NE Svalbard and East Greenland), and a Primary Origin for the Upper

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Precambrian Research 206–207 (2012) 137–158

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Precambrian Research

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Cryogenian glaciations on the southern tropical paleomargin of Laurentia

(NE Svalbard and East Greenland), and a primary origin for the upper

Russøya (Islay) carbon isotope excursion

a,b,∗ c,d e f

Paul F. Hoffman , Galen P. Halverson , Eugene W. Domack , Adam C. Maloof ,

f c,d

Nicholas L. Swanson-Hysell , Grant M. Cox

Department of Earth & Planetary Sciences, Harvard University, 20 Oxford Street, Cambridge, MA 02138, USA

School of Earth and Ocean Sciences, University of Victoria, Box 1700, Victoria, BC V8W 2Y2, Canada

Centre of Tectonics, Resources and eXploration (TRaX), School of Earth and Environmental Sciences, The University of Adelaide, Adelaide, SA 5005, Australia

Department of Earth and Planetary Sciences/GEOTOP, McGill University, 3450 University Street, Montréal, QC H3A 0E8, Canada

Department of Geosciences, Hamilton College, 198 College Hill Rd., Clinton, NY 13323, USA

Department of Geosciences, Princeton University, Guyot Hall, Princeton, NJ 08544-1003, USA

a r t i c l e i n f o a b s t r a c t

Article history: Neoproterozoic successions in NE Svalbard and East Greenland host a pair of glacigenic formations, the

Received 5 May 2011

younger of which are correlated with the terminal Cryogenian (Marinoan) glaciation based on their

Received in revised form 17 February 2012

lithologically and isotopically diagnostic cap dolostones. A deep negative carbon isotope excursion (CIE)

Accepted 29 February 2012

occurs stratigraphically beneath the older glacigenic formations in both areas, but no CIE is preserved

Available online 13 March 2012

beneath the younger glacial horizon in either area. This led to uncertainty over the age of the CIE and

87 86

the paleoenvironmental signiﬁcance of the units separating the glacigenic formations. Sr/ Sr ratios in

Keywords:

Sr-rich limestone associated with the CIEs are ≤0.7064 in East Greenland and ≤0.7068 in NE Svalbard,

Neoproterozoic

Cryogenian consistent with early Cryogenian values globally and inconsistent with late Cryogenian ratios, which

≥

are exclusively 0.7071. The CIEs are tentatively correlated with the Islay excursion in the Scottish and

Carbon isotopes

Islay excursion Irish Caledonides, and potentially correlative subglacial CIEs in northwestern Canada, northwestern Tas-

ı Snowball Earth mania and the southwestern United States. In NE Svalbard, newly-acquired Corg data covary with

ı13

Jormungand Ccarb across the CIE, but the organic excursion is roughly half the amplitude of the inorganic excursion.

ı13 18 13

Ccarb and ı Ocarb do not covary, nor does ı Corg vary as a function of total organic carbon content.

A primary origin for the CIE is supported and the accompanying proliferation of stromatolites suggests

rising carbonate saturation (falling pCO2) prior to glaciation. New data suggest that the hiatus responsi-

ble for the missing Trezona CIE occurs at the top, not the bottom, of the Bråvika member sandstone in

NE Svalbard. A dramatic decline in regional average thickness of older Cryogenian glacial deposits from

the paleomagnetically-determined subtropics (British Isles) to the deep tropics (Svalbard) is consistent

with meteoric net accumulation and sublimation zones on a snowball Earth, and inconsistent with the

Jormungand climate zonation.

1. Introduction 1982; Fairchild and Hambrey, 1984, 1995; Dowdeswell et al., 1985;

Fairchild et al., 1989; Fairchild and Spiro, 1990; Harland et al.,

Inspired by the ambient glaciers, a distinguished parade of geol- 1993; Harland, 1997; Halverson et al., 2004; Halverson, 2011).

ogists has examined the Neoproterozoic glacial deposits exposed Two discrete glacigenic units (Fig. 4) occur within the Polarisbreen

in the Caledonian (Siluro-Devonian) thrust-fold belt (Figs. 1–3) of Group (Hambrey, 1982; Halverson, 2011), a succession of carbonate

NE Svalbard (Kulling, 1934; Harland and Wilson, 1956; Wilson and and mainly ﬁne-grained siliciclastic rocks of Cryogenian through

Harland, 1964; Chumakov, 1968, 1978; Edwards, 1976; Hambrey, early Ediacaran age that sits atop the Akademikerbreen Group,

a Neoproterozoic carbonate-dominated succession 1350–2500 m

thick (Wilson, 1961; Knoll and Swett, 1990; Fairchild et al., 1991;

∗ Halverson, 2006; Halverson et al., 2007b). Distinct stratigraphic

Corresponding author at: 1216 Montrose Ave., Victoria, BC V8T 2K4, Canada.

nomenclature evolved on the islands of Spitsbergen and Nor-

Tel.: +1 250 380 0059; fax: +1 250 592 5528.

E-mail address: [email protected] (P.F. Hoffman). daustlandet (Kulling, 1934; Winsnes, 1969; Harland, 1997), but for

138 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 1. Neoproterozoic glacigenic deposits on the Laurentian (closed triangles) and Baltic (open triangles) paleomargins of the Caledonian (Siluro-Devonian) orogen in a pre-

Eocene continental restoration (Bullard et al., 1965), modiﬁed after Spencer (1975). The 718 Ma (early Cryogenian) paleoequator is based on paleomagnetic poles from the

Franklin Large Igneous Province of Arctic North America (Evans and Raub, 2011) and applies only to the Laurentian paleomargin. Devonian left-slip on the Billefjorden (BFZ)

and other orogen-parallel faults displaced NE Svalbard from its Cryogenian location in the Greenland Sea area (Harland and Gayer, 1972; Harland, 1997). Early Cryogenian

glacigenic deposits are unknown on the Baltic paleomargin, which did not exist in the position shown (in grey) before the Devonian.

simplicity, Halverson et al. (2004) and this paper elect to use the the GSSP (Knoll et al., 2006) and globally (Halverson et al., 2004;

Spitsbergen nomenclature throughout the 160-km-long outcrop Hoffman et al., 2011).

belt. The two glacigenic sequences are separated by a non-glacial

The older glacigenic unit is a relatively thin (0–52 m) glacial interval comprising the upper Elbobreen Formation (Fig. 4). The

marine interval named the Petrovbreen Member (Hambrey, 1982) MacDonaldryggen Member (Hambrey, 1982) is a 200-m-thick

of the Elbobreen Formation (Fig. 4), the oldest formation in the highstand sequence of dark-coloured, parallel-laminated siltstone

Polarisbreen Group. It consists of mainly well-stratiﬁed, carbonate- with thin allodapic dolostone beds and stellate pseudomorphs

clast diamictite with glacially faceted and striated clasts (Hambrey, of calcite (Fig. 5a), possibly after ikaite, that predate compaction

1982; Harland et al., 1993). The glacial marine sequence onlaps an (Halverson et al., 2004). The Slangen Member (Hambrey, 1982)

erosion surface with at least 50 m of stratigraphic relief, developed is a 20–30-m-thick regressive sequence of microbialaminated

on stromatolitic dolostone of the upper Russøya Member (Fairchild and evaporitic dolarenite and quartz-dolarenite (Halverson et al.,

and Hambrey, 1984; Halverson et al., 2004; Halverson, 2011) of the 2004). The informal Bråvika member (Halverson et al., 2004) is a

Elbobreen Formation. northward-thickening, coarsening-upward wedge of well-sorted

The younger glacigenic sequence is the Wilsonbreen Formation and well-rounded quartz-chert arenite with large-scale crossbed-

(Fig. 4), which is thicker on average (∼117 m) and composed pre- ding and thin intercalations of ﬂaggy and stromatolitic dolostone.

dominantly of non-stratiﬁed, polymictic diamictite with faceted Its reciprocal relationship with the northward-tapered Wilson-

and striated clasts, and subordinate channelized sandstone and breen diamictite and its northward-directed paleocurrents led

conglomerate (Edwards, 1976; Hambrey, 1982; Harland et al., Halverson et al. (2004) to suggest that the Bråvika member rep-

1993; Halverson et al., 2004; Halverson, 2011). A thin but resents a terrestrial proglacial outwash (sandur) facies of the

widespread band of glacio-lacustrine limestone and dolostone beds Wilsonbreen Formation. In this interpretation, the compositional

is sandwiched between diamictite and associated facies of the and textural maturity of the arenite implies recycled sand. The base

Ormen and Gropbreen members (Fairchild and Hambrey, 1984; of the Bråvika member is conformable with the Slangen Member

Fairchild et al., 1989; Fairchild and Spiro, 1990; Fairchild, 1993; Bao (Fig. 5b) and the Bråvika member appears to have been unlithi-

et al., 2009; Halverson, 2011). Calcite-hosted trace sulfate in these ﬁed when overidden by the Wilsonbreen ice sheet, based on the

beds is extraordinarily depleted in O, indicating a high atmo- unfractured nature of the contact and abundant reworked quartz

spheric CO2 concentration, potentially consistent with a long-lived sand (but no sandstone clasts) in the overlying basal Wilsonbreen

global or near-global glaciation (Bao et al., 2009). The Wilsonbreen diamictite. The Bråvika member was formerly included in the

diamictite is conformably overlain by the Dracoisen Formation, the Wilsonbreen Formation (Halverson et al., 2004; Halverson, 2011)

base of which has many of the lithologic, stratigraphic and iso- but here it is placed in the Elbobreen Formation (Figs. 3 and 4)

topic hallmarks of the basal Ediacaran cap-carbonate sequence at because of new evidence to be discussed.

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 139

Fig. 2. Bedrock geology of the heavily-glaciated Caledonian thrust-fold belt in northeastern Spitsbergen and adjacent Nordaustlandet, Svalbard, showing the locations of

measured sections (MS 1, 7, 8 and 9) of the lower Elbobreen Formation (Russøya Member), from which the paired organic and inorganic carbon isotope data were obtained

across the ‘Islay’ excursion (Fig. 6). The box indicates the area of Fig. 3. ESZ = Eolusletta Shear Zone; LFZ = Lomfjorden Fault Zone.

140 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 3. Bedrock geology of the Caledonian thrust-fold belt on Gotiahalvøya, northwestern Nordaustlandet. Note the pair of synclines in Polarisbreen Group strata south of

Sveanor, the intervening thrust tip, the system of conjugate transcurrent faults of late Caledonian age that accommodate WSW-ENE shortening and NNW-SSE extension (i.e.,

horizontal plane strain), and the overall increase in exposed stratigraphic depth from west to east. Locations of measured sections MS7 and MS8 (Fig. 6), P7601 and P7609

(Fig. 13) are indicated. The area was earlier sketch-mapped by Kulling (1934) and Flood et al. (1969).

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 141

Fig. 4. Composite stratigraphic column of the Polarisbreen Group of NE Svalbard interpreted (a–e) according to the single glaciation ‘sikussak-oasis’ model (Halverson et al.,

2004) and (f–j) the standard double glaciation model preferred in this paper. Crossed circle in sea glacier indicates motion directed into the page.

1.1. The Russøya carbon isotope excursion (CIE) permanent shorefast sea ice (Fig. 4b), also known as sikussak, and

the evaporitic Slangen Member formed when tropical sea-surface

Halverson et al. (2004) documented a remarkable carbonate temperatures reached the melting point, allowing open water in

ı13

Ccarb excursion in the upper Russøya Member, directly beneath basins with sufﬁcient inﬂow of water to balance evaporation, yet

the older glacial horizon (Fig. 6a). Over a stratigraphic interval of physically protected from sea glacier invasion (Fig. 4c). There are

ı −

<50 m, Ccarb values fall steadily from +5‰ to 6‰ (VPDB), before two reasons why this model has become obsolete. First, deep neg-

partially recovering to −4‰ beneath the Petrovbreen diamictite. ative CIEs now are known to exist beneath both Cryogenian glacial

The CIE is preserved in areas of positive relief on the subglacial horizons in NW Canada (Halverson, 2006; Macdonald et al., 2010)

erosion surface; it is truncated or absent in erosional depres- and the Irish-Scottish Caledonides (McCay et al., 2006; Prave et al.,

sions, where the diamictite is thickest (Halverson et al., 2004). 2009; Sawaki et al., 2010). The older subglacial CIE is named the

The stratigraphic relationship demonstrates that the CIE predates Islay excursion (Prave et al., 2009) after the Islay Limestone in SW

the subglacial erosion surface. In contrast, the Slangen Mem- Scotland, where the CIE was ﬁrst documented (Brasier and Shields,

87 86

ber beneath the younger glacial horizon is uniformly enriched in 2000). Second, Sr/ Sr values between 0.70672 and 0.70677 in

Ccarb and no negative CIE has been found anywhere, including Sr-rich limestone of the upper Russøya Member (Halverson et al.,

the dolostone intercalations within the Bråvika member. This led 2007a) are consistent with older Cryogenian (<0.7068) but not with

Halverson et al. (2004) to hypothesize that both the Petrovbreen younger Cryogenian (>0.7071) values globally (Brasier and Shields,

and Wilsonbreen diamictites belong to the terminal Cryogenian 2000; Halverson et al., 2007a, 2010). Parsimony (Fig. 4f–j) now

global glaciation, thereby allowing the upper Russøya CIE to be cor- demands correlation of the upper Russøya CIE with the Islay excur-

related with the late Cryogenian Trezona excursion (McKirdy et al., sion, the Petrovbreen Member with the older Cryogenian (Rapitan)

2001; Halverson et al., 2002, 2005; Hoffman and Schrag, 2002), glaciation, the Wilsonbreen Formation with the terminal Cryoge-

while the Dracoisen cap dolostone marks the base of the Ediacaran. nian (Marinoan) glaciation, and the base of the Dracoisen Formation

To account for the MacDonaldryggen and Slangen members with the Ediacaran GSSP. The MacDonaldryggen and Slangen mem-

as products of a global or near-global glaciation, Halverson et al. bers represent the clement intermission between the pan-glacial

(2004) created a time-dependent ice dynamic model (Fig. 4a–e), intervals (Fig. 4g–h), and the absence of the Trezona excursion

in which the Macdonaldryggen Member accumulated beneath implies a signiﬁcant hiatus beneath the Wilsonbreen Formation.

142 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 5. Sedimentary features in the Elbobreen Formation in Murchisonfjorden, Nordaustlandet, Svalbard: (a) glendonite nodules (white) on bedding planes of parallel-

laminated siltstone, upper MacDonaldryggen Member, 3.0 km southeast of Sveanor (Fig. 3); (b) conformable contact between dolostone ribbons (beneath pen) of the Slangen

Member, exhibiting dolomitized molar-tooth structure (rare post-Sturtian example), and quartz-chert-arenite (above pen tip) of the Bråvika member, at Sveanor; (c) stereo-

pair of columnar stromatolite bioherms (form-genus Collenia, Raaben, 1969) in the upper Russøya Member, near Kinnberget, north side of Murchisonfjorden, and (d) helical

columnar stromatolite biostrome (form-genus Kussiella, Raaben, 1969) in the uppermost Russøya Member, near Sveanor. (c) and (d) illustrate the proliferation of columnar

stromatolites associated with the Islay ı Ccarb excursion. Coin 2 cm diameter; pens 15 cm long.

There is accordingly no need for the sikussak-oasis model in NE a southward Cryogenian migration is compatible with a moderately

Svalbard, but it remains conceptually valid in principle. The recent reliable paleopole from the basal Ediacaran cap dolostone (Mirassol

correction (Li and Pierrehumbert, 2011) to the sea glacier ﬂow Formation) in SW Brazil (Trindade et al., 2003), assuming a con-

model (Goodman and Pierrehumbert, 2003), yielding a substan- ventional Amazonia-southern Laurentia palinspastic connection. If

tially smaller pole-to-equator ice thickness gradient, makes the correct, this transit is of particular interest because the deep tropics

sikussak-oasis model more widely applicable, in theory, because and subtropics are the climatic zones that offer the clearest dis-

it magniﬁes the difference in thickness between the tropical sea tinction between the snowball Earth and Jormungand (Abbot et al.,

glacier and co-zonal sikussak ice. 2011) climate states. The Jormungand state has a narrow strip of

open water in the deep tropics, broad zones of ablative, relatively

dark ice (albedo ≤0.55) in the subtropics and white, snow-covered

1.2. Paleogeographic setting

ice poleward of the subtropics (Abbot et al., 2011). In the Jor-

mugand state, precipitation exceeds evaporation–sublimation in

Maﬁc igneous rocks of the ca 718 Ma Franklin Large Igneous

the deep tropics, while the subtropics experience net sublimation

Province (LIP) have been sampled all across Arctic Canada and

(Abbot et al., 2011). In a snowball Earth, the hydrologic vectors

northwest Greenland, yielding a high-quality, compaction-proof,

are switched around: sublimation exceeds precipitation in the

grand mean, paleomagnetic pole that centers Laurentia on the

deep tropics, while the subtropics experience net precipitation

paleoequator when the older Cryogenian Rapitan glaciation began

∼ (Pierrehumbert et al., 2011). If the Cryogenian glaciations were Jor-

717 Ma (Denyszyn et al., 2009; Macdonald et al., 2010; Evans and

mungand states, the Petrovbreen glaciation would have occurred

Raub, 2011). In a pre-North Atlantic (pre-Eocene) reconstruction

in a zone of net precipitation and the Wilsonbreen glaciation in

(Fig. 1), NE Svalbard restores offshore North Greenland, exactly

a zone of net sublimation. If they were snowball Earth states, the

on the 718 Ma paleoequator (Evans and Raub, 2011). However, it

Petrovbreen glaciation occurred in a zone of net evaporation and

moved northward relative to cratonic Laurentia during Devonian

the Wilsonbreen glaciation in a zone of net precipitation.

left-slip faulting and its Cryogenian location may have lain several

degrees south of the paleoequator in the Greenland Sea (Fig. 1) area

(Harland and Gayer, 1972; Harland, 1997; Maloof et al., 2006). 1.3. Present study

Preliminary paleomagnetic data from the ca 615 Ma Long Range

Dykes of western Newfoundland and Labrador (McCausland et al., In 2004, the ﬁrst four listed authors studied three sections of

◦ ◦

2009) suggest that Laurentia had drifted 17 ± 12 southward by the Tillite Group (Haller, 1971; Hambrey and Spencer, 1987) in the

early Ediacaran time, carrying the Greenland Sea margin of Lau- Fjord Zone of East Greenland (Fig. 1). Our goals were (1) to com-

rentia from the deep tropics during the during the Petrovbreen pare the broadly correlative Tillite and Polarisbreen groups (Fig. 7),

glaciation into the subtropics after the Wilsonbreen glaciation. Such and the underlying carbonates, as formerly adjacent parts of the

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 143

Fig. 6. Litho- and chemostratigraphy of the upper Backlundtoppen Formation (Akademikerbreen Group) and the Russøya Member of the Elbobreen Formation (Polarisbreen

ı13 13 18

Group) in NE Svalbard: (a) Ccarb, (b) ı Corg and (c) ı Ocarb. The stratigraphic column (left) combines section MS1 (Backlundtoppen Formation) and MS7 (Russøya Member)

of Halverson et al. (2004). DD stands for Dartboard Dolomite Member and P stands for Petrovbreen Member. Stable isotope data containing the Islay excursion are shown from

sections MS1 (circles), MS7 (diamonds), MS8 (triangles) and MS9 (squares). Three data points from MS9 (squares) that appear to lie within the basal Russøya Member actually

lie within the Backlundtoppen Formation in strata above the Dartboard Dolomite Member but below the sequence boundary. Similarly, a data point from MS8 (triangles)

that appears to lie within the Petrovbreen Member in the columnar section (MS7) actually represents the uppermost Russøya Member where less erosional truncation occurs

ı13 13 13 beneath the diamictite. Cross-plots illustrate: (d) correlation between Ccarb and ı Corg, (e) lack of a clear relationship between ı Corg and total organic carbon (TOC)

18 13 2

content, and (f) relative invariance of ı Ocarb during large changes in ı Ccarb. Squared correlation coefﬁcients (R ) and the slope in (d) were calculated without the outlying

data point (grey cross).

Laurentian paleomargin (Harland and Gayer, 1972; Harland, 1997); 1987; Moncrieff and Hambrey, 1989, 1990; Moncrieff, 1988;

(2) to compare the basal Canyon Formation (Fig. 7), lithologically Fairchild and Hambrey, 1995; Stouge et al., 2011), as are the car-

and isotopically, with basal Ediacaran sequences globally (Knoll bonates of the upper Eleonore Bay Group (Bed groups 18–20)

et al., 2006); (3) to compare the pairs of glacigenic sequences in East equivalent to the Russøya Member (Herrington and Fairchild, 1988;

Greenland and NE Svalbard as more proximal and distal deposits, Fairchild et al., 2000). Our work ampliﬁes but does not alter the

respectively, with respect to the interior of Laurentia (Fig. 4), and main conclusions of these authors.

also in terms of drift from the deep tropics to the subtropics over We examined the Tillite Group and subjacent strata in three

Cryogenian and earliest Ediacaran time; and (4) to test the sikussak- areas over a strike length of 92 km (Fig. 8)—on the south side of Bro-

oasis model as applied to the Arena Formation (Fig. 7) of the Tillite getdal in Strindberg Land, at Tillitekløft (Fig. 9b) near Kap Weber

Group. In 2007, the ﬁrst three authors made additional sedimento- in Andrée Land, and Kap Oswald (Fig. 9a) on Ella Ø in Kong Oscar

logical and isotopic observations on westernmost Nordaustlandent Fjord. The third area is located behind Lauge Koch’s mid-20th cen-

in Svalbard (Figs. 1–3). As well, the ﬁfth listed author measured tury Expedition Base in Solitärbugt (Poulsen and Rasmussen, 1951).

bulk organic carbon isotopes (ı Corg) in 60 samples selected from Coordinates for the base of each section are given in Fig. 10. These

three of the sections used to deﬁne the upper Russøya carbonate areas were described by Hambrey and Spencer (1987).

CIE (Halverson et al., 2004). This paper summarizes and discusses

the resulting ﬁeld observations and isotopic data.

2.1. Upper Eleonore Bay Group

2. Tillite Group, Fjord Zone, East Greenland At Kap Oswald on Ella Ø (Poulsen and Rasmussen, 1951;

Hambrey and Spencer, 1987; Herrington and Fairchild, 1988;

The Tillite Group and its glacigenic formations (Fig. 7) are Fairchild et al., 2000), a ridge-forming dolostone unit in Bed group

described well in the existing literature (Hambrey and Spencer, 18 of the upper Eleonore Bay Group is characterized by tepee

144 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 7. Stratigraphic correlations between the upper Eleonore Bay Group (El. B. Gp) and Tillite Group of East Greenland (left) and Polarisbreen Group of NE Svalbard (right).

87 86

structures and ﬁbrous-isopachous cements (Fig. 10). It crops out Sr-rich limestones in Bed group 20 have Sr/ Sr ratios of

behind the Expedition Houses in Solitärbugt (Fig. 9a) where it is 0.70634–0.70638 (Fairchild et al., 2000), less radiogenic than any

overlain by 80 m of thin-bedded wavy calcilutite with black chert sample from the preceding upper Russøya CIE (Bed group 19 equiv-

nodules, and a thicker sequence (Bed group 19) of greenish-grey alent) in Svalbard (0.70675; Halverson et al., 2007a).

and brick-red argillites punctuated by turbidites and debrites of

tan dolostone and grey limestone. The lower part of Bed group 19

is covered by the braid-plain of the lower Storeelv (Store River). 2.2. Ulvesø Formation

The Eleonore Bay Group ends at a sharp erosion surface, overlain

by massive diamictite of the Ulvesø Formation (Fig. 10). The lower diamictite unit of the Tillite Group, the Ulvesø For-

Northwest of Tillitekløft (Fig. 8), Bed group 18 includes mation (Hambrey and Spencer, 1987; Moncrieff and Hambrey,

repetitive 5-m-thick cycles comprised of wavy dololutite and cross- 1990) consists mainly of massive to weakly stratiﬁed, clast-poor

bedded dolarenite, capped by well-developed tepees and tepee diamictite carrying angular to subrounded clasts of dolostone,

breccias, cemented by ﬁbrous-isopachous dolospar while inﬁl- limestone, chert and quartzite in decreasing order of abundance.

trated from above by dolomicrite. Bed group 19 is overlain by a Striated clasts (Fig. 11b), azimuthal clast fabrics (Hambrey and

shoaling-upward sequence of organic-rich to intraclastic, crossbed- Spencer, 1987) and grooved pavements (Moncrieff and Hambrey,

ded limestone (Bed group 20) that is absent on Ella Ø (Hambrey 1989) connote dynamic grounded ice. However, waterlain (rainout)

and Spencer, 1987; Herrington and Fairchild, 1988; Stouge et al., diamictite predominates over basal (ice-contact) tillite, which led

2011). At Brogetdal, black argillaceous limestone succeeds thicker- Moncrieff and Hambrey (1990) to infer an extensive ice shelf carry-

bedded, shallower-water limestone of Bed group 20 (Hambrey and ing englacial debris. Various bodies of well-sorted quartz-arenite,

Spencer, 1987). The sub-Ulvesø disconformity cuts down section some crossbedded and others disturbed, occur below, within, and

from north to south. above the diamictite sequence (Fig. 10). A submarine toe-of-slope

Fairchild et al. (2000) give Ccarb values for Bed groups 19 and debris-apron environment (Eyles, 1993) is contraindicated by these

13 13

ı ı

20, while Buchardt et al. (2008) obtained Ccarb and Corg data bodies, and also by the uniformly abrupt appearance and disap-

from Bed group 19 and the upper part of Bed group 18. Ccarb is pearance of diamictite within an otherwise unremarkable basinal

enriched in Bed group 18 (mean value +7‰) but variably depleted succession (Fig. 10), and lack of discrete ﬂow units within diamictite

− 13

(0 to 10‰) in Bed group 19. The Corg values were obtained from bodies 10 s of meters in thickness.

13 13

strata with depleted Ccarb values. The difference ( C) between The type section at Kløftelv (Cliff River) near Kap Oswald (Fig. 9a)

13 13

ı ı ∼

paired Ccarb and Corg values is 30‰, consistent with typ- consists of three diamictite bodies (U1-3), the ﬁrst of which is

ical fractionation of primary biomass (Buchardt et al., 2008). Bed in knife-sharp contact with argillite of Bed group 19 (Fig. 10).

group 20 exhibits a return to enriched Ccarb values of +3 to +9‰ U1 is a friable, pale olive-coloured, clast-poor diamictite with a

(Fairchild et al., 2000). The return to positive values is consistent sandy mudstone matrix and clasts of limestone and stromatolitic

with CIEs in the Coates Lake Group beneath the Rapitan glacial hori- dolostone that coarsen upward to boulder size. U2 is a more

zon in NW Canada (Halverson, 2006) and in the Islay and Lossit consolidated, weakly stratiﬁed, clast-poor diamictite with an anas-

formations beneath the Port Askaig Formation in Scotland (Prave tomosing foliation (feuilletage of Deynoux, 1985) and pebble-size

et al., 2009). The isotopic data lend further credence to correlations clasts. Pebble-lined, ﬂat-topped channels containing crossbedded

between East Greenland and NE Svalbard (Fig. 7), assuming the CIEs quartz-dolomite-arenite and ﬁne conglomerate occur near the top

in Bed group 19 and the upper Russøya Member are correlative. of the U2 diamictite, which is overlain by a bed of quartz-arenite

The return to positive Ccarb values is not observed in East Sval- grading upward to polymictic conglomerate. The arenites have

bard (Fig. 6) and could have been removed by karstic and/or glacial westward-directed palaeocurrents. U3 is a highly-consolidated,

erosion beneath the Petrovbreen diamictite (Halverson et al., 2004). clast-poor, planar-foliated, upward-coarsening diamictite that

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 145

Fig. 8. Geological sketch map of the central Fjord Zone (after Bengaard, 1992), East Greenland, showing locations (dotted circles) of measured sections in Fig. 10.

changes colour at the top from greenish-grey to maroon. The top dune migration, interpreted as aeolian (Moncrieff and Hambrey,

of U3 is a sharp, load-casted (ball-and-pillow) surface on which 1990). The arenite is sharply overlain by a massive, clast-poor,

quartz-arenite with a thin basal pebble lag is downfolded into the sandy diamictite with small carbonate clasts that is just 1.4 m

ı13

underlying diamictite (Hambrey and Spencer, 1987). In Bastion- thick. Four of the clasts have Ccarb values between +3.6‰ (lime-

bugt (Fig. 9a), polygonal sandstone wedges occur near the top of stone) and −1.7‰ (dolostone). The diamictite is erosively overlain

the Ulvesø diamictite. The absence of dyke-in-dyke or cross-cutting by a clast-supported pebble conglomerate of the same thickness,

relations (Fig. 11d) suggests that the wedges are not clastic dykes. followed abruptly by 3.0 m of subaqueous mega-rhythmite. The

At Tillitekløft (Fig. 10), the Ulvesø Formation is thinner but mega-rhythmite consists of 0.4-m-thick couplets of stratiﬁed sandy

otherwise similar to Kap Oswald, 70 km to the south. The lower diamictite and current-rippled argillaceous siltstone. It is con-

diamictite is weakly-stratiﬁed and the clasts become smaller formably overlain by dark-grey argillite of the Arena Formation.

upward. The upper diamictite contains faceted and striated clasts

(Fig. 11b), and ﬂat-topped lenses of stratiﬁed quartz-wacke, vari- 2.3. Arena Formation

ably deformed by intrastratal shear parallel to their upper surfaces

(Fig. 11a, c). The stratiﬁed quartz-wacke lenses suggest the pro- The Arena Formation (Hambrey and Spencer, 1987) is a recessive

duction and ﬂowage of meltwater beneath stagnant grounded ice, sequence of dark-coloured argillite with distal turbidites (Bouma

followed by renewed ice ﬂowage and attendant subglacial defor- DE and CDE) of ﬁne-grained sand- and siltstone (Fig. 11e). Tur-

mation (Domack et al., 2008). We interpret the lower diamictite as bidity ﬂow was directed broadly northward, based on ripple-drift

a waterlain rainout deposit and the upper one as a subglacial tillite, orientation. Transgressive and regressive quartz-arenites occur at

associated with intermittently stagnant ice. the base and top of the argillite-turbidite sequence in the southern

At Brogetdal, 20 km north-northeast of Tillitekløft, the Ulvesø sections (Fig. 10), consistent with a northward deepening of the

Formation is signiﬁcantly different (Fig. 10). It begins with 32 m of basin (Moncrieff, 1988).

subfeldspathic quartz-arenite featuring granule ripples and large- Between the Kløftelv and Storeelv (rivers), near Kap Oswald

scale tabular crossbedding resulting from northeastward-directed (Fig. 9a), the transgressive tract consists of ﬂat-bedded and tabular

146 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 9. Key sections of the Tillite Group in East Greenland: (a) Kap Oswald on Ella Ø, type sections of the glacigenic Ulvesø and Storeelv formations (Hambrey and Spencer,

1987; Stouge et al., 2011). Arrows indicate stratigraphic way up and dashed lines are faults. Photograph taken from Bastion Mountain overhang. (b) Tillitekløft (Tillite Canyon),

near Kap Weber in Andrée Land, type section of the Canyon Formation and reference sections of the Ulvesø and Storeelv formations (Hambrey and Spencer, 1987). Asymmetric

apophyses of quartz-arenite, projecting downward into diamictite, are associated with small, normal-sense, growth faults (yellow lines, ‘d’ indicates down-thrown side).

Note onlap relationship (small black arrow) at the base of the cap dolostone (basal Canyon Formation), which thins southward and is absent at Kap Oswald. (For interpretation

of the references to colour in this ﬁgure legend, the reader is referred to the web version of the article.)

crossbedded, medium to ﬁne-grained, well-sorted, sub-feldspathic ﬂow regime. In both the above locations, regressive quartz-arenite

quartz-arenite, with southeastward-directed ﬂow inferred from at the very top of the Arena Formation was impregnated before

crossbedding. On the shoreline of Bastionbugt (Fig. 9a), the basal lithiﬁcation by boulders of syenogranite and dolostone, the upper

Arena Formation is a ﬁning-upward sequence of ﬂat-bedded, surfaces of which are faceted and grooved at a common horizon,

small-scale trough-crossbedded and current-rippled, quartz-chert which is overlain by the basal Storeelv diamictite. This surface

arenites with westward-directed (tilt-corrected) paleocurrents. is interpreted as a subglacial boulder pavement (Moncrieff and

Hambrey and Spencer (1987) illustrate “wave ripples” here (their Hambrey, 1989).

Fig. 17a) but unidirectional foreset laminae within the morpholog- At Tillitekløft (Fig. 10), the Ulvesø diamictite is over-

ically symmetrical ripples suggests a combined wave and current lain by quartz-arenite composed of well-rounded, medium- to

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 147

Fig. 10. Measured sections of the Tillite Group at Kap Oswald, Tillitekløft and Brogetdal (Fig. 8), and the upper Eleonore Bay Group at Kap Oswald (lower right). Thicker

sections of the Canyon Formation are telescoped.

148 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 11. Sedimentary features in the Tillite Group, Fjord Zone, East Greenland: (a) ﬂat-topped (arrows) quartz-arenite lenses (stained dark maroon) within the upper diamictite

of the Ulvesø Formation at Tillitekløft; (b) glacially-striated dolostone clast from the upper diamictite of the Ulvesø Formation at Tillitekløft; (c) detail of dashed box in (a),

showing recumbent folds in quartz-arenite (highlighted) related to intrastratal shear at the upper bounding surface (arrow); (d) polygonal sandstone wedge (S) in massive

diamictite (D) of the upper Ulvesø Formation, Bastionbugt, Kap Oswald; (e) argillite with distal turbidites characteristic of the Arena Formation at Brogetdal; (f) possible

boulder pavement between massive clast-poor diamictites in the lower Storeelv Formation, Storeelv, Kap Oswald. Hammer is 0.33 m long, coin 0.03 m in diameter, and staff

1.5 m long with 0.5-m divisions. (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of the article.)

coarse-grained sand with paleohorizontal granule lags and large- 2.4. Storeelv Formation

scale tabular crossbedding, indicating southeastward-directed

ﬂow. The entire argillite-turbidite sequence is well exposed on the The Storeelv Formation (Hambrey and Spencer, 1987) is thicker

south side of Brogetdal (Figs. 10 and 11e), where the turbidites and lithologically more diverse than the Ulvesø Formation (Fig. 10).

are grouped in three coarsening-upward sets. Climbing-ripple ori- It contains granitoid and metamorphic basement clasts as well

entations indicate that the turbidity currents were northwest- to as intrabasinal debris. Like the Ulvesø, the Storeelv Formation

northeast-directed. The argillite contains disc-shaped concretions includes both waterlain diamictite and basal tillite (Moncrieff and

of Fe-rich carbonate (siderite, ankerite or Fe-dolomite). At Broget- Hambrey, 1988, 1990). They are associated with subglacial melt-

dal, the argillite-turbidite sequence is disconformably overlain by water deposits, subaqueous outwash, debrite and glaciomarine

well-sorted crossbedded quartz-arenite that is lithologically indis- sediments including ice-proximal diurnal rhythmite (Moncrieff

tinguishable from that found above the basal Storeelv diamictite and Hambrey, 1988, 1990).

(Fig. 10). The Arena Formation is devoid of glacial indicators, aside In the type section (Fig. 10), the lower Storeelv Formation

from cm-scale calcite pseudomorphs at the Arena/Storeelv con- comprises a series of heterolithic, greenish to mainly maroon, non-

tact, possibly after gypsum (Moncrieff and Hambrey, 1990) or ikaite stratiﬁed diamictites, punctuated by thin, tabular and channelized

(Fig. 7). pebbly sandstone and conglomerate. At the base, a clast-supported

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 149

conglomerate with large boulders of syenogranite is pressed down

into the underlying Arena Formation sandstone. The top of this

conglomerate is a striated boulder pavement of subglacial ori-

gin (Moncrieff and Hambrey, 1989). A second boulder pavement

(Fig. 11f) occurs 40 m above the ﬁrst (Fig. 10). Striations on this

pavement are oriented parallel to the preferred long-axis orienta-

tions of elongate clasts in the overlying diamictite (mean azimuth

◦

020 ) and plucked faces indicate ice-ﬂowage toward the north-

northeast, consistent with aqueous palaeoﬂow directions inferred

from climbing ripples in a sandstone ∼1.0 m below the boulder

pavement (Moncrieff and Hambrey, 1989). The middle Storeelv

Formation is a complex of mainly maroon sandy diamictites with

tabular beds and channels of climbing-rippled sandstone and

conglomerate with north-northeastward directed palaeocurrents.

There are several planar erosion surfaces of possible subglacial ori-

gin (Fig. 10). The upper Storeelv Formation in the type section

consists of massive and stratiﬁed diamictite, clearly distinct and

thinner debris-ﬂow deposits, and a thin interval diurnal siltstone

rhythmite (Moncrieff and Hambrey, 1990).

At Tillitekløft (Fig. 9b), the Storeelv Formation comprises a set

of maroon (lower) to olive-green (upper), clast-poor diamictites,

each with a sandy mudstone matrix and clasts of stromatolitic,

oolitic and intraclastic dolostone and limestone, black chert, silt-

stone and porphyritic quartz-syenite. The thin basal diamictite is

separated from the overlying more friable diamictite by a boul-

der pavement. The diamictites are mostly massive but the middle

two contain deformed and inclined lenses of very ﬁne-grained

maroon quartz-arenite and siltstone. The upper diamictites are

separated by sheets of ﬁne- to medium-grained quartz-arenite.

Large-scale, low-angle, inclined surfaces within the thicker quartz-

arenite dip to the southeast and northeast, after tilt-correction, and

could represent point-bar accretion surfaces related to northeast-

and northwest-directed channel ﬂow. Asymmetric apophyses of

Fig. 12. Litho- and chemostratigraphy of the basal Canyon Formation (Tillite Group)

quartz-arenite projecting downward into diamictite are associated

cap dolostone at Brogetdal (Figs. 8 and 10), East Greenland. Note scale offset between

with small, normal-sense, growth faults (Fig. 9b). The youngest 13 ı18 13

ı Ccarb (black dots) and Ocarb (grey circles), decline in ı Ccarb with stratigraphic

diamictite has a current-rippled upper surface, overlain by a thin height characteristic of basal Ediacaran cap dolostones, and lack of covariation

13 18 2

between ı C and ı O (R = 0.0009).

pebble lag and <2.0 m of recessive, dark greenish-grey argillite that carb carb

is conformable with the basal Canyon Formation cap dolostone

(Figs. 9b and 10). Coarser peloids are better sorted. Towards the base and top, macro-

The section at Brogetdal is lithologically similar to Tillitekløft, peloids disappear and the dolostone becomes thin bedded with

but quartz-arenite is proportionally more prominent and a unit partings of red or maroon argillite. The base is generally cov-

of parallel-laminated siltstone with sandy turbidites occurs in the ered but appears to be sharp; the top is gradational with maroon

middle of the unit (Fig. 10). Crossbedding indicates northeasterly- argillite lacking dolostone beds. The cap dolostone is isotopically

ı and southerly-directed palaeocurrents in the arenites above and depleted and Ccarb decreases systematically up-section (Fig. 12).

below the siltstone, respectively. Conglomerate-ﬁlled wedges This Ccarb pattern is a distinguishing feature of basal Ediacaran

occur beneath the coarse-grained quartz-arenite at the top of the cap dolostones generally (Kennedy et al., 1998; Hoffman et al.,

formation. 2007; Rose and Maloof, 2010). At Tillitekløft, the cap dolostone is

Moncrieff and Hambrey (1990) interpret the Storeelv Formation thinner (8.0 m) and low-angle crossbedding was not observed. The

as having accumulated largely beneath a broad ice shelf, which was local substrate is different—argillite at Tillitekløft and quartz aren-

fed by wet-base outlet ice streams at the periphery of a continental ite at Brogetdal—but the gradational upper contact with maroon

ice sheet in a region of subdued topography. argillite is the same. At Kap Oswald (Fig. 9a), the cap dolostone is

missing entirely and its place between the topmost (20-m-thick)

2.5. Canyon Formation diamictite and maroon argillite of the lower Canyon Formation is

occupied by 0.6 m of quartz-arenite with a basal granule lag. The

The Canyon Formation (Haller, 1971) is a thick postglacial arenite contains ﬁve beds, the tops of which are current-rippled

depositional sequence dominated by greenish-grey argillite with with westward- and southward-directed paleoﬂows. Local atten-

dolomitic turbidites and concretions. A pale yellowish-weathering uation of basal Ediacaran cap dolostone close to point sources of

peloidal dolostone with characteristic basal Ediacaran features clastic input has previously been observed in the Nuccaleena For-

(Knoll et al., 2006; Hoffman et al., 2011) forms the trans- mation of South Australia (Rose and Maloof, 2010).

gressive tract (cap dolostone) of the depositional sequence The multicoloured argillaceous sequence above the cap dolo-

(Figs. 9b, 10 and 12). The highstand tract culminates with shallow- stone (Fig. 9b) is remarkably similar to the lower Dracoisen

water dolarenite (Fig. 10). Formation of NE Svalbard (Fig. 4) and the lower Sheepbed For-

◦

The transgressive cap dolostone is thickest (14.0 m) at Brogetdal mation of the Canadian Mackenzie Mountains west of 130 W

(Fig. 12), where it is strongly laminated with abundant low-angle (Halverson, 2011; Hoffman and Halverson, 2011). Maroon mud-

crossbedding. The lamination is deﬁned by oscillatory covariation stone overlying the cap dolostone passes upward into rhythmic

in the size and sorting of palimpsest macro- and micropeloids. alternations of greenish marlstone and dark maroon argillite, which

150 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

pass upward in turn into dark grey argillite interpreted as the an accompanying change in ı Corg. In contrast, a diagenetic origin

13 13

maximum ﬂooding interval. The parallel stratigraphies suggest for the Ccarb CIE predicts that no coeval Corg excursion should

paleoceanographic symmetry north and south of the paleoequa- occur in the upper Russøya Member.

tor at the dawn of the Ediacaran Period on the windward tropical Herein, we report new ı Corg data (see Supplementary Infor-

margin of Laurentia (Fig. 1). mation for methods and data) for the bulk organic carbon of

carbonate rocks in measured Sections 1, 8 and 9 (Fig. 2), which

3. Polarisbreen Group, westernmost Nordaustlandet, NE are among the sections where Ccarb data deﬁning the upper

Svalbard Russøya CIE were obtained (Halverson et al. (2004). The average

ı13

Ccarb value for the Kinnvika Member, Dartboard Dolomite Mem-

Our new observations refocus on the upper Russøya CIE (Fig. 6), ber and lower Russøya Member below the CIE is 4.6‰ (Fig. 6a). Over

situated beneath the older Cryogenian Petrovbreen subglacial dis- the same stratigraphic interval the ı Corg values have an average

− conformity, and on the absence of a Trezona-equivalent CIE beneath of 26.6‰ (Fig. 6b). In the uppermost Russøya Member, where

13 13

− ı

the terminal Cryogenian Wilsonbreen Formation. We describe the Ccarb values dip below 4‰, Corg values have an average

−

CIEs as occurring beneath, not before, the glacial deposits, so as not value of 30.7‰. If a single outlier (crossed in Fig. 6d) is excluded,

to appear judgemental over a possible diagenetic origin (Swart and the square of the correlation coefﬁcient (R ) for the combined

−12

Kennedy, 2012). Contact relations, described in Section 1.1, indi- dataset for all three sections is 0.60 with a p-value of 3.1536e (or

−11

cate that the Russøya CIE predates the subglacial erosion surface. 0.55 and 4.8941e if the outlier is included), which demonstrates

13 13

ı ı Analogous contact relations between the Trezona CIE and younger a strong correlation between Ccarb and Corg. In contrast, there

13 18

ı Cryogenian (Marinoan) subglacial erosion surfaces have been doc- is no covariance between Ccarb and Ocarb (Fig. 6c, f).

umented in Namibia (Halverson et al., 2002) and South Australia The magnitude of the change in ı Corg across the upper Russøya

(Rose et al., 2012). The age relations do not, in fact, rule out a syn- excursion is slightly less than half that of the Ccarb change

glacial diagenetic origin—the ultimate subglacial erosion surface (Fig. 6d). In addition to the reduced magnitude of the excursion,

could form late in the glacial period—but they do eliminate an origin there is more scatter in the Corg records (particularly when

through deep burial diagenesis (Derry, 2010). We test for synglacial comparing data from MS8 with that from MS1 and MS9). In con-

meteoric diagenesis (Swart and Kennedy, 2012) with isotopes in trast to carbonate carbon isotope records, whose values should

Section 3.2. reﬂect a nearly constant small fractionation from primary DIC, the

photosynthetic fractionation (εp) between primary biomass and

3.1. Upper Russøya stromatolite proliferation DIC is signiﬁcantly inﬂuenced by the biological speciation of pro-

duction, sea-surface temperature and aqueous CO2 concentration

On Nordaustlandet, stromatolites (including columnar form- (Freeman and Hayes, 1992), phosphate concentration at the depth

genera Baicalia and Kussiella) are exceptionally proliﬁc in strata of organic production (Pagani et al., 2002), and cellular growth

hosting the upper Russøya CIE. In measured Sections 7–9 (Fig. 2), rate (Bidigare et al., 1997), geometry (Popp et al., 1998) and size

stromatolites make up 52% (44 m) of strata in all three sections (Henderiks and Pagani, 2008). Accordingly, even with minimal

ı13

having C < 0‰ (Halverson et al., 2004). In the lower and middle post-depositional alteration, it is expected that organic carbon iso-

Russøya Formation, where C > 0‰, microbialaminite makes up tope records should display more scatter than coeval carbonate

only 9% of the total thickness and columnar stromatolite is absent records (e.g., Maloof et al., 2010). The reduced magnitude of the

13 13

altogether (Halverson et al., 2004). Stromatolite proliferation is less Corg excursion relative to the Ccarb excursion, resulting in

evident to the south on Spitsbergen, where the CIE is truncated in a decrease in the difference between the organic and inorganic

measured Sections 3–5 (Halverson et al., 2004) and has columnar isotopes (ı C), could be explained in multiple ways. Given the

stromatolite only near the base in measured Section 1 (Fig. 6). strong dependence of εp on CO2(aq) concentrations (Freeman and

Stromatolite bioherms occur in the Islay Limestone beneath the Hayes, 1992), changes in ı C are often leveraged as a rough pCO2

glacigenic Port Askaig Formation on Garbh Eileach (Spencer, 1971). proxy (e.g., Kump and Arthur, 1999; Sansjofre et al., 2011). Accord-

It was in this formation that the older Cryogenian Islay CIE was ﬁrst ingly, the decrease in ı C during the upper Russøya CIE could

recognized and subsequently veriﬁed (Brasier and Shields, 2000; reﬂect a decrease in pCO2, consistent with impending glaciation.

McCay et al., 2006; Prave et al., 2009; Sawaki et al., 2010). The However, this explanation is non-unique and the decrease in C

87 86

Sr/ Sr ratio of 0.7064 in the Islay Limestone supports an older could also reﬂect increased growth rate, phosphate concentration

Cryogenian age (Brasier and Shields, 2000; Sawaki et al., 2010). or sea-surface temperature, smaller cell size, or a change in micro-

On the northern paleomargin of Laurentia, stromatolite bioherms bial community composition to one more dominated by organisms

occur in strata hosting the Trezona CIE (upper Wynniatt and Kil- such as cyanobacteria that obtained their carbon from the bicar-

lian formations) on Victoria Island in Arctic Canada (Jones et al., bonate pool. Another explanation for the reduced magnitude of

2010), but not in carbonates hosting the Islay CIE in the Macken- the ı Corg excursion is that it reﬂects an exogenous input of non-

zie and Ogilvie mountains (Coates Lake and Lower Mount Harper contemporaneous organic carbon (Jiang et al., 2010, 2012; Johnston

groups) of NW Canada (Macdonald et al., 2010). Deeper water facies et al., 2012). This carbon could be derived either from recycling of

and active (extensional) tectonics may have been unfavourable for old sedimentary organic carbon or from an oceanic pool; in either

stromatolites in those areas. case, it would damp the record of primary Corg. We are unable

to discriminate between these explanations with existing data but,

3.2. Paired organic and carbonate carbon and oxygen isotopes in regardless of the reason for the smaller amplitude, the negative

13 13

ı the upper Russøya CIE Corg excursion is both signiﬁcant and coeval with the Ccarb

excursion.

If the carbon that is associated with organic matter in a car-

bonate rock is derived from the same dissolved inorganic carbon

(DIC) pool as the carbonate carbon, then large-scale changes to the 3.3. Where is the Trezona CIE?: Is the Bråvika member synglacial

isotopic composition of DIC should be reﬂected both in carbonate or preglacial?

ı13 13

ı ( Ccarb) and organic carbon ( Corg) isotope records. The inter-

ı13

pretation that the Ccarb CIE in the upper Russøya Member is due If the upper Russøya CIE underlies the older Cryogenian glacial

to a perturbation in the isotopic composition of DIC thus predicts horizon, where is the Trezona CIE that should underlie the

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 151

Wilsonbreen Formation (Fig. 4)? Carbonates of the Slangen Wilsonbreen Formation), which were not strongly lithiﬁed in Cryo-

Member are enriched in Ccarb by 2–5‰ relative to VPDB on Spits- genian time. This is supported by the ubiquity of silica sand grains

bergen and by 5–7‰ to the north on Nordaustlandet (Halverson ‘ﬂoating’ in the mud-rich matrix of diamictites, and by the dis-

et al., 2004). The Bråvika member (Halverson et al., 2004) is proportionately small fraction of arenite clasts in the diamictites

a coarsening-upward southward-tapered wedge of crossbedded (Hambrey and Spencer, 1987; Harland et al., 1993), relative to

quartz-chert-arenite and sandy conglomerate that stratigraph- their overwhelming abundance in the source material accessible to

ically overlies the Slangen Member and underlies the basal glacial erosion given the presence of igneous-metamorphic (base-

diamictite of the Wilsonbreen Formation. Locally the Bråvika mem- ment) debris (Nystuen et al., 2008).

ber contains lenses of stromatolitic or ribbon-bedded dololutite.

Halverson et al. (2004) interpreted the Bråvika member as a ﬂu-

3.4. Molar-tooth structure

vial proglacial (sandur) facies of the Wilsonbreen Formation. Four

observations prompted this interpretation. First was the recipro-

The term ‘molar-tooth’ refers to ptygmatically-folded veins of

cal thickness relationship between the Bråvika member and the

uniform calcite microspar of early diagenetic origin found in ﬁne-

Wilsonbreen diamictite, the arenite thickening northward at the

grained sublittoral carbonates of Proterozoic age (Frank and Lyons,

expense of the diamictite. Second was the northward-directed

1998; James et al., 1998). Most examples of well-developed molar-

paleoﬂow based on large-scale crossbedding in the Bråvika mem-

tooth predate the Sturtian glaciation (Shields, 2002): James et al.

ber. Third was the appearance of carbonate-clast conglomerate in

(1998) illustrate only one younger example, from the Wonoka For-

the northernmost section of the Bråvika member, on the coast of

mation (Ediacaran) of South Australia. In NE Svalbard, molar-tooth

Franklinsundet, which we interpreted to be reworked from Wilson-

is well-developed in the upper Russøya Member below the Petro-

breen diamict. Fourth was the northward (down-slope) increase

vbreen Member (Halverson et al., 2004), just as it is in Bed group

in grain size in the Bråvika member, implying progradation that

20, beneath the Ulvesø Formation in East Greenland. This distri-

we attributed to a northward advancing ice margin. Accordingly,

bution of molar-tooth in Svalbard and East Greenland is consistent

the Bråvika member was assigned to the Wilsonbreen Formation

with a Sturtian age for the Petrovbreen-Ulvesø glaciation. However,

(Halverson et al., 2004; Halverson, 2011). A disconformity was pos-

molar-tooth also occurs in the upper Slangen Member at Sveanor

tulated between the Slangen and Bråvika members (Fig. 4), and this

(Fig. 5b), deposited well after the Ulvesø-Petrovbreen glaciation.

hiatus could account for the missing Trezona CIE.

Although replaced by fabric-retentive dolomite, neither its iden-

In 2007, stable isotope data were obtained from the thickest

tity as molar-tooth nor its early diagenetic origin is in doubt. Thus,

section of the Slangen Member revealed by mapping (Fig. 13b),

the Sveanor occurrence represents a second post-Sturtian example

and from a 2.6-m-thick intercalation of dololutite at a relatively

of molar-tooth structure and the ﬁrst of Cryogenian age.

high stratigraphic level within the Bråvika member in the same

area (Fig. 13a). The Ccarb values in both units are 5–7.5‰. The

Trezona CIE is still missing. The contact between the Slangen and 4. Discussion

Bråvika members at Sveanor (Fig. 3), which was covered by a snow-

bank in previous expeditions, was perfectly exposed. Although the 4.1. Primary or diagenetic origin of the upper Russøya CIE?

contact is abrupt on the scale of meters, dololutite ribbons of the

Slangen Member are interlayered with basal Bråvika quartz-chert- The standard interpretation of Neoproterozoic CIEs in terms of

arenite on the scale of centimeters. The contact itself is conformable change in the isotopic composition of DIC in contemporaneous sea-

(Fig. 5b) with no evidence of signiﬁcant hiatus. Finally, we revis- water (e.g., Halverson and Shields-Zhou, 2011) has been challenged

13 18

ı ı

ited the northernmost section of the Bråvika member, previously on the grounds that Ccarb covaries with Ocarb in some cases

inferred to be the most proximal facies based on the presence of (Knauth and Kennedy, 2009; Derry, 2010), and that certain CIEs

carbonate-clast conglomerate (Halverson et al., 2004). More careful (e.g., Trezona and Shuram) lack coincident excursions in the ı C

examination shows the Franklinsundet coastal section to be a more of contained organic matter (Fike et al., 2006; McFadden et al.,

ı distal, coastal-plain and marginal-marine facies complex, in which 2008; Swanson-Hysell et al., 2010; Johnston et al., 2012). Ccarb

the carbonate clasts are limestones of strictly intraformational ori- and Ocarb covariation is predicted if CIEs were the product of

gin. Like the dololutite intercalations to the south (Fig. 13a), the diagenesis under the inﬂuence of meteoric ﬂuids, charged with

13 13

limestone intraclasts are enriched in C. C-depleted DIC stemming from the respiration of organic mat-

The conformable nature of the Slangen-Bråvika contact at ter (Knauth and Kennedy, 2009), or alternatively as the product

Sveanor (Fig. 5b), the similarity of their Ccarb values (Fig. 13), of deep burial diagenesis at elevated temperatures (Derry, 2010).

and the invalidation of evidence that carbonate clasts in the Bråvika Glacioeustatic changes drive meteoric groundwater ﬂows capa-

member have a glacial origin, all point in the same direction: that ble of producing globally synchronous CIEs diagenetically, given

the Bråvika member is more closely related to the Slangen Mem- sufﬁcient organic matter and terminal electron acceptors (Swart

ber than to the Wilsonbreen glaciation. We therefore reassign the and Kennedy, 2012). Additionally, the lack of covariance between

Bråvika member to the Elbobreen Formation (Figs. 3, 4, 7, 13) and organic and carbonate ı C could reﬂect the presence of an enor-

tentatively suggest that the hiatus responsible for missing Trezona mous pool of dissolved organic carbon (Rothman et al., 2003;

CIE lies between the Bråvika member and the Ormen Member of Fike et al., 2006; Swanson-Hysell et al., 2010), attenuation of the

the Wilsonbreen Formation (Fig. 3). photosynthetic fractionation (εp) accompanying export production

It might be assumed that a glacioﬂuvial sand sheet could be dis- (Sansjofre et al., 2011), or the presence of detrital organic matter

tinguished from a nonglacial tropical river deposit with a hand lens. unrelated to syndepositional DIC (Jiang et al., 2010, 2012; Johnston

Not so for the Cryogenian glacigenic formations of East Greenland et al., 2012).

and NE Svalbard, in which massive diamictite including subglacial The Islay CIE (Table 1) has been a non-participant in the above

tillite is intercalated with bodies of chemically- and texturally- discussions for want of data. With the present results (Fig. 6), dis-

ı18 mature, quartz- and quartz-chert arenite (Hambrey and Spencer, cussion can begin. The ﬁrst observation is that Ocarb does not

ı13

1987; Harland et al., 1993; Halverson et al., 2004; Fig. 10). A covary with Ccarb in the upper Russøya Formation (Fig. 6c, f),

simple explanation for this phenomenon is that the arenite was contrary to the predictions of a meteoric or burial diagenetic origin

recycled from older non-glacial sandstones (Eleonore Bay and Lom- for the CIE (Knauth and Kennedy, 2009; Derry, 2010). What if mete-

fjorden supergroups, and the Bråvika member in the case of the oric diagenesis occurred on a snowball Earth, where groundwater

152 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Fig. 13. Litho- and chemostratigraphy of the Bråvika (a) and Slangen (b) members of the Elbobreen Formation (Polarisbreen Group) on Gotiahalvøya, Nordaustlandet, Svalbard

13 18 13 18 2

(Fig. 3). Note scale offset between ı Ccarb (black dots) and ı Ocarb (grey circles), covariation between ı Ccarb and ı Ocarb in the Slangen Member (R = 0.646), and continuity

in ı Ccarb between the members.

ﬂows were recharged by basal ice-sheet melting? On a snowball would have accumulated in the central and higher areas of the ice

Earth, the water vapour responsible for the growth and mainte- sheets on a snowball Earth. It is this isotopically depleted ice that

nance of dynamic ice sheets originates from the sublimation of would have been advected to the area of basal meltwater produc-

equatorial sea ice (Donnadieu et al., 2003; Pollard and Kasting, tion (Pattyn, 2010). The magnitude of the isotopic depletion in the

2004; Pierrehumbert et al., 2011). Sublimation, unlike evapora- basal ice of the Cryogenian ice sheets on paleo-equatorial Laurentia

tion, imparts no isotopic fractionation because of the low rate of has not been investigated in a model. But only if the magnitude of

ı18

diffusion of oxygen in ice. The ice–water equilibrium fractionation the Oice anomaly in the melting region was near-zero would the

ı18 ı13 18

∼ ı ( Oice-water) is 3‰ (O’Neil, 1968), so the steady-state composi- absence of Ccarb– Ocarb covariation (Fig. 6a–c, f) be consistent

ı18 tion of sea ice ( Oice) on a snowball Earth would have been at least with a meteoric diagenetic origin for the upper Russøya CIE.

13 13

ı ı

1.5‰ heavier than mean (non-glacial) ocean water, assuming that The observed covariation between Corg and Ccarb (Fig. 6a,

roughly half of the water was sequestered in the form of ice sheets b, d) is also inconsistent with a diagenetic origin for the upper

and the sea glacier. In contrast, the North American and East Antarc- Russøya CIE. As discussed in Section 3.2, there are many possible

tic ice sheets at the Last Glacial Maximum were ∼30‰ and ∼60‰ explanations for the smaller magnitude of the ı Corg excursion, but

lighter, respectively, than standard mean ocean water (Sima et al., none of them allow for a plausible diagenetic explanation for the

2006; Lorius et al., 1985). This resulted from Rayleigh distillation observed isotopic covariation. Although muted in comparison with

13 13

ı ı in the atmosphere accompanying progressive precipitation. If all of the Ccarb excursion, the Corg excursion is signiﬁcant (Fig. 6d)

the precipitation fell as snow, Rayleigh distillation would have no and coincident with the Ccarb excursion. The simplest explana-

direct effect on the bulk ice composition, but isotopically lighter ice tion is that a substantial fraction, at least, of the organic C was drawn

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 153

Table 1

ı13

Cryogenian subglacial negative C excursions (CIEs).

Paleocontinent Stratigraphic host Glaciation Reference

Trezona CIE (Marinoan/late Cryogenian)

Australia Trezona Fm Elatina McKirdy et al. (2001)

Swanson-Hysell et al. (2010)

Rose et al. (2012)

Congo Ombaatjie Fm Ghaub Halverson et al. (2002)

Hoffman (2011)

Laurentia-N Keele Fm Ice Brook (Stelfox) Halverson et al. (2005)

Johnston et al. (2012)

Baltica Grasdal Fm Smalfjord Halverson et al. (2005)

Laurentia-S Craignish/Ardrishaig Stralinchy/Reelan Prave et al. (2009)

Islay CIE (Sturtian/early Cryogenian)

Laurentia-N Beck Spring Fm Surprise Horodyski and Knauth (1994)

Corsetti and Kaufman (2003)

Tasmania Black River Fm Julius River Calver (1998)

Laurentia-S Bed group 19 Ulvesø Fairchild et al. (2000)

Buchardt et al. (2008)

Laurentia-S Islay (Lossit) Fm Port Askaig Brasier and Shields (2000)

McCay et al. (2006)

Prave et al. (2009)

Sawaki et al. (2010)

Laurentia-S Russøya Mb Petrovbreen Halverson et al. (2004)

This paper

Laurentia-N Copper Cap Fm Rapitan Halverson (2006)

Laurentia-N Lower Mt Harper Gp Rapitan Macdonald et al. (2010)

Laurentia-N Killian Unnamed Jones et al. (2010)

Formations with CIE-associated stromatolite proliferation in upper case.

Laurentia north (-N) and south (-S) of the Cryogenian paleoequator.

from the same DIC pool as the carbonate C. Taken together, the prevents microbial ‘bridging’ between structures (Hoffman, 1976).

13 13

ı ı observed covariation between Corg and Ccarb, and the lack In those areas, microbial mats would be unable get a foothold with-

of covariance in Ocarb, make a diagenetic interpretation of the out early lithiﬁcation.

upper Russøya CIE implausible. If truly correlative, other examples In the Cryogenian, evaporation need not have been the driver

of the Islay CIE (Table 1) are likely to be primary in origin as well. of carbonate oversaturation and no protection was needed against

grazing by animals (Butterﬁeld, 2011). At the beginning of each

Cryogenian ice age, the saturation state of surface waters should

4.2. Signiﬁcance of CIE-associated stromatolite proliferation

logically have risen in response to the drawdown of sea level

(Berger, 1982; Opdyke and Walker, 1992; Ridgwell et al., 2003)

Of the 13 Cryogenian subglacial CIEs known to us, six are asso-

and atmospheric CO (Zeebe and Wolf-Gladrow, 2001), and also

ciated with stromatolite proliferations (Table 1). By proliferation 2

because increased aeration of deep waters due to temperature

we mean increased stromatolite abundance, size and diversity in

change would raise the saturation state of surface waters at the

form, relative to underlying strata of comparable water depth. In

expense of deep waters (Higgins et al., 2009). The ﬁrst explana-

the upper Ombaatjie Formation, for example, columnar stromato-

tion, a hypsometric effect, cannot apply to the Islay and Trezona

lites of the form-genera Tungussia and Conophyton (Raaben, 1969)

CIEs because the stromatolites and CIEs formed on the tops of

are prominent; in the upper Russøya Formation, columnar form-

shallow-water platforms before sea-level fall (e.g., Russøya and

genera Baicalia and Kussiella dominate. In other areas, stromatolitic

Ombaatjie formations). Those low-latitude platforms would have

bioherms are developed (Black River, Islay and upper Wynniatt-

become subaerially exposed once sea-level began to fall. On the

Killian formations). In most of the cases where stromatolites do

other hand, atmospheric pCO could have fallen signiﬁcantly before

not proliferate (e.g., Keele, Grasdal, Bed group 19 and Copper Cap 2

any continental ice-sheets had developed (Voigt et al., 2011), given

formations), the subglacial CIEs are associated with deeper water

a paleogeography in which polar continents were absent (Evans,

facies unsuitable for stromatolite development.

2000; Trindade and Macouin, 2007; Hoffman and Li, 2009; Evans

The environmental range of microbial mats responsible for the

and Raub, 2011). The paleogeography allowed the decline in pCO

construction of modern stromatolites in coastal marine settings 2

and temperature, inextricably linked through climate sensitivity,

is deﬁned by animal grazing, principally by snails, and suitable

and consequently the increase in the saturation state of surface

substrates (Garrett, 1970; Logan et al., 1974; Kendall et al., 2002;

waters, to be expressed in the form of stromatolite prolifera-

Maloof and Grotzinger, 2012). The exceptional sublittoral and

tion on platforms before any signiﬁcant glacioeustatic drawdown.

lower intertidal stromatolites of Shark Bay, Western Australia,

∼ Moreover, the ﬂux of alkalinity from continental weathering was

owe their existence to net evaporation (salinity 70 ppm), which

relatively unimpaired by the advancing sea ice in the absence

excludes mat-grazing cerithid gastropods and increases carbon-

of polar continents. The alkalinity ﬂux governed the quantity of

ate saturation. The latter speeds early lithiﬁcation (e.g., beachrock)

carbonate produced; sea-surface temperature governed only the

and makes lithiﬁed crusts available for microbial colonization.

location of (non-skeletal) carbonate production.

Once established, the microbial mats provide their own substrate

What remains to be explained is why the inferred Cryogenian

through binding–cementation feedback. The columnar and high-

pCO and temperature declines were associated with CIEs. The

relief stromatolites occur in areas of persistent wave action, which 2

154 P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158

Myr timescale of Cryogenian CIEs justiﬁes their interpretation in

terms of changes in the composition of global carbon inputs ( Cin)

and fractional organic burial rates (forg) (Kump and Arthur, 1999).

Kaufman et al. (1997) postulated that Cryogenian subglacial CIEs

resulted from enhanced upwelling of anoxic, alkalinity-charged,

13 13

C-depleted deepwaters. The postulate requires a ı CDIC gradi-

ent much larger than that in the present ocean (<2‰), given the

magnitude of the CIEs. It also requires a means of suppressing

export production, expected to rise in response to enhanced nutri-

ent upwelling, driving surface-water ı CDIC heavier. Hoffman et al.

(1998) attributed the Trezona CIE to reductions in productivity and

forg in response to increased sea-ice cover. As the tropical ocean was

still ice-free when the host carbonates were deposited, the mag-

nitude of the CIEs, equivalent to steady-state forg = 0 at their nadir

ı ≈ −

assuming Cin 6‰ or less (Halverson and Shields-Zhou, 2011),

is far in excess of what declining productivity alone would predict.

Schrag et al. (2002) related the same CIE to methane release from

submarine gas hydrate, which destabilized the climate through

steady-state greenhouse gas substitution. This concept rested on

≥ high rates of organic burial (forg 0.4), in tropical deltas, inferred

ı ≥

from the time-averaged Ccarb 5‰ for most of the 180 Myr

preceding the Islay CIE (Halverson and Shields-Zhou, 2011). The

challenge is to explain how such high rates of forg could be main-

tained, let alone how submarine methane releases could invade the

atmosphere, given the oxidizing potential generated by the organic

burial itself. Remineralization of organic matter is another source

of C-depleted DIC, producing CO2 in the presence of O2, or HCO3

anion where respiration is achieved anaerobically through micro-

bial sulfate and iron reduction (MSR and MIR). Tziperman et al.

(2011) modelled a 1.0-Myr, large-amplitude (ı C = −15‰) CIE,

accompanied by a transient fall in atmospheric pCO2 from 260 to

160 ppm, forced by transient change to the ratio of anaerobic to

aerobic remineralization. Their model requires a large DOC pool

and ample supplies of Fe(III) and sulfate. The modest attendant

rise in seawater pH accords with the observed stromatolite pro-

liferations (Tziperman et al., 2011). They acknowledge, however,

Fig. 14. Regional average stratigraphic thickness of the (a) younger and (b) older

the same essential challenge as with the Schrag et al. (2002) con- Cryogenian glacial sequences in Svalbard (Harland et al., 1993; Halverson et al.,

cept: oxygenic ﬁxation of organic matter, followed by anaerobic 2004), the Fjord Zone of East Greenland (Hambrey and Spencer, 1987; this study),

and the British Isles (Spencer, 1971; McCay et al., 2006). n = the number of measured

remineralization, amounts to an oxygen source (Tziperman et al.,

sections in each region (n = 80). Paleolatitudes for the older Cryogenian deposits

2011). This source is large in the model scenario and must be regu-

are based on the recalculated grand mean paleopole from the contemporaneous

lated by appropriate sinks in order to maintain anoxic deepwaters,

Franklin LIP (Denyszyn et al., 2009; Macdonald et al., 2010; Evans and Raub, 2011).

which are essential to the model. For the Islay CIE, volcanism and There are no reliable paleopoles from Laurentia contemporaneous with the younger

weathering of the contemporaneous Franklin LIP affords a possi- Cryogenian glaciation. The marked decrease in regional average thickness of the

older glacial deposits from the subtropics (British Isles) to the deep tropics (Svalbard)

ble sink for O2 and source of Fe and S (Macdonald et al., 2010).

parallels the decline in net meteoric accumulation minus sublimation on a hard

Erupted across the paleoequator on the windward margin of Lau-

snowball Earth.

rentia (Evans and Raub, 2011), its location was ideal for chemical

weathering, although weathering rates would have been less than

for a younger Phanerozoic analogue because of the absence of land the situation would have been reversed: net sublimation would

plants (Tziperman et al., 2011). A role for the Franklin LIP as a critical have been greatest at the equator and the subtropics would have

sink for O2 and CO2 at the onset of the older Cryogenian (Rapi- experienced the largest net accumulation (Pierrehumbert et al.,

tan) glaciation might appear more convincing if a comparable sink 2011). Given that the British Isles, Fjord Zone and Svalbard provide

existed at the time of the Trezona CIE. However, a model of silicate a paleomagnetically-constrained transect of the southern tropics

weathering and drawdown of atmospheric CO2 in the aftermath on the Laurentian margin during the older Cryogenian glaciation

of an older Cryogenian snowball Earth implies that no indepen- (Fig. 1), can the nature of the respective glacial deposits be used

dent trigger was needed for the younger glaciation, merely a return to discriminate between the two contrasting climate states, Jor-

to the conditions responsible for the previous refrigeration (Mills mungand and snowball Earth? Are there systematic differences

et al., 2011). On the other hand, if the Islay CIE is a manifestation of between the older and younger Cryogenian glacial sequences in the

the proximal cause of glaciation, the Trezona CIE implies that cause three areas that could be used to infer polar wander (plate motion)

was repeated. or alternate climate states?

We computed the regional average thickness of the total glacial

4.3. Snowball Earth or Jormungand? sequence for both glaciations in each area from 80 measured

sections in the literature and the present study (Fig. 14). The

If a lane of open water existed along the paleoequator—the numbers are evenly divided between the older and younger

Jormungand climate state—net sublimation would have been great- Cryogenian glaciations. We assume no correlation within a glacial

est in the subtropics and net precipitation would have peaked period: each section could represent a small and different fraction

at the equator (Abbot et al., 2011). On a hard snowball Earth, of the total glacial period. We compare only net regional sediment

P.F. Hoffman et al. / Precambrian Research 206–207 (2012) 137–158 155

accumulation over each glacial period. For the older Cryogenian Large Igneous Province (∼718 Ma) in Arctic North America, syn-

glaciation, paleolatitude is tightly constrained by paleomagnetic chronous with the onset of the early Cryogenian Rapitan glaciation

data from the contemporaneous (718 Ma) Franklin LIP (Denyszyn in NW Canada, restore NE Svalbard and East Greenland to the inner

et al., 2009; Macdonald et al., 2010; Evans and Raub, 2011). The tropics on the southern windward paleomargin of Laurentia.

regional average glacial sequence thickness increases by a factor of Deep negative Ccarb excursions (CIEs) underlie the older

30 from the deep tropics to the subtropics (Fig. 14b). This could be glacial horizons in both areas, returning to positive values in East

related to tectonics, ice-sheet drainage, preservational accident, Greenland. Sr isotopes in the same strata support their correlation

ı or climatic zonation. On a snowball Earth, the change from net with the global Islay CIE. The absence of a coeval Ocarb excursion

meteoric accumulation in the subtropics to net sublimation (at sea combined with the presence of a coeval ı Corg excursion verify a

level) in the deep tropics corresponds to an equatorward rise of syndepositional seawater DIC origin for the Islay excursion in NE

the equilibrium-line altitudes (ELA) on the ice sheets. The ﬂux of Svalbard. The ı Corg excursion is less than half the magnitude of

ice through the subtropical ice sheets is greater than for those in the Ccarb excursion, but more data are needed to shorten the list

the deep tropics, and the lower ELAs in the subtropics route more of potential explanations for this behaviour. The CIE is accompanied

of the ice-borne sediment to the marine environment, improving by a proliferation of stromatolites, possibly in response to rising

its preservation potential. In Jormungand, translating meteoric carbonate saturation (falling pCO2) and consequent early cemen-

balance into ice-sheet dynamics is more complicated because tation.

some of the precipitation falls as rain. Nevertheless, ice sheets in No CIE occurs beneath the younger glacial horizon in either area.

the deep wet tropics should have substantially lower ELAs than In NE Svalbard, crossbedded quartz-chert-arenite of the Bråvika

those in the dry subtropics, just as they do in the modern and LGM member, previously interpreted as a proglacial facies of the Wilson-

climates (Broecker and Denton, 1990). We expect maximum sed- breen diamictite, has Ccarb values in carbonate intercalations

iment delivery by ice sheets in a Jormungand climate to be in the (+6‰) indistinguishable from conformably underlying marine car-

deep tropics, with minima in sediment delivery in the subtropics. bonates of the Slangen Member (Elbobreen Formation). The hiatus

This is exactly the opposite of what is observed (Fig. 14b). Glacial responsible for the missing global Trezona CIE is now considered to

sequence thickness variation for the older Cryogenian glaciation, be a disconformity between the Bråvika member and the Wilson-

given the strong paleomagnetic constraints, is consistent with a breen diamictite.

hard snowball Earth and inconsistent with a Jormungand state. The regional average thickness of older Cryogenian glacial

Because the climatic interpretation of regional thickness variation sequences (n = 41) declines dramatically from the British Isles

is non-unique, the analysis needs to be extended globally, so far as through the Fjord Zone of East Greenland to the Svalbard

paleomagnetic constraints exist. Archipelago. Reliable paleomagnetic constraints from the contem-

Average thickness variation for the younger glaciation is less poraneous Franklin LIP (∼718 Ma) place the British Isles in the

dramatic (Fig. 14a). There are no direct paleomagnetic constraints subtropics and Svalbard in the deep tropics, along the paleo-

for Laurentia at 635 Ma. Preliminary paleomagnetic results from austral margin of Laurentia. Equatorward attenuation in regional

the Long Range (615 Ma) and Grenville (590 Ma) dyke swarms sug- average thickness is consistent with a hard snowball Earth, on

gest net southward motion of Laurentia during Cryogenian and which meteoric net accumulation peaks in the subtropics and is

early Ediacaran time (McCausland et al., 2009; Evans and Raub, strongly negative (net sublimation) in the deep tropics. It is incon-

2011). Assuming a conventional Laurentia-Amazonia restoration sistent with the Jormungand climate state, which is dryest in the

within Rodinia, a moderately reliable paleopole from Amazonia at subtropics and most wet at the equator. Regional average thick-

635 Ma (Trindade et al., 2003) implies that much of that south- ness variation is less useful for the younger Cryogenian glaciation

ward motion occurred in the Cryogenian. Assuming such motion, because there are no paleomagnetic constraints for Laurentia of

the increased thickness of younger versus older Cryogenian glacial that age (655–635 Ma).

sequences in the Fjord Zone and Svalbard could reﬂect the apparent

northward migration of the subtropical zone of maximum meteoric Acknowledgements

net accumulation from one snowball Earth to the next. The appar-

ently thin (“meters to tens of meters”) Stralinchy-Reelan sequence Fieldwork for this study was supported by U.S. National Science

(McCay et al., 2006) could represent the dry mid-latitudes of a Foundation collaborative grants OPP-0352694 to PFH and OPP-

snowball Earth (Pierrehumbert et al., 2011), far from the deep trop- 035273 to EWD. A Geological Society of America research grant to

ical sources of water vapour. The same zonal migration would shift NLSH and a Sloan Research Fellowship to ACM supported organic

the dry subtropics in the Jormungand state to the Fjord Zone and carbon stable isotope measurements. We thank VECO Resources,

Svalbard, inconsistent with the relatively thick average sections in for logistical help in East Greenland, and Heinrich Eggenfeller of

those areas (Fig. 14a). This evidence in favour of snowball Earth the yacht S/Y ARCTICA of Longyearbyen, for travel in NE Svalbard.

over Jormungand for the younger glaciation carries little weight, PFH acknowledges support from the Earth System Evolution Pro-

however, because of the lack of contemporaneous paleomagnetic gram of CIFAR (Canadian Institute for Advanced Research). Linda

constraints from Laurentia. Godfrey is thanked for technical support of stable isotope measure-

ments at Rutgers University. Francis Macdonald commented on an

5. Conclusions early draft and constructive suggestions of Daniel Le Heron and an

anonymous reviewer improved the manuscript.

Two discrete glaciations are represented in the Polarisbreen

Appendix A. Supplementary data

Group of NE Svalbard and the Tillite Group of the Fjord Zone in

East Greenland. The younger Wilsonbreen and Storeelv diamictites

Supplementary data associated with this article can be found, in

are terminal Cryogenian (Marinoan) in age, based on their litho-

87 86 the online version, at doi:10.1016/j.precamres.2012.02.018.

logically and isotopically diagnostic cap dolostones. Sr/ Sr ratios

in Sr-rich limestones underlying the older Petrovbreen and Ulvesø

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