member are less 13C-enriched(d13C �5%) than the underlying stromatolites.Above thisinterval is another prominent £ooding surface. In mostsections, the overlying strata form a ¢ning- then coarsen-ing-upward parasequence (15^401m) of predominantlyshale, siltstone and ¢ne sandstone. Lesser green shalesand dolomite breccia lenses occur within the siliciclasticinterval.

At Br�vika (measured section 10), the northernmostmeasured section of this interval, a stromatolitic pinnaclereef, whose upward growth apparently kept up with the£ooding of the underlying carbonates, protrudes fromthe lower carbonates well into the silty interval. This45-m-high reef fortuitously documents marine carbon-isotopic composition of 5^7% through the lowerKinnvik-ka Member (Fig. 4b). Wedges of stromatolites, intraclastbreccia, and muddy debris £ows with dolomite clasts ex-tend laterally into the siltstones, capping m-scale coarsen-ing-upward cycles. Signi¢cant relief on the upper contactof the Kinnvikka dolostones in the Murchisonfjord areasuggests that such reefs may be a common feature.Theseisolated stromatolite buildups likely supplied the dolomi-tic detritus in other lower Kinnvikka sections and maypartly account for the highly variable thickness of the over-lying siliciclastic interval. The inter-reef and upperKinnvikka Member sediments are typically recessive-weathering, but based on £oat appear to consist of red ar-gillite with minor sandstone, dolomitic breccia and grain-stone. The nature of the contact with the overlyingDartboard Dolomite is uncertain due to poor exposure.

Dartboard Dolomite Member

TheDartboardDolomiteMember (Fig. 3) forms the top ofthe Akademikerbreen Group (Wilson, 1961; Wilson &



Harland,1964) and occurs throughout the outcrop belt ex-cept at its northernmost exposure in Ny Friesland (Kluft-dalen). Elsewhere in Ny Friesland and Olav V Land, theDartboard Dolomite is �25m thick, commonly stroma-tolitic, typi¢ed by lobate^cuspate laminations, and perva-sively recrystallized and vuggy (Fairchild & Hambrey,1984). In some exposures, the stromatolites are laterallylinked and steep-walled with synoptic relief of up to 3mand, in planview, de¢ne concentric patterns of laminations(hence the name ‘Dartboard’), reminiscent of Conophyton(Knoll et al., 1989). In other exposures, the cuspate struc-tures are nearly vertical, disrupted and brecciated at theapex (Fairchild & Hambrey, 1984), and linear to irregularin plan view.These structures more closely resemble peri-tidal tepees (cf. Kendall & Warren, 1987) than coniformstromatolites, and indicate deposition on a shallow, episo-

dically desiccated platform (Fairchild & Hambrey, 1984).The upper contact of the Dartboard Dolomite is brec-ciated at measured section1and in several other Ny Fries-land sections (Fairchild &Hambrey, 1984).

TheDartboardDolomiteMember is super¢cially simi-lar in Nordaustlandet exposures, but sedimentologicallydistinct and not as pervasively recrystallized. At Kinnvik-ka (measured section 10, Fig. 4b), the entire unit is 16mthick and its base consists of tightly packed peloids andbreccias with abundant, isopachous, void- ¢lling cements(Fig. 5a). Up section, macropeloids and microbial laminaeare increasingly interbedded. As microbial laminae be-comes predominant, it forms broad domes with interven-ing, peloid- ¢lled, and occasionally bridges troughs. Inthe upper half of the unit, continuous microbial laminaeformwave-like bedforms with meter-scale amplitude and

Fig.4. Stratigraphic columns and corresponding d13C data (in per mil units vs.VPDB) for the Kinnvikka through PetrovbreenMembers in Spitsbergen (a) andNordaustlandet (b). Each column is hung on a d13C5 0 datum (dashed line) where data is available(approximatedwhere no or insu⁄cient data are available) for the pre-Petrovbreen negative anomaly.The choice of this datum impliesthat the time at which d13C equaled 0% was the same everywhere in the basin, although this assertion is not proven. Hanging thestratigraphic columns on this datum demonstrates whatwe interpret to be variable erosional truncation of the upper Russ�yaMember,andwe note that the implied depth of erosion roughly correlateswith the thickness of the PetrovbreenMember diamictite (e.g.measuredsection 3). Solid lines show correlation of the base of the KinnvikkaMember, top of the Dartboard DolomiteMember, and the base ofthe upper parasequence of the Russ�yaMember. Gaps in stratigraphic coverage represent unexposed section. See the Supplementarymaterial for tabulated d18O and d13C data.



wavelength. Geometrically, the structures resemble te-pees, with broad troughs and pinched crests that intersectto form a quasi-orthogonal grid in plan view. However,distinct from the Ny Friesland examples, these structuresare neither disrupted nor brecciated, and no other evi-dence points to a shallow-water or intermittently emer-gent environment. Enigmatically, whereas the laminationappears more consistent with sediment stabilized by mi-crobial mats, the bedforms are more compatible with trac-tion transport of granular sediment. The microbiallamination £attens, and the dolomite becomes ferroan to-ward the top of the unit, which in Nordaustlandet appearsgradational with the overlying Russ�yaMember.

d13C through the Dartboard DolomiteMember consis-tently ranges from 3% to 6%, but shows no discernibletrend up section (Fig. 4).Nor does d13C vary noticeably be-tween individual sections, despite facies variations and theevidence for extensive alteration in Ny Friesland.

Russ�ya Member

The poorly exposed lower part of the Russ�ya Member isvariable in thickness and facies, but forms a distinct

upward-shoaling sequence 45^135m thick. At Backlund-toppen (measured section 1), the brecciated contact of theupperDartboard is overlain by 3m of black shale, followedby 26m of bluish-grey to black limestone rhythmites withincreasing proportions of thin dolomite beds upwards.The carbonates grade upward into dark, marly shales withpeculiar, crinkly microbial laminae. The sequence iscapped by a 7.5m-thick, poorly sorted, coarse-grainedsandstone. Elsewhere in Spitsbergen, the upper part ofthe sequence is more organic-rich and its top surface isnot easily identi¢ed (Fig. 4a). InNordaustlandet, the lowerRuss�ya Member appears continuous with the DartboardDolomite, where microbial laminae below pass upwardinto dolomitic, then heavily silici¢ed rhythmites beneatha poorly exposed interval of marly black shale, rhythmites,and limestone ribbonites with molar tooth structures.This thick (up to 110m), relatively deep-water intervalshoals upward into grainstones and microbialaminites,with an occasional thin sandstone bed.

The upper Russ�ya Member is well exposed at Back-lundtoppen (measured section 1), Dracoisen (measuredsection 4) and in several locations inNordaustlandet (Fig. 4).In Ny Friesland (Fig. 4a), it comprises a series of 5^10-m

Fig.4. (Continued)



thick cycles that variable consist of shale, black shale, orblack limestone at the base and sandstone or dolomitegrainstone at the top. Molar tooth structures (Fig. 5b) arecommon in the black limestone and are locally extremelydense, comprising up to 25% of the total rock mass. Large(up to 5-cm diameter) pyrite nodules occur locally in theblack shales. Interbedded dolomites are typically iron-stained and vuggy and in places contain gypsum castsand silici¢ed anhydrite nodules (Fairchild & Hambrey,1984; Harland et al., 1993). Up section, the carbonate phaseis exclusively dolomite. The uppermost Russ�ya consistsof an assortment of facies, including m-scale biohermalstromatolite mounds (Fairchild & Hambrey, 1984), and ishighly variable between sections. At Ditlovtoppen, theupper Russ�ya is condensed ( �10m) and is capped by adistinctive columnar stromatolite biostrome. In Nordaus-tlandet (Fig. 5b), the upper Russ�ya is thinner (o40m)and lessvariable than inNyFriesland. It consists of two para-sequences, the lower of which is topped by biohermal stro-matolites or microbialaminites.The upper parasequence iscapped by a 2^6m-thick biostrome (Kussiella sp.; Fig. 5c),nearly identical to that atDitlovtoppen, butwith collectivelyspiraling columns. The top of this parasequence is brec-ciated andvariably truncated on the outcrop-scale.

The lower and upper Russ�ya have very di¡erent d13Ctrajectories (Figs. 3 and 4). The lower Russ�ya hoversaround 5% through most of its thickness, following an in-itial rise from �1.5% near the base of the member atBacklundtoppen (measured section 1). In contrast, theupper Russ�ya declines from 5% to � 6% just below thePetrovbreen diamictite. InNordaustlandet, the decline be-gins beneath the penultimate parasequence (Fig 4b), andspans multiple facies, indicating that the anomaly doesnot record a surface-to-deep d13C gradient. In some Nor-daustlandet sections, regarded here as the most complete,an upturn towards less negative values is preserved in theupper parasequence. In contrast, the anomaly is missingaltogether in some Spitsbergen locations (measured sec-tions 3 and 5; Fig. 4).

Petrovbreen Member diamictite

The Petrovbreen Member diamictite is typically thin( �10m on average; Harland et al., 1993) and presentnearly everywhere above the Russ�ya Member (Fig. 4).The PetrovbreenMember is generally thicker in Ny Fries-land, with the thickest section (52m) at Klofjellet (mea-sured section 3).The PetrovbreenMember had previously

Fig. 5. Distinctive lithofacies in theDartboardDolomiteMember of the Backlundtoppen Formation and the Russ�ya and Petrovbreenmembers of the Elbobreen Formation. (a) coarse peloids in a microbially laminated framework (the coin, for scale is the same size as apenny) at measured section 9, (b, c) Lithofacies of the upper Russ�yaMember: (b) Pervasive molar tooth structure in black limestoneribbonites at measured section 4, (c)Kussiella sp. biostrome at the top of the uppermost parasequence at measured section 9.(d) PetrovbreenMember:massive diamictite overlain by ripple cross-laminated dolarenite at measured section 5.Note predominance ofangular dolomite (light grey) and chert (black) clasts, typical of the PetrovbreenMember (Harland et al., 1993).



been reported from only one locality on Nordaustlandet(measured section 11; Hambrey, 1982; Harland et al., 1993),but a thin and patchily preserved diamictite is found inmost well exposed sections in Nordaustlandet.

In most sections, the bulk of the Petrovbreen Memberconsists of tan andyellow-weatheringdolomitic rhythmites,¢nely laminated and rippled wackestone, and poorly sortedto massive diamictite (Fairchild & Hambrey, 1984). At Dra-coisen (measured section 5), a ripple cross-laminated grain-stone with outsized clasts comprises the uppermost part ofthe diamictite (Fig. 5d). The clast composition in the dia-mictite is consistently dominated by subangular, pale greydolomite (grainstones and stromatolites) and black chert(Fig. 5d) (Hambrey, 1982; Harland et al., 1993). Sandstoneand siltstone clasts are less common andvolcanic fragmentshave been reported (Fairchild &Hambrey, 1984) from a sin-gle region (measured sections1^2).Harland etal. (1993) sug-gested that up to10%of the clasts are striated, but probablythe percentage is much lower, as neither dolomite nor chert

clasts commonly preserve unequivocal striations. Rhyth-mites, interpreted by Hambrey (1982) as varvites, occurthroughout the Petrovbreen Member and commonly con-tain ¢ne dropstones.The lithological characteristics of thePetrovbreen Member suggest deposition in a glacimarineor glaciolacustrine setting (Hambrey, 1982; Fairchild &Hambrey, 1984; Harland et al., 1993).

Figure 6 presents a d13C^d18O cross plot of dolomiteclasts, matrix, and varvites compared to carbonate samplesfrom three sections of theRuss�yaMember in Spitsbergen.Although d13C of two of the clasts is negative, as is expectedfrom the composition of the immediately underlying car-bonates, most of the samples lie between 3% and 5%.Thecarbon-isotopic compositions of these clasts are consistentwith derivation from lower in the Russ�yaMember, as pro-posed byFairchild&Hambrey (1984), but anAkademikerb-reen source cannot be ruled out. Similarly, the chert clastsare most likely derived from the Russ�ya Member, wheredark cherts commonly occur at the top of parasequences.

MacDonaldryggen Member

The MacDonaldryggen Member consists dominantly ofterrigenous mudstone, but includes ¢nely laminated car-bonates o1.0m thick that occur directly above the Pet-rovbreen diamictite in several locations in Olav V Land[and were interpreted by Kaufman et al. (1997) as cap car-bonates].These carbonates are dark grey (tan weathering)limestones or dolomites inwhichwispy organic matter de-¢nes an anastomosing lamination (Harland et al., 1993);they resemble the lean (o40 cm) cap carbonate locallyoverlying the Gaskiers diamictite in Newfoundland (Myr-ow&Kaufman, 1998). However, d13C in the basal MacDo-naldryggen carbonates (Fig.7) is highly variable (� 17% to0%) and in some cases clearly a¡ected by the reminerali-zation of organic matter (cf. Irwin etal., 1977).This isotopicpattern (Fig. 7) is distinct from typical Neoproterozoic capcarbonates whose isotopic compositions are uniformglobally (e.g. Ho¡man & Schrag, 2002). Nor do these car-bonates, which are in some cases interbeddedwith the un-derlying diamictite,mark an abrupt facies change as wouldbe expected if theywere deposited immediately following aNeoproterozoic glacial episode.

Fig. 6. d13C^d18O cross-plot of Russ�ya carbonates andPetrovbreenMember matrix and clasts.The ¢eld of both thediamictite matrix and clasts generally overlap with the valuesfrom the underlying Russ�yaMember, although d18O values aremore spread out [which Fairchild et al. (1989) interpret torepresent a spectrum from least-altered (d18O40%) to most-altered (d18Oo� 5%) rock £our], and indicate that the Russ�yaMember was the predominant source of clasts to the Petrovbreendiamictite. 1Data from Fairchild et al. (1989).

Fig.7. d13C^d18O cross-plot of baseMacDonaldryggen carbonates, calcitepseudomorphs, and thin dolomite bedswithintheMacDonaldryggenMember shales. Notethe extremely low d13C values for some of thecalcite samples, which indicate thatremineralized organic matter was a signi¢cantsource of alkalinity (cf. Irwin etal.,1977). Boxeddata are from a sample at measured section 8that contains the pseudomorph shown in theinset, interpreted here to be a glendonitenodule.



The bulk of theMacDonaldryggenMember consists of¢nely laminated silty mudstone, dark olive green in colour.These uniformly ¢ne-grained and parallel-laminateddeposits are up to 200m thick, contain minor sandyinterbeds, and are virtually devoid of current-generatedbedforms.AtSlangen (measured section 2), the uppermostMacDonaldryggen mudstones grade upward intopure limestone ribbonites and grainstones with d13C

values of 2^4% (Fig. 8). The top of the sequencehere is an exposure surface, indicated bymudcrackswithinthe limestones reported by Fairchild & Hambrey(1984), who also cite mudcracks in several other sections.However, most sections lack both carbonates andmudcracks, and the mudrocks instead pass gradation-ally into marly ribbonites of the lower SlangenMember.

Fig. 8. Stratigraphic columns and isotopic data spanning theWilsonbreen Formation in (a) northeastern Spitsbergen and(b) northwestern Nordaustlandet, with the top and bottom of theWilsonbreen Formation as datums. Note the variable scale of theWilsonbreen Formation and Br�vikaMember. Grey lines in (a) show correlation of theMiddle Carbonate inMember in Spitsbergensections (absent in Nordaustlandet). Rose diagram shows palecourrent data (n519) for the Br�vikaMember at measured sections 9^11.(c) Schematic cross-section (vertically exaggerated) of theWilsonbreen diamictites and Br�vikaMember, illustrating the northwardthinning of the diamictites and reciprocal southward taper of the Br�vika sandwedge. See Fig. 4 for legend and the Supplementarymaterial for tabulated d18O and d13C data.



The mudstone is speckled with rosette-shaped, calcitenodules, 2^10mm in diameter (Fig. 7) that formed authi-genically before the sedimentwas fully compacted.The no-dules appear to be pseudomorphs of radial clusters oftightly nested crystals. They resemble glendonite pseudo-morphs, which have long been used as a cold paleoclimateindicator (Boggs, 1972; Kemper & Schmitz, 1981; De Lurio&Frakes,1999; Swainson&Hammond, 2001).TheMacDo-naldryggen pseudomorphs have not been studied in detail,but their external projections include both canted pyrami-dal and squared prismatic habits, which are diagnostic ofthe unusual crystal habit (Swainson & Hammond, 2001) ofikaite, monoclinic CaCO3 � 6H20. Quaternary and modernikaite tufa formswherewaters rich in orthophosphate (a cal-cite and aragonite inhibitor) and Ca are discharged into al-kaline waters that are close to their freezing point (Council& Bennet, 1993; Buchart et al., 1997). Authigenic ikaite crys-tals, precursor glendonites, have been found at shallowdepths (2^8m) in organic-rich sediments of the Antarcticshelf (Suess et al., 1982) and in estuaries of the southwestBeaufort Sea in Alaska (Kennedy et al., 1987). Most ancientglendonites are strongly depleted in 13C (Boggs, 1972; Suesset al., 1982; De Lurio & Frakes, 1999), typically �15% to� 22%, and remineralized organic matter or anaerobicallyoxidized methane is widely cited as sources of the isotopi-cally depleted carbonate. The MacDonaldryggen pseudo-morphs are depleted in 13C and 18O, relative to dolomitesforming discontinuous beds and concretions in the samemudrocks that are assumed to better represent seawatercompositions (Fig. 7), but are not nearly as 13C-depleted astypical glendonites.We emphasize that we have not studiedthe pseudomorphs in detail and that our interpretation thatthey are glendonites is tentative; the original mineral couldhave been gypsum or anhydrite, as is common in the Rus-s�yaMember (Fairchild &Hambrey,1984).

Slangen Member

TheSlangenMember (Fig.8) is a regressive parasequence,20^30m thick, composed of silty to marly ribbons cappedby pale grey, cherty dolomite. Much of the dolomite con-sists of cross-strati¢ed grainstone, variably composed ofooids, tabular intraclasts and rounded dolomite grains(Fairchild & Hambrey, 1984; Harland et al., 1993). Tabularcross-bedding, well-de¢ned by trains of bedding-parallel,chert- ¢lled vugs (Harland et al., 1993), is common in thelower half of the grainstone.Dark microbialaminite occursas thin lenses within the grainstone in some sections andas thicker deposits above the grainstone.The microbiala-minite has a well-developed laminoid fenestrate fabric,with void- ¢lling sparry calcite, length-slow chalcedony(Fairchild & Hambrey, 1984) and, in rare instances, eitheranhydrite or gypsum.The evidence for evaporite mineralsand extremely high Na concentration in dolomite (2000^4000 p.p.m. at measured section 5) suggests that this unitwas deposited in an intermittently restricted environment(Fairchild & Hambrey, 1984). The general facies and se-quence are similar to those observed beneath the Recent

coastal sabkhas (supratidal salt £ats) of the southern Ara-bian Sea (e.g. Kendall & Skipwith, 1969), as discussed indetail in Fairchild &Hambrey (1984).

The upper contact of the Slangen Member typically issharp, except at Ditlovtoppen (measured section 4), wherethe grainstone passes transitionally upwards into 2.5m ofmudcracked red siltstone with thin dolomite interbeds.Elsewhere, the top 5^20 cm below the contact is typicallysilici¢ed and broken into angular and tabular clasts thatshow no evidence of reworking (Fairchild & Hambrey,1984). Despite the evidence for exposure, no erosional orkarstic relief is evident at the outcrop scale. Instead, thesurface typically appears to be £at.

In Spitsbergen sections, d13C varies in the grainstonefrom 2.5% to 4.5% (Fig. 8a). In measured section 5, thesection sampled in greatest detail, d13C rises smoothlythrough the grainstone by 2%. d18O also increases by thesame amount, lending credence to the interpretation thatthe grainstone was deposited in a progressively restrictedenvironment.The Nordaustlandet sections (Fig. 8b) exhi-bit a similar pattern of �2% enrichment up sectionthrough the grainstone, but are on average 2^3% higherthan the Spitsbergen sections (Fig.9).

Spatially variable diagenesis on carbonate platformscommonly generates lateral di¡erence in d13C of a similaror greater magnitude (e.g. Allan &Mathews, 1982; Chafetzet al., 1999) than those in the Slangen Member. However,the reproducibility of the upward positive d13C trend in allsections and de¢ciency in organic matter in the Slangendolomites (Fairchild & Hambrey, 1984) suggest that theSlangen Member preserves a signal of evolving seawaterd13C. Insofar as this hypothesis is true, the d13C di¡erencebetween Spitsbergen and Nordaustlandet sections may beinterpreted in one of two ways: either it records a roughlynorth-south isotope gradientwithin theSvalbard basin, orthe grainstone is diachronous within the Slangen Mem-ber, and in any given section captures only a portion of alonger term rise in marine d13C (Fig. 9). The former hy-pothesis implies a lateral isotope gradient of �2% (Fig.9a), which is similar to d13C gradients between the innerand outer shelf in the Otavi Group of northern Namibia(Halverson et al., 2002) and is consistent with depositionof the grainstone in a restricted environment.On the otherhand, simple stacking of theNordaustlandet sections atopthe Spitsbergen sections yields a nearly continuous d13Cincrease of 4% up section (Fig.9b), consistentwith progra-dation of the grainstone from south to north.

Wilsonbreen Formation

Br�vikaMember

In Nordaustlandet and northernmost Ny Friesland (He-clahuken), a clean, medium-grained sandstone composedpredominantly ofwell rounded quartz and chert grains se-parates the Slangen Member from the Wilsonbreen dia-mictites (Fig. 8). We informally refer to this unit as theBr�vika Member and include it within the WilsonbreenFormation.TheBr�vikaMember thickens northward from



8m at Ditlovtoppen (measured section 4) to 233m at thenorth coast of Nordaustlandet (measured section 11).TheBr�vika Member in measured sections 6^8 appears ¢nergrained and recessive near the base, although the natureof the contactwith the SlangenMember is obscured.Thinlenses of stromatolitic dolomite occur within the lowersandstones. In these locations, the top of the sandstonecoarsens with the addition of angular dolomite clasts.Thecontact with the overlying diamictite is a sharp scour sur-face with a thin pebble lag, and the lower diamictites arerich in sand evidently reworked from the Br�vikaMember.

In the northernmost exposure (measured section 11), theBr�vika Member comprises three distinct parasequences, alower one dominated by sandstone (whichwe correlate withthe entire Br�vika Member in measured sections 6^8), amiddle one in which the sandstone contains angular, out-sized, dolomite clasts, and an upper one composed of argil-lite andmarly limestone ribbonites, capped by a0.5-m-thicklimestone-clast diamictite (beneath the basal Dracoisen capdolostone). d13C in these carbonates �7%, falling on thetrajectory of increasingd13Cde¢ned by theSlangenMember(Fig 9b). Planar and trough cross-strata throughout mea-sured section 11 indicate a mean paleo£ow direction towardthe north (Fig. 8), the direction of stratigraphic thickening.

The stratigraphic relations and geometry of the Br�vikaMember (Fig. 8c) and overlying diamictite suggest that it isa diachronous unit, with sands being deposited in thenorthern part of the basin while Wilsonbreen diamictitewas being deposited to the south.The coarsening upwardnature of the sandstone sequence could re£ect the ap-proach of theWilsonbreen ice sheet.

Diamictite

The Wilsonbreen Formation in Ny Friesland and Olav VLand (Fig. 8) is divided into three members (Harland

et al., 1993). The basal Ormen Member consists predomi-nantly of poorly strati¢ed diamictite with abundant lensesand interbeds of conglomerate, coarse sandstone, andbreccia (Fairchild & Hambrey, 1984; Harland et al., 1993).Clasts in the diamictite are dominantly carbonate, but sand-stone, chert, gneiss and coarse-grained pink granite clastsalso are abundant (Harland et al., 1993). Striated and frac-tured stones are common (Harland et al., 1993).The sourceof the metamorphic clasts is unknown. OrmenMember fa-cies indicate deposition in a terrestrial setting, whereas theexotic clasts suggest deep erosion and glacial transport overlong distances (Fairchild&Hambrey,1984).Because glacierswere in contact with basement, carbonate and siliciclasticclasts in the diamictite could have been derived from anystratigraphic level in the sub-glacial succession.

The Middle Carbonate Member consists mostly ofmedium-grained, dolomitic sandstone, but is distin-guished by thin carbonate beds found only inNyFrieslandand Olav V Land sections. On the Ormen nunatak (mea-sured section 2), the carbonate is a 0.5-m-thick, silty, dolo-mitic microbial laminae. At Dracoisen (measured section5), the carbonate occurs as multiple, thin (o0.1m), discon-tinuous, and often scoured, isopachous limestone stroma-tolites encased within medium-grained sandstone. AtDitlovtoppen the carbonate occurs in 1.5m of dolomiteand limestone ribbonites and rhythmites. The d13C of allsampled carbonates, including those published by Fair-child et al. (1989), lie in a range of 1^5% (Fig. 8). Fairchildetal. (1989) argued that the carbonateswere deposited in anevaporitic environment (during a period of glacial retreat;Fairchild & Hambrey, 1984) based on highly enriched d18Ocompositions (up to110.8%) of some dolomites, as well asthe occurrence of gypsum casts and anhydrite relics insome samples.Whereas new analyses presented here yieldmuch lower d18O values in limestone samples collected in

Fig.9. d13C and d18O data from the Slangen grainstone (dolomite) and the Dracoisen cap dolostone, which bracket theWilsonbreenFormation. Both intervals are scaled to a mean thickness. (a) If the grainstone was deposited everywhere synchronously, then the spreadbetween the Spitsbergen andNordaustlandet data indicate an �2%d13C gradientwithin the basin. (b) If the grainstone was depositeddiachronously, such that it is younger to the north, the data indicate a gradual rise in d13C of 4% through the SlangenMember. Incontrast to the variable isotopic values for the SlangenMember, both d13C and d18O through the Dracoisen cap dolostone are nearlyidentical in all sections, implying synchronous deposition throughout the basin, despite variations in thickness (3-18m).



measured section 5 (Fig. 8a), high sulfate and Na(3500p.p.m.; unpublished data) concentrations supportthe hypothesis that the Wilsonbreen carbonates formedin evaporitic waters, analogous to carbonate forming todayin the Dry Valleys of Antarctica (Fairchild et al., 1989), butwith locally a large component of 18O-depleted water, po-tentially derived from glacial runo¡ (Fairchild et al., 1989).

The Gropbreen Member is variable in thickness and issimilar lithologically to the Ormen Member, consistingmostly of poorly strati¢ed, maroon diamictite (Fig. 10b).Clast composition is also similar, though with a greaterproportion of basement material. A striated boulder pave-ment within Gropbreen diamictite at Ditlovtoppen hasbeen reported (Chumakov, 1968, 1978), and sand wedgepolygons occur in various sections (Chumakov, 1968, 1978;Harland et al., 1993). AtGropfjellet, near Backlundtoppen,the sand wedges are developed in the uppermost diamic-tite layer and are ¢lled by sand identical to that of an over-lying, ripple cross-laminated sandstone.

OnNordaustlandet the diamictite is lithologically simi-lar but has a higher concentration of volcanic clasts.Volca-nic and granite clasts in Nordaustlandet are lateGrenvillian in age (Johannson et al., 2000), similar tothe underlying basement (Gee et al., 1995). The Middle

Carbonate Member is absent on Nordaustlandet and thethickness of the diamictite is highly variable, reciprocatingwith the underlying Br�vika Member.These relations in-dicate that the diamictite on Nordaustlandet may beequivalent in age only to the Gropbreen Member in NyFriesland, as would be expected from a northward advanceof theWilsonbreen ice front.

Dracoisen Formation (cap-carbonatesequence)

Like other Neoproterozoic caps worldwide (Williams,1979; Aitken, 1991; Kennedy, 1996; Ho¡man & Schrag,2002), the Dracoisen cap dolostone (Fig. 10a) forms thetransgressive base of a cap-carbonate sequence (cf. Ho¡-man & Schrag, 2002) that sharply overlies glacigenic dia-mictite without evidence of reworking or hiatus.Comparable in thickness to other cap dolostones (Wil-liams, 1979; Ho¡man & Schrag, 2002), the Dracoisen capis thinnest (3m) in Nordaustlandet (measured section 6)and thickest (18m) at Ditlovtoppen in Ny Friesland (Fig.10a). It weathers to a yellowish-tan colour and is a usefulmarker bed (Wilson andHarland, 1964). It is lithologicallysimilar in all sections, composed of ¢nely laminated,

Fig.10. (a) Field photograph of the upperWilsonbreen Formation and overlyingDracoisen Formation at Ditlovtoppen(measured section 4), illustrating the uniformcap dolostone (18-m-thick), overlain by red,green, then black shales of the cap-carbonatesequence. (b) Polished slab slows low-anglecross-lamination, reverse-graded peloids,and black, interpeloid cement typical of theDracoisen cap dolostone. (c)Field photographwith line work and (d) idealized sketch oftypical megaripple structures, whichexclusively occur in the upper half of theDracoisen cap dolostone and other capdolostones globally and demonstrate growthof the bedform by vertical accretion underoscillatory £ow.



micro- to macropeloidal dolostone with variable amountsof black, interpeloidal cement (Fairchild & Spiro, 1987).The peloids form discontinuous, normal- and reverse-graded laminae or overlapping pockets (Fig. 10b).The lar-ger peloids are commonly truncated, £attened, or fused toone another, implying that they were not rigid particles(Fig. 10b).The preservation of such delicate structures at-tests to rapid cementation.

Very low-angle cross-strati¢cation is common in thecap dolostone, particularly towards the top, where accre-tionary oscillation megaripples (Fig. 10c,d) also occur.The megaripples have characteristic wavelengths of 3^6mand amplitudes of 20^40 cm with linear, parallel creststrending NW^SE (present-day), perpendicular to the in-ferred direction of tradewinds. Individual sets of megarip-ples are up to1.4m thick.

The most complete section of theDracoisen Formation(Fig. 11) is found at Ditlovtoppen (measured section 4) insouthern Ny Friesland (Fairchild & Hambrey, 1984; Har-land et al., 1993). Here, the transgressive upper dolostoneis transitional over a few meters through red siltstone andoverlain by variegated red and green shales. Up section,the green shales pass into black shales, which comprisethe remainder of the cap-carbonate sequence (Fig. 10a).Mud cracks mark the ¢rst evidence of subaerial exposureat172m above the base of theDracoisen Formation. Abovethis level, the remainder of the Dracoisen consists mostly

of mudcracked and ripple cross-laminated red and greensiltstoneswith minor sandstone and interbedded dolomitebeds, the most prominent ofwhich is a 9-m-thick, micro-bialaminite dolomitic unit with cross-cutting cauli£owerchert layers (Fairchild & Hambrey, 1984) �230m abovethe base of the Dracoisen Formation. No Ediacaranmacrofossils occur in Svalbard, indicating a considerablehiatus (430Myr) between the top of the Dracoisen andthe overlying Tokammane Formation of Cambrian age(Knoll & Swett, 1987).

d13C and d18O data for the Dracoisen cap dolostone areplotted in Fig. 8. In all sampled sections, d13C begins at�� 3% at the base and gradually decreases by about 1%through the dolostone. This trend is consistent through-out northeastern Svalbard, as seen in Fig. 9, in which acomposite plot of all data, normalized to a constant capdolostone thickness, show variation of less than �1%.Locally (e.g. measured section 4) carbonate interbeds oc-cur above the cap dolostone, and d13C in these records acontinued decline to a low of � 5%. d18O also exhibits arough negative trend through the dolostone, falling about1% from a low of �� 5% at the base. These values arelower on average by �4% than typical dolomites in Sval-bard, but typical of other Neoproterozoic cap dolostones(Halverson, 2003).

Sparse carbonate beds in the middle and upperDracoi-sen record a pronounced rise in d13C to highly enriched

Fig.11. Stratigraphic and d13C pro¢le for theDracoisenFormation atDitlovtoppen(measuredsection 4). See Fig. 4 for the legend and theSupplementary material for tabulated data.



values (4110%) in the dolomite beds in the middle of theFormation.This rise coincides with sedimentary features(mud cracks and evaporates) that suggest at least intermit-tent restriction of the basin, and therefore, may not repre-sent the evolution of open marine d13C.

Elemental chemistry of the Dracoisen cap dolostone

Elemental data from three sections of the Dracoisen capare summarized inFig.12.The relative18O-depletion, veryhigh concentration of Fe and Mn, and high Mn/Sr of theDracoisen cap are characteristic of digenetic alteration bymeteoric £uids (Brand&Veizer,1980,1981; Banner&Han-son, 1990; Jacobsen & Kaufman, 1999). However, strati-graphic and geochemical characteristics are inconsistentwith extensive meteoric diagenesis. First, the cap dolos-tone is bounded above by a £ooding surface with no evi-dence for exposure, and therefore, it seems unlikely that itwas subjected to extensive meteoric £uid £ow after deposi-tion. Second, the meteoric diagenesis model (Banner &Hanson,1990; Jacobsen&Kaufman,1999) predicts that gi-ven the very high Mn/Sr ratios in the dolomites (10^40mol mol�1), both d13C and d18O should have equili-brated with the meteoric £uids (Banner & Hanson, 1990).The strong reproducibility in d13C and the preservation ofa distinct trend in all sections contradicts this model. Sec-ond, the high Na concentrations (1200^3800p.p.m.) is theopposite of what one would predict had the cap dolostoneseen a large £ux of freshwater, andwould seem to indicatethat if the dolostone was diagenetically altered, it would

have been under restricted, hypersaline conditions, forwhich there is no other evidence. Therefore, we regardthe unusual geochemistry of the Dracoisen cap dolostonenot as a record of pervasive alteration, but as a manifesta-tion of the unusual hydrological and geochemical condi-tions in the aftermath of glaciation (Ho¡man et al., 2002).

DISCUSSION

The Pre-Petrovbreen d13C anomaly

The most salient feature in the isotopic compilationthrough the Polarisbreen Group is the large negativeanomaly ( �11%) in the upper Russ�ya Member, just be-neath the Petrovbreen diamictite. d13C declines smoothlythrough shallow-water sediments in Nordaustlandet sec-tions, which rules out a huge depth-dependent isotopegradient (cf. Shields et al., 1997; Calver, 2000; Fairchildet al., 2000) as the source of the anomaly. Furthermore, inorder for the full isotopic decline to have been recorded inshallow-water carbonates, it must have preceded any rapidsea-level drawdown related to the buildup of continentalglaciers. This buildup of glaciers would have lowered sealevel and exposed the Polarisbreen platform to erosionprior to deposition of thePetrovbreen diamictite.We inter-pret the absence of the anomaly inmeasured sections 3 and5 as a consequence of this pre-glacial erosion and note thatthe thickness of the Petrovbreen diamictite in each sectionbroadly coincides with the apparent depth of erosion, asinferred from the extent of truncation of the d13C anomaly

Fig.12. Isotopic and elemental cross-plotsfor three sections of the Dracoisen capdolostone.



(Fig. 4). The truncation of the upper Russ�ya Member isvariable, and, along with the preservation of isotopicallynegative clasts in the Petrovbreen, rules out the possibilitythat the anomalywas the result of extreme diagenetic over-printing beneath the Petrovbreen exposure surface (Hal-verson et al., 2002).

The stratigraphic constraints on the pre-glacial d13Canomaly are homologous to those for a pre-glacial anomalyof identical magnitude in the upper Ombaatjie Formation(Otavi Group), beneath theMarinoan-aged Ghaub glacialdeposits in northernNamibia (Halverson et al., 2002) (Fig.13). In both cases, stratigraphic evidence indicates that thefull negative shift occurred prior to the onset of continen-tal glaciation, thus removing glaciation as a possible causefor the anomaly. Similar pre-glacial negative d13C anoma-lies (Fig. 13) occur globally, including in the MackenzieMountains in northwest Canada (Keele Formation; Ho¡-man & Schrag, 2002), the Adelaide Rift Complex, SouthAustralia (Enorama and Trezona Formations;Walter et al.,2000; McKirdy et al., 2001), and Scotland (Lossit Forma-tion, Brasier & Shields, 2000). In all cases, the decline ap-pears gradual, punctuates trends of enriched composition(45%), and is followed by a trend towards more positivevalues directly beneath the overlying glacial deposits. InNamibia, Canada, and Australia, the anomalies precedethe younger of two discrete glaciations (Fig. 13), and nocomparable anomaly is known from beneath the older,Sturtian, glaciation. In these three successions, the Stur-tian glacial deposits host signi¢cant volumes of bandediron formation and are overlain by organic-rich cap-car-bonate sequences with similar carbon-isotopic signatures(Kennedy et al., 1998; Ho¡man & Schrag, 2002). Likewise,the cap-carbonate sequences to the younger, Marinoan,glaciations in these three areas are similar to one anotherstratigraphically, sedimentologically, and isotopically(Ho¡man & Schrag, 2002).We assume that the older andyounger glaciations in Namibia, Canada and Australiawere coeval (Kennedy et al., 1998). Since pre-glacial nega-tive d13C anomalies occur only beneath theMarinoan gla-cial deposits, we propose that this anomaly is a usefulglobal correlation tool that can be used where obviousSturtian^Marinoan glacial pairs are absent.

The Marinoan cap carbonate

The sedimentary record of deglaciation is contained in thetransgressive Marinoan cap-carbonate sequence world-wide (Kennedy etal.,1998;Ho¡man&Schrag, 2002). Com-parative studies have shown the Marinaon cap carbonatesto be lithologically variable, with particular facies- speci¢cstructures and textures, including megaripples, peloidsand sea£oor-encrusting cements (Aitken,1991; Grotzinger&Knoll,1995;Kennedy,1996; James etal., 2001;Ho¡man&Schrag, 2002; Nogueira et al., 2003).The most widespreadfeature of these sequences is the basal cap dolostone,which is typically 3^20m-thick and consists of pale yellowto grey, ¢nely laminated, £inty dolomite (Ho¡man& Schrag, 2002), bounded above by a £ooding surface.

Megaripples, which we interpret to re£ect rapid deposi-tion on a severe, storm-dominated shelf, are characteristicof most cap dolostones and consistently occur in theupper part of the unit. The homology of the Dracoisencap dolostone and the Ravensthroat cap dolostone innorthwest Canada (Fig. 13) is particularly striking,but virtually all of the distinctive features observedin the Dracoisen cap dolostone are shared by other capdolostones in outer shelf settings in Australia (Kennedy,1996), Namibia (Ho¡man, 2002) and South America(Nogueira et al., 2003) as well. Moreover, the carbonisotope pro¢le through the Dracoisen cap dolostoneis indistinguishable from other Marinoan cap dolostones(Fig.13).

The Polarisbreen glaciation

Unlike Namibia, Canada and Australia, the pre-glacialcarbon-isotope anomaly in Svalbard precedes the older ofthe two diamictite-dominated units. If the PetrovbreenandWilsonbreen Formations are regarded as separate gla-ciations (Hambrey, 1982; Fairchild & Hambrey, 1984; 1995;Kennedy, 1998), then correlation of the Russ�ya anomalywith the overseas successions (Fig.13) implies that the Pet-rovbreen is Marinoan and the Wilsonbreen represents athird, post-Marinoan, glaciation (Ho¡man & Schrag,2002). However, this interpretation is di⁄cult to reconcilewith the striking similarity of theDracoisen cap dolostonewithMarinoan caps elsewhere, andwith the absence of anycap carbonate on top of the Petrovbreen diamictites.Thesecorrelation problems disappear if the Petrovbreen andWilsonbreen diamictites, alongwith the intervening inter-val, encompass a single glacial episode equivalent to theMarinoan. For simplicity, we call this the Polarisbreen gla-ciation. If we are to accept the evidence from elsewherethat theMarinoanwas a snowball Earth episode (Ho¡man& Schrag, 2002), then the sedimentary succession repre-senting the Polarisbreen glaciation must also be compati-ble with the snowball hypothesis.

In the paragraphs that follow,we use stratigraphic infor-mation from East Svalbard (given above) and East Green-land (from the literature) to develop a possible scenario(Fig. 14) in which the entire succession from the base ofthe Petrovbreen diamictites up through the Dracoisen capdolostone (Ulves� Formation through Canyon cap dolos-tone in East Greenland; Fig. 1) is accounted for as phasesof a single, snowball-type, climatic cycle.We propose thatthe older andyounger diamictites record the creation (Fig.14, Phase a) and destruction (Fig 14, Phase d) of the snow-ball conditions, but that the fully established snowball epi-sode ^ de¢ned as having thick (4100m) sea ice coveringthe open ocean globally (Warren et al., 2002) ^ is repre-sented by the ¢ne clastics of the MacDonaldryggen For-mation (Fig. 14, Phase b) and the regressive carbonate-evaporite sequence of the SlangenMember (Fig.14, Phasec).Although our paleoclimatic interpretation of the overallsuccession is unconventional in several respects, webelieve it is consistent with (1) our ¢eld observations and



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isotope data, (2) previously published isotopic data, (3) ourproposed inter-regional correlation criteria and (4) recenttheoretical advances in the glacial dynamics of snowballEarth (Warren et al., 2002; Donnadieu et al., 2003; Good-man & Pierrehumbert, 2003).

Phase a: PetrovbreenMember (Ulves� Formation): buildup ofsubtropical ice domes

The ¢rst physical sign on the EGES platform of the com-ing ice age was a �50-m sea level drop related to the

Fig.14. TheMarinoan glaciation in east Svalbard andEastGreenland, interpreted in terms of a ¢ve-phase (aê) snowball-type climaticcycle.The cross-section is drawn along a north-south line (in present coordinates) across the EGES platform. Formation/Membernames are derived from Spitsbergen (Fig.1).The stratigraphic section to the left (scaled to thickness) is a schematic version of thatdeveloped in the text and ¢gure Fig. 3 and is plotted to correspond to the cross sections on the right. Phase (a): An ice dome advancedover the coastal highlands of Laurentia/East Greenland, heralding the onset of local glaciation prior to full snowball conditions.Theadvancing ice led to deposition of PetrovbreenMember glacimarine diamictites above a karstic erosional surface within thestromatolitic Upper Russ�yaMember. Phase (b): Full snowball conditions were marked byMacDonaldryggenMember mudstones,whichwere deposited beneath sikussak ice by sediment laden, subglacially derived, meltwater plumes. Pseudomorphs are abundant inthe mudstone, and the top of the unit in more glacioproximal EastGreenland sections often contains dropstones related to the breakupof the sikussak ice. Phase (c): As equatorial temperatures rose, the sikussak ice disintegrated over the EGES platform. However, seaglaciers were still fed by ice £ow from polar latitudes, therefore an oasis of open ocean developed on the EGES platformwhile ice stillcovered much of the world’s oceans.Trade winds would have fed the ice dome with moisture from the oasis and kept glacially derivedicebergs andwinter sea ice stacked along the East Greenland side of the platform. Dolomite grainstones of the SlangenMember weredeposited on the East Svalbard side of the platform. Phase (d): Reinvigoration of the hydrologic cycle due to the opening of snowballoases on tropical shelves led to the growth of ice domes. Br�vikaMember sandstones were reworked into glacio£uvial outwash andeventually overridden by the advancing ice.Terrestrial diamictites and glaciolacustrine carbonates of theWilsonbreen Fmweredeposited proximal to the ice dome. Phase (e):The basal Dracoisen Fm cap dolostone was deposited in the immediate aftermath of thesnowball Earth, when tropical temperatures ¢nally reached levels su⁄cient to melt the sea glacier.The cap dolostone is not preserved inthe deepest part of the platform. See the Discussion section of the text for a full explanation of each phase.



growth of high latitude and high altitude glaciers. Semi-arid conditions prevailed on the exposed EGES platform,where tens of meters of karstic relief developed (Fig. 4).Trade winds deposited aeolian sandstones preserved inthe present-day northwest portion of the platform (Ham-brey & Spencer, 1987; Moncrie¡ and Hambrey, 1990). Asthe subtropics cooled, moisture delivered from the openocean (in the present-day east) began to accumulate as icein paleo-highlands now situated in the hinterland ofsouthern East Greenland (Higgins et al., 2001). An icedome nucleated in these highlands (Moncrie¡, 1989; Phil-lips &Fredricksen,1981;Henriksen,1981; Smith&Robert-son,1999b) and descended to sea level (Fig.14a), depressingtheEGESplatform glacioisostatically.Massive diamictitesas thick as 300m (Ulves� Formation) accumulated at thegrounding line of the ice sheet in the southwestern partof the platform (Higgins, 1981; Hambrey & Spencer,1987), while ice-rafted debris rained-out over present-day northern and eastern regions (Fairchild & Hambrey,1984;Hambrey&Spencer,1987;Harland etal.,1993), ¢llingtopography and leaving a thin veneer of glacimarine dia-mictite (Petrovbreen Formation).

Clasts in the Petrovbreen diamictites are subrounded tosubangular (Fig. 5d), suggesting that theywere transportedat least partially as subglacial debris and that sources ofangular englacial debris such as nunataks were not impor-tant (Dowdeswell etal.,1985;Hambrey&Spencer,1987, andreferences therein). Clasts were derived entirely from thedirectly underlying sedimentary units indicating that theice sheet was likely cold-based and unable to accomplishextensive bedrock quarrying (Cu¡ey et al., 2000; Atkinsetal., 2002).The abundance of dropstones in distally depos-ited glacimarine sediments also points to a cold-based icesheetwith little or no ice shelf (o70 kmwide) thatwas cap-able of calving cold, debris-rich icebergs (Carey &Ahmed,1961; Orheim & Elverhoi, 1981). Together, these observa-tions indicate that by the end of Petrovbreen/Ulves� times,the platformwas dominated by a cold-based ice dome thatwas grounded in the present-day southwest of the plat-form andwhich distributed debris-rich icebergs to the restof the platform. At least two sand wedge horizons in theUlves� Formation (Spencer, 1985; Hambrey & Spencer,1987; Moncrie¡ & Hambrey, 1990), indicative of subaerialexposure and thermal contraction cracking of permafrost,imply a £uctuating ice line that left some regions of thewestern platform intermittently ice-free.

Phase b:MacDonaldryggenMember (Arena Formation):formation of sikussak ice

Across the entire EGES platform, glacimarine diamictitesterminate abruptly at the top of the Petrovbreen/Ulves�Formation (Hambrey & Spencer, 1987; Harland et al.,1993). The succeeding MacDonaldryggen/Arena Forma-tion consists of dropstone-free, ¢nely laminated, siltyclaystone and ¢ne sandstone. In the grounding-line-prox-imal regions of SW East Greenland, the Arena Formationis signi¢cantly coarser and contains two distinctive drop-

stone horizons indicating that some form of glaciationpersisted. More distal sections in NE Svalbard consist ofthick, ¢nely laminated deposits that lack current-gener-ated bedforms, but are speckled with glendonite nodules.We propose that the abrupt termination of ice-rafted deb-ris at the top of the Petrovbreen/Ulves� Formation doesnot mark the end of a glaciation, but rather the abrupttransition into a snowball Earth (Ho¡man & Schrag,2002). Howdo we rationalize this interpretation?We must¢rst step back and consider some basic features of ice dy-namics and their probable role in the development andmaintenance of a snowball Earth.

Theory suggests that ice dynamics play a critical role inthe advance of sea ice to the Equator (Warren et al., 2002;Goodman & Pierrehumbert, 2003; Lewis et al., 2003). Insimulations lacking ice dynamics, vertical mixing at theice lines halts the advance of sea ice, preventing a snowballEarth (Poulsen et al., 2001; Poulsen, 2003).The dynamicsof sea ice hasten the ice-line advance by transporting latentheat poleward and fresh water equatorward. The physicalmechanisms are varied, from dispersal of drift ice bywindsand currents (Lewis etal., 2003), to gravitational £owage ofthick (4100m) sea ice toward the zone of melting (Good-man & Pierrehumbert, 2003). The latter phenomenon isaptly described as a ‘sea glacier’ (Warren et al., 2002).Goodman & Pierrehumbert (2003) ¢nd that at dynamicsteady state, a sea glacier poised at the brink of ice-albedorunaway would be �0.2-km thick in polar and middle la-titudes, but would taper to zero thickness at the 0 1C iso-therm. Ice £ow velocities peak at 2 kmyr�1 near the iceedge and result in the deposition there of the meltwaterequivalent of 1.5m of additional precipitation annually(Goodman & Pierrehumbert, 2003).This freshwater inputnot only raises the freezing point, but also suppresses ver-tical mixing at the ice line.

Simulations with ice dynamics revive the classic conceptof ice-albedo instability, wherein the ¢nal invasion of seaglaciers through the tropics occurs rapidly ( �1.6 kyr ifmelting ceased) due to runaway ice-albedo feedback, afterwhich the ice slowly thickens for a few thousand years asthe ocean loses its reservoir of solar heat (Warren et al.,2002). Once the ice fronts meet at the equator, maximumice £ow velocities slow to just 55myr�1 in the simulation(Goodman & Pierrehumbert, 2003), but this in£ow ratemaintains �0.45 km of ice at the equator (vs. 0.2 km for sta-tic ice) and reduces ice thickness in the middle and polar la-titudes to �1.0 km (vs. 2 km at the poles in the static case).Equatorward £owage of the sea glacier is balanced by pole-ward vapor transport and by melting and freezing at thebase of the ice in the tropics and extratropics, respectively.

If a snowball glaciation took place, the synglacial sedi-mentary record on continental margins should re£ect thedynamic interplay between three types of ice mass (Fig.14): (1) the mobile sea glacier, (2) shorefast, multi-annual,sea ice (known as sikussak; Koch, 1945) on continentalshelves, shallow platforms and inland seas and (3) terres-trial glaciers and ice domes. Given the proximally dee-pened bathymetric pro¢le characteristic of glaciated



shelves (Anderson,1999), we may expect the shear zone be-tween the sea glacier and the shorefast sikussak to occur atthe outer edge of the shelf (Fig. 14b).The maximum thick-ness of sikussak ice, given adequatewater depth, should ap-proximate that of ‘static’ ice in the simulations describedabove. In reality, sikussak ice will be thinner due to subgla-cial meltwater runo¡ from adjacent ice domes, analogousto the thin ice on the present Antarctic Dry Valley lakes(e.g.McKay etal.,1985).Atmiddle and high latitudes, sikus-sak ice will be thicker than the adjacent sea glacier andwillreadily become entrained in the glacial £ow. In the subtro-pics and tropics, however, sikussak ice will be thinner thanthe bounding sea glacier, which will tend to encroach ontothe shelf and buttress the shorefast sikussak. Counterintui-tively then, abrupt invasion of sea glaciers into the tropicswould have the e¡ect of blocking the £ow of terrestrial gla-cier ice onto the platform (Fig. 14a, b).With the platformcovered by sikussak ice, the terrestrial ice sheet could nolonger calve and the source of dropstones would be re-moved (Funder et al., 1998; Dowdeswell et al., 2000).Thus,the abrupt termination of the Petrovbreen/Ulves� diamic-tites and associated ice-rafted debris need not signify theend of the glaciation. But is the MacDonaldryggen-ArenaFormation compatible with sedimentary processes activeon the platform during a snowball glaciation?

The sea glacier and shorefast sikussak would backstoptidewater glacial £ow (Fig. 14b), causing the terrestrial icesheet to slowly thicken (De Angelis & Skvarca, 2003), de-spite the severe reduction in precipitation rate when theocean is frozen. Net sublimation at the sea surface in thesubtropics would provide moisture borne by trades thatwould precipitate as ice in the coastal highlands borderingthe platform (Fig.14b). Numerical simulations by Donna-dieu et al. (2003) imply that a dynamic polythermal icesheet could develop in as little as180 kyr after snowball on-set from sublimation-derived moisture alone. If the icedome became even partially warm-based, ice loss wouldhave been accommodated by subglacial meltwater streamsrather than calving, and basal melting at the groundingline would have rendered shelf-ce free of debris, furtherreducing the chance of generating ice-rafted debris basin-ward of the grounding line (Carey &Ahmed,1961; Drewry&Cooper, 1981; Orheim&Elverhoi,1981; Alley et al., 1989).Net evaporative mass loss from the surface of the sikussakwould have caused di¡usive cooling through the ice col-umn and freezing of seawater to its base. Persistent upwardadvection of sikussak ice would have prevented any wind-blown, glacier-derived, or anchor ice-derived (Reimnitzet al., 1987) debris from raining out as ice-rafted debris.The notable absence of ice-rafted debris in the MacDo-naldryggen/Arena laminites distant from the groundingline is compatible with a frozen ocean.

Sublimation at the surface of the sikussak would causefreezing of sea water at its base, generating dense brinesthat would sink to the bottom and £ow basinward as den-sity currents. When ampli¢ed by tidal pumping, densitycurrents may reach velocities high enough to transportsand and cut sinuous channels 40m deep (Elverh�i et al.,

1989; Dowdeswell et al., 1998). During the Pleistocene, la-minated facies in Antarctica (Orheim & Elverh�i, 1981)and Greenland (Dowdeswell et al., 1998) are restricted tointerglacial periods when open water is available for seaice formation, and the resulting brine-driven traction cur-rents are able to deliver coarse sediment to distal, other-wise shale-dominated regions. On the Barents shelftoday, brine-driven gravity currents formed during seaice formation account for as much as 70%of the annual se-dimentation in distal environments (1.4^3.5 cm kyr�1), theother 30% coming mostly from ice-rafted debris en-trained in sea ice delivered from the Russian Arctic andsubordinately from the edges of buoyant, subglacially de-rived meltwater plumes (Elverh�i et al., 1989). The scour-ing, winnowing and rippling visible in grounding-lineproximal sands of the Arena Formation (Hambrey &Spencer, 1987) are consistent with brine-driven densitycurrents generated by freezing seawater at the base of thesikussak in response to surface sublimation.

Phase c: SlangenMember: sikussak melts away and a snowballoasis appears

Atmospheric CO2 concentrations and greenhouse forcingmust slowly rise on a snowball Earth because silicateweathering cannot balance volcanic outgassing. After mil-lions of years (Caldeira & Kasting, 1992), tropical sea-sur-face temperatures reach 0 1C and sea ice begins to melt.Whereas the stagnant sikussak could melt o¡ completely,creating an oasis of open water on the EGES platform(Fig. 14c), the tropical ocean would remain sealed o¡ be-cause of glacial in£ow from middle and high latitudes(Goodman & Pierrehumbert, 2003).We propose that theupper part of the MacDonaldryggen/Arena Formationand the Slangen regressive carbonate-evaporite sequence(Fig. 8) were deposited in a snowball oasis over the EGESplatform. Carbonate sediment was produced under eva-porative conditions in the oasis and quickly prograded toform a vast supratidal sabkha as base level fell. The exis-tence of extensive open water and warmer temperatureswould have intensi¢ed vapor transport and snow accumu-lation on downwind ice domes. Expansion of these andother ice sheets globally would have caused glacioeustaticlowering, consistent with seaward (northward) prograda-tion and subaerial exposure (o¥ap) of the Slangen carbo-nate bank. Weathering of carbonate detritus of glacialorigin could provide the alkalinity £ux necessary to preci-pitate new carbonate within the basin (Fairchild et al.,1989).

Kennedy et al. (2001) reported high d13C compositionsfrom purportedly syn-glacial (Marinoan) carbonates andinterpreted them to demonstrate relatively high fractionalburial of organicmatter during the glaciation.Whereas ourinterpretation of the 13C-enriched Slangen Member assyn-glacial is consistent with their ¢ndings, its isotopiccomposition may re£ect other factors, including the d13Cof the carbonate available for weathering during the glacia-tion (Higgins & Schrag, 2003). Additionally, the limited



dissolved organic carbon reservoir in the oasis would bedominated by the huge atmospheric reservoir (0.1^0.2 barCO2; Caldeira & Kasting, 1992). The latter would not beisotopically equilibrated with the ocean, but instead withvolcanic input with d13C �� 5% to � 7% (DesMarais &Moore, 1984). Consequently, the ��13% equilibriumfractionation between CO2 and calcite for seawater at 0 1C(Friedman &O’Neil, 1977) could result in carbonate beingprecipitated in the oasis with d13C as heavy as 6^8% (Hig-gins&Schrag, 2003), irrespective of any contribution fromcarbonate dissolution or organic matter burial.The rapidrise in d13C observed in the Slangen carbonates couldre£ect equilibration with the mantle-dominated atmo-spheric carbon reservoir.

On the landward part of the platform, no equivalent tothe SlangenMember is preserved.Glacial erosion and gla-cioisostatic loading had probably created a landward-dip-ping inner shelf region (consistent with SW-directedpaleocurrents in upper Arena Formation and youngerunits of EastGreenland;Hambrey&Spencer,1987), wheredeeper waters were maintained and where onshore windstended to stack seasonal sea ice and icebergs (Fig. 14c).The uppermost Arena Formation contains a sudden reap-pearance of ice-rafted debris, including boulder-sizedgranite dropstones (Moncrie¡ & Hambrey, 1988) that mayre£ect the ¢nal disintegration of the local sikussak, mobi-lizing outlet glaciers from the inland ice dome. The pre-sence of granite clasts indicates that ice streams haderoded down to basement in the hinterland.

Phase d:Wilsonbreen Fm: inland ice dome advances ontoa sand sea

Increased precipitation in the hinterland caused the icedome to expand and £owmore actively.The ice margin ad-vanced across the EGES platform (Fig. 14c, d), creating agrooved and striated peneplane (Moncrei¡ & Hambrey,1988) on which polymictic diamictites and sandstones oftheWilsonbreen/StoreelvFormationwere deposited. Sub-glacial channel sands and tills (containing abundant meta-morphic basement clasts) indicate that the ice sheet was atleast partly warm-based (Chumakov, 1968; Dowdeswell,1985; Harland et al., 1993). Sands that had collected on theexposed Slangen bank were reworked by northward-£ow-ing proglacial streams. These deposits (Br�vika Memberof the Wilsonbreen Formation) thicken northward at theexpense of the diamictites (Fig.8c), and coarsen upward re-lated to the approach of the glacier terminus (Fig. 14d).Sand wedges indicative of subaerial exposure are devel-oped on various horizons within the Wilsonbreen (Chu-makov, 1968; Harland et al., 1993) and Storeelv (Spencer,1985;Hambrey&Spencer,1987;Moncrie¡,1990) diamictites.Thin carbonate horizons in the middle Wilsonbreen For-mation of southern East Svalbard (Fig. 9) precipitated inevaporative, dry-valley lakes fed by glacial meltwater (Fair-child & Spiro, 1987; Fairchild et al., 1989).

As with the initial Petrovbreen/Ulves� glaciation, thehinterland ice front on the EGES platform was unstable

during the terminal Wilsonbreen/Storeelv reglaciation.Orbital forcing, sea-ice switching (Gildor & Tziperman,2000), binge-purge cycling (MacAyeal, 1993), ice-streammigration (Payne & Donglemans, 1997), or a host of otherfactors could account for the observed instability. Evensmall changes in glacial mass balance or local base levelcould dramatically change the paleogeography of the plat-form. As greenhouse forcing continued to rise, the hinter-land ice sheet ultimately expanded close to the presentnorthern exposed limits of the platform.

Phase e: Dracoisen cap dolostone: sea glaciers collapse andcataclysmic deglaciation ensues

With CO2 levels continuously rising, the melting rate ofthe tropical sea glacier eventually overtook the in£ow rateof ice from higher latitudes. The equatorial ice rapidlythinned and open water appeared, greatly intensifying thereverse ice-albedo andwater-vapor greenhouse feedbacks.The sea glacier quickly collapsed,melting away in a periodof �100 yrs (R.T. Pierrehumbert, pers. comm.).The timescale for melting the terrestrial ice sheets and the conse-quent sea-level rise is more uncertain, at least a millen-nium if the ice was 10 times thicker (2 km) than theterminal sea glacier (0.2 km, Goodman & Pierrehumbert,2003) to begin with. Following millions of years of deepfreeze, deglaciation was cataclysmic.With a pCO2 of �0.2bar needed to deglaciate a Neoproterozoic snowball Earth(Caldeira & Kasting, 1992), tropical sea surface tempera-tures were raised to �50 1C (R.T. Pierrehumbert, pers.comm.) in the greenhouse aftermath. Intense carbonate(dolomite) weathering delivered alkalinity to the oceanand drove the precipitation of cap dolostone on the timescale of the post-glacial transgression (Higgins & Schrag,2003). Rapid melting and surface warming imposed severedensity strati¢cation on ocean waters, impeding mixingbetween the oxic, brackish, surface water and euxinic, sal-ine, hydrothermally dominated, deep water.

Sturtian glaciation in Svalbard?

If the Polarisbreen diamictites jointly represent the Mar-inoan glaciation in Svalbard (Fig.14), thenwhere in the un-derlying Hecla Hoek stratigraphy is the older Sturtianglobal glaciation (Knoll, 2000)? No older glacial depositshave been found (and may never be given the limited areaof exposure), but perhaps this is not surprising consider-ing that the vast conformable succession of shallow-watercarbonates would have provided no topography for glacialnucleation. Glacial deposits are largely absent from theMarinoan-age carbonate platform in Namibia for thatreason, but are found on the upper continental slope(Ho¡man et al., 1998b; Ho¡man, 2002), a facies notencountered in Svalbard. Post-glacial cap carbonates, onthe other hand, are transgressive and normally extend farbeyond the limits of the related glacial deposits (Ho¡man& Schrag, 2002).

Sturtian cap carbonates have characteristic features inAustralia, Canada and Namibia (Kennedy et al., 1998),



which should facilitate its recognition in Svalbard in theabsence of Sturtian glacial deposits. They are generallydark grey, regressive sequences that begin with argillac-eous limestone containing allodapic beds or, less com-monly, sublittoral microbial laminae.They are condensed,compared with Marinoan cap carbonates, and the trans-gressive tracts that host the declining arm of the d13Canomaly are generally missing. Consequently, d13C riseswith stratigraphic height from � 2% or lower at the baseto highs of15% (Kennedy etal.,1998;McKirdy etal., 2001;Yoshioka et al., 2003) (Fig. 13). Ho¡man & Schrag (2002)speculated that Sturtian seawater did not reach criticaloversaturation with respect to CaCO3 until after the post-glacial transgression, and that the weathering £ux ofalkalinity was lower than in the Marinoan because lesscarbonate rock was available globally for submarine dis-solution andweathering.

Even though negative d13C values occur within theAka-demikerbreen carbonates (Knoll et al., 1986; Kaufmanet al., 1997; Halverson et al., 2003), we argue that the lowerRuss�ya Member is the closest analog beneath the Pet-rovbreen diamictite to the Rasthof Formation cap carbo-nate. Like the Rasthof (Ho¡man et al., 1998; Kennedy,1998), it comprises a single, shoaling-upward half-se-quence that is relatively rich in organic matter and beginswith deep-water, bluish-grey to black, argillaceous lime-stone, intercalated in some sections with thin dolomitebeds. In the most carbonate-rich section (measured sec-tion 1) d13C rises from1.5% to 5% in the lower half of thesequence. Allowing for poor exposure and lack of carbo-nate in the lowermost Russ�ya, the d13 C pro¢le matchesthat of the Rasthof Formation (Yoshioka et al., 2003).

If the lowerRuss�yaMember is aSturtian cap carbonate,then based on characteristics of other equivalent cap carbo-nates, the glacial surface should be at the top of the Dart-board Dolomite (Fig. 4). The contact is poorly exposed,and sedimentologically, the Dartboard itself is as strangeas its name. In fact, many of its distinctive features (isopa-chous sea- £oor cements, macropeloids, wave-like bed-forms, sublittoral microbial laminae) are similar toMarinoan cap dolostones. Most Sturtian cap carbonateslack transgressive cap dolostones and begin at the maxi-mum £ooding stage or later (Ho¡man & Schrag, 2002), sowhy should a 13C-enriched Sturtian cap dolostone appearin Svalbard (no equivalent is known in East Greenland)? Itis tempting to speculate that platformwaters over Svalbardreached critical oversaturationwith respect to the dolomiteprecursor earlier than elsewhere because of a local source ofalkalinity, which in this case would have been the strongly13C-enriched Akademikerbreen carbonate platform.

If the DartboardDolomite were the Sturtian cap dolos-tone, by analogy to Marinoan cap dolostones, its baseshould be the glacial surface. Unfortunately, the base ofthe Dartboard is no better exposed than the top, and noobvious sequence boundary has been identi¢ed at thisstratigraphic level.The variable depth of deposition, fromsublittoral (Knoll et al., 1989) to shallow/intermittently ex-posed, implied by the range in facies in the Dartboard

Dolomite, suggests signi¢cant relief on this contact. Onthe other hand, the top of the Dartboard Dolomite islocally brecciated (Fairchild & Hambrey, 1984), and absentaltogether in northernmost Spitsbergen, leaving openingthe possibility that theDartboardDolomite, while litholo-gically similar to Marinoan-aged cap dolostones is notpart of the Sturtian cap-carbonate sequence.

Irrespective of the origin and timing of the Dartboarddolomite, it is clear that the EGES platform experienceda profound change in depositional style in the uppermostAkademikerbreen Group (Herrington & Fairchild, 1989).Fairchild & Hambrey (1995) attributed the sudden in£uxof ¢ne siliciclastic sediment in the underlying KinnvikkaMember (Fig. 4) and local deepening on the EGES plat-form to tectonic uplift.While tectonic uplift mayhave beena factor, the evidence for high-amplitude £uctuations inbase level, including the development in measured section10 of a 45m-high, pinnacle reef with multiple debrisaprons, is highly suggestive of a glacioeustatic oscillation.An alternative explanation for the in£ux of ¢ne clasticsediment in the Kinnvikka Member could be transientlandscape disequilibrium associated with climatic dete-rioration during the incipient Sturtian glaciation (Zhanget al., 2001).

The hypothesis that the base of the PolarisbreenGroup(Russ�ya Member) is the cap carbonate to a cryptic Stur-tian glaciation is speculative, and we regard the problemof the Sturtian glaciation on the EGES platform as unre-solved.We cannot rule out that the Sturtian glaciation isrepresented either in the underlying AkademikerbreenGroup or by the Petrovbreen Member, but global correla-tions that will be developed in a separate paper renderthese options less likely.The potential key interval (Kinn-vikka^Dartboard^lower Russ�ya) is exposed in an unex-plored area directly west of the Vestfonna Ice Dome onNordaustlandet (Fig.1).

CONCLUSIONS

The Neoproterozoic Polarisbreen Group of northeasternSvalbard contains two diamictite-rich units, the olderPetrovbreen Member and the younger Wilsonbreen For-mation, that are conventionally interpreted to representtwo distinct glaciations. The Petrovbreen diamictites areoligomictic, derived from directly underlying carbonatesof the Russ�yaMember, and thinner overall than theWil-sonbreen diamictites, which are polymictic and includeextrabasinal basement material. Awell-de¢ned, 10%, ne-gative d13C anomaly occurs in the Russ�ya Member di-rectly beneath the erosion surface on which thePetrovbreen diamictites were deposited. This anomaly ishomologous (Fig. 13) with those observed beneath theElatina glacials in South Australia (McKirdy et al., 2001),the Ghaub glacials in northern Namibia (Halverson et al.,2002), and the Stelfox glacials in northwestern Canada(Ho¡man & Schrag, 2002), all of which are correlated withthe Marinoan glaciation in the South Australian type sec-tion (Kennedy et al., 1998; Ho¡man & Schrag, 2002). The



Petrovbreen diamictites have no well-developed post-gla-cial cap carbonate, but the younger Wilsonbreen diamic-tites have a cap dolostone (basal Dracoisen Formation) thatis similar compositionally, sedimentologically and isotopi-cally (Fig. 13) to Marinoan-age cap dolostones in Australia(Kennedy, 1996), Canada (James et al., 2001) and Namibia(Ho¡man & Schrag, 2002). Thus, from the perspective ofthe bounding strata and global correlations, the two dia-mictite-rich units in Svalbard appear to represent the be-ginning and end of a singleMarinoan glacial cycle.

We propose a new interpretation of the glacial history ofthe EGES platform that is consistent with these correla-tions andwith the snowball Earth hypothesis for theMar-inoan glaciation. In Polarisbreen times, the EGES marineplatform was situated in the subtropics of the southernhemisphere on the south-facing margin of Laurentia, re-presented by the hinterland of present-day East Green-land. Carbonate sedimentation of the Russ�ya Memberendedwith a450m base-level drop that heralded the gla-ciation. Adjacent to the platform in East Greenland, an icedome built up with cold-based ice descending to tide-water. The ice dome deposited thick tills at the ice frontand released debris-laden icebergs that rafted the veneerof Petrovbreen glacimarine diamictite across the platform.Overlying the ice-rafted debris is a thick unit (MacDonal-dryggenMember) of ¢nely laminated suspension depositsthat lack dropstones but are speckledwith calcite pseudo-morphs that we interpret to be glendonites.We regard thePetrovbreen^MacDonaldryggen contact as the onset ofthe full-blown snowball Earth, when ‘sea glaciers’ (thick£oating marine ice that £ows toward the zone of meltingor ablation) invaded the tropical oceans (Goodman&Pier-rehumbert, 2003). The £ux of ice-rafted debris disap-peared when ‘sikussak’ ice (multi-annual, shorefast, seaice) choked the EGES platform.The sikussak backstoppedthe ice dome and prevented calving, and may itself havebeen pinned to the platform by the sea glacier impingingon the shelf edge.The MacDonaldryggen suspension de-posits accumulated beneath the sikussak ice from hypo-pycnal plumes and density £ows issuing from ice streamsbeneath the slowly growing ice dome.

As the snowball Earth gradually warmed (becauseweathering could not keep pace with CO2 outgassing), va-pour transport increased and the global sea glacier £owedmore rapidly. In£ow of sea ice from higher latitudes main-tained thick ice on the tropical ocean even after sea-sur-face temperatures reached the melting point. Theshorefast sikussak ice, however,melted away, leaving an oa-sis of open water over the EGES and other subtropicalmarine platforms and inland seas, while the world oceanremained ice covered.We interpret the Slangen Member,a regressive sabkha-like evaporitic dolomite sequenceoverlying the MacDonaldryggen Member, to be the sedi-ment deposited in this oasis on the snowball Earth. Open-ing of the EGES oasis led to rapid expansion of thedownwind ice dome, which advanced across the shelf anddeposited the Wilsonbreen ice-contact deposits. In frontof the advancing ice, sands of the Br�vika Member were

redeposited by outwash streams.TheWilsonbreen ice sheethad advanced almost to the present northern limit of thePolarisbreen outcrop belt when, in our interpretation, thesea glaciers ¢nally collapsed, leading to cataclysmic degla-ciation and attendant sea level rise. The basal Dracoisencap dolostone is the signature of this transgression. Ironi-cally, our interpretation associates the full-blown snowballEarth with two units that are not overtly glacial in origin,the MacDonaldryggen Member suspension deposits andthe SlangenMember sabkha cycle.

If both diamictite-rich units in Svalbard belong to theMarinoan glaciation, the older Sturtian glaciation shouldbe represented in the underlying strata (Knoll, 2000). Noglacial deposits are known beneath the Petrovbreen/Ulve-s� diamictites.This is perhaps not surprising consideringthe conformable nature of the vast, shallow-marine, car-bonate platform (Akademikerbreen Group), which wouldprovide no topography on which terrestrial glaciers mightnucleate. However, a possible Sturtian cap-carbonate se-quence is found in the lowerRuss�yaMember, which is se-dimentologically similar to the Rasthof Formation inNamibia (Ho¡mann and Prave, 1996; Ho¡man et al., 1998;Kennedy et al., 1998; Yoshioka et al., 2003).The basal Rus-s�yaMember lacks a negative d13C anomaly, but it shows apositive d13C trend like that of other Sturtian cap carbo-nates (Kennedy et al., 1998).

Evidence for open water on shallow platforms duringthe Marinoan glaciation has been reported in other re-gions (e.g., Williams, 1996; Prave, 1999; Williams &Schmidt, 2000; Kennedy et al., 2001; Condon et al., 2002).Since Kirschvink (1992), this evidence has always been ci-ted in opposition to the snowball hypothesis.We concludethe opposite, that shorefast sea ice in the relatively ablativesubtropics should melt away, exposing seawater on shallowshelves and inland seas, well before the global sea glacier¢nally collapsed. Oases should develop as a natural conse-quence of the snowball cycle, and theyhighlight the role ofsea glaciers in the critical phases of the cycle (Goodman &Pierrehumbert, 2003).More generally, they remind us thatwhile the Neoproterozoic stratigraphic records we rely onare epicontinental, the snowball Earth cycle is fundamen-tally an oceanographic phenomenon.

SUPPLEMENTARYMATERIAL

The following material is available from:http://www.blackwellpublishing.com/products/journals/suppmat/BRE/BRE234/BRE234sm.htm

Appendix A1Data table of d13C and d18O data presentedin this paper.

ACKNOWLEDGMENTS

This workwas supported by theNational Science Founda-tion (Arctic Science Program grant OPP-9817244),the NASA Astrobiology Institute, the Canadian Institute



for Advanced Research (Earth System Evolution Project),and a GSA graduate research grant. Norsk Polarinstituttprovided logistical support in Svalbard, and WinfriedDallman helped us greatly in organizing our ¢eld work.We thank Dan Schrag for use of his laboratory and manydiscussions about the data in this paper. Ian Fairchild, Pe-ter DeCelles, and an anonymous referee provided excel-lent reviews that greatly improved this paper, as didcomments by Andy Knoll on an earlier draft. We thankSam Bowring, Dan Condon, Jason Goodman, John Grot-zinger, Matt Hurtgen, Stein Jacobsen, Jay Kaufman, PeterMoore, Ray Pierrehumbert, Susannah Porter, PascalePoussart, and Steve Warren for valuable discussions. CrisCarman, Anne Estoppey, John Higgins, Leslie Hsu, BriceJones, Peter Moore, and Meg Smith, provided ¢eld assis-tance. Ethan Goddard and Greg Escheid supervised thelaboratory work.

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Manuscript received 20 August 2003;Manuscript Accepted:21April 2004



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BasinResearch The Marinoanglaciation …late Neoproterozoic Polarisbreen Group (Fig.1) sits atop 2km of conformable carbonate strata, the Akademikerb-reenGroup,andwithin its mixed