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11
Chemical Geology 237 (2007) 23 – 45
www.elsevier.com/locate/chemgeo
Stratigraphy and geochemistry of a ca 800 Ma negative carbon
isotope interval in northeastern Svalbard
Galen P. Halverson a,⁎, Adam C. Maloof b , Daniel P. Schrag c ,
Francis Ö. Dudás d , Matthew Hurtgen e
a
UMR 5563, LMTG-IRD-CNRS, Université Toulouse-3, 14 avenue Edouard Belin, 31400 Toulouse, France
b
Department of Geosciences, Princeton University, Guyot Hall, Princeton, NJ 08544, USA
c
Department of Earth and Planetary Sciences, Harvard University, 20 Oxford St., Cambridge, MA 02138, USA
d
Department of Earth, Atmospheric, and Planetary Sciences, MIT, Building 54-1124, Cambridge, MA 02139, USA
e
Department of Geological Sciences, Northwestern University, 1850 Campus Dr., Evanston, IL 60208, USA
Accepted 8 June 2006
Editor: P. Dienes
Abstract
The Neoproterozoic Akademikerbreen Group in northeastern Svalbard comprises 2 km of nearly pure carbonate section. The
carbonates are generally highly 13C-enriched (δ13C N 5‰), but this trend is interrupted by an ∼ 325 m interval of low δ13C values
(−4 to 0‰) in the upper Grusdievbreen and lower Svanbergfjellet formations. An abrupt negative isotopic shift at the onset of this
low δ13C interval is reproduced in detail in multiple sections along the length of the outcrop belt (125 km) and everywhere
coincides with a prominent sequence boundary and change in lithology. Likewise, the return to positive δ13C values coincides with
a second exposure surface. Correlation of the lower Akademikerbreen Group δ13C record with a nearly identical isotopic profile in
the Bitter Springs Formation of Central Australia suggests an age of ∼ 800 Ma for the low δ13C interval and confirms that it is a
global seawater signal. The coincidence of the negative and positive δ13C shifts with major stratigraphic perturbations in the
otherwise conformable succession suggests that both episodes of transient sea level change were related to global phenomena.
87
Sr/86Sr ratios rise transiently from an average of 0.7063 to 0.7066 within this interval. Whereas large negative δ13C anomalies in
the Neoproterozoic are commonly associated with episodes of widespread glaciation, the Akademikerbreen low δ13C interval
precedes the oldest (Sturtian) of the known Neoproterozoic glacial events, and no other evidence suggests an ice age at this time.
We propose instead that the negative δ13C interval is related to a pair of inertial interchange true polar wander (TPW) events.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Neoproterozoic; Svalbard; Chemostratigraphy; δ13C;
87
Sr/86Sr; TPW
1. Introduction
The carbon-isotopic composition of the Neoproterozoic oceans was generally high, but fluctuated by at
⁎ Corresponding author.
E-mail address: [email protected] (G.P. Halverson).
0009-2541/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.chemgeo.2006.06.013
least 15‰ (Shields and Veizer, 2002) coincident with
episodes of widespread glaciation (Knoll et al., 1986;
Kaufman and Knoll, 1995; Kaufman et al., 1991, 1997).
The highly 13C-enriched intervals typical of interglacial
times have received only minor attention, but it is
widely assumed that the enrichment of the oceans was a
result of unusually efficient burial of organic matter in
24
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
Neoproterozoic ocean basins, perhaps as a consequence
of high erosion rates related to widespread orogenesis
and abundant nutrient availability (Derry et al., 1992;
Kaufman et al., 1995). The negative δ13C anomalies
have received much greater attention, due in large part to
their intimate relationship to glaciation (Hoffman et al.,
1998a,b). Three Neoproterozoic glaciations and associated δ13C anomalies have been identified (Halverson et
al., 2005; Xiao et al., 2004): the oldest (Sturtian) postdates ∼ 750 Ma and is perhaps as young as ∼ 710 Ma
(Brasier et al., 2000), the middle (Marinoan) glaciation
likely ended ∼ 635 Ma (Hoffmann et al., 2004) and the
youngest (Gaskiers) was short-lived and occurred at
∼ 580 Ma (Bowring et al., 2003; Knoll et al., 2004). The
negative δ13C anomalies are best known from cap–
carbonate sequences to the Sturtian and Marinoan
glaciations (Williams, 1979; Kennedy, 1996; Kaufman
et al., 1997; Kennedy et al., 1998; Hoffman and Schrag,
2002) and had long been regarded as solely post-glacial
phenomena. This assumption underlies popular causeand-effect hypotheses for the presumed linkage between
climate change and δ13C patterns, but it is now apparent
that the negative δ 13 C anomalies are variable in
magnitude and pattern for each of the Neoproterozoic
glacial episodes (Halverson et al., 2005). Furthermore,
at least one glaciation is preceded by a global negative
δ13C anomaly (Hoffman et al., 1998a; McKirdy et al.,
2001; Halverson et al., 2002), implying a complex
relationship between global climate and δ13C in the
Neoproterozoic (Schrag et al., 2002).
The absence of a simple cause-and-effect relationship
between negative δ13C excursions and glaciation raises
the possibility that carbon isotope anomalies could occur
independent of glaciation. Hill et al. (2000a,b) found
evidence for a decoupling between Neoproterozoic δ13C
anomalies and glaciation when they documented a salient
interval of low δ13C in pre-Sturtian (∼800 Ma) carbonates
(Walter et al., 1995; Hill and Walter, 2000) of the Bitter
Springs Formation in central Australia. We have documented what we interpret to be the same low δ13C interval
in the Akademikerbreen Group in northeastern Svalbard
where it is defined by a sharp decline of ∼8‰ in the
middle Grusdievbreen Formation and a rise of similar
magnitude in the lower Svanbergfjellet Formation. Both
isotope shifts are associated with regionally persistent,
conformable exposure surfaces and appear to represent
switches between contrasting steady state conditions of
carbon cycling, related to global events that caused
transient fluctuations in sea level. In this paper we present
detailed stratigraphic, isotopic, and elemental data
spanning the Grusdievbreen and Svanbergfjellet formations to constrain possible geological processes that may
11
have contributed to the carbon isotopic and stratigraphic
patterns spanning the low δ13C interval in Svalbard.
2. Geological background
The Svalbard archipelago consists of three tectonic
terranes that were juxtaposed during the Silurian–
Devonian Ny Friesland orogeny (Harland and Gayer,
1972; Harland et al., 1992; Gee and Page, 1994; Lyberis
and Manby, 1999). The thick, Neoproterozoic–Ordovician
middle–upper Hecla Hoek Succession (Fig. 1) underlies
the northern part of the eastern terrane and is nearly
identical to the Neoproterozoic succession in the southern
East Greenland Caledonides, suggesting that both were
deposited in a contiguous basin (East Greenland–East
Svalbard — EGES platform) that was dissected during
the Caledonian orogeny (Harland and Gayer, 1972).
The origin of the EGES platform remains obscure, but
presumably it was associated with the fragmentation of
Rodinia, which in this region involved the separation of
four cratons including Laurentia, Baltica, and Amazonia
(Torsvik et al., 1996; Hartz and Torsvik, 2002). Paleomagnetic data indicate that the EGES platform remained
in tropical latitudes throughout deposition of the Akademikerbreen Group (Maloof et al., 2006).
Fig. 1. Generalized lithostratigraphy of the Neoproterozoic successions in northeast Svalbard after Harland et al. (1993), Fairchild and
Hambrey (1984), Knoll and Swett (1990), Halverson et al. (2005). See
text for discussion of age constraints (shown in ovals).
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
25
The north–south trending outcrop belt of Neoproterozoic rocks comprises up to 7 km of strata, traceable
for 120 km from northern Spitsbergen to western
Nordaustlandet in northeast Svalbard (Fig. 2). A
different stratigraphic nomenclature is used in Spitsbergen and Nordaustlandet, but due to the unambiguous
correlation of the strata between the two islands
(Fairchild and Hambrey, 1995) and for the sake of
simplicity, only the nomenclature from Spitsbergen is
used here (Fig. 1). The ∼ 4 km-thick Veteranen Group
forms the base of the Neoproterozoic succession and
consists predominantly of siliciclastic sediments, but
contains three distinct carbonate units (Wilson, 1958;
Knoll et al., 1986). The Veteranen Group is constrained
to be younger than ∼ 950 Ma based on U–Pb dates from
magmatic zircons in basement granites in the Nordaustlandet region (Gee et al., 1995; Johannson et al., 2000)
and detrital zircons in the Planetfjella Group in Ny
Friesland (Larianov et al., 1998).
The Veteranen Group siliciclastic sediments are
transitional with the ∼ 2 km of nearly pure carbonate
section of the Akademikerbreen Group. These carbonates comprises the Grusdievbreen, Svanbergfjellet,
Draken, and Backlundtoppen formations, all of which
are conformable across the length of the outcrop belt
(Fig. 2) and generally thin from south to north (Wilson,
1961), indicating that the cratonic margin lay north of
the EGES basin (in present Svalbard coordinates) at the
time of deposition. The Grusdievbreen Formation
exceeds 600 m in thickness and is separated into two
informal members (upper and lower) at one of only two
major sequence boundaries in the Akademikerbreen
Group. Whereas mid-shelf limestone rhythmites dominate the Grusdievbreen lithology, the overlying
Svanbergfjellet contains abundant microbialaminites
and stromatolites in additional to limestone ribbon
rock. The Svanbergfjellet Formation is ∼ 600 m thick in
the south (Olav V Land), but thins dramatically to the
north (Wilson, 1961). It is separated into four informal
members (Wilson, 1961; Knoll and Swett, 1990), the
lower two of which are separated by a second sequence
boundary (Fig. 3). The ∼ 200 m-thick Draken Formation consists predominantly of dolomitic intraformational conglomerates, microbialaminites, and
grainstones deposited in a tidal flat/lagoonal setting
Fig. 2. Geological map of the north–south trending Neoproterozoic
outcrop belt in northeastern Svalbard: Numbered boxes key stratigraphic
sections measured and sampled for this study. Location names: 1)
Svanbergfjellet, 2) Golitsynfjellet (north and south), 3) Dracoisen, 4)
Raudberget, 5) Glintdalen, 6) Faksegåven, 7) Murchisonfjord. ESZ =
Eolussletta Shear Zone; LFZ = Lomfjorden Fault Zone.
26
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
Fig. 3. Composite stratigraphic section and δ13C data (closed circles) through the Grusdievbreen and Svanbergfjellet formations at measured section 7
(Murchisonfjord). Open circles are δ13C data from measured section 6 to fill in limited data coverage through poorly exposed section at measured section 7.
(Knoll et al., 1991). The upper Draken Formation is
gradational with the overlying Backlundtoppen Formation, which is ∼ 500 m-thick and comprises black
limestone ribbon rock, oolitic and pisolitic grainstone,
and pale gray dolomite clastic grainstone and stromatolites (Wilson, 1961; Knoll et al., 1989). An influx of
silt and fine sand into the EGES basin during
deposition of the uppermost Akademikerbreen Group
(Kinnvikka Member) marks the end of the long-lived
carbonate platform (Fairchild and Hambrey, 1995;
Halverson et al., 2004).
The overlying mixed carbonate-siliciclastic Polarisbreen Group contains two separate diamictites units.
Halverson et al. (2004) interpreted these glacials as
representing the beginning and end of a Marinoan
snowball cycle (Fig. 1), and the lowermost Polarisbreen
Group (Russøya Member of the Elbobreen Formation)
as the chronostratigraphic equivalent of the Sturtian
cap-carbonate sequence (Halverson et al., 2005). If this
hypothesis is correct, it implies a minimum age of
~ 700 Ma for the top of the AkademikerbreenPolarisbreen contact (cf. Brasier et al., 2000, though
note the timing of the Sturtian glaciation in other
successions remains poorly resolved). Alternatively,
the glacigenic Petrovbreen Member (Fig. 1) may
represent the Sturtian glaciation (e.g. Kennedy et al.,
11
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
27
Table 1
Isotopic and elemental data for 28 samples on which 87Sr/86Sr was determined from the Grusdievbreen and Svanbergfjellet formations
Sample
Formation Scaled stratigraphic height
87
G341.464.5
G341.445.3
G19.71.8
G19.57.2
G19.25.7
G341.300
G341.294.5
G341.282.5
G341.251.6
G341.215.6
G341.205.1
G341.178
M9.125.6
G341.146.6
M9.83.4
M9.55.0
G33.145.9
G148.158.7
G33.131.7
G148.149.2
G148.140.3
G33.112.1
G148.136.3
G148.126
G33.91
G33.62.2
G33.14.5
G336.281.7
Svan.
Svan.
Svan.
Svan.
Svan.
Svan.
Svan.
Svan.
Svan.
Svan.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
Grus.
0.70659
0.70661
0.70625
0.70629
0.70758
0.70681
0.70698
0.70635
0.70673
0.70634
0.70644
0.70642
0.70660
0.70637
0.70641
0.70640
0.70643
0.70635
0.70644
0.70640
0.70699
0.70884
0.70624
0.70625
0.70652
0.70628
0.70626
0.70627
917.3
898.1
851.8
843.5
805.7
752.8
747.3
735.3
704.4
668.4
657.9
630.8
617.3
599.4
575.1
546.7
488.9
476.8
474.7
467.3
458.4
455.1
454.4
444.1
434
405.2
357.5
281.7
Sr/86Sr δ13C
δ18O
Mg/Ca (mol/mol) Mn/Sr (mol/mol) Mn (ppm) Sr (ppm)
4.88 − 6.32 0.799
4.34 − 6.08 0.016
5.95 − 6.35 0.075
5.93 − 6.79 0.028
2.73 − 6.91 0.031
−2.98 − 5.05 0.276
−1.97 − 4.65 1.092
−0.02 − 5.53 0.321
−1.40 − 6.25 0.187
−0.70 − 6.90 0.015
−2.12 − 7.28 0.019
−0.31 − 7.26 0.025
0.57 − 8.49 0.009
0.08 − 6.92 0.034
−0.31 − 8.51 0.007
−0.75 − 9.17 0.015
−1.09 − 8.32 0.018
−1.25 − 8.14 0.017
−1.01 − 8.33 0.018
−1.46 − 8.40 0.024
−0.22 −9.85 0.010
−0.44 −11.75 0.020
0.33 − 8.85 0.011
6.42 − 7.01 0.024
5.79 − 10.60 0.021
7.27 − 7.07 0.009
6.55 − 6.26 0.026
6.97 − 5.67 0.010
1.598
0.046
0.116
0.099
1.050
0.338
0.796
0.113
0.322
0.030
0.123
0.027
0.006
0.020
0.004
0.002
0.007
0.029
0.041
0.121
0.677
1.249
0.16
0.012
0.022
0.006
0.006
0.012
105.8
29.6
72.0
60.0
77.6
49.3
63.3
38.1
38.0
17.5
28.4
13.8
4.3
18.4
5.2
2.8
5.2
12.5
13.7
45.4
40.9
104.7
28.6
13.2
7.2
6.3
9.7
17.1
105.6
1025.5
993.5
964.0
117.9
232.7
126.9
536.3
188.4
930.2
368.3
805.9
1241.8
1481.4
2029.7
2283.0
1236.9
695.9
594.1
271.3
96.3
133.7
281.8
1719.9
513.6
1692.4
2525.5
2278.5
Stratigraphic heights are all scaled to the composite section from measured section 7 (Fig. 3). 87Sr/86Sr data in italics are regarded as secondary based
on Sr concentrations. See Supplementary materials for a complete tabulation of isotopic and elemental data.
1998), in which case the top of the Akademikerbreen
Group is N 700 Ma. An inferred cryptic unconformity
beneath the Cambrian Tokammane Formation of the
Oslobreen Group truncates the Polarisbreen Group
(Knoll and Swett, 1987), and the absence of Ediacaran
fossils beneath this contact (Knoll and Swett, 1987)
suggests a minimum age of 575 Ma (Bowring et al.,
2003) for the top of the Neoproterozoic succession in
Svalbard (Fig. 1).
Previous studies revealed a series of δ13C excursions
in the Neoproterozoic succession of Svalbard (Knoll et
al., 1986; Derry et al., 1992; Kaufman et al., 1997);
negative δ13C anomalies in the Polarisbreen Group are
clearly associated with glacial episodes (Knoll et al.,
1986; Fairchild and Spiro, 1987; Fairchild et al., 1989;
Kaufman et al., 1997; Halverson et al., 2004). Carbonates in the Akademikerbreen Group were shown to be
generally very 13C-enriched, with the exception of
negative δ13C values in the upper Grusdievbreen and
lower Svanbergfjellet formations (Knoll et al., 1986;
Kaufman et al., 1997), which are the focus of this
paper.
3. Materials and methods
3.1. Sample collection and preparation
All samples used in this study were collected in the
course of measuring stratigraphic sections. Samples were
chosen so as to minimize veining, fractures, and
weathered surfaces. All samples were slabbed, polished
and subsampled with 1–5 mm dental drill bits. Splits of
sample powder were used for all geochemical analyses
described below.
3.2. Stable Isotopes
δ13C and δ18O isotope data (VPDB) were acquired
on a VG Optima dual inlet mass spectrometer attached to
a VG Isocarb preparation device in the Harvard
University Laboratory for Geochemical Oceanography.
External error (1σ) better than ± 0.1‰ was achieved for
both δ13C and δ18O. See Halverson et al. (2004) for a
description of methods. See Supplementary materials for
tabulated data.
Fig. 4. Stratigraphic, δ18O, and δ13C profiles spanning the G1 sequence boundary in the Grusdievbreen Formation in multiple sections. The 0-m datum is the G1 sequence boundary, which is easily identified
in section along the length of the outcrop belt (Fig. 5a). The second column from the left is a composite, with the upper Grusdievbreen Formation derived from measured section 5, and the lower
Grusdievbreen Formation from measured section 6.
28
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
29
Fig. 5. Field photographs of the Grusdievbreen and Svanbergfjellet formations. a) G1 sequence boundary (white line) separating the upper and lower
members of the Grusdievbreen Formation at measured section 3b. Note the relief on this contact (2 m), and the sharp change in lithology in the
parasequence above the boundary. b) Karst pipe consisting of heavily recrystallized dolomite packstone penetrating limestone grainstone ∼10 m
beneath the G1 contact at measured section 6. c) Edgewise limestone breccias within the silty member above the G1 sequence boundary. Note the
breccia bundles are continuous with coherent layers from which the clasts are derived. d) Aragonite pseudomorphs (“crystal fans”) forming a
continuous seafloor cement in limestone ∼20 m above the G1 boundary at measured section 5. The crystal fans occur along bedding surfaces and the
space between them is filled by onlapping micrite, indicating that they are primary depositional features. e) The S1 sequence boundary at measured
section 1 (the type section of the Svanbergfjellet Formation). The hammer, for scale (outline in white circle), is lying on the brecciated, silicified, and
ferruginized zone beneath black shales of the Middle Limestone Member. Despite heavy brecciation and silicification of the exposure surface, no
evidence of erosional relief is apparent on outcrop scale. f) Columnar Minjaria stromatolites in the Lower Limestone Member form a distinctive
biostrome 10–15 m above the S1 boundary throughout northeast Svalbard (Knoll and Swett, 1990). The hammer and pen (for scale in b, c, e, and f )
are 13 cm and 33 cm long, respectively.
3.3. Elemental chemistry
Elemental analyses (Ca, Mg, Sr, Mn, Fe) were
performed on a Jovin Yvon 46P ICP-AES mass
spectrometer in the Harvard University Laboratory for
Geochemical Oceanography. All samples were prepared
by dissolving ∼4 mg of carbonate powder in ∼4 ml of
2% nitric acid. SCP multielement and single element
standards were used for element-specific calibration at the
beginning of each run. External error (1σ), determined by
repeat analyses, was b 7% for all elements, with the best
precision achieved for Ca, Mg, and Sr.
30
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
3.4. Strontium isotopes
New 87Sr/86Sr data presented in Table 1 were
acquired at the MIT Radiogenic Isotope Laboratory.
5–10 mg aliquots of powdered carbonate were leached
sequentially 3–5 times for 15–45 min in 0.2 M
ammonium acetate to remove loosely bound Sr cations
(cf. Bailey et al., 2000). Carbonate was dissolved in
0.5 N acetic acid and insoluble residue was removed by
centrifuging. Sr was isolated via standard column
chemistry techniques and analyzed on a VG Sector 54
(TIMS). All samples are referred to a standard value of
0.710250 for NBS 987, whose long-term average value
on the Sector 54 is 7.10246, with 2-σ internal precision
of 0.000016.
4. Stratigraphy and geochemistry of the lower
Akademikerbreen Group
4.1. Grusdievbreen Formation
The lower Akademikerbreen Group is well exposed at
measured sections 4, 5, 6, and 7. Fig. 3 presents a composite section from measured section 7 (Murchisonfjord),
where we have the most detailed stratigraphic and isotopic
coverage. We use this section as a reference for discussion
of the stratigraphy and carbon isotope record through the
Grusdievbreen and Svanbergfjellet formations.
The Grusdievbreen Formation is ∼600 m thick,
relatively uniform across the length of the outcrop belt,
transitional with the underlying Veteranen Group (Wilson, 1961), and informally separated here into an upper
and lower member (Fig. 3). The base of the Grusdievbreen Formation is identified by the first dolomite-rich beds
above finely laminated, hummocky cross stratified
siltstones of the Oxfordbreen Formation (Wilson, 1961).
The lower part of the Lower Grusdievbreen Member
consists of intercalated siltstone, nodular (dolomite and
limestone) rhythmites ribbon rock (Knoll and Swett,
1990), and minor isolated stromatolite mounds. Sampling
through this interval is sparse due to the abundance of silt
and poor exposure in some sections, but available data
indicate δ13 C values near 1‰ at the base of the
Grusdievbreen Formation and variable and more 13Cenriched values through the succeeding 200 m (Fig. 3).
A colorful 10 m-thick unit of orange and pinkweathering intraformational breccia with interbedded
green and red shale overlies a flooding surface in the
middle of the Lower Grusdievbreen Member and
coincides with a spike to δ13C values as low as 1.7‰
(Fig. 3). δ13C values rise gradually and smoothly up to
8‰ through ∼ 50 m of strata above the flooding surface,
11
then remain high through most of the remaining 200+ m
of the Lower Grusdievbreen Member, which consists
almost exclusively of monotonous mid-shelf black
limestone ribbon rock (Knoll and Swett, 1990) with
interbedded storm-generated intraclast breccias and
minor grainstone. δ13C values begin to decline just
below a prominent sequence boundary (G1) that
separates the upper and lower members at ∼450 m
above the base of the Grusdievbreen Formation (Fig. 1).
The G1 sequence boundary and bracketing stratigraphy, discussed in more detail below, coincide with a sharp
decline in δ13C (Fig. 4). The lower 20 to 40 m above the
G1 boundary comprise a brick-red, upward-shoaling
parasequence that stands out amidst the background
stratigraphy (Fig. 5). δ13C values remain low (b0.5‰)
through the remainder of the Grusdievbreen Formation.
Limestone ribbon rock and clastic grainstones dominate
the lithology of the Upper Grusdievbreen Member, but
stromatolite bioherms are also locally abundant. Molar
tooth structures are common in the ribbon facies and
molar tooth breccias form the base of some ribbon rock
beds. Styolites are common throughout the Upper
Grusdievbreen Member and demonstrate significant
compaction of the sediment column.
4.1.1. The G1 boundary
TheG1sequenceboundaryisbest exposedontheeastern
side of the Golitsynfjellet nunatak, alongside the AkademikerbreenGlacierinOlavVLand(Fig.2),wheretwosections
(3a,b) 2 km apart have been measured (Fig. 5a). There, the
uppermost 1.3 m beneath the G1 sequence boundary is a
coarse, dolomitic intraformational conglomerate. Erosional relief developed on the exposure surface is visible on the
outcrop and cuts down as much as 2 m, in places removing
the dolomitic conglomerate bed (Fig. 5a) and demonstrating that at least one phase of dolomitization preceded
erosional truncation. To the north, in the Lomfjord region
(Fig. 2, measured sections 5 and 6), a heavily recrystallized, yellow, vuggy dolomitic packstone directly underlies
the G1 boundary, forms a bed up to 4 m thick, and fills karst
pipes that developed in the underlying limestones. At
measured section 6, the karst is pervasive just below the
boundary and individual pipes in contact with the exposure
surface extend as much as 10 m deep. Grikes up to 0.5 m
wide are developed along joints and occur as much as 20 m
below the G1 surface (Fig. 5b). Since dolomitization of the
packstone must have occurred after the development of G1
exposure surface in this region, it appears that dolomitization of the Akademikerbreen platform was ongoing
spanning the sequence boundary. No evidence of
karstification or erosion related to the G1 sequence
boundary has been identified around Murchisonfjord,
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
although it should be noted that this surface is not well
exposed in this region.
At Golitsynfjellet, a bed of green silt overlies the
sequence boundary and varies in thickness up to 2 m,
filling erosional relief on the G1 sequence boundary. The
green silt passes transitionally into red silt, which is
generally finer grained but contains coarse-grained lenses
and scattered carbonate intraclasts. Scour surfaces and
hummocky cross stratification are common in the red silt
facies. The first coherent carbonate lenses appear at 3.8 m
above the boundary. These white to pink limestone lenses
alternate with silt layers, vary in thickness from 1 to 6 cm,
are discontinuous, and are intermittently brecciated,
suggesting early lithification on the seafloor. Large tabular
clasts from these beds are commonly bundled edgewise in
packages filling pockets scoured into underlying silt
layers and draped by red silt (Fig. 5c). In places, these
clasts are imbricated, whereas others bundles resemble
“beach rosettes,” (Fig. 5c) suggesting deposition under
oscillatory flow. Whereas edgewise conglomerates are
often associated with shallow-water conditions, such as a
beaches or tidal flats (Demicco and Hardie, 1994, p. 42),
the association of the breccias above the G1 boundary
with hummocky cross stratification suggests the lowermost carbonate beds were deposited between fair weather
and storm weather wave base and that storm waves may
have caused the disaggregation of the carbonate lenses.
Up section, the percentage of silt gradually decreases
and discrete silt layers disappear by ∼20 m above the
exposure surface. At the same level, crystal fan
pseudomorphs up to ∼3 cm in diameter occur within a
1 m-thick horizon. Individual crystals have square
terminations, indicating that the fans were originally
aragonitic. The crystal fan horizon is best developed at
measured section 5, where the fans are tightly packed in
continuous layers (Fig. 5d) and account for N 25% of total
sediment volume. The crystal fan layers are infilled and
draped by micrite, indicating that they grew directly on the
seafloor, analogous to seafloor cements in some Marinoan
cap–carbonate sequences (Hoffman and Schrag, 2002).
Above the crystal fan horizon, recrystallized, chalky
limestone ribbon rock passes upward into small (2–4 cm
diameter), columnar stromatolites, grainstones and
microbialaminites, up to a flooding surface that marks
the top of the basal parasequence of the upper member
of the Grusdievbreen Formation (Fig. 4). At measured
section 2, the parasequence is 47 m thick, whereas in
Murchisonfjord (measured section 7), it is only half as
thick. This difference in thickness is controlled, at least
partly, by the abundance of siliciclastic sediment in the
parasequence, which increases from north to south. This
pattern of southward thickening is at odds with the
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31
hypothesis that the paleo-cratonic margin lies northward
of the Svalbard outcrop belt and suggests either that the
Murchisonfjord region was a zone of sediments bypass
or that another siliciclastic sediment source existed to
the south or southwest. Despite the differences in
thickness, the parasequence above the G1 everywhere
stands out starkly against the background stratigraphy of
predominantly black limestone ribbon rock (Fig. 5a).
In all measured sections (Fig. 4), the δ13C trend is
identical across the G1 boundary, beginning to decline
from a high of 7‰ directly beneath the exposure surface,
then dropping sharply across the boundary, and gradually
declining to a low of ∼−1.5‰ roughly coincident with the
crystal fan horizon (Fig. 4). However, whereas δ13C values
in the lowermost limestone above the boundary in our two
Golitsynfjellet sections (3a,b) are already negative in the
northern two sections, δ13C values are slightly positive
directly above the sequence boundary and gradually
decline to b0‰ over the next 5 m. Allowing for
differences in sedimentation rates, the δ13C values across
the G1 boundary in all sections are otherwise identical
within 0.5‰ (Fig. 4). Therefore, the northern sections
appear to preserve a more complete record of the evolution
of marine δ13C values spanning the sequence boundary,
perhaps because the first carbonates to precipitate were not
as diluted by incoming silt.
Just as the δ13C trends in all sections are virtually
identical across the G1 boundary, so also are the patterns
in δ18O values reproduced in multiple sections (Fig. 4).
Beneath the exposure surface, δ18O values rise abruptly
from ∼− 11 to − 6‰. In the lowest limestones above the
surface, δ18O values are − 9 to − 12‰. δ18O values then
rise gradually over the next10 to 20 m of section to an
average of − 8‰ (Fig. 4), which is a typical value for
limestones in the Akademikerbreen Group (Derry et al.,
1992; Halverson, 2003). The perturbation in the
unusually smooth and reproducible δ18O trend is clearly
related to the G1 sequence boundary, but does it
represent a primary seawater signal?
The concentration of dolomite beds in the Grusdievbreen Formation at the G1 boundary suggests that
dolomitization was related to the sea level fall that
exposed the Akademikerbreen platform. Variable truncation of the dolomite unit at measured sections 2a,b
indicates that one phase of dolomitization must have
preceded development of the G1 sequence boundary. On
the other hand, at measured section 6, dolomitized
packstone fills karstic cavities, indicating that dolomitization also proceeded after the platform had been
subaerially exposed. Therefore, the similar increase in
δ18O values beneath the exposure surface in both
regions, despite the apparent difference in timing of
Fig. 6. Stratigraphic, δ18O, and δ13C profiles spanning the S1 sequence boundary in the Svanbergfjellet Formation from 5 measured sections. The 0-m datum is the S1 sequence boundary, which is
everywhere brecciated and silicified (Fig. 5e). Dashed line above the S1 boundary shows correlation of the flooding surface at the top of the Minjaria biostrome.
32
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
dolomitization, suggests that the positive upward trend
in δ18O values records a secondary rather than primary
seawater signal. A cross-plot of δ18O vs. Mg/Ca for all
samples spanning the sequence boundary shows a clear
positive, linear trend, the least-squares best fit of which
is the line δ18O = 3.77(Mg/Ca) − 10.3‰ (Fig. 7). The
3.77‰ difference between stoichiometric calcite (Mg/
Ca ≈ 0) and dolomite (Mg/Ca ≈ 1) is consistent with the
estimated equilibrium fractionation between these two
carbonate phases (Land, 1980; Kah, 2000) and suggests
that the positive deflection in δ18 O is purely an
equilibrium isotope effect, with the most 18O-enriched
values corresponding to most heavily dolomitized
samples.
The y-intercept (Mg/Ca = 0) of the Lower Grusdievbreen Member (LGM) dolomitization trend (−10.3‰) is
significantly lower than average δ18O values in the
Akademikerbreen Group but is only slightly more
negative than the average δ18O composition (−9.6‰) of
the best-preserved limestones (Mg/Ca, Mn/Sr b 0.1) in the
20 m beneath the G1 boundary (Fig. 8). In contrast, the
average δ18O value in the best-preserved limestones from
20 to 120 m beneath the G1 boundary is −7.4‰ (Fig. 8).
Therefore, rather than a trend of increasing δ18O directly
below the G1 boundary, as seen in the bulk data (Fig. 4),
δ18O data from the best preserved samples reveal a more
gradual decline of ∼2‰ that begins before the downward
shift in δ13C values (Fig. 8).
In all measured sections δ18O values directly above the
sequence boundary are deflected negative of the background trend (Fig. 4). A combination of field observations
and geochemical data suggest that meteoric diagenesis
contributed to this signal. Firstly, the limestones in the
lower 30 m above the G1 boundary are generally
recrystallized, and despite evidence for primary aragonite
precipitation, have an average Sr concentration of
∼150 ppm, compared to N 1000 ppm in limestones well
below and above the boundary. Based on the predominance of red silt in this stratigraphic interval, the
diagenetic fluids were oxidizing. A cross-plot of Sr and
Mn concentrations in Fig. 7b shows a distinct dog-leg
trend with [Sr] decreasing precipitously with increasing
[Mn] in samples with [Mn]b 50 ppm, before flatting out in
samples with [Mn] N 50 ppm. This trend is characteristic
of carbonate systems that have experienced a variable
degree of fluid–rock interaction during meteoric diagenesis (Brand and Veizer, 1980, 1981; Banner and Hanson,
1990; Jacobsen and Kaufman, 1999). This diagenesis also
affected the oxygen isotopic composition of limestones
above the G1 boundary as seen in the δ18O–Mn/Sr crossplot in Fig. 7c. Whereas significant variation in δ18O
occurs independent of changes in Mn/Sr, samples from
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33
the lower part of the Upper Grusdievbreen Member
(UGM) plot along a logarithmically decreasing trend
(Fig. 7c inset) as predicted by models of meteoric
diagenesis (Banner and Hanson, 1990; Jacobsen and
Kaufman, 1999).
Considering this evidence for alteration by meteoric
waters, it is not surprising that the δ18O anomaly above
the G1 boundary (Fig. 4) is absent in a plot of δ18O data
from only the best-preserved limestones (Fig. 8).
Therefore, despite the reproducibility of the low-δ18O
trend above the G1 boundary in all sections (Fig. 4), it is
regarded as purely a secondary alteration signature. In
contrast, δ13C in the same samples appears insensitive to
variation in Mn/Sr (Fig. 7d), suggesting the smoothly
declining δ13C trend (Fig. 8) is a primary seawater signal.
4.2. The Svanbergfjellet Formation
The Svanbergfjellet Formation, which is well
exposed in all sections except measured section 4
(Figs. 2, 6), is informally divided into four members
(Knoll and Swett, 1990), and thins dramatically from
over ∼ 600 m thick in southern exposures to ∼ 100 m in
the northernmost exposure (Wilson, 1961), suggesting a
northward cratonic margin to the EGES basin at this
time (Halverson et al., 2005). The contact with the
Grusdievbreen Formation is poorly defined and difficult
to pinpoint in the field, but is generally identified by a
switch from limestone to dolomite and concomitant
increase in closely spaced flooding surfaces and
abundant shallower-water facies. The Lower Dolomite
Member (Fig. 3) is distinguished by the concentration of
faults and folds in this unit, presumably due to the more
brittle behavior of the dolomite than surrounding
limestone units during Caledonian shortening.
Microbialaminites and small, laterally-linked stromatolites are particularly common facies in the Lower
Dolomite Member (Fig. 3). In contrast, in the Upper
Algal Dolomite Member (Fig. 3), larger stromatolites
comprising discrete biostromes (Knoll and Swett, 1990)
are more common and variably include Baicalia, Colonella, Tungusia, and Conophyton. The Lower and
Upper Limestone members consist predominantly of
black, ribbony, and often hummocky cross-stratified
limestone with abundant molar tooth structures. However, stromatolites also occur in the Lower Limestone
Member and include a prominent Minjaria bed (Figs. 5, 6)
that makes a useful stratigraphic marker across the outcrop
belt (Knoll and Swett, 1990). All four members of the
Svanbergfjellet Formation thin to the north, although
some of the variability between sections may be
attributable to tectonic thickening.
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G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
Fig. 7. Isotopic and elemental data from the Grusdievbreen and Svanbergfjellet formations. a) δ18O–Mg/Ca cross-plot shows two distinct
dolomitization trends whose slopes are consistent with the approximate equilibrium isotope fractionation between dolomite and calcite (e.g. Land,
1980). The upper trend is from dolomites in the Lower Dolomite Member (LDM) of the Svanbergfjellet Formation, and the lower trend is from the
Lower Grusdievbreen Member (LGM). The LDM trend is about 4‰ heavier than the LGM trend, implying a higher initial δ18O composition in the
former and consistent with data from the best preserved limestones that show higher δ18O values beneath the S1 boundary than beneath the G1
boundary (Fig. 8). b) Sr–Mn concentration cross-plot shows a typical pattern for meteoric diagenesis of originally aragonitic limestones (e.g. Banner
and Hanson, 1990). c) δ18O–Mn/Sr cross-plot shows significant variability, but three distinct patterns are visible. An inferred meteoric diagenesis
trend of logarithmically decreasing δ18O with increasing Mn/Sr (inset) in limestones from the lower part of the Upper Grusdievbreen Member
(UGM), high Mn/Sr values (N2) corresponding to dolomitized samples in the LDM and upper UGM, and a wide range (6‰) in δ18O independent of
Mn/Sr. d) δ13C–Mn/Sr cross-plot shows a distinct diagenetic trend of decreasing δ13C with increasing Mn/Sr in both dolomite and limestone samples
in the lower part of the Lower Limestone Member (LLM), which suggests that primary δ13C values should be slightly heavier (1 to 2‰) in these
samples. e) 87Sr/86Sr–Sr cross-plot shows a typical diagenetic pattern in which 87Sr/86Sr is nearly invariant in samples with Sr concentrations
N250 ppm (dashed line), but sharply rising 87Sr/86Sr in samples with Sr b250 ppm. Therefore, only samples with N250 ppm Sr are regarded as
primary. f) 87Sr/86Sr–δ18O cross-plot shows no systematic variation in Sr isotope ratios in samples with δ18O N −9‰.
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
11
35
Fig. 8. Composite isotopic and paleomagnetic records through the Grusdievbreen and Svanbergfjellet formations (data in Supplementary materials). The
δ13C record represents all data from all measured and sampled sections. 87Sr/86Sr data are from samples with [Sr] N 250 ppm in measured sections 3, 4, and
7. Grey line represents the best approximation of seawater 87Sr/86Sr spanning the Bitter Springs isotope stage. δ18O is from all measured sections with
elemental data for which Mg/Ca and Mn/Srb 0.1. The paucity of data spanning the S1 boundary reflects the preponderance of dolomite in this stratigraphic
interval, whereas the lack of data in the upper Svanbergfjellet Formation is the artifact of the absence of elemental data for samples from this interval. The
right-hand column is a composite plot of paleomagnetic data from the Grusdievbreen and Svanbergfjellet formations, summarized from Maloof et al.
(2006) and demonstrating an ∼55° discordance in paleodeclination across the G1 boundary and a return to pre-G1 paleodeclination in the Svanbergfjellet
Formation. All opposite polarity intervals are reversed for simplicity, and inclination data is omitted as it varies little through the plotted interval.
Just as the G1 sequence boundary in the Grusdievbreen
Formation delineates the sharp decline in δ13C values at
the onset of the low δ 13 C interval in the lower
Akademikerbreen Group, a second sequence boundary
(S1) at the contact between the Lower Dolomite Member
and Lower Limestone Member marks the abrupt return to
positive δ13C values in the Svanbergfjellet Formation
(Figs. 3, 6). Following the return to 13C-enriched values in
the Lower Limestone Member, δ13C values vary around a
mean of 5‰ with a pair of discernible fluctuations of
3–4‰ in amplitude that do not correspond to any
perceptible change in sedimentation (Fig. 3). Spitsbergen
sections are slightly 13C-enriched (0–1.5‰) relative to
Nordaustlandet, but otherwise, the isotopic trend is
virtually identical in the two regions (Fig. 6).
4.2.1. The S1 boundary
The S1 sequence boundary in the lower Svanbergfjellet
Formation separates dominantly shallow-water, cyclic
dolomites of the Lower Dolomite Member (LDM) below
from deeper-water sediments of the Lower Limestone
Member (LLM) above. This sequence boundary occurs
along the length of the outcrop belt, but like the G1 boundary, is better developed in Ny Friesland, where the
uppermost 0.5 m of the Lower Dolomite Member is typically heavily silicified and brecciated (Fig. 5e) and sharply
overlain by 5 to 8 m of black shale (green shale in
Nordaustlandet). The boundary is also ferruginized in most
sections, but this iron enrichment could have been a secondary effect related to leaching from the overlying black
shales. No erosional relief is apparent on outcrop scale, and
the cyclic pattern of the underlying stratigraphy precludes
determining the extent, if any, of regional scale truncation.
The black shales comprise the base of an upwardshoaling parasequence and grade upward into hummocky cross-stratified and marly dolomite ribbon rock.
This storm-dominated facies passes transitionally into a
regionally extensive biostrome, which consists of the
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G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
stromatolite Minjaria (Fig. 5f) in most locations, but
also contains Conophyton near the top of the bed in the
Lomfjord region and Nordaustlandet. The top of the
biostrome is a flooding surface and in all measured
Spitsbergen sections is overlain by a thick stack
(N100 m) of bituminous rhythmites and ribbon rocks
of the Lower Limestone Member. In some sections, the
rhythmites contain crinkly laminae, presumably the
remnants of microbial mats growing on the seafloor
beneath storm wave base. In Nordaustlandet, intermittent stromatolites persist above the Minjaria biostrome
and may be correlative with the microbially-influenced
sediments in Spitsbergen sections.
δ13C values in the upper part of the Lower Dolomite
Member vary from 0 to −5‰ in both limestone and
dolomite samples, but generally hover near −3‰ and
show a slight rise beneath the S1 boundary. δ13C ≈ 2‰ in
the lowermost carbonates above the sequence boundary,
and continues to rise gradually upsection to 5 to 7‰,
typical values for the upper Svanbergfjellet Formation.
This trajectory of rising δ13C in the lower part of the
Lower Limestone Member (Figs. 6, 8) is reproduced in
detail in all sections, and analagous to the negative δ13C
shift at the G1 sequence boundary, this trend is clearly
associated with the S1 sequence boundary. However,
δ13C values in the both dolomite and limestone samples in
the lower part of the Lower Limestone Member show a
distinct, logarithmically decreasing trend with increasing
Mn/Sr ratios (Fig. 7d). This decrease in δ13C with inferred
increase in the extent of fluid–rock interaction may be in
part related to mobilization and oxidation of organic
matter from the black shales at the base of the Lower
Limestone Member. In any case, this pattern suggests that
the preserved δ13C values in the lower part of the Lower
Limestone Member may have been depleted in 13C by as
much as 2‰ during diagenesis. However, even if these
samples were disregarded, the S1 sequence boundary
would still delineate a sharp rise in δ13C values.
The δ18O composition of carbonates bracketing the
S1 boundary is heavily influenced by the abundance of
dolomite in the lower Svanbergfjellet Formation. Like
the dolomitized samples beneath the G1 boundary, these
Svanbergfjellet dolomites define a coherent swath in a
cross-plot of δ18O versus Mg/Ca (Fig. 7a). The best fit
to these data is the line δ18O = 3.06(Mg/Ca) − 6.31. This
dolomitization trend is offset 4‰ higher than the trend
through the Grusdievbreen dolomites. Similarly, the
δ18O composition of the best-preserved limestones in
the Svanbergfjellet Formation is on average heavier than
that in the Lower Grusdievbreen Member (Fig. 8). If this
difference reflects the primary depositional environment, then either the waters of the EGES basin were
11
more restricted and saline or significantly cooler prior to
the S1 positive δ13C shift than before the G1 negative
shift. The Svanbergfjellet dolomitization trend shows
more scatter than the Grusdievbreen trend (Fig. 7a), but
this could be due to the fact that the dolomite in the
Lower Dolomite Member spans ∼150 m, and the
original δ18O of carbonates deposited in this interval
may have varied.
4.3. Strontium isotopes
We have analyzed 28 carbonate samples spanning the
G1 and S1 boundaries for 87Sr/86Sr (Table 1). The density
of data is greatest in the Grusdievbreen Formation, where
limestones are very pure and have Sr concentrations of
1000–3000 ppm. The Sr concentration of ancient
carbonates is very sensitive to the degree of fluid–rock
interaction experienced during meteoric diagenesis (Brand
and Veizer, 1980) and declines sharply with increasing Mn
concentration (Fig. 7b). Although meteoric waters have
low Sr concentrations, they are highly radiogenic, such that
during meteoric diagenesis, the 87Sr/86Sr ratio of carbonates increases. This effect is negligible where the fluid–
rock interaction is minimal and Sr concentrations remain
high, but becomes very large where Sr concentrations are
low (Banner and Hanson, 1990; Jacobsen and Kaufman,
1999). A cross-plot of 87Sr/86Sr versus Sr concentration
(Fig. 7e) clearly shows this pattern, with a very steep
radiogenic trend occurring in samples with [Sr]b 250 ppm.
On the other hand, data from samples with [Sr] N 250 ppm
define a flat, linear trend, showing no systematic increase
in 87Sr/86Sr with decreasing Sr concentration. No coherent
correlation between 87Sr/86Sr and δ18O exists (Fig. 7f),
and as independent evidence indicates significant variation
in the primary δ18O composition of the Akademikerbreen
carbonates, we consider Sr concentrations as the most
useful diagnostic tool for determining the fidelity of Sr
isotope signatures. Samples with [Sr] N 250 ppm are
regarded here as faithful proxies for original seawater
composition.
Insofar as this interpretation is correct, then our Sr
isotope data through the Svanbergfjellet and Grusdievbreen formations (Fig. 8) show a limited variation from
0.70625 to 0.70660, consistent with the least radiogenic
values documented in this stratigraphic interval by Derry
et al. (1989). Scatter over a range of ∼0.00025 is apparent, and may be the consequence of leaching of
radiogenic Sr from clastic components during diagenesis
or sample preparation (Derry et al., 1989). Bearing this
intrinsic scatter in mind, we have identified subtle but
distinct trends in the evolution of seawater 87Sr/86Sr
through the Lower Akademikerbreen Group. First,
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
87
Sr/86Sr remains nearly constant at 0.70625 in the lower
Grusdievbreen Formation, then rises to as high as 0.70660
above the G1 boundary (Fig. 8). This slight rise broadly
coincides with a pulse of siliciclastic sediment input into
the EGES basin (Fig. 4), raising the question of whether
this signal could be the result of Sr in-growth from
radiogenic Rb bound in clay minerals. However, these
elevated 87Sr/86Sr ratios are also found in very pure, Srrich limestones in the upper part of the upper Grusdievbreen Formation, suggesting that this increase in 87Sr/86Sr
is a primary signal. 87Sr/86Sr then declines gradually back
to values near 0.70625 beneath the S1 boundary, and
remains at this level into the lowermost limestones above
the boundary. Limited data show a sharp rise up section in
the Lower Limestone Member, but this trend needs to be
confirmed with additional analyses.
5. Discussion
Previously published isotopic data from the Neoproterozoic succession in Svalbard showed low δ13C values
in the middle Akademikerbreen Formation (Knoll et al.,
1986; Derry et al., 1989, 1992). We have confirmed
these values and documented in detail the evolution of
seawater δ13C through the Grusdievbreen and Svanbergfjellet formations (Fig. 8). Several striking features
of this record deserve mention. First, the negative
isotope interval appears symmetrical to the first order,
interrupting an overall trend of high δ 13 C that
characterizes the Akademikerbreen Group (Knoll
et al., 1986; Halverson et al., 2005). The isotope interval
is delineated by a sharp drop of 8‰ in δ13C in the
Grusdievbreen Formation and a similarly sharp rise of
the same magnitude in the lower Svanbergfjellet
Formation (Fig. 8). The most impressive feature of the
δ13C record is that both shifts correspond precisely to
transient fluctuations in sea level that exposed the
Akademikerbreen platform to subaerial erosion and
karstification. No other major sequence boundaries
occur within the Akademikerbreen Group.
Negative δ13C anomalies are a hallmark feature of the
Neoproterozoic. Whereas these anomalies were formerly
ascribed solely to post-glacial cap carbonates (e.g.
Kennedy, 1996; Kaufman and Knoll, 1995; Kaufman et
al., 1997), recent studies have demonstrated that a N 10‰
negative δ13C anomaly preceded the Marinoan glaciation
(McKirdy et al., 2001; Halverson et al., 2002, 2004).
Likewise, it appears that a post-Marinoan glaciation is
preceded by the most extreme drop in δ13C in the
Neoproterozoic (Halverson et al., 2005; Xiao et al., 2004).
These findings have revealed that the relationship between
the negative δ13C anomalies and glaciation is not
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37
straightforward and is different for each event. Nevertheless, the association between the anomalies and glaciation
remains. For this reason, we previously speculated that the
negative isotope shift at the G1 boundary was a proxy for a
cryptic glaciation (Halverson and Maloof, 2001). This
hypothesis was reinforced by the similarity of the basal
parasequence of the upper Grusdievbreen Formation to
typical Marinoan cap–carbonate sequences. This upwardshoaling parasequence stands out amidst the background
stratigraphy (Fig. 5a), contains reddened siliciclastic sediments (otherwise rare in the Akademikerbreen Group), and
has a regionally extensive layer choked with seafloor
cements, which are uncommon in the Neoproterozoic
outside of Marinoan cap carbonates (Groztinger and Knoll,
1995). The combined isotopic and stratigraphic patterns
above the G1 boundary are reminiscent of Marinoan cap–
carbonate sequences, albeit without a cap dolostone.
On the other hand, the sequence above the S1
boundary resembles typical Sturtian cap–carbonate
sequences, which lack a basal cap dolostone, are relatively
rich in organic matter (or disseminated sulfides), and
contain abundant sub-littoral, microbially-influenced
sediments. The δ13C trend in Sturtian cap carbonates
includes negative values at their very base (Kennedy et al.,
1998; Hoffman and Schrag, 2002; Yoshioka et al., 2003),
but the anomaly is commonly base-truncated due to
delayed deposition of the lower cap carbonate (Hoffman
and Schrag, 2002; Halverson et al., 2005). Therefore, the
Sturtian cap carbonate is most consistently distinguished
isotopically by a sharp rise in δ13C in the lower part of the
cap–carbonate sequence. These stratigraphic and isotopic
features are found in the strata overlying the S1 boundary,
and it is possible that the black shales directly above the
exposure surface could mask a basal negative δ13C
anomaly. However, for reasons discussed below, we
believe this S1 boundary is not related to glaciation.
5.1. Global correlations
The G1 boundary clearly cannot correspond to the
Marinoan glaciation as the Marinoan cap–carbonate
sequence comprises the lower Dracoisen Formation of
the Polarisbreen Group (Fig. 1; Halverson et al., 2005). But
can the S1 surface correspond to the Sturtian glaciation? If
our interpretation that both glacial diamictites in the
Polarisbreen Group belong to the Marinoan glaciation
(Halverson et al., 2005) is correct, then the Sturtian
glaciation must be represented, if not directly by glacial
deposits, then by a sedimentological and geochemical
proxy for the cap carbonate (Knoll, 2000) somewhere in
the stratigraphic section beneath the Petrovbreen Member
diamictite (Fig. 1). Due to its association with a major δ13C
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G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
anomaly, it is tempting to speculate that the S1 boundary
could be temporally equivalent to Sturtian glacial deposits
found in other basins. No geochronological data is
available to test directly this correlation. However, much
indirect chemostratigraphic and stratigraphic evidence
militates against this interpretation. For example, in no
succession where unambiguous Sturtian age glacial
deposits are found is the glaciation preceded by a
prolonged interval of negative δ13C values (Halverson et
al., 2005) as is the S1 boundary. It cannot be ruled out that
the negative isotope interval actually corresponds to the
glaciation itself (implying that it was not a global
glaciation) and has hitherto been undocumented. However,
Sturtian glacial deposits commonly fill major erosional
unconformities (Young, 1992), and the attendant fall in sea
level would have precluded deposition of the Lower
Dolomite Member, which was deposited in shallow water.
Also, average 87Sr/86Sr (0.7063) in the Grusdievbreen and
Svanbergfjellet formations is significantly lower than the
value of ∼0.7068 commonly associated with the Sturtian
glaciation (Kaufman et al., 1997; Kennedy et al., 1998;
Jacobsen and Kaufman, 1999; Shields, 1999; Walter et al.,
2000; Yoshioka et al., 2003).
Perhaps the most compelling argument that the S1
boundary does not correspond to the Sturtian glaciation
is the apparent correlation of the carbon isotopic record
of the Grusdievbreen–Svanbergfjellet with that of the
Bitter Springs Formation (Halverson et al., 2005), which
lies stratigraphically well below the Sturtian age glacial
deposits of the Aralka Formation in the Amadeus Basin
of central Australia (Walter et al., 1995). Fig. 9 shows
the Bitter Springs δ13C data (Hill et al., 2000a) plotted
along with the Svalbard data, scaled such that the
negative and positive shifts align. Both the positive and
negative shifts are equal in magnitude, and the isotopic
pattern between the G1 and S1 boundaries is virtually
identical to the δ13C profile through the lower half of the
Loves Creek Member (Hill et al., 2000a). 87Sr/86Sr data
from the Bitter Springs Formation are rather variable
and do not definitively corroborate this correlation.
However, 87Sr/86Sr values as low 0.7057 occur in the
Gillen Member (Hill et al., 2000a), and one test of this
correlation will be whether similarly unradiogenic
values occur in the lowermost Akademikerbreen Group.
This correlation enables a rough estimate of the age of
the lower Akademikerbreen Group since age constraints
in Australia are tighter than those in Svalbard. Hill et al.
(2000a,b) previously argued that the upper Bitter Springs
Formation is ∼830 Ma based on correlation of volcanics
in the upper Loves Creek Member with the 827 ± 6 Ma
Gairdner Dyke Swarm (Wingate et al., 1998) on the
Gawler craton. However, a younger age seems more
11
Fig. 9. Overlay of δ13C data from the Bitter Springs Formation (Hill et al.,
2000a) and data from the Grusdievbreen and Svanbergfjellet formations
(from Fig. 3), scaled so as to match the negative and positive δ13C shifts
across the G1 and S1 boundaries. Note the nearly perfect match of δ13C
datawithintheBitterSpringsstage.Theexceptionisapositivespikeinthe
Bitter Springs Formation that corresponds to an interval of non-marine
deposition (Hill et al., 2000a) and may not reflect a primary seawater
signal. Approximate ages of G1 and S1 boundaries are taken from
correlation of the Bitter Springs stage with the Callanna Group in the
Adelaide Geosyncline, which is constrained by U–Pb ages on the Rook
Tuff(Fanningetal.,1986)atthebaseoftheCallannaGroup(Preiss,2000)
and the Boucaut Volcanics at the base of the overlying Burra Group
(Preiss, 2000).
likely since the Bitter Springs Formation appears to
correlate with the isotopically light Curdimurka Subgroup
(Hill and Walter, 2000) in the Adelaide geosyncline
(Preiss, 2000). If this correlation is correct, then the
beginning and end of the negative δ13C interval are
broadly constrained to ca 800 Ma by the 802 ± 10 Rook
Tuff (Fanning et al., 1986) in the lower Curdimurka
Subgroup and the 777 ± 7 Ma Boucaut Volcanics (Preiss,
2000) in the lower Burra Group (Fig. 9).
Rainbird et al. (1996) correlated the Bitter Springs
Formation with the upper part of the their “Succession
B” in northern Canada based on lithological and
stratigraphic similarities between the two continents.
This correlation implies that the negative δ 13 C interval
should occur within the Little Dal Group in the
Mackenzie Mountains and in the Shaler Supergroup
and equivalent strata in the Amundsen Basin (Rainbird
et al., 1996). A preliminary study has shown an interval
of 13 C-depleted carbonates in the Shaler Supergroup
(Asmeron et al., 1991), while new data from the
Mackenzie Mountains reveal a pronounced negative
anomaly in the middle Little Dal Group that is most
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
likely correlative with the Bitter Springs anomaly in
Svalbard and central Australia (Halverson, 2006).
5.2. The Neoproterozoic
87
Sr/86Sr record
The combination of a distinct δ13C pattern with nearly
constant Sr-isotopic composition make the low δ13C
interval in the lower Akademikerbreen Group particularly
well suited for global correlations. It is well established
that seawater 87Sr/86Sr increased dramatically from
unradiogenic values b 0.7060 in the early Neoproterozoic
to N0.7070 by the Marinoan glaciation (Veizer et al.,
1983; Asmeron et al., 1991; Derry et al., 1992; Jacobsen
and Kaufman, 1999; Shields, 1999; Walter et al., 2000;
Melezhik et al., 2001; Thomas et al., 2004). However, as
the comprehensive compilations of Neoproterozoic
87
Sr/86Sr data by Melezhik et al. (2001) and Thomas
et al. (2004) have revealed, the current database is
contradictory and suggests highly variable seawater
87
Sr/86Sr, particularly in the middle Neoproterozoic.
Several factors contribute to the confusing Neoproterozoic 87Sr/86Sr record. The combination of a lack of
radiometric ages for this time interval, poorly constrained
correlations, and inconsistent sample quality all result in
artificial variation seen in composite Sr isotope records.
Most compilations indicate that 87Sr/86Sr values as low as
those in the Grusdievbreen and Svanbergfjellet formations (~0.7063) are restricted to carbonates older than ca
770 Ma (Melezhik et al., 2001; Thomas et al., 2004).
However, Kuznetsov et al. (2003) suggest instead that
87
Sr/86Sr remained below 0.7063 until ~650 Ma.
In their model on Neoproterozoic sedimentary cycling,
Derry et al. (1992) argued that both the generally
increasing 87Sr/86Sr and high δ13C seawater compositions
through the latter half of the Neoproterozoic reflected high
rates of continental erosion and marine sedimentation,
resulting in increased input of radiogenic Sr into the
oceans and high fractional burial of organic carbon
(forg = Corganic-buried/Ctotal-buried). Even with the much
greater residence time of Sr in the oceans compared to
C, given the apparent longevity of the negative δ13C
interval in the Akademikerbreen Group, one would
anticipate a decrease in 87Sr/86Sr if the drop in δ13C was
the result of decreased rates of continental erosion.
Instead, 87Sr/86Sr rises slightly by ∼0.0002 directly
above the G1 sequence boundary and remains elevated
(up to 0.70660) through most of the negative δ13C
interval (Fig. 8). Although it cannot be ruled out that the
change in 87Sr/86Sr was related to a decrease in the input
of hydrothermal Sr, the concurrence of the increase in
87
Sr/86Sr with the input of siliciclastic sediments into the
EGES basin and the decay of this subtle anomaly (Fig. 8)
11
39
with the return to pure carbonate deposition suggest that it
records a change in continental weathering patterns.
However, since silicate weathering rates can only increase
over long time scales with a concomitant increase in CO2
outgassing rates, this pulse of radiogenic Sr into the
oceans must have been related either to a transient
elevation of pCO2 or change in the average 87Sr/86Sr
composition of weathered continental crust.
5.3. The δ18O record
The significance of the δ18O composition of ancient
marine carbonates is a subject of much debate. On the one
hand, the best-preserved Proterozoic limestones are
consistently depleted by 5–8‰ when compared to
Cenozoic values (Veizer and Hoefs, 1976; Burdett et al.,
1990; Kah, 2000; Frank and Lyons, 2000). Similarly,
δ18O values of marine calcite, aragonite, and phosphate
shells suggest a gradual rise in the δ18O composition of
the oceans from −8 to 0‰ through the Phanerozoic
(Veizer et al., 1999). This argument is countered by
modeling results (e.g. Muehlenbachs and Clayton, 1976;
Muehlenbachs, 1998) and δ18O analyses of silicate
minerals in ophiolites of various ages (Gregory and
Taylor, 1981; Holmden and Muehlenbachs, 1993), which
suggest that seawater δ18O is buffered against long-term
change by hydrothermal circulation through mid-ocean
ridges. However, these arguments disregard possible
variations in the temperature at which crustal alteration
occurs, which can drive long-term changes in marine
δ18O (Wallman, 2001) and also exert a dominant control
on the δ18O composition of altered oceanic crust.
Although the purpose of this paper is not to try to resolve
this controversy, our data are consistent with long term
variation of marine δ18O over geological time.
The oxygen isotopic trends spanning the G1 boundary
underline the potential pitfalls in interpreting the δ18O
record of Neoproterozoic carbonates. Both the rise in
δ18O beneath the G1 boundary and the negative anomaly
above it are highly reproducible in all measured sections.
However, a combination of petrologic and geochemical
evidence suggests that these two trends are the consequence of dolomitization and meteoric diagenesis,
respectively. The fact that the dolomitization trend is
preserved (Fig. 7a) despite the evidence for widespread
groundwater flow reveals that these dolomite beds were
relatively resistant to isotopic equilibration. The reproducibility of the meteoric signal in all measured sections
(Fig. 4) is also impressive and suggests that the silty
parasequence was at one time a basin-wide aquifer. Since
this aquifer would have been disconnected by faulting and
folding during Caledonian deformation, it is reasonable to
40
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
surmise that the meteoric signal is pre-Caledonian in age.
The preservation of the discordant isotopic signals across
the G1 boundary (Fig. 4) further suggests that the δ18O
composition of these rocks has not been appreciably
modified in the past 400 Ma. These observations imply a
high preservation potential for Svalbard limestones,
despite isolated instances of extensive, pre-Caledonian
meteoric diagenesis.
Dolomitization and meteoric diagenesis left distinct
imprints on the elemental composition of Akademikerbreen carbonates (Fig. 7). Fig. 8 shows a compilation of the
δ18O composition of all samples spanning the G1 and S1
boundaries for which elemental data indicate minimal
alteration (Mg/Ca, Mn/Sr b 0.1). Notably, both the positive
deflection beneath the G1 boundary and the negative
anomaly above it apparent in raw δ18O data (Fig. 4) are
absent from this compilation of least altered data.
Interestingly, δ18O in the best-preserved samples above
the G1 boundary is roughly equal to or slightly higher than
average δ18O in the lower Akademikerbreen Group,
whereas a decline in δ18O of ∼2‰ occurs below the
exposure surface (Fig. 8). If this decline in δ18O values
represents a primary signal, then either it reflects a transient
warming episode (∼8 °C) or a freshening of the EGES
basin. Given the geological evidence for the emergence
and dolomitization of the Akademikerbreen platform, the
latter hypothesis is unlikely. Since the downturn in δ13C
values also begins beneath the G1 boundary, it seems that
the mechanism responsible for perturbing the global
carbon cycle, and perhaps the warming of the EGES
basin, preceded the nadir in sea level during which the G1
sequence boundary developed.
5.4. Evidence for True Polar Wander events?
The association of large shifts in the carbon-isotopic
composition of Neoproterozoic seawater with regionally
persistent sequence boundaries insinuates glaciation.
However, indirect evidence for a pre-Sturtian age implies
that the Bitter Springs isotope stage precedes any of the
known Neoproterozoic ice ages. A ca 800–780 Ma
glaciation cannot be ruled out, but no supporting evidence
has been found in any basin of this age. Therefore, a
different mechanism must be sought that can simultaneously account for the δ13C pattern, the pair of relative
sea level fluctuations, and enhanced continental weathering above the G1 sequence boundary, as indicated by both
the influx of red silt and rise in 87Sr/86Sr.
New paleomagnetic data from limestone and siltstone
samples spanning the G1 boundary in the Grusdievbreen
Formation and from the Upper Algal Dolomite Member
in the Svanbergfjellet Formation add a key piece to this
11
puzzle. These data (Maloof et al., 2006) will be discussed
in greater detail in a separate paper, and only the results are
summarized here (Fig. 8). A stable, high temperature
magnetic component has been found in about 85% of a
suite of samples collected from three different sections
spanning the G1 boundary. This remnant magnetization is
reproduced in all sections, is distinct from regional
overprints of younger age from Svalbard, and passes a
Caledonian fold test. Whereas the paleomagnetic inclination remains nearly constant (∼15°), the declination
changes by ∼55° across the sequence boundary (Fig. 8).
Paleomagnetic data from the Svanbergfjellet Formation,
which pass a syn-sedimentary fold test, preserve a
paleomagnetic declination similar to that beneath the G1
boundary, only with a slightly higher inclination.
The evidence for meteoric diagenesis above the G1
boundary raises the possibility that the paleomagnetic
signature is an overprint signal. However, if primary, and
assuming the geocentric axial dipole hypothesis is valid
for the middle Neoproterozoic, then the 55° discordance
in declination can only be accounted for by a large rotation
of the EGES platform with respect to the earth's spin axis.
Plate tectonics can produce such rotations, but a minimum
of ∼30 Ma is required to account for a rotation of this
magnitude (Maloof et al., 2006). Given the conformable
nature of the strata (Figs. 3, 5) and continuity of the δ13C
pattern (Fig. 8) across the sequence boundaries, this time
span is inconsistent with the data.
Another mechanism capable of producing rotations
of this magnitude is true polar wander (TPW), where the
entire solid earth rotates with respect to the spin axis,
which remains fixed in a celestial reference frame.
Inertial interchange true polar wander (IITPW) is a
variety of TPW in which mantle heterogeneities are
arranged in such a way that a small redistribution of
mass causes the solid earth to rotate by up to 90° (Fisher,
1974). During an IITPW event, continents furthest from
the IITPW rotational axis will undergo large latitudinal
shifts (paleoinclination), while those closest to the axis
will undergo large rotations (paleodeclination) and little
latitudinal change. The rate of rotation of the silicate
earth during an IITPW event is limited by the relaxation
time of Earth's hydrostatic bulge, and a full 90° rotation
is estimated to require 3–20 Ma (Steinberger and
O'Connell, 1997).
An IITPW event would also generate transient sea
level changes related to the delayed response of the solid
earth with respect to the ocean under gravitational load
(Mound and Mitrovica, 1998; Mound et al., 1999). A
continent traversing the equator would experience a
relative rise in sea level as it moved to lower latitude and a
relative fall in sea level as it moved to higher latitude, with
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
maximum effects at 45° from the IITPW axis. (Mound
et al., 1999). Based on the position of the EGES platform
with respect to a calculated IITPW axis, Maloof et al.
(2006) estimated a relative fall in sea level change of up to
25 m during both the G1 and the S1 IITPW events. Since
the transient fall in sea level would coincide with only a
fraction of the total IITPW event, the low in sea level
would endure less than ∼3–20 m.y. (Steinberger and
O'Connell, 1997). Therefore, a pair of IITPW events,
which is consistent with the paleomagnetic data, could
account for the G1 and S1 sequence boundaries.
Kirschvink et al. (1997) proposed an early Cambrian
IITPW event and speculated that it contributed to the early
radiation of metazoa. Kirschvink and Raub (2003) further
hypothesized that a Cambrian IITPW event would trigger
episodic releases of methane from marine clathrate
reservoirs as a consequence of changes in ocean
circulation (and therefore temperature) and sea level
(pressure). Due to the highly 13C-depleted composition of
methane (∼−60‰; Kvenvolden, 1993), Kirschvink and
Raub (2003) argued that these methane bursts could
explain high order fluctuations in the early Cambrian
δ13C record. Outbursts of this potent greenhouse gas to
the oceans and atmosphere may help explain the δ18O
data beneath the G1 boundary, which suggest warming of
the EGES basin prior to exposure of the carbonate
platform. The pulse of increased silicate weathering could
also be related to this warming event, but another possible
explanation is that during the IITPW rotation, Laurentia
and other low-latitude continents traversed new climatic
regimes, which could have changed the loci of continental
weathering and destabilized the landscape (cf. Zhang
et al., 2001) over vast areas. The input of methane-derived
carbon and decrease in biological removal of carbon from
the oceans could account for the increased alkalinity
implied by the formation of seafloor cements above the
G1 boundary. However, methane alone cannot account
for the low δ13C values between the G1 and S1 boundaries
as the volume of methane required to suppress marine
δ13C by ∼8‰ through ∼325 m of carbonate far outstrips
the potential global reservoir of methane clathrate (Buffet,
2000; Dickens, 2001). Another more fundamental change
to global carbon cycling is required.
The most straightforward way to alter seawater δ13C for
extended intervals of time is to change the fractional burial
of organic carbon in marine sediments (forg; Kump and
Arthur, 1999). The ∼8‰ decline across the G1 boundary
could be accounted for by a shift in forg from ∼0.35 to 0.15.
Organic carbon burial is mainly determined by a
combination of sedimentation rates and primary productivity, the latter of which is controlled by nutrient supply
and both of which are tied to continental weathering
11
41
(Schrag et al., 2002). Today, approximately 70‰ of total
organic carbon removed from the biosphere is buried in
large tropical river deltas (Hedges and Keil, 1995). Schrag
et al. (2002) postulated that a preponderance of continents
in low latitudes would have maintained high forg and
therefore the high δ13C characteristic of much of the
Neoproterozoic (Knoll et al., 1986) as a consequence of a
large number of tropical rivers and more restricted
equatorial ocean basins in which bottom-water anoxia
would facilitate burial of organic matter and nutrient
recycling. An IITPW event in the middle Neoproterozoic
(G1) would have rotated those equatorial continents and
basins far from the IITPW axis into higher latitudes,
reducing the number of large tropical rivers (Maloof et al.,
2006) and exposing previously restricted basins to zonal
ocean currents, thereby increasing mixing and inhibiting
the development of anoxic conditions. The consequence of
such an event would be a reduction in global carbon burial
rates and a decrease in δ13C to a new steady state
controlled by the new distribution of the continents.
Likewise, a second (S1) IITPW event during which the
continents returned to their equatorial positions would
boost organic carbon burial and drive δ13C back up to the
typical high Neoproterozoic values. Therefore, at least
qualitatively, a pair of IITPW events could account for the
large, symmetric δ13C shifts at the G1 and S1 boundaries.
A proposed connection between the negative δ13C
interval and an IITPW event is admittedly speculative.
However, Li et al. (2004) have independently proposed a
TPW event occurring sometime between ∼800 and
750 Ma to explain disparate paleopoles from South
China, and suggest that it drove a mantle superswell and
associated ca. 800 Ma flood basalts from polar to tropical
latitudes. Determining whether the Li et al. (2004) TPW
event is synchronous with the Akademikerbreen low δ13C
interval and consistent with the proposed G1 and S1
IITPW events will require additional geochronological
and chemostratigraphic constraints. However, our IITPW
hypothesis is straightforward to test because the relative
sea level changes and paleomagnetic disparities associated with an IITPW event are different but predictable for
every continent (Mound et al., 1999; Evans, 2003), given
a particular paleogeography.
6. Conclusions
We have presented detailed stratigraphic and chemostratigraphic data spanning a salient interval of low δ13C in
the middle Akademikerbreen Group in northeast Svalbard. We estimate this interval to be ca 800 Ma based on
the correlation with the lower Loves Creek Member of the
Bitter Springs Formation in central Australia, which has a
42
G.P. Halverson et al. / Chemical Geology 237 (2007) 23–45
virtually identical δ13C trend (Fig. 9; Hill et al., 2000a).
Insofar as this correlation is correct, then this isotopic
anomaly is global and preceded the Sturtian glaciation.
The symmetry of the δ13C anomaly suggests that the
positive and negative shifts that define its beginning and
end represent transitions between distinct stable states of
global organic carbon burial. A slight rise in seawater
87
Sr/86Sr above the G1 boundary implies that forg and
erosion rates were not directly coupled. However, the
coincidence of the δ13C shifts with major sequence
boundaries point to a connection to global phenomena
that were capable of affecting regional sea level, patterns
of organic carbon burial, and weathering patterns on the
continents.
Given the association between large negative δ13C
anomalies and glaciation in the Neoproterozoic (Knoll
et al., 1986), it is tempting to associate the Akademikerbreen negative δ13C interval to a pre-Sturtian glaciation
that has hitherto been undocumented. However, the
absence of any direct evidence for glaciation in Svalbard,
central Australia, and other presumably coeval sedimentary successions (e.g. in northern Canada) demands an
alternative explanation. New paleomagnetic data from
Svalbard (Maloof et al., 2006) may provide the answer.
These data indicate a rapid 55° rotation of the EGES
platform spanning the G1 boundary and a counter rotation
of approximately the same magnitude across the S1
boundary. We propose that a pair of IITPW events can
account for both the paleomagnetic data and the transient
changes in relative sea level at the G1 and S1 sequence
boundaries. The low δ13C values that occur between these
two sequence boundaries could be accounted for by
dramatically altered patterns of organic carbon burial (cf.
Schrag et al., 2002) between the two IITPW events
resulting from changes in nutrient delivery and ocean
circulation patterns. Although this hypothesis is speculative, it is straightforward to test since the Akademikerbreen δ13C anomaly should be easy to identify in
carbonates of equivalent age, and the paleomagnetic and
eustatic imprints of an IITPW event are variable but
predictable for other continents, based on their position
relative to the IITPW axis (Mound et al., 1999; Evans,
2003). If subsequent paleomagnetic studies support the
evidence from Svalbard for a pair of IITPW events ca
800 Ma, the associated shifts in steady state marine δ13C
composition will be testament to the profound effect of
paleogeography on global carbon cycling.
Acknowledgments
This work was supported by the National Science
Foundation (Arctic Science Program grant OPP-9817244
11
to Paul Hoffman, Harvard University), the NASA
Astrobiology Institute, the Canadian Institute for Advanced Research (Earth System Evolution Project), and a
GSA graduate research grant to GPH. We thank Alcides
Sial for inviting GPH to submit this manuscript.
Constructive reviews by Ján Veizer and Andrey Bekker
greatly improved the quality of the manuscript. Paul
Hoffman supervised and funded the field project in
Svalbard, which was carried out as part of GPH's and
ACM's PhD research. Sam Bowring supervised strontium isotope analyses at MIT. Andy Knoll provided much
insight and many suggestions on fieldwork and related
manuscripts. Norsk Polarinstitutt provided logistical
support in Svalbard. Winfried Dallman and Ken Petersen
helped organize fieldwork. Ethan Goddard and Greg
Eischeid provided assistance and supervision in stable
isotopic and elemental analyses.
Appendix A. Supplementary data
Supplementary data associated with this article can
be found, in the online version, at doi:10.1016/j.
chemgeo.2006.06.013.
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