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Continental carbonate facies of a Neoproterozoic panglaciation, north-east Svalbard

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The Marinoan panglaciation (ca 650 to 635 Ma) is represented in north-east Svalbard by the 130 to 175 m thick Wilsonbreen Formation which contains syn-glacial carbonates in its upper 100 m. These sediments are now known to have been deposited under a CO2-rich atmosphere, late in the glaciation, and global climate models facilitate testing of proposed analogues. Precipitated carbonates occur in four of the seven facies associations identified: Fluvial Channel (including stromatolitic and intraclastic limestones in ephemeral stream deposits); Dolomitic Floodplain (dolomite-cemented sand and siltstones, and microbial dolomites); Calcareous Lake Margin (intraclastic dolomite and wave-rippled or aeolian siliciclastic facies); and Calcareous Lake (slump-folded and locally re-sedimented rhythmic/stromatolitic limestones and dolomites associated with ice-rafted sediment). There is no strong cyclicity, and modern analogues suggest that sudden changes in lake level may exert a strong control on facies geometry. Both calcite and dolomite in stromatolites and rhythmites display either primary or early diagenetic replacive growth. Oxygen isotope values (-12 to +15‰VPDB) broadly covary with δ13C. High δ13C values of +3·5 to +4·5‰ correspond to equilibration with an atmosphere dominated by volcanically degassed CO2 with δ13C of -6 to -7‰. Limestones have consistently negative δ18O values, while rhythmic and playa dolomites preserve intermediate compositions, and dolocretes possess slightly negative to strongly positive δ18O signatures, reflecting significant evaporation under hyperarid conditions. Inferred meltwater compositions (-8 to -15·5‰) could reflect smaller Rayleigh fractionation related to more limited cooling than in modern polar regions. A common pseudomorph morphology is interpreted as a replacement of ikaite (CaCO3·H2O), which may also have been the precursor for widespread replacive calcite mosaics. Local dolomitization of lacustrine facies is interpreted to reflect microenvironments with fluctuating redox conditions. Although differing in (palaeo)latitude and carbonate abundance, the Wilsonbreen carbonates provide strong parallels with the McMurdo Dry Valleys of Antarctica.
Crystal fabrics in FA5. (A) Transmitted light. Millimetre-scale microbial calcite laminites separated by thinner ice-rafted laminae and locally containing till pellets (P). Calcite laminites display peloidal clots and local fenestrae and have variably bulbous tops. W2, Reinsryggen, 79Á8 m. (B) Transmitted light. Similar horizon to (C) displaying clastic lamina overlain by clotted and fenestral microbial lamina with bulbous top. W2, East Andromedafjellet, 13Á5 m. (C) Stained thin section, transmitted light. Dolomitic microbial laminite containing dolomicrite and dolomicrospar laminae and floating detritus (white). Conspicuous fenestrae are filled by pinkstained calcite. W2, South Ormen, 31Á4 m. (D) Transmitted light. W2, Reinsryggen, 79Á8 m, microbial fabrics. Faintly clotted (peloidal) micrite (examples arrowed) and microspar laminae with intervening calcite-filled fenestra . Dark patches are micro-till pellets (some are labelled 'P'), now partly silicified. (E) Transmitted light, stained thin section. Enlargement of the ikaite pseudomorphs of Fig. 18B showing relic crystal outlines in dark micrite and replacive calcite (micro-)spar mosaic of zoned euhedral non-ferroan calcite overgrowth by ferroan calcite. W3, South Klofjellet, 116Á5 m. (F) Paired transmitted light (left) and CL (right) images. Fenestral microbial fabric similar to (D) displaying consistent crystal zonation: brighter to duller in micritic matrix and sharper zones within fenestral cements. Fabric is consistent either with primary calcite growth within extracellular polymeric substance or replacement of ikaite, followed by cementation of fenestrae. W2, Reinsryggen, 80Á6 m.
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Continental carbonate facies of a Neoproterozoic panglaciation,
north-east Svalbard
IAN J. FAIRCHILD*, EDWARD J. FLEMING*
1
, HUIMING BAO, DOUGLAS I.
BENN§, IAN BOOMER*, YURI V. DUBLYANSKY, GALEN P. HALVERSON**,
MICHAEL J. HAMBREY††, CHRIS HENDY‡‡, EMILY A. MCMILLAN*,
CHRISTOPH SP
OTL, CARL T. E. STEVENSON* and PETER M. WYNN§§
*School of Geography, Earth and Environmental Sciences, University of Birmingham, Birmingham
B15 2TT, UK (E-mail: i.j.fairchild@bham.ac.uk)
Department of Geology, The University Centre in Svalbard (UNIS), N-9171 Longyearbyen, Norway
Department of Geology and Geophysics, Louisiana State University, E235 Howe-Russell Complex,
Baton Rouge, LA 70803, USA
§School of Geography and Geosciences, University of St Andrews, St Andrews KY16 8YA, UK
Institut f
ur Geologie, Leopold-Franzens-Universit
at Innsbruck, Innrain 52, A-6020 Innsbruck, Austria
**Department of Earth & Planetary Sciences/Geotop, University St. Montreal, St. Montreal, Quebec,
Canada H3A 0E8
††Department of Geography and Earth Sciences, Aberystwyth University, Aberystwyth, Ceredigion
SY23 3DB, UK
‡‡Department of Chemistry, University of Waikato, Private Bag 3105, Hamilton, 3240, New Zealand
§§Lancaster Environment Centre, University of Lancaster, Lancaster LA1 4YQ, UK
Associate Editor – Guy Spence
ABSTRACT
The Marinoan panglaciation (ca 650 to 635 Ma) is represented in north-east
Svalbard by the 130 to 175 m thick Wilsonbreen Formation which contains
syn-glacial carbonates in its upper 100 m. These sediments are now known
to have been deposited under a CO
2
-rich atmosphere, late in the glaciation,
and global climate models facilitate testing of proposed analogues. Precipi-
tated carbonates occur in four of the seven facies associations identified:
Fluvial Channel (including stromatolitic and intraclastic limestones in
ephemeral stream deposits); Dolomitic Floodplain (dolomite-cemented sand
and siltstones, and microbial dolomites); Calcareous Lake Margin (intraclastic
dolomite and wave-rippled or aeolian siliciclastic facies); and Calcareous
Lake (slump-folded and locally re-sedimented rhythmic/stromatolitic lime-
stones and dolomites associated with ice-rafted sediment). There is no strong
cyclicity, and modern analogues suggest that sudden changes in lake level
may exert a strong control on facies geometry. Both calcite and dolomite in
stromatolites and rhythmites display either primary or early diagenetic repla-
cive growth. Oxygen isotope values (12 to +15&
VPDB
) broadly covary with
d
13
C. High d
13
C values of +35to+45&correspond to equilibration with an
atmosphere dominated by volcanically degassed CO
2
with d
13
Cof6to
7&. Limestones have consistently negative d
18
O values, while rhythmic
and playa dolomites preserve intermediate compositions, and dolocretes pos-
sess slightly negative to strongly positive d
18
O signatures, reflecting signifi-
cant evaporation under hyperarid conditions. Inferred meltwater
compositions (8to155&) could reflect smaller Rayleigh fractionation
1
Present address: CASP, West Building, 181A Huntingdon Road, Cambridge CB3 0DH, UK.
443©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists
Sedimentology (2016) 63, 443–497 doi: 10.1111/sed.12252
related to more limited cooling than in modern polar regions. A common
pseudomorph morphology is interpreted as a replacement of ikaite
(CaCO
3
H
2
O), which may also have been the precursor for widespread repla-
cive calcite mosaics. Local dolomitization of lacustrine facies is interpreted
to reflect microenvironments with fluctuating redox conditions. Although dif-
fering in (palaeo)latitude and carbonate abundance, the Wilsonbreen carbo-
nates provide strong parallels with the McMurdo Dry Valleys of Antarctica.
Keywords Carbon isotopes, Cryogenian, ikaite pseudomorphs, lacustrine,
oxygen isotopes, Snowball Earth.
INTRODUCTION
The second of two Neoproterozoic panglacia-
tions, in which ice sheets reached sea-level in
the tropics, terminated 635 Myr ago at the base
of a transgressive cap carbonate defining the
Cryogenian–Ediacaran System boundary
(Table 1). Deposits of Cryogenian ice ages are
preserved on most continents and are commonly
interpreted as glacimarine (Arnaud et al., 2011).
However, in north-east Svalbard, the Marinoan-
aged, 130 to 175 m thick Wilsonbreen Formation
(Halverson, 2011) uniquely contains non-marine
carbonates as well as subglacial tillites (Fig. 1).
These evince hyperarid terrestrial environments
(Fairchild et al., 1989) and an atmosphere rich
in carbon dioxide during glaciation (Bao et al.,
2009), the latter conclusion fulfilling a predic-
tion of the Snowball Earth hypothesis (Kirsch-
vink, 1992; Hoffman et al., 1998). Because
Wilsonbreen Formation outcrops are restricted
to remote icefield nunataks on Spitsbergen
(Fig. 2) and coastal eastern Nordaustlandet, they
have rarely been visited, and hence current
knowledge of the sedimentary architecture has
been incomplete. Wilsonbreen Formation car-
bonates contain the highest carbonate and sul-
phate d
18
O values and lowest sulphate D
17
O
signatures so far discovered in the geological
record, features which evoke one of the most
extreme climatic events in Earth history (Bao
et al., 2009; Benn et al., 2015). This article char-
acterizes a range of non-marine environments in
which the carbonates were precipitated using a
combination of field, petrographic and stable
isotope evidence, and scrutinizes their analogy
to the extreme terrestrial environments of the
modern McMurdo Dry Valley region of Antarc-
tica (Fairchild et al., 1989).
The study area in the Svalbard mainland of
Spitsbergen (Figs 1 and 2) and the basin contin-
uation to the north-east have long been recog-
nized as classic areas for late Precambrian
glaciation (Kulling, 1934). The first detailed
description of the Wilsonbreen Formation was
by Wilson & Harland (1964), although carbo-
nates were discussed only in terms of its bound-
ing dolomites which serve as stratigraphic
markers. A later sedimentological synthesis
(Hambrey, 1982; Fairchild & Hambrey, 1984)
showed that evidence of glacial activity was
confined to two glacial units in the Polarisbreen
Group: the Wilsonbreen Formation, and a newly
discovered, thin, older unit in the Elbobreen
Formation (Petrovbreen Member also known as
E2; Table 1). Several distinctive forms of carbo-
nate occur in association with the glacial depo-
sits. Dolomitic glacial rock flour was first
demonstrated in ultra-thin sections (Fairchild,
1983). Subsequent stable isotope studies showed
the presence of glacimarine precipitates in the
Petrovbreen Member and glacilacustrine depo-
sits in the Wilsonbreen Formation (Fairchild &
Hambrey, 1984; Fairchild & Spiro, 1987; Fair-
child et al., 1989). It was also shown that these
carbonates contrast with a distinctive marine
transgressive ‘cap carbonate’ (following Wil-
liams, 1979) over the Wilsonbreen Formation.
Halverson et al. (2004) provided a much more
detailed chemostratigraphic framework for the
Polarisbreen Group and postulated that both
diamictite units belonged to the same, Marinoan
glaciation. However, new chemostratigraphic
data led to a reversion to the previous two-fold
glaciation interpretation of older literature
(Halverson, 2006; Hoffman et al., 2012). No
direct geochronological constraints exist on
Svalbard, but the low
87
Sr/
86
Sr values at the
base of the Polarisbreen Group correlate well
with those associated with the first evidence of
Neoproterozoic glaciation elsewhere (Halverson
et al., 2010). Also, the transgressive–regressive
succession above the Wilsonbreen Formation
closely resembles the basal Ediacaran facies suc-
cession globally (Halverson et al., 2004). The
end of Marinoan glaciation, and hence the top
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
444 Ian J. Fairchild et al.
Table 1. Neoproterozoic chronostratigraphy and north-east Svalbard lithostratigraphy (after Halverson, 2011, updated by unpublished data). Glacial units
are highlighted in red. This paper deals with the Wilsonbreen Formation, representing the younger of the two Cryogenian glaciations
Geological
System Group Formation Member
Thickness
(m) Lithologies Interpreted environment
Ediacaran Polarisbreen Dracoisen D4 to D7 265 Sandstones and mudstones (D4
and D6); dolomite D5 and D7
Playas (D4 to D6); Coastal (D7)
D1 to D3 200 Cap carbonate (D1) transitional
(D2) to black shale D3)
Transgressive coastal (D1) to
offshore (D2 to D3)
Cryogenian Wilsonbreen W3 (Gropbreen) 65 to 95 Diamictites and sandstones
with minor limestone
Glacilacustrine; subglacial at base
W2 (Middle
Carbonate)
20 to 30 Three main intervals of carbonate-
bearing sandstones and siltstones
with intervening diamictites
Carbonate intervals are fluvial and
lacustrine with glacial influence;
diamictites are glacilacustrine
rainout deposits
W1 (Ormen) 55 to 85 Brecciated underlying dolomite
locally overlain by sandstones and
conglomerates passing up into
diamictites and sandstones with
local rhythmites
Basal periglaciated surface, locally
succeeded by fluvial deposits, then
glacilacustrine rainout deposits and
sediment-gravity flows
Elbobreen E4 (Slangen) 20 to 30 Oolitic dolomite Regressive peritidal
E3 (Macdonaldryggen) 200 Finely laminated dolomitic silty
shale
Offshore marine
E2 (Petrovbreen) 10 to 20 Dolomitic diamictites, rhythmites
and conglomerates
Glacimarine
(base to be
defined)
E1 (Russøya) 75 to 170 Dolomites overlain by limestone
with molar tooth structure, black
shale and dolomite
Shallow marine
Tonian Akademiker-
breen
Backlund-
toppen
530 Limestone and dolomite Carbonate platform
Draken 350 Intraclastic dolomite Peritidal
Svanberg-
fjellet
425 Limestone and dolomite Carbonate platform
Grusdiev-
breen
700 Limestone and dolomite Carbonate platform
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 445
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
446 Ian J. Fairchild et al.
of the Wilsonbreen Formation is 635 Ma (Roo-
ney et al., 2015, and references therein). Palaeo-
magnetic constraints suggest that north-east
Svalbard lay in the subtropics throughout the
Cryogenian (Li et al., 2013) as part of Laurentia,
which spanned the equator in the centre of the
fragmenting Rodinia supercontinent.
The closely similar stratigraphy in central East
Greenland (now officially redesignated as North-
east Greenland; Hambrey & Spencer, 1987;
Moncrieff & Hambrey, 1990) indicates that it
represents a basin continuation (Hambrey, 1983;
Knoll et al., 1986; Fairchild & Hambrey, 1995),
subsequently offset to the south by left-lateral
strike-slip faulting (Harland, 1997). The distinct
Neoproterozoic succession in western Svalbard
cannot easily be correlated with its counterpart
in north-east Svalbard and was probably depo-
sited in a separate basin (Harland et al., 1993).
Cryogenian events and panglaciation
Harland (1964) championed the idea of low-
latitude late Proterozoic glaciation based on
globally distributed diamictites that were con-
sidered to be glacigenic and correlative. How-
ever, firmly establishing the glacial affinity of
these units, determining their low-latitude ori-
gin, and confirming their correlation required
another four decades (Fairchild & Kennedy,
2007). Despite the complexity of individual suc-
cessions, as summarized in Arnaud et al. (2011),
it is now widely recognized that two panglacial
events occurred during the Cryogenian period
(ca 720 to 635 Ma), referred to as the Sturtian
(or early Cryogenian) and Marinoan (or late
Cryogenian) glaciations (Halverson et al., 2005;
MacDonald et al., 2010; Hoffman et al., 2012;
Calver et al., 2013; Lan et al., 2014; Rooney
et al., 2014, 2015). The traditional geological
interpretation of Cryogenian glaciation invoked
low atmospheric CO
2
levels, but it was not until
whole-Earth energy balance models emerged (for
example, to study a future ‘nuclear winter’) that
it became understood that a frozen, high-albedo
planet represents a stable climatic state (Budyko,
1969). Kirschvink (1992) first applied this model
to interpreting the geological record, suggesting
a Neoproterozoic Snowball Earth that was facili-
tated by low palaeolatitudes of the continents at
this time and that the key to escape from the
Snowball state was the build-up of volcanically
derived CO
2
(Caldeira & Kasting, 1992), which
at sufficiently high levels would overcome the
control of the albedo of Earth’s icy surface on
global energy balance and trigger deglaciation.
Hoffman et al. (1998) and Hoffman & Schrag
(2002) subsequently elaborated upon the Snow-
ball Earth model and provided the first sedimen-
tological, stratigraphic and geochemical
evidence in its support. Snowball Earth theory
stimulated data-collection and modelling which
have significantly clarified the criteria for attri-
bution of a suite of phenomena to Snowball
Earth conditions. Neoproterozoic glaciations are
associated with negative carbon isotope anoma-
lies (Knoll et al., 1986; Kaufman et al., 1997)
and Hoffman et al. (1998) argued that the ano-
maly apparently encompassing the Marinoan
glacial period reflected low organic productivity
as a result of glacial climate and an ocean that
became isolated from the atmosphere. Addition-
ally, the post-glacial cap carbonates and the
associated negative carbon isotope anomalies
were regarded as a positive test of the Snowball
Earth hypothesis, reflecting rapid meltdown and
re-equilibration of oceanic and atmospheric CO
2
during a post-Snowball, ultra-greenhouse envi-
ronment. However, subsequent work implies a
more nuanced scenario and the interpretations
Fig. 1. Reconstruction of the sedimentary architecture and palaeoenvironments of the Wilsonbreen Formation
and its constituent members (W1 to W3). Precipitated carbonate is present in the palaeoenvironmental group
called ‘Carbonate Fluvial and Lacustrine’ throughout W2 (except locally at the top and base), also in W3 and at
one of the locations in W1 as listed in Table 2. The Svalbard archipelago is shown bottom right and Spitsbergen
is the main island of the group, while Nordaustlandet is the island to the north-east. The rectangle on the Sval-
bard map shows the study area as enlarged upper right with Wilsonbreen Formation outcrops (red) within nuna-
taks (grey) rising from the highland snowfield. From north to south, study locations are: DRA (Dracoisen); here
the main section is located on a nunatak informally known as Multikolorfjellet with some additional sampling
from W2 at a second nunatak termed Tophatten 1 km to the north; DIT (Ditlovtoppen); AND (East Andromedafjel-
let); REIN (a ridge on South Andromedafjellet informally known as Reinsryggen); KLO (South Klofjellet); with
some additional observations from a partial W2 section 1 km away, at North Klofjellet; McD (MacDonaldryggen);
GOL (Golitsynfjellet, intermediate between McD and BAC) a partial W2 section was illustrated by Fairchild
et al. (1989); BAC (Backlundtoppen-Kvitfjellet ridge); PIN (an unnamed nunatak informally termed Pinnsvin-
ryggen); SLA (Southeast Slangen) and ORM (South Ormen).
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 447
A B
C
D
E
1
2
3
D1
W3
W
2
W1
W2
W2
W1
W3
W1
E4 W2
W3 D1
D1
W3
W2 W1
E4
BAC
REIN
KLO
ORM
DRA
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
448 Ian J. Fairchild et al.
of both the pre-glacial and post-glacial anoma-
lies remain disputed (Kennedy et al., 2001a,b;
Halverson et al., 2002; Hoffman & Schrag, 2002;
Schrag et al., 2002; Trindade et al., 2003; Hoff-
man et al., 2007, 2012; Le Hir et al., 2008, 2009;
Kennedy & Christie-Blick, 2011).
Early articulation of the Snowball Earth model
lacked a framework for interpreting the glacial
formations and their relevance for the theory. In
the original ‘hard’ Snowball, a universally thick,
ice cover on the oceans, composed of an upper
zone of glacier ice and lower zone of frozen sea-
water was envisaged (Pierrehumbert, 2005; Pier-
rehumbert et al., 2011). In this scenario, sea-
level should have been greatly lowered and mar-
ine ice margins should have occurred far below
the shallow-water pre-glacial sediments. Evi-
dence to support this hypothesis was found in
Namibia. On platform tops, glacigenic deposits
are rare, whereas on platform margins tidewater-
glacier grounding-line phenomena are inferred
to be some hundreds of metres lower topograph-
ically (Hoffman, 2011; Domack & Hoffman,
2013). On the other hand, most glacial sedimen-
tologists opposed the Snowball theory, com-
monly citing evidence of repeated glacial
advances and retreats in marine environments,
and wave-generated and storm-generated struc-
tures indicating open water inconsistent with
the hard Snowball scenario (Xiao et al., 2004;
Etienne et al., 2007; Allen & Etienne, 2008; Le
Heron et al., 2011, 2013). Apparent support
from models showing equatorial open water (e.g.
Hyde et al., 2000) faced the problem that the
simulated climate solutions were unstable.
Hence, Hoffman (2011) responded that: “counter
arguments [to the Snowball model] based on
temperate-type glacial sedimentology fail to
grasp that the preserved glacial sedimentary
record reflects the end of the Snowball Earth,
when melting was bound to emerge triumphant”.
New modelling that showed stable, long-lived
ice-free regions (‘Waterbelts’) in the tropics has
motivated rethinking the Snowball glacial record
(Pierrehumbert et al., 2011). Abbot et al. (2011)
proposed the term ‘Jormangund’ state to
describe a stable, ice-free equatorial fringe
achieved as the result of the low albedo of low-
latitude sea ice. Alternatively, Rose (2015)
argued for an entirely new climatic state in
which convergence of ocean heat transport at
tropical ice edges stabilizes a belt of open water
straddling the equator. Hoffman et al. (2012)
noted that although the Jormangund state pre-
served the pattern of modern low-latitude cli-
mate belts, with a moister equatorial region, the
hard Snowball climatic pattern (Pierrehumbert
et al., 2011) would result in higher precipitation
minus evaporation in the subtropics and an
extremely arid equatorial zone. This important
distinction draws attention to the need to study
low-latitude continental glacial deposits such as
the Wilsonbreen Formation where more direct
evidence of climatic conditions can be obtained.
The surviving essential predictions of Snow-
ball Earth theory can be summarized as follows:
(1) globally, glaciations must occur syn-
chronously, (2) they must be long-lasting
(>1 Ma) to allow (3) the build-up of atmospheric
CO
2
to high levels when (4) sedimentation
occurs in a brief period prior to termination.
Although earlier geochronological compilations
had legitimate doubts about (1) and (2) (Allen &
Etienne, 2008), they are now more firmly esta-
blished (Rooney et al., 2015), although the dura-
tion of the Marinoan glaciation (5 to 15 Myr) is
imprecisely known. The Wilsonbreen Formation
has now permitted positive tests of (3) and (4).
Compared to the Cenozoic, Neoproterozoic
strata offer limited options for determining
atmospheric CO
2
concentrations. Bao et al.
(2008) proposed a bold new approach based on
processes occurring during stratospheric ozone
formation which results in an enrichment in the
isotope
17
O in ozone and carbon dioxide and
depletion in oxygen (O
2
). This non-mass-depen-
Fig. 2. Examples of studied section outcrops. (A) Member W2 at Multikolorfjellet, Dracoisen (DRA) illustrating the
three groups of glacial retreat facies beds (numbered ‘1’ to ‘3’) separated by diamictites which also make up mem-
bers W1 and W3. (B) Member W2 at ORM (South Ormen) with pale sand-dominated units and dark red finer units.
Bedding is inverted and dips steeply away from the photographer. (C) REIN (Reinsryggen) section on the south
flank of Andromedafjellet illustrating Wilsonbreen Formation members overlain by cap carbonate member D1. Per-
sons for scale are ca 18 m tall. (D) KLO (South Klofjellet) section of the entire Wilsonbreen Formation on steep
slopes cut by minor faults, one of which is highlighted. The thickness of the Wilsbonbreen Formation, W1 to W3, is
135 m. (E) BAC (Backlundtoppen-Kvitfjellet ridge) photomontage from a helicopter hovering above the glacier Wil-
sonbreen. The visible section is vertical (thrust fault shown) and the accessible part defines a narrow ridge between
the cliff and a snowbank on the ridge crest. The thickness of the Wilsbonbreen Formation, W1 to W3, is 180 m.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 449
Table 2. Summary of characteristics of carbonates in members W1 and W3. Analyses are of calcite except where
italicized (dolomite): SD =standard deviation; n=number of samples
Section
Member,
m above
member base
(from member
top)
Facies
association Lithology
d
18
O&d
13
C&
nMean SD Mean SD
AND W3 30 (15) 5 Intraclastic stromatolitic rhythmites
(carbonate layers 5 to 10 mm thick)
with intervening diamictite. Some
calcite-filled pseudomorphs
922 020 027 033 5
AND W3 38 (7) 5 Stromatolitic limestone (laminae
5 to 10 mm thick) with diamictite
laminae, variably broken up within
diamictite. Observed over 100 m
laterally
816 055 023 070 6
AND W3 405(45) 5 Brecciated limestone ––
REIN W3 56 (65) 5 Limestone with scattered sand 892 044 1
KLO W3 25 (18) 5 Carbonates interlaminated with
diamictite at base of bed. EFK15
closely resembles the middle
carbonate layer at section AND
797 183 1
503 104 1
KLO W3 28 (15) 5 Laminated carbonates in diamictite.
Some crystal pseudomorphs both
within sediment and growing upward
886 040 038 010 3
404 116 1
KLO W3 31 (12) 5 Laminated carbonates in ice-rafted
sediments with prominent upward
growing crystal pseudomorphs (in
dolomite) of ikaite
944 238 031 109 4
KLO W3 35 (8) 5 Similar to horizon 4 m lower in section ––
PIN W3 n/a (26) 5 Slump folded and brecciated thickly
laminated stromatolitic limestone in
red silty sandy diamictites. Traceable
laterally for 30 m
1253 028 004 029 3
ORM W1 35(205) 5 Slump-folded rhythmites with distinct
millimetre-scale stromatolitic limestone
laminae, some with bulbous tops,
separated by diamictite
125032 2
ORM W3 28 (30) 3 Sandstone with convolute bedding
with nodular dolocrete (floating quartz,
micro-nodules and cracks) along specific
laminae
+678 560 2
ORM W3 33 (25) 4 Diamictite rests on massive stromatolitic
limestone with pseudomorphs and
slump folds or intraclasts over greenish
cross-laminated sandstone with ooids
1076 112 220 156 6
ORM W3 465(115) 5 Red, thickly laminated and slump-
folded stromatolitic limestone in
diamictite. Locally massive with a lot
of early cement
944 238 031 109 4
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
450 Ian J. Fairchild et al.
dent effect does not influence
18
O abundances.
The
17
O signal can be preserved in the oxygen
atoms of sulphate in rocks if atmospheric oxy-
gen is used to oxidize sulphides on the land sur-
face. Sulphate has the fortuitous property of not
exchanging oxygen atoms with other species
over 1000 Myr timescales at surface conditions,
provided it is not subject to microbial redox
cycling. Bao et al. (2008) showed that at time
periods when atmospheric PCO
2
was enhanced,
sulphate in the geological record was
17
O-
depleted. The most significant anomaly was dis-
covered in barite crystal fans within the carbo-
nate succession overlying Marinoan glacial
diamictites in South China.
Bao et al. (2009) subsequently found more
profound
17
O-deficiencies in carbonate-asso-
ciated sulphate (CAS) in lacustrine limestone of
the central part (member W2) of the Wilson-
breen Formation, consistent with very high
atmospheric PCO
2
actually during glaciation.
More recent studies in other geographic regions,
coupled with process modelling approaches,
have supported evidence for high PCO
2
during
deposition of Marinoan cap carbonates (Bao
et al., 2012; Cao & Bao, 2013; Killingsworth
et al., 2013; Bao, 2015), but the Wilsonbreen
Formation remains the only unit where PCO
2
can be estimated during glaciation.
Benn et al. (2015) applied the same approach
to a much larger data set of limestones from
members W2 and W3 to argue that similar high
PCO
2
values (estimated at 10 to 100 mbar atmo-
spheric CO
2
) occurred throughout the deposition
of the Wilsonbreen Formation. Because it would
have taken a long time to accumulate CO
2
, these
authors inferred that the bulk of the Wilson-
breen Formation was deposited in a relatively
short period near the end of the glaciation. In
turn, this implies an extended hiatus early in
the glaciation, which manifested as a perma-
frosted horizon at the base of the Wilsonbreen
Formation.
Coupled ice sheet and atmospheric general
circulation model results in Benn et al. (2015),
using Snowball Earth boundary conditions at
20 mbar atmospheric PCO
2
, display thick gla-
ciers on the continents coexisting with extensive
areas of bare ground and widespread hyperari-
dity. Under these conditions, the climatic conse-
quences of orbital precession cycles is to force
ice margins to migrate by at least hundreds of
kilometres, and the present authors link this
cyclicity to the presence of distinct ice advances
within the Wilsonbreen Formation (Fig. 1). The
presence of an arid tropical continent during the
Marinoan ice age is also consistent with obser-
vations of periglacial structures, aeolian, fluvial
and evaporative floodplain facies in the Aus-
tralian Marinoan successions (Williams, 2008;
Ewing et al., 2014; Retallack et al., 2015).
Although conclusions based on the Marinoan
glaciation should not necessarily apply to the
much longer Sturtian glaciation, the new results
provide a way to reconcile the opposed posi-
tions stated in Allen & Etienne (2008), of tem-
perate glacial conditions during panglaciation,
and Hoffman (2011), of sediment deposition
occurring rapidly during meltdown (Allen,
2015). The purpose of this contribution is to
provide a detailed sedimentological analysis of
the Wilsonbreen Formation as template for gla-
cial sedimentation in the terrestrial realm during
Snowball glaciation and to investigate the geo-
morphic-climatic system in which the Wilson-
breen carbonates were deposited. This analysis
begins by examining the most plausible modern
analogue.
The Antarctic McMurdo Dry Valleys as
analogues for the Wilsonbreen Formation
carbonates
Walter & Bauld (1983) were the first to attempt
to compare the frigid, arid Dry Valleys region
with carbonates and associated facies in Neopro-
terozoic successions, but the Australian exam-
ples chosen were either postglacial or had an
uncertain relationship with glaciation. Later,
Fairchild et al. (1989) made a more specific
comparison with member W2 of the Wilson-
breen Formation. Conversely, leading workers
on the Dry Valleys (Lyons et al., 2001) com-
mented that they might compare with Protero-
zoic Snowball Earth; they have also been used
as analogues for the Martian surface (Marchant
& Head, 2007; Dickson et al., 2013). The local
‘Alpine’ glaciers of the Dry Valleys are cold-
based, but outlet glaciers such as the Taylor Gla-
cier are warm-based, albeit because of sitting in
a saline basin (Mikucki & Priscu, 2007). Given
the criteria for recognition of a glacial thermal
regime presented by Hambrey & Glasser (2012),
warm-based analogues are needed to understand
much of the glacigenic facies of the Wilsonbreen
Formation (Fleming et al., 2016). Nevertheless,
for the carbonate facies, the context of the Dry
Valley region provides unique parallels for the
following evidence provided by Fairchild et al.
(1989): (i) a record of extreme evaporation with
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 451
an interpreted up to 20&difference in d
18
O
between input waters and those responsible for
the precipitation of the carbonates with heaviest
d
18
O signatures; and (ii) alternating deposition
of glacial sediment and continental deposits
indicating ice advance and retreat, specifically
including rhythmic and microbial, presumed
lacustrine carbonates and evidence of former
evaporites.
Figure 3 illustrates the geographical context of
the Dry Valleys region, which has been studied
extensively over several decades by groups led
from New Zealand, Japan and the USA, with
Taylor Valley being the focus since 1993 of the
McMurdo Dry Valleys Long-term Ecosystem
Research Program of the National Science Foun-
dation. The Dry Valleys occupy 4800 km
2
in the
Transantarctic Mountains between the East
Antarctic Ice Sheet (EAIS) and the Ross Sea.
The individual valleys mostly trend eastwest,
are up to 80 km long and 15 km wide and are
internally drained, commonly with several dis-
crete lake basins in each valley. Two outlet gla-
ciers from the EAIS (Wright Upper and Taylor
glaciers) only just cross the regional bedrock
altitude divide into the eastward-draining catch-
ments, while Ferrar Glacier flows all the way to
the coast (Fig. 3). Local valley glaciers extend
from the mountains into dry valleys, for exam-
ple, Canada Glacier in Taylor Valley (inset in
Fig. 3) and ephemeral streams develop in sum-
mer. A variety of lakes occur, including those
occupied by ice frozen to the bed in the rela-
tively high-altitude Victoria Valley, highly saline
lakes with no ice cover (for example, Don Juan
Pond, Upper Wright Valley) and finally lakes
with a 3 to 5 m ice cover (for example, Lakes
Brownsworth and Vanda in Wright Valley, and
Lakes Bonney, Hoare and Fryxell in Taylor Val-
ley), which melt only in a marginal moat in
summer (Green & Lyons, 2009; Dickson et al.,
2013). Microbe-dominated biotas flourish wher-
ever and whenever liquid water is present,
including as mats on stream and lake beds. Pho-
tosynthesizing algae are important in lakes dur-
ing the spring season (Fountain et al., 1999).
Microbial mats consume nutrients rapidly from
stream water, and biogeochemical cycling in
these systems is strongly influenced by exchange
of hyporheic waters with stream water
(McKnight et al., 1999).
Microclimates in the Dry Valleys comprise a
coastal zone, with soils that transiently thaw in
summer, and a stable cold upland zone with
particularly low relative humidity (Marchant &
Head, 2007; Marchant et al., 2013). The inter-
vening valleys have a mean annual temperature
of 16 to 21°C and the maximum average
daily temperature is below zero throughout the
year (Fountain et al., 1999). The valleys receive
less than 10 mm water-equivalent of precipita-
tion per year, almost always as snow. Because of
the prevailing low relative humidities (for exam-
ple, between 50% and 60% in Taylor Valley,
Fountain et al., 1999), ablation (mostly as subli-
mation) greatly exceeds precipitation. Wilson
(1981) described the consequences of this geo-
graphic configuration and climatology using
physical and chemical principles. The amount
of precipitation rises with altitude, but falls
inland further from the Ross Sea. The snowline
marks the boundary where ablation exceeds pre-
cipitation and it rises inland as precipitation
declines. Wilson (1981) attempted to explain the
distribution of salts based on an understanding
of downslope-increasing aridity, but it seems
that variable meteorology confounds the predic-
tions in detail. Nevertheless, it is the case that
deliquescent salts flow downhill in the subsoil
above a permanently frozen layer. Lakes with a
lid of ice maintain mass balance between abla-
tion at the surface and freezing of lake water at
the base of the lid, replenished seasonally by
stream inflow. Once snow is removed by abla-
tion, sunlight penetrates through vertical ice
crystals and significant solar heating of lakes
can occur. The present configuration of salts
allows deductions of both long-term and short-
term history. Specifically, spatial variability in
salts requires a long-term (>10
5
to 10
6
years) sta-
bility of the ice-free subaerial valley sides. On
the other hand, some lakes have a basal brine
layer, which diffusion modelling shows origi-
nated ca 1200 years ago when some lakes were
ice-free shallow brines, before being re-filled by
fresh meltwater.
Wilson (1981) did not address the effect of
wind on aridity and salt distribution. Katabatic
winds flowing down the East Antarctic Ice Sheet
strongly enhance aridity regionally. As this air
warms adiabatically, humidity decreases, parti-
cularly in winter (Nylen et al., 2004). It is now
clear that the episodically strong summer winds
are actually warm foehn winds, which arise
from strong pressure gradients that arise during
cyclonic development over the Ross Sea, gov-
erned by hemispheric climatic anomalies. These
topographically enhanced and channelled winds
typically flow at >5m s
1
westerly along the
Dry Valleys and cause very large intra-annual
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
452 Ian J. Fairchild et al.
163°
E
160°
E
77·2°
S
77·6°
78°S
McMurdo
Sound
EAIS
EAIS
Miers Valley
Taylor Valley
Wright Valley
Ice movement
at LGM
Victoria Valley
Ross Ic e Shelf
EA
WA
RIS
Ferrar Glacier
Tay lo r G lac ier
18
km
A
B C
Fig. 3. (A) Oblique aerial view of the McMurdo Dry Valleys (01/01/1999 imagery from US Geological Survey via
Google Earth) with location arrowed in inset of Antarctica (upper right). EAIS =East Antarctic Ice Sheet.
LGM =Last Glacial Maximum. Abbreviations on upper right inset: EA =East Antarctica, WA =West Antarctica,
RIS =Ross Ice Shelf. The lower left inset shows an oblique aerial photograph looking west up Taylor Valley with
cold-based valley glaciers on hillside on left, ice-covered Lake Bonney (valley floor, right) with the tip of the Tay-
lor Glacier beyond. (B) Despite sub-zero temperatures, runoff occurs from the Lower Wright Glacier (an outlet gla-
cier of the coastal Wilson Piedmont Glacier) and feeds the Onyx River. This stream flows inland to the west,
eventually to Lake Vanda, visible in (A) in the central part of the valley. Note the aeolian sands banked against
the glacier. (C) Oblique aerial view of Victoria Lower Glacier (eastern end of Victoria Valley) which feeds a stream
flowing inland, westward to Lake Vida, seen in the distance. Aeolian dunes tranverse to the stream are visible on
the right-side of the the valley.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 453
and inter-annual increases in meltwater produc-
tion and streamflow (Doran et al., 2008; Speirs
et al., 2013). The high incidence of these winds
in summer 2001/2002, when positive degree
days increased by an order of magnitude, led to
rises in lake level of 05to10 m, effectively
wiping out the previous 14 years of lowering of
lake level in a period of three months (Doran
et al., 2008). These details effectively demon-
strate the sensitivity of the environment to cli-
matic changes, which in turn will strongly
influence sedimentary facies.
The Dry Valleys region and the adjacent off-
shore area, seaward of Ferrar Glacier, demon-
strate a 34 Myr history of glaciation. Cores from
several offshore drillholes indicate that the first
evidence of glaciation was at the Eocene/Oligo-
cene boundary, and that the thermal characteris-
tics of the glaciers changed from temperate
(associated with abundant flora), through poly-
thermal conditions to the predominantly cold
glaciers of today (Hambrey & Barrett, 1993; Ham-
brey et al., 2002; Barrett, 2007; Naish et al.,
2008). The McMurdo region has generated con-
siderable controversy concerning the timing of
the switch to cold, arid conditions, specifically
whether the switch took place in Miocene or
Late Pliocene time; these views have been
reviewed and re-evaluated by Barrett (2013). Evi-
dence for landscape evolution based on the
position of Miocene volcanic ash deposits
clearly demonstrates that, after the establish-
ment of the current large, cold, and stable East
Antarctic Ice Sheet in the Miocene, only surfi-
cial landscape modification has occurred (Sug-
den et al., 1995; Lewis et al., 2007). Importantly,
the cold, dry climates of the Dry Valleys
remained stable, even during significant warm-
ing events recorded in the Ross Sea during the
Pliocene. This long-term climatic stasis can be
compared with the predicted long-term hydro-
logical inactivity anticipated on a Snowball
Earth as carbon dioxide levels slowly rose.
Large lakes formed during the Last Glacial
Maximum in all the major Dry Valleys this
required two conditions: (i) expansion of the Ross
Sea Ice sheet to block the marine margins of the
Dry Valleys (Hendy, 1980; Hall et al., 2013); and
(ii) increased meltwater production in the valley
(by wind-induced melting, Doran et al., 2008),
despite significantly colder conditions. Con-
versely, dating of aragonitic lacustrine deposits
shows that Taylor Glacier (with characteristically
light oxygen isotope values) expanded during
interglacial periods, as did local valley-side gla-
ciers, which have heavier isotope values (Hendy,
1979, 1980). These phenomena presumably
reflect higher snow accumulation on the EAIS,
enhanced meltwater production and partial
blockage of the valley by ice. Importantly, these
inferences draw attention to the potential anti-
phasing of global temperature and local glacier
advance, and the greater importance of regional
humidity controls on Milankovitch timescales.
In the current paper, a wealth of new data on
the Wilsonbreen Formation carbonates are pre-
sented which allows the analogies previously
made to be tested and evaluated in depth. The
interpretation of these data is assisted by mod-
ern analogues and new computer simulation
studies of Neoproterozoic climates (Pierrehum-
bert et al., 2011; Benn et al., 2015).
METHODS
Fieldwork (2010/2011) was supported by heli-
copter and by snow scooter. Sections, including
six entirely new locations, were measured by
30 m tape, orientated by compass and Abney
level, and linked to bedding dips, with total
thickness checked by GPS with uncertainty of
around 5%. The best available sections in each
region were logged (Figs 1 and 2), although they
vary in quality from almost complete exposure,
to intermittent outcrop separated by ground-
level frost-shattered regolith. The carbonate
rocks are chemically fresh, but laboratory study
of cut or sectioned samples was necessary to
identify facies in many cases. Over 350 samples
were sawn, 210 of which were thin-sectioned;
60 of these were stained with Alizarin Red-S
and potassium ferricyanide, and over 30 pol-
ished sections were studied by cold-cathode
cathodoluminescence (CL) at 15 kV.
Carbon and oxygen stable isotope data are pre-
sented here as d
13
Candd
18
O in parts per thou-
sand with respect to the VPDB (Vienna Pee Dee
Belemnite) standard. Differences between labora-
tories are insignificant in relation to the wide
range of isotope values (29&in d
18
Oand8&in
d
13
C) in this study. Supplementary data in Bao
et al. (2009) included methods and all data col-
lected to that date. New data were obtained at the
University of Birmingham using a continuous-
flow Isoprime IRMS (isotope ratio mass spectro-
meter; Isoprime Limited, Stockport, UK), with a
multiflow preparation system. Samples of
between 80 to 250 lg powdered carbonate were
reacted with phosphoric acid at 90°C for at least
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
454 Ian J. Fairchild et al.
90 min. Results were calibrated using IAEA
(International Atomic Energy Agency) standards
NBS-18 and NBS-19 (National Bureau of Stan-
dards) and repeatability on an internal standard
was better than 01&for d
13
C and 015&for
d
18
O. A fluid inclusion study is reported in the
supplementary information. Sulphate oxygen
and sulphur isotopes are presented in Benn et al.
(2015). The present study draws on previously
presented trace element data from the carbonate
fraction soluble in dilute nitric acid (Fairchild &
Spiro, 1987; Fairchild et al., 1989; Bao et al.,
2009), while new trace element and other isotope
analyses will be presented elsewhere.
In figure captions, the location and strati-
graphic position are given in a standard format.
For example, ‘W2, Dracoisen, 70 m’ refers to a
sample from member W2 from the Dracoisen sec-
tion, 70 m above the base of the Formation (or
the base of the section where the base is not
seen). The context and oxygen isotope composi-
tion of the carbonate from that horizon can be
found in the stratigraphic section diagrams:
Fig. 4 for Dracoisen and supplementary Figs 1 to
6 for the other sections. Carbonates in members
W1 and W3 are summarized only in Table 2, but
full stratigraphic logs of the Wilsonbreen Forma-
tion are given in Benn et al. (2015, fig. S1).
WILSONBREEN FORMATION
ARCHITECTURE, COMPOSITION AND
POST-DEPOSITIONAL HISTORY
The Wilsonbreen Formation, dominated litho-
logically by sandy diamictites, contains a wealth
of evidence indicating that it is largely glaci-
genic, and was deposited in both aqueous and
subglacial settings (Hambrey, 1982; Fairchild &
Hambrey, 1984; Dowdeswell et al., 1985; Har-
land et al., 1993). New work (Benn et al., 2015;
Fleming et al., 2016), including five previously
undescribed sections, has resulted in a coherent
stratigraphic-facies reconstruction (Fig. 1). This
paper focuses on beds (sandstones, rhythmites
and mudrocks) containing precipitated carbo-
nate; the presence of such strata was originally
used to define member W2 (Hambrey, 1982). In
the northernmost section (Dracoisen), W2 is
readily distinguished by three such carbonate-
bearing beds separated by diamictites (Fig. 2A),
and overall a similar pattern applies in other
sections (Fig. 1). However, thin rhythmite units
also occur in member W1, one containing pre-
cipitated carbonate, and most of the thin non-
diamictite (sandstone and rhythmites) beds in
member W3 contain such carbonate facies
(Table 2). Although glacigenic rocks continue
NNE to the coast and to Nordaustlandet, these
were not included in this study because diamic-
tites are thinner and less well exposed, and pre-
cipitated carbonates are absent (Halverson et al.,
2004; Hoffman et al., 2012). Neither do such car-
bonates occur in the equivalent Storeelv Forma-
tion of Northeast Greenland, representing the
south-western continuation of the basin (Ham-
brey & Spencer, 1987; Moncrieff & Hambrey,
1990; Hoffman et al., 2012; Fleming, 2014).
Terrigenous detritus
Clarifying the nature of reworked (terrigenous)
clastic material is important for the petrological
study of the Wilsonbreen Formation. Harland
et al. (1993) summarized information from five
sites, determining that carbonate clasts make up
40 to 85%, igneous and metamorphic clasts 5 to
33%, and sandstones and quartzites 10 to 20%
of stones (i.e. gravel-sized debris); the basement
lithologies are primarily granitoids and gneisses,
with some basalt in which ferromagnesian min-
erals are commonly chloritized. In the wider
context of terranes affected by the Caledonian
orogen, detrital zircon studies indicate prove-
nance from Archaean and Palaeoproterozoic
rocks of the cratonic interior, and Mesoprotero-
zoic detritus derived from the eroded remnants
of the Grenville–Sveconorwegian orogen
(Cawood et al., 2007). A systematic study of car-
bonate clast compositions will be presented
elsewhere, but in summary 80% of pebbles are
dolomite rock and 20% are limestone. Many of
these carbonates resemble units from the under-
lying succession (Table 1) in lithology and iso-
tope composition, but no stratigraphic trends in
clast type or isotope composition were found.
The sand fraction is dominated by quartz and
feldspar with subordinate dolomite and lime-
stone and the mud fraction likewise comprises
quartzo-feldspathic debris and dolomite (Fig. 5).
The dominance of silt in the mud fraction is
shown in graded rhythmites (Fig. 5C and D);
notably detrital calcite occurs in the mud frac-
tion only locally and the proportion of clay min-
erals is very low. The dolomite matrix of
Polarisbreen Group diamictites is predominantly
the result of glacial transport and comminution,
as shown by the presence of sub-micron relic
rock flour (Fairchild, 1983) in which clay mine-
rals are virtually absent (Fig. 5D). This detrital
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 455
matrix contains a complex mixture of both
bright and dully luminescent red dolomite when
viewed in CL (Fig. 5A). The mean carbon and
oxygen isotope compositions of dolomitic matrix
in diamictites and wackes (n=43) are
+2413&and 3518&, respectively,
which are slightly lower values than for dolo-
mite pebbles (Fig. 6A). This matrix composition
forms a useful reference point for comparison
with precipitated carbonates in the Wilsonbreen
Formation.
Burial diagenetic modification
The Wilsonbreen Formation is overlain by a
marine transgressive succession but, above the
transgressive–regressive cycle of the lower
Dracoisen Formation, most Ediacaran strata
(Table 1) are non-marine playa lake facies, in
turn overlain by 950 m of Cambro–Ordovician
platform carbonates (Harland, 1997), indicating
a minimum burial of 15 km. Late Silurian Cale-
donian folding and local thrusting followed
(Fig. 2), but small-scale folding is generally
absent, as is penetrative deformation. The Wil-
sonbreen Formation lies outside the thermal
aureoles of Devonian granite plutons in the fold
belt (Harland, 1997). Pores in the cap carbonate
are filled with traces of bitumen, indicating that
the sediments passed through the oil window,
but good preservation is indicated by the ability
to remove individual striated clasts from diamic-
tite matrix (Hambrey, 1982) and some unusually
high oxygen isotope compositions of dolomites
(Fairchild et al., 1989; Bao et al., 2009). A fluid
inclusion study (see supplementary information)
indicates that initial pore fluids were advec-
tively replaced by a typical meteoric fluid with
d
2
H in the range 60 to 100&and that
18
Ois
slightly enriched by exchange with the solid
phase. However, fluid inclusion volumes are
orders of magnitude too small to have affected
the bulk solid composition. Physical compaction
effects are not noticeable because of the low clay
content of the sediments and early cementation
of the carbonates, but very locally lamina
boundaries are styolitic. Uncemented sandstones
exhibit straight to concavo-convex boundaries
between quartz grains.
Many Wilsonbreen Formation outcrops are
reddened, and even in unreddened facies,
organic carbon contents are low (below detec-
tion levels of 02% on three W2 carbonates).
Fairchild & Hambrey (1984) regarded the hema-
tite pigment as post-depositional based on the
occurrence of Liesegang band structures. Ongo-
ing palaeomagnetic studies may further eluci-
date the timing of this hemitization. Reddening
is less pervasive in clean sandstones, implying
that the source of iron in reddened zones is the
fine fraction. The low preservation of organic
carbon may reflect low abundance of clay (Ken-
nedy et al., 2006) as well as a low-productivity
continental setting. In either case, it permitted
the diagenetic reddening, which contrasts with
the dark-coloured early Cryogenian glacimarine
deposits in the study area (Member E2, Table 1).
Burial cements are minor and readily distin-
guished from syn-depositonal phenomena. Lim-
ited quartz overgrowths are common in
sandstones and on sand grains floating in dolo-
mites, while some sandstones display partial
poikilotopic calcite cement. Ferroan saddle
dolomite and ferroan calcite occlude larger pri-
mary pores, in dolomites and limestones,
respectively. These carbonate phases also occur,
locally together with quartz and white mica in
crystal pseudomorphs. Saddle dolomite averages
12215&in d
18
O and 1815&in d
13
C
and calcite spar 10516&and +2015&,
respectively (Bao et al., 2009).
FACIES ANALYSIS
Facies Associations
Following Benn et al. (2015) seven facies associa-
tions (FAs) are recognized in the Wilsonbreen
Formation. Table 3 presents summary descrip-
tions and interpretations, whereas Fig. 7 illus-
trates inferred environments. The glacial and
Fig. 4. Profile of member W2 at Dracoisen. The lithological log emphasizes physical characteristics (see key,
upper right) while the assignment to facies associations (centre) also draws on petrological and stable isotope
information. FA1 to FA6 represent an environmental continuum (cf. Fig. 7) and FA7 is grouped with FA1 as the
terrestrial glacial end-member. Each datapoint (bar on thick black line) represents the stratigraphic position of a
studied sample or observed lithological boundary in the field. The oxygen isotope composition of precipitated cal-
cites and dolomites (the latter adjusted by 3&) are shown. Mixed-mineralogy samples (<90% calcite or dolomite
in calcite–dolomite mixtures) are not plotted. The same conventions are used for profiles of the other studied sec-
tions, which are presented as supplementary figures (Figs S1 to S6).
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
456 Ian J. Fairchild et al.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 457
periglacial facies associations (FA1, 6 and 7) are
discussed in Hambrey (1982), Fairchild & Ham-
brey (1984), Benn et al. (2015) and Fleming et al.
(2016), and only a few salient points are men-
tioned here. Two distinct glacier advances from
the south are recognized, just below and above
member W2, by occurrences of FA1: sheared
diamictite, sandstones and gravels, locally resting
on highly deformed rhythmites (glacitectonites)
or striated clast pavements (Fig. 1, labelled ‘FA1
subglacial’). The bulk of the Wilsonbreen Forma-
tion is, however, composed of weakly stratified
diamictites (FA6) with decimetre to metre-scale
lenses of rhythmites (with precipitated carbonate
in members W2 and W3). Local presence of drop-
stones or isolated gravel clasts (lonestones), till
pellets and stratification, and the absence of sub-
glacial shearing phenomena, point to a subaque-
ous ice-rafted origin. Local lensing conglomerates
are interpreted as sediment-gravity flows which,
at Dracoisen, comprise a conspicuous low-angle
cross-stratified unit (Fig. 1, labelled ‘FA6 proxi-
mal’), interpreted as a grounding-line fan (Benn
et al., 2015; Fleming et al., 2016).
Gravel with ventifacts overlying shattered
dolomite represents the Periglacial Facies Asso-
ciation (FA7) at the base of the Wilsonbreen For-
mation, interpreted as a multi-million year
continental hiatus (Benn et al., 2015). Facies
Association 7 also occurs at the base of W2
where a periglacial exposure horizon with sand-
stone wedges penetrating subglacial till is found
(Fig. 1) at South Ormen, South Klofjellet and
Ditlovtoppen and also at the Golitsynfjellet sec-
tion (Fairchild et al., 1989), which is between
McDonaldryggen and the Backlundtoppen–Kvitf-
jella Ridge (BAC). The periglacial horizon is
overlain by sandstones correlated with those at
the base of W2 at Dracoisen (DRA).
The other four facies associations contain pre-
cipitated carbonate and are described in detail
below with summaries presented in Table 3 and
Fig. 7. Each of the facies associations occurs as
units 01to40 m in thickness (Fig. 4 and sup-
plementary information Figs S1 to S6).
Facies Association 2 (Facies 2S and 2T)
Description
This facies association is represented by erosion-
ally based, fining-upward tabular sandstone beds,
05to40 m in thickness. Facies 2T typically
A B
100 µm
1 mm
10 µm
D
C
Fig. 5. Detrital textures and
minerals. (A) and (B) Paired images
of a polished thin section under CL
and in transmitted light,
respectively, of the matrix of a silty
sandstone (W2, 812 m, Dracoisen)
illustrating the abundance of faint
blue-luminescing feldspar and
varied fragments of both bright and
dull red-luminescing dolomite in
the mud fraction. (C) Thin section,
transmitted light. Siltstone graded
rhythmites (W3, 604 m, East
Andromedafjellet) displaying
several sand-sized till pellets. (D)
Crossed polars. Matrix at the top of
a graded silt layer illustrating
quartzo-feldspathic debris and
micron-sized dolomite, but no clay
minerals (W1, 10 m Slangen).
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
458 Ian J. Fairchild et al.
–4
–3
–2
–1
0
1
2
3
4
5
6
–15 –10 –5 0 5 10 15
13
C
V-PDB
(‰)
18
O
V-PDB
(‰)
FA2 calcite FA4 calcite
FA4 dolom ite FA3 dolom ite
FA5 calcite FA5 dolom ite
Dolomite detrital matrix mean
A
FA2 calcite: y = 0·37x + 5·31 R² = 0·58
FA4 calcite:R² = 0·0006
FA4 dolo mite : y = 0·35x + 3·27 R² = 0·46
FA3 dolo mite : y = 0·10x + 3·41 R² = 0·58
FA5 calcite: R² = 0·0011
FA5 dolo mite : y = 0·31x + 2·73 R² = 0·71
–2
–1
0
1
2
3
4
5
6
15 –10 –5 0 5 10 15
13CV-PDB (‰)
18OV-PDB (‰)
FA5 dolomite
FA5 calcite
FA4
dolomite
FA4 calcite
B
–4
–2
0
2
4
6
8
10
–8 –6 –4 –2 024
13
C
VPDB
, (‰)
18
O
VPDB
(‰)
Dolomite cl asts Limestone cla sts Matrix
Fig. 6. (A) Summary of stable isotope compositions of precipitated carbonate. Inset gives detail of compositions
of clasts and fine dolomite matrix. Dolomite and calcite groupings are separated and only samples with >90% of
either dolomite or calcite in the carbonate fraction are plotted. (B) Stable isotope fields illustrating the degree of
covariance of the sample groups.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 459
Table 3. Facies and facies associations of the Wilsonbreen Formation. The facies associations are numbered to
represent an environmental continuum as depicted in Fig. 7
Facies association Constituent facies Environmental interpretation
FA1: Deformed
Diamictite
Detailed facies
analysis in
Fleming et al.
(2016)
Dominated by massive diamictites, locally with deformed
stratification with lenticular sand and gravel bodies
(1 to 5 m thick) with disrupted margins. Locally
associated with upward-increasing shear of underlying
sediment and striated boulder pavements
Subglacial till and channel
bodies; glacitectonites
FA2: Fluvial
Channel
2S:Beds of 05to40 m very fine to medium-grained,
locally cross-stratified sandstone with erosional
conglomeratic base. Some variants have low-angle
accretion surfaces with silt-dominated beds.
Stromatolitic limestone is either absent or abundant,
forming decimetre-scale packages with millimetre to
centimetre-scale laminae separated by sand laminae
and laterally eroded into trains of intraclasts
2T: As above, but lacking carbonate
precipitates
Ephemeral stream channel
sands with strongly seasonal
flows. In some cases there is a
widespread carpet of calcified
microbial mats
FA3: Dolomitic
Floodplain
All facies contain dolomite with a positive d
18
O
composition that actively cements, displaces and/or
upwardly accretes the sediment
3D: Nodules, centimetre to decimetre-scale of dolomite-
cemented silts and sands transitional to structureless
decimetre to metre-scale beds with floating silt and
sand in dolomite, internal nodules and calcite-
cemented fractures
3S: Stromatolitic dololaminites, decimetre-
scale and millimetre-laminated
Dolocretes representing soil
horizons, typically above fluvial
deposits, with locally raised
water table
3N: Nodular dolocrete
3D: Bedded dolocrete
3S: Subaerial stromatolites
FA4: Calcareous
Lake Margin
4R: Rhythmites with millimetre-scale dolomite, limestone
or mixed-mineralogy laminae, commonly desiccated,
alternating with 1 to 10 mm wavy cross-laminated silty
sandstones. Carbonate laminae are often stromatolitic.
Almost invariably reddened
4I: Intraclastic, decimetre-scale, very fine to medium-
grained sandstones. Very locally contain ooids
4S: Indistinctly horizontally stratified well-sorted
fine to medium-grained sandstone, with well-rounded
grains, locally with centimetre-scale dolomite laminae
4R: Shallow lake to playa
environment with wave-
reworked sediment
alternating with microbial
mat accretion
4I: Discrete storm horizons
in shallow lake/playa reworking
sand and stromatolitic carbonate
4S: Aeolian sandflat deposits
FA5: Carbonate
Lake
5R: Rhythmites with millimetre-scale limestone, dolomite
or mixed-mineralogy laminae, usually with stromatolitic
microstructure and locally building decimetre
stromatolites with centimetre-scale relief. Alternate with
01to50 mm laminae of diamictite/wacke, commonly
with till pellets and occasional lonestones. Slump folding
and brecciation common
5D: Discrete gravel to diamictite units with significant
intraclastic debris in addition to terrigenous gravel
5R: Carbonate microbial mat
deposits in lacustrine
environment subject to seasonal
ice-rafting and slope instability
5D: Slope-related resedimentation
FA6: Glacilacustrine
Detailed facies
analysis in Fleming
et al. (2016)
Dominated by massive and stratified diamictites
with common lonestones and till pellets.
Decimetre to metre-scale intervals of silty rhythmites
occur locally, especially in member W1. May contain
lenses of conglomerate, sometimes channelled, and
lenses or thin beds of sandstone forming packages
which can be inclined at up to 20°
Ice-rafted glacilacustrine
sediment locally reworked
as sediment-gravity
flows. Inclined packages
form grounding-line fans
(termed proximal glacilacustrine
on Fig. 1) and rhythmites
likewise reflect ice-margin retreat
FA7: Periglacial
See Fairchild &
Hambrey (1984) and
Benn et al. (2015)
Decimetre-wide sandstone wedges penetrating up to 2 m
into underlying sediment from a discrete surface that may
have a gravel lag. At base of Wilsonbreen Formation,
gravel with ventifacts overlying shattered dolostone
Exposed periglaciated bedrock
and sand wedges from exposed
periglacial environment
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
460 Ian J. Fairchild et al.
comprises moderately sorted very fine to med-
ium-grained sandstone with a basal erosional sur-
face capped by a gravel lag. Internal cross-
stratification is tabular, typically with low-angle
accretion surfaces and set thickness of 05to
10 m. This facies is seen at the base of the W2
section at Ditlovtoppen (Fig. S1), South Ormen
(Fig. S2), East Andromedafjellet (Fig. S3) and
Reinsryggen (Fig. S4), and is also prominent near
the top of the member at Dracoisen (Fig. 8H) and
at Ditlovtoppen, where thin siltstone beds define
low-angle accretion surfaces (Fig. 8G).
Facies 2S is distinguished from Facies 2T by
the presence of limestone laminae and intraclasts,
and reduced silt content. The best-exposed exam-
ple is a 4 m thick unit that forms the basal part of
member W2 at Dracoisen (Figs 2A and 4), resting
on massive diamictite with decimetre-scale ero-
sional relief. A thin basal pebble conglomerate
passes up into 07 m of pebbly sandstone,
whereas the main, central part of the bed is very
fine to medium-grained sandstone with local tab-
ular cross-stratification, with set thickness up to
15 cm and current ripples (Fig. 8D). At both
levels, there are abundant (ca 25 to 50%) layers of
limestone, with individual laminae typically 1 to
3 mm thick (Fig. 8A). Universal features of the
limestone are domed growth morphologies
(Fig. 8A) and differentiated microstructures
(Fig. 8B) with well-developed slightly irregular
laminae of micrite, microspar and detritus-rich
carbonate with regular millimetre-scale fenestrae
(Fairchild, 1991; Riding, 2000). In one case a per-
vasive vertical structure reminiscent of cyanobac-
terial or algal filament moulds (Knoll et al., 1993)
was observed (Fig. 9B). All limestone occur-
rences are laterally discontinuous, and grade lat-
erally to trains of intraclasts that are locally
stacked at high angles. In places, extensive crusts,
up to 05to60 mm thick, of radiaxial calcite
Fig. 7. Summary cartoon of facies associations. Facies Associations 2 to 5 are colour-coded here and in isotope
plots. The margin of an ice sheet is depicted, terminating partly on land and, in the foreground, in the lake. Facies
Association 1 is shown in a subglacial environment (subglacial sediments are coloured brown); periglacial phe-
nomena are also grouped in this facies association. Facies Associations 2 and 3 both occur in a fluvial setting
(sediments coloured yellow), while FA5 and FA6 were deposited in lakes (sediments coloured light brown).
Facies Association 4 includes both shallow lacustrine and coastal sediments.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 461
cement develop, also broken to form intraclasts
(Fig. 8C).
Under CL, intraclasts fluoresce uniformly
brightly, whereas later vein cements show more
variable characteristics (Fig. 9A). This pattern is
mirrored geochemically in high Mn contents of
several thousand ppm and exceptionally high
Mn/Fe of 1 (Bao et al., 2009) of the intraclasts.
Strontium contents are 150 to 300 ppm, while
Mg is 4000 to 6000 ppm (Bao et al., 2009),
equivalent to ca 2 mole % MgCO
3
. The radiaxial
fabrics appear pristine and microdolomite inclu-
sions are absent. Sulphate content is high (2000
to 5000 ppm, Bao et al., 2009), whereas the
preservation of a negative D
17
O anomaly demon-
strates that the sulphate has not undergone
redox cycling (Bao et al., 2009; Benn et al.,
2015). Stable isotope data from stromatolitic
limestones define a coherent field (Fig. 6A) with
d
18
O ranging from 105to34&and d
13
C
from +09to+46&, weighted towards higher
values, with overall isotopic covariation
(Fig. 6B). A micromill traverse through syndepo-
sitional calcite reveals ‘lamina to lamina’ varia-
tions of 20&in d
18
O and 05&in d
13
C, without
strong covariation (Fig. 8E and F). The clear pet-
rographic distinction between syn-depositional
and later calcite spar cement is reflected also in
the low d
18
O signature of the spar of 10 to
12&(Fig. 8F).
Interpretation
Facies Association 2 is distinguished by trac-
tional sediment transport. The consistent pre-
sence of laterally extensive basal erosion
surfaces imply channelization, while for Facies
2T the low-angle accretion surfaces with thin
silt beds suggests migrating point-bar deposits
(Davies & Gibling, 2010), although confirmation
would require more extensive exposures. The
presence of high-angle cross-stratification, good
sorting and disrupted intraclasts in Facies 2S
are characteristics found in either tidal sand
flats or in low sinuosity fluvial channels. The
limestones contain two features typical of Neo-
proterozoic microbial deposits: macroscopically
domed laminae and differentiated microstruc-
tures formed by periodic variations in phenom-
ena such a sediment trapping, gas generation
and carbonate precipitation, fulfilling criteria for
stromatolites (Fairchild, 1991; Riding, 2000).
Limestones were lithified, as demonstrated by
broken crusts and dispersed and stacked intra-
clasts, for example along foresets, as expected
for significant tractional flows in shallow-water.
The later disruption and discontinuity of the
limestone units indicates highly variable flow
conditions, whereas the locally highly regular
microbial lamination, including clastic layers,
indicates periodically fluctuating flows.
Facies 2S resembles tidal sandflat deposits,
consistent with the occurrence of well-preserved
radiaxial cements with relatively high d
18
O com-
positions (Fairchild & Spiro, 1987). However,
there is an absence of herringbone cross-stratifi-
cation and reactivation surfaces which com-
monly occur in such facies (Fairchild, 1980;
Fairchild & Herrington, 1989). There is a clear
contrast with the more regular macrostructures
of marine stromatolites elsewhere in the basin
(Fairchild & Herrington, 1989; Knoll & Swett,
1990; Halverson et al., 2004) and Neoproterozoic
deposits more generally (Grotzinger & Knoll,
1999). This may be attributable to a more hostile
environment, with highly variable rates of sedi-
mentation. Furthermore, unlike the Facies 2S
stromatolites, Neoproterozoic peritidal deposits
are invariably dolomitized (Knoll & Swett,
1990), probably a feature of high Mg/Ca in sea-
water (Hood & Wallace, 2012).
The present authors favour a freshwater, fluvial
context for this facies association. It is recognized
that radiaxial fabrics are not diagnostic of marine
waters, but also occur in speleothems (Neuser &
Fig. 8. Facies Association 2 (Fluvial Channel). (A) to (F) are Facies 2S (W2, Dracoisen, at around the 60 m level),
whereas (G) and (H) are Facies 2T. (A) Sandstone with microbial laminites and low-domed stromatolites variably
broken into intraclasts. (B) Stained thin section in transmitted light of stromatolite microstructure with micrite
‘M’, microspar ‘S’, fenestral ‘F’ and detrital ‘D’ laminae. (C) Stained thin section in transmitted light showing the
broken edges of stromatolitic intraclasts with radiaxial calcite cement crusts in a calcareous sandstone matrix. (D)
Cross-stratified sandstone (set 30 cm high) with stromatolite intraclasts overlain by current ripple forms. (E) Pol-
ished rock slice with arrow denoting micromilled traverse shown in (F). (F) Micromill isotope traverse across two
micrite/microspar lamina and a central zone of calcite spar which has a much lower isotope signature. (G) Sand-
stone body (with pebbly base) showing accretionary surfaces (for example, dashed line). Palaeohorizontal shown
by solid black line. (W2, Ditlovtoppen, 119 m; tape has markings at 10 cm intervals). (H) Photomontage of tabular
sandstone unit of FA2 with a pebble horizon near its top (arrowed). It rests erosively on floodplain (FA3) silts and
is overlain by red lake margin (FA4) sediments; ruler is 25 cm long (W2, Dracoisen, 83 m).
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
462 Ian J. Fairchild et al.
–14
–12
–10
–8
–6
–4
–2
0
0
1
2
3
4
5
13
C (left axis)
18
O
(right
axis)
A B
C D
E
G
F
H
Spar
FA2
FA3
FA4
3 mm
1 mm
0·5 mm
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 463
Richter, 2007). Furthermore a non-marine envi-
ronment is consistent with the relatively low Sr
content of the calcite (Fairchild & Baker, 2012).
Hence, the radiaxial calcite is interpreted here as
the pristine original low-Mg calcite phase.
Micrite and microspar fabrics have similar che-
mistries and are also considered to reflect deposi-
tional conditions. Following the same arguments
made for stable isotopes, the implication is that
the Mn content of FA2 calcites is also primary
and reflects low contemporary atmospheric PO
2
(cf. Hood & Wallace, 2014), but not anoxic condi-
tions under which sulphate reduction would
have occurred, disturbing the D
17
O systematics
and producing sulphides (Bao et al., 2009).
Importantly, the Mg composition of Facies 2S
limestone is similar to modern speleothem depo-
sits in a cool Scottish cave depositing from waters
with Mg/Ca controlled by dolomite dissolution
(Fairchild et al., 2001). The lamina thickness of
the stromatolites is very similar to those of mod-
ern fluvial microbial tufas which are similarly
complex deposits containing both biologically
mediated and inorganic precipitates (Andrews &
Brasier, 2005; Andrews, 2006). Interestingly, the
lamination of these modern deposits reflects a
strong annual variation in discharge; Facies 2S
likewise possesses physical sedimentological
characteristics consistent with those of ephem-
eral streams. In this interpretation, the stable
isotope compositions, which are similar for radi-
axial and microsparry calcites, can also be inter-
preted as primary, in which case the isotopic
covariation (Fig. 6B) reflects an evolutionary
trend towards more evaporated equilibrated solu-
tions (Talbot, 1990), rather than variations associ-
ated with recrystallization. Modern Antarctic
streams lack the calcite mineralization, but micro-
bial mats are well-developed and are adapted to
ephemeral flow conditions, readily reactivating
even after being dry for many years (McKnight
et al., 2007). The fluvial interpretation will be
developed later in the light of the relationship of
FA2 to other facies associations.
Facies Association 3 (Facies 3D and 3S)
Description
Facies 3D is marked by discrete zones of pro-
nounced dolomite cementation within a sand-
stone or siltstone, in some cases pervasive.
Where dolomite is most abundant, detritus floats
in a displacive mass of dolomite crystals
(Fig. 10B, D, E and F), but other dolomite-
cemented silty sands are still clast-supported
(Fig. 10C). Locally, distinct dolomite-cemented
nodules are visible (Fig. 10A) or a structureless
dolomite bed occurs with a low content of float-
ing silt and sand. The most characteristic struc-
tures are millimetre-scale nodular dolomicrite
structures within massive dolomite-cemented
layers and associated with calcite-filled frac-
tures. These phenomena are found at one hori-
zon in member W3 (Fig. 10D), as well as in
A
500 µm
MS
M
R
V
1 mm
B
Fig. 9. Photomicrographs of stromatolitic limestones from FA2. (A) Paired transmitted light (left) and CL (right)
micrographs. Stromatolitic laminae of orange-luminescing microspar (MS) with subhedral authigenic quartz and
clastic layer (M) including quartz and feldspar grains (the latter luminesces dark blue). Large fenestra is filled by
radiaxial calcite (R) seen in both transverse and basal sections and displaying brighter earlier growth and duller
later growth. These fabrics are cut by a vein (V) filled with bright to dull luminescing calcite. W2, Dracoisen,
585 m. (B) Stained thin section, transmitted light. Alternating micrite, microspar and clastic laminae with promi-
nent irregular vertical ‘filamentous’ structure of clear calcite. W2, Dracoisen, 585m.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
464 Ian J. Fairchild et al.
several locations in member W2 (for example,
Fig. 10F). A rarer phenomenon is the presence
of equant centimetre-scale cauliflower-shaped
pseudomorphs, filled by ferroan saddle dolomite
cement (Fig. 11A) occurring at the top of a con-
glomerate-based fining-upward cycle (Fig. 8G).
Since saddle dolomite is a burial phase (Radke
& Mathis, 1980), the pseudomorphs must have
been occupied with soluble crystals that dis-
solved during burial prior to cementation.
Facies 3S refers to dolomitic laminites, with
broad centimetre-scale domed macrostructure,
with an aspect ratio of 10:1; they are found
uniquely in a single complex bed, in association
with Facies 3D, and overlying Facies 2S (at
70 m, Dracoisen, Fig. 4). It has been studied on
the ‘Multikolorfjellet’ cliffs and the ‘Tophatten’
nunatak 1 km to the north. The bed is around a
metre in thickness, and has an internally vari-
able structure. Most commonly, the base of the
bed shows minor erosion of underlying diamic-
tite and begins with crudely laminated, very fine
to medium-grained green sandstone, locally with
millimetre-scale limestone layers that are partly
disrupted into intraclasts (Facies 2S). In places,
the limestone passes upward into intensively
dolomite-cemented sand in which the rock fab-
ric appears to have expanded. In this sand dolo-
mite corrodes quartz detritus and there are
cavities lined with dolomicrospar and occluded
by calcite. At the ‘Tophatten’ locality, a chaotic
breccia unit a few decimetres thick is locally
found at the base of the bed instead of sand-
stone. Everywhere, the top of the bed is marked
by 10 to 20 cm of dololaminites with a complex
microstructure, which alternate on a centimetre-
scale with displacively cemented sands
(Fig. 10E). The laminae are variably composed
of dolomicrite or dolomicrospar and contain
common fenestrae (Fig. 9E and H), while weath-
ered surfaces reveal a finely textured micro-
topography (Fig. 10G). Locally, slightly lower in
the bed, limestone laminites (FA5) form a 10 cm
horizon overlying a 20 to 30 cm chaotic carbo-
nate breccia (Fig. 10I) and gradually become dis-
rupted downward.
Dolomite from FA3 is characteristically bright
under CL (Fig. 11B to D). In Facies 3S, dolomi-
crite clots are uniformly bright, while adjacent
dolomicrospar displays duller growth filling
small fenestrae, and larger fenestrae are filled by
dolospar with more variable properties
(Fig. 11B). Manganese (3000 to 4000 ppm), Fe
(10 000 to 15 000 ppm), Na (2000 ppm) and Sr
(250 to 350 ppm) concentrations are all unusu-
ally high (Bao et al., 2009), and unpublished
electron microsope images and microanalyses
from the present authors also show enrichments
in many transition metals and rare earths, as
well as a consistent chemical zonation within
crystals of dolomicrite mosaics. In Facies 3D, a
difference in mean CL brightness, and hence
timing of growth, is observed between some
nodules and surrounding matrix (Fig. 11D),
while zonation within individual crystals grow-
ing between siliciclastic sand grains is observed
locally (Fig. 11C). Sulphate is enriched
(4000 ppm); there is no D
17
O anomaly but sul-
phate d
18
O is exceptionally high (Bao et al.,
2009; Benn et al., 2015).
Carbon and oxygen isotope values are corre-
lated in Facies 3D (Fig. 6), but Fig. 12 illus-
trates that the two analyses from member W3
lie about 1&higher in d
13
C than expected from
this trend. Facies 3D has a range of d
18
O from
19to+114&, but the mean value is biased
upward by multiple analyses from a single
sample which passes up into Facies 3S with
even higher values (Fig. 12). Facies 3S is nota-
ble for possessing possibly the heaviest oxygen
isotope values of carbonate rocks so far
recorded in the geological record (Bao et al.,
2009), with values up to +147&(VPDB;
Fig. 12). A micromill traverse (Fig. 10H and J)
demonstrates that these extreme high values
are maintained on the millimetre-scale, but that
over petrographic boundaries, d
18
O values can
vary by as much as 6&.
Interpretation
For Facies 3D, the dolomicritic, syndepositional,
passive to displacive growth with nodular struc-
ture and cracks is characteristic of calcretes in
which precipitation is driven by evaporative
concentration at or above a water table.
Although rare, the spar-filled pseudomorphs
(Fig. 11A), interpreted to have originated as
anhydrite (Fairchild et al., 1989), attest to eva-
porative conditions. The displacive growth, nod-
ules and cracks identify these carbonates as
alpha calcretes (Wright, 1990), which in
Phanerozoic examples tend to occur on non-car-
bonate substrates and in more arid conditions
than the more common beta calcretes that are
influenced by higher plants (Wright & Tucker,
2009). The strongly positive d
18
O values and
covariation with carbon isotopes require evapo-
ration (Fairchild et al., 1989), which at the
higher end of the spectrum necessitates an extre-
mely arid environment. Dolocretes are less com-
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 465
mon than calcareous calcretes and tend to be
better developed when originating from ground-
water than when pedogenic, as in Triassic strata
of the Paris Basin (Sp
otl & Wright, 1992). In this
example, pedogenic and groundwater dolocretes
had a similar range of stable isotope composi-
tions to each other, but their covariance slope
(1:1) was steeper than in FA3. Overall, the
absence of any signal of light carbon from oxida-
tion of organic matter in FA3 is consistent with
the extremely low organic carbon contents of the
rocks. The Wilsonbreen Formation dolocretes
are interpreted as pedogenic, primarily because
the extremely high d
18
O values would require
ground surface conditions for such extremely
effective evaporation to occur. As will be dis-
cussed later, this interpretation is also consistent
with the vertical facies relationships.
The differentiated microstructures in Facies
3S are again typical of microbialites, as are the
distinctive surface textures (Fig. 10G; Callow
et al., 2011). Such laminites are found in asso-
ciation with soils and intermittently flooded
subaerial surfaces (Alonso-Zarza, 2003),
although younger examples include root mats
from higher plants that are clearly inapplicable
here. Klappa (1979) ascribed centimetre-scale
laminated deposits of ‘hard pan’ on calcretized
limestone substrates as originating from the
activities of lichen which colonize, bore into
and form accretionary deposits on surfaces. The
lichen-formed deposits exhibit features such as
fenestrae, sediment incorporation and variable
crystal size which are consistent with the Wil-
sonbreen Formation example. However, no evi-
dence of alteration of underlying cemented
material has been found, and the Wilsonbreen
Formation microbial laminae are more distinct
and are noticeably domed, in contrast with
laminar calcretes. In fact, the microbial layering
is indicative of active upward accretion, rather
than slow pedogenetic alteration. Accretion
occurred through both growth of carbonate-
mineralized microbial mats and deposition of
sand laminae. A shallow depression on a flood-
plain/playa margin is an apposite environment.
A combination of a high water table from
which evaporation could occur and very shal-
low-water inundation followed by drying out
and sediment addition, perhaps by aeolian
action, is indicated. At Dracoisen (Fig. 9I), the
gradational relationship between laminated
carbonate and underlying chaotic breccia is a
classic characteristic of evaporite dissolution
breccias. The calcite-cemented nature of the
breccia is consistent with removal of one or
more horizons of calcium sulphate evaporites
either during deposition of the bed or soon
afterwards following resumption of glacial con-
ditions. Brasier (2011) presented a similar
Mesoproterozoic example from Ontario in
which stromatolites, associated with collapse
breccias and calcretes, are inferred to form at a
playa lake margin. Likewise, the modern
McMurdo Dry Valleys contain a record of many
shallow saline lakes and salt pans (Wilson,
1981).
Dolomite is known to precipitate as a primary
phase or by replacement of a CaCO
3
precursor
in a range of surface environments (Warren,
Fig. 10. Facies Association 3 (Dolomitic floodplain). Facies 3D is shown in (A) to (D), Facies 3S in the other
panels, and both facies in (E). All of the Facies 3S images come from W2, Dracoisen, 70 m. (A) Nodular dolocrete
with calcite-lined vugs in siltstone with scale in millimetres (W2, East Andromedafjellet, 35 m). (B) Stained thin
section in plane polarized light showing matrix-supported fabric of dolomite cementation of sandy siltstone.
Dolomicrospar lines a fenestra which is occluded by calcite. W2, Dracoisen, 70 m. (C) Thin section in plane polar-
ized light. Grain-supported dolomicrite cement of silty sandstone. The dolomite has a d
18
O composition of +27&
and has a uniform texture in contrast to clastic dolomite of Fig. 5D. W2, South Klofjellet, 57 m. (D) Displacive
dolomite cement supporting silt and sand grains. Well-developed structure of dark nodules which show different
CL characteristics from surrounding dolomite from which they are separated by curved cracks. W3, South Ormen,
78 m. (E) Stained thin section of interlaminated dolomite-cemented sand and microbial laminae with fenestrae,
some occluded by ferroan dolomite (turquoise arrow) or ferroan calcite (purple arrow). (F) Stained thin section
illustrating similar fabric to (D), but with calcite cementation of cracks and larger pores (W2, Dracoisen, 83 m) (G)
Field photograph of textured bedding surface of dololaminite identified as microbial mat texture (W2, Dracoisen,
70 m). (H) Polished rock chip of microbial dololaminites with arrow marking position of 52 mm micromill tra-
verse [shown in (J) below]. (I) Microbial laminites draping downward into underlying laminate whose brecciation
is attributed to evaporite dissolution collapse. Outlined ruler is 20 cm long. (J) Stable isotope profile (in &with
respect to V-PDB) of microbial dololaminites along line illustrated in (H) The isotopes covary over a magnitude of
6&for d
18
O and 1&for d
13
C.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
466 Ian J. Fairchild et al.
B
200 μm
1 cm
D
C
A
H
G
F
E
IJ
1 mm
1 mm
100 μm
5 mm
7
8
9
10
11
12
13
14
15
3·2
3·4
3·6
3·8
4
4·2
4·4
4·6
4.8
5
d13C left axis
d18O righ t ax is
1 mm
δ
13C
δ
18O
δ13C le axis
δ18O right axis
J
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 467
2000), although the initial crystals (protodolo-
mite) may lack well-developed ordering reflec-
tions, which increase during ageing (Gregg
et al., 1992). The petrographic characteristics of
FA3 dolomicrite are consistent with very early
diagenetic replacement of a precursor carbonate
or of primary growth of (proto)dolomite; the lat-
ter is clearly the case for zoned dolomicrospar
cavity-linings (cf. Hood & Wallace, 2012). The
presence of euhedral crystals within displacive
fabrics (Fig. 11C) is distinctive. Although Tan-
don & Friend (1989) interpreted euhedral growth
zones in displacive calcite in calcretes as evi-
dence of recrystallization, it is more logically
interpreted as a primary growth fabric, as argued
by Sp
otl & Wright (1992) in a dolocrete
example.
The extremely high d
18
O values rule out post-
depositional modification. Interpretation of the
values as reflective of the depositional environ-
ment is also consistent with the trace element
chemistry and preserved crystal growth zones.
The high Mn content and absence of pyrite
implies low pO
2
, but not anoxia, although the
sulphate oxygen isotope systematics are indica-
tive of more redox variability than in Facies 2S.
A
2 mm
B
C D
2 mm
100 µm
0·5 mm
Fig. 11. FA3 (dolomitic floodplain). (A) Transmitted light, stained thin section. Sandy dolocrete (FA3) containing
equant nodule cemented by ferroan saddle dolomite (turquoise), with local late calcite (red), interpreted as a fill of a
small anhydrite nodule. W2, Backlundtoppen-Kvitfjellet ridge, 747 m. (B) Facies 3S stromatolite with fenestrae.
Paired transmitted light (left) and CL (right) images. Brightly luminescing dolomite may be primary or an early
replacement of a precursor. W2, Dracoisen 6995 m. (C) Paired transmitted light (left) and CL (right) images. Dolo-
crete showing very fine-grained quartz sand grains floating in dolo(micro-)spar with crystals displaying a common
zonation of bright to dull CL. Displacive primary dolomite growth is the preferred interpretation. W2, Ditlovtoppen,
1185 m. (D) Paired transmitted light (left) and CL (right) images. Dolocrete, similar to Fig. 10D and F, with sparse
floating quartz and feldspar (black and blue, respectively, in CL) and calcite-filled cracks (centre) and pores (base).
Uniformly luminescing dolomicrite crystals differ in brightness within nodules presumably forming at different
stages. Calcite-filled pores show CL zonation (base) or no CL (cracks, centre). W2, Dracoisen, 829m.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
468 Ian J. Fairchild et al.
Specifically, bacterially mediated electron shut-
tling by Mn-species can catalyse repeated transi-
tions between sulphate and sulphite, which
would erase any D
17
O anomaly and lead to high
d
18
O in sulphate (Bao et al., 2009). Such pro-
cesses could trigger dolomite nucleation given
the evidence from other field and experimental
studies on the catalytic role of sulphate reduc-
tion (Vasconcelos et al., 2005; Zhang et al.,
2012). The inferred redox variations may be
related to a supply of brine primarily from
within the sediment, contrasting with the sur-
face waters from which Facies 2S precipitated.
The occurrence of the highest d
18
O values in
laminated dolomites of Facies 3S is consistent
with their formation by very near-surface eva-
poration, whereas abrupt variations in d
18
O
(Fig. 10J) imply occasional inundations by less
evolved waters. In summary, FA3 provides
examples of facies that stretch the boundaries of
earth surface phenomena and indicate deposi-
tion in unusually arid terrestrial environments.
Facies Association 4 (Facies 4I, 4R and 4S)
Description
Facies 4R is the most common facies in this
association and consists of rhythmic alterna-
tions of carbonate and sorted terrigenous sedi-
ment, which occur in association with
structures implying shallow-water to intermit-
tently emergent conditions, such as wave-ripple
lamination and desiccation structures. The fine
carbonate layers are usually dolomitic, or
mixed dolomitic–calcitic, but include some
limestone (Figs 7A and 12). The coarser sedi-
ment layers are universally composed of coarse
silt to fine sand with evidence of tractional
sorting, which distinguishes this facies associa-
tion from FA5. Wherever laminae are suffi-
ciently thick, undulatory cross-lamination is
displayed (Fig. 13B) which can be confidently
identified as wave-generated. Locally, symme-
trical ripples are preserved in cross-section
(Fig. 13A) or on bedding planes (Fig. 13C). Dry-
-4
-3
–2
–1
0
1
2
3
4
5
6
15 –10 –5 0 5 10 15
13
C
V-PDB
(‰)
18
O
V-PDB
(‰)
Facies 3S dolomite, member W2 Facies 3D dolomite, member W2
FA4 calcite, member W2 FA4 dolomite, member W2
FA5 calcite, member W2 FA5 dolomite, member W2
Facies 3D dolomite, member W3 FA4 calcite, member W3
FA5 calcite, member W3 FA5 dolomite, member W3
Fig. 12. Stable isotope plot, differentiating facies within FA3 and (in purple and larger symbols) samples from
member W3.
©2015 The Authors. Sedimentology ©2015 International Association of Sedimentologists, Sedimentology,63, 443–497
Continental glacial carbonates 469
ing out is commonly indicated by desiccation
structures with associated small intraclasts
(Fig. 13B and H) or salt pseudomorphs
(Fig. 13D), although such structures are not
present in the majority of samples. Four exam-
ples of apparently non-evaporitic crystal pseu-
domorphs have been found, but these are
much better developed in FA5 and are
described in that section. Carbonate laminae
are micritic in texture and typically uniform,
although differentiated clotted microstructures
also occur, similar to those described below in
FA5, consistent with precipitation beneath ben-
thic microbial mats (Riding, 2000). This facies
was locally highly affected by subsequent glaci-
tectonic deformation at the top of W2 at Ditlov-
toppen, as described by Fleming et al. (2016).
The isotope traverse of Fig. 13J reveals a shift
in isotopes from the sandy layers (with detrital
dolomite) into dolomicrite, implying that the lat-
ter is authigenic, as confirmed by CL observa-
tions (Fig. 14A and B). The limestone laminae in
this facies association display a range of d
18
O
values from 119to32&with a mean of
8