Precambrian Research 137 (2005) 167–206
Evidence for Paleoproterozoic cap carbonates in North America
A. Bekker a,b,∗ , A.J. Kaufman c , J.A. Karhu d , K.A. Eriksson a
a
Department of Geological Sciences, Virginia Polytechnic Institute and State University, Blacksburg, VA 24061, USA
b Department of Earth and Planetary Sciences, Harvard University, Cambridge, MA 02138, USA
c Department of Geology, University of Maryland at College Park, College Park, MD 20742, USA
d Department of Geology, University of Helsinki, FIN-00014, Finland
Accepted 1 March 2005
Abstract
The early Paleoproterozoic Snowy Pass Supergroup of the Medicine Bow Mountains and Sierra Madre, Wyoming, USA and
Huronian Supergroup, Ontario, Canada were deposited along the present-day southern flank of the Wyoming and Superior cratons.
Whereas three discrete levels of glacial diamictite are developed in both successions, carbonate strata are known only directly
above the middle diamictite (Vagner and Bottle Creek formations in Medicine Bow Mountains and Sierra Madre, respectively,
and Espanola Formation in southern Ontario) in these thick correlative siliciclastics-dominated strata. The carbonates from each
succession record negative ␦13 C values (−4.0 to −0.8‰, V-PDB) and attenuated carbon isotopic difference between organic
and inorganic phases. Oxygen in carbonates is strongly depleted in 18 O suggesting exchange with hot fluids, which is consistent
with pervasive recrystallization of carbonates and remobilization of elements. However, the stratigraphic coherence of carbon
isotopic compositions and the general lack of correlation between ␦13 C and either ␦18 O values or trace element concentrations
supports a primary origin for 13 C-depleted carbonates, which are interpreted here to reflect anomalous oceanic compositions.
The intimate association of thick carbonate units containing abundant carbonate debris flows with immediately underlying
glacial strata indicates that chemical precipitation resulted from a rapid flux of carbonate alkalinity onto ocean margins during
post-glacial transgression. Although these early Paleoproterozoic carbonates are similar to Neoproterozoic ‘cap dolomites’ in
stratigraphic position and carbon isotopic compositions, the older post-glacial accumulations begin with limestone and lack
many of the sedimentary structures typical of Neoproterozoic deposits. Furthermore, it is not understood why carbonates only
occur above the middle of the three glacial horizons whereas these deposits are ubiquitous above Neoproterozoic diamictites.
The differences might reflect lower overall carbonate saturation in early Paleoproterozoic oceans which contrasts sharply with
Archean and later Paleoproterozoic intervals and higher siliciclastic inputs in rift environments, which shut down carbonate
deposition.
Geological and geochemical indicators suggest a stepwise increase in atmospheric oxygen across the Paleoproterozoic glacial
epoch. The tempo and mode of atmospheric oxygen rise has significant consequences for the abundance of the important
greenhouse gases CH4 and CO2 and hence for oceanic acidity. If we accept that atmospheric oxidation of methane to carbon
∗ Corresponding author at: Geophysical Laboratory, Carnegie Institution of Washington, 5251 Broad Branch Rd., N.W., Washington, DC
20015, USA. Fax: +1 202 478 8901.
E-mail address: [email protected] (A. Bekker).
0301-9268/$ – see front matter © 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.precamres.2005.03.009
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A. Bekker et al. / Precambrian Research 137 (2005) 167–206
dioxide resulted in each of the three discrete glaciations, it implies that atmospheric CH4 remained high throughout the interval
and that pulsed oxidation events, plausibly linked to higher primary productivity and lower hydrothermal activity, led to surface
refrigeration. If correct, the unique presence of cap carbonate above the middle Paleoproterozoic diamictite may reflect an
appropriate balance of CO2 and CH4 sufficient to provide enough alkalinity to seawater through silicate weathering, but not so
high that carbonate preservation would be inhibited by enhanced acidity.
© 2005 Elsevier B.V. All rights reserved.
Keywords: Chemostratigraphy; Paleoproterozoic; Cap carbonates; Huronian Supergroup; Snowy Pass Supergroup; Rise of atmospheric oxygen
1. Introduction
At the beginning and end of the Proterozoic Eon
(2500–543 Ma), the Earth was covered on several separate occasions by large continental ice sheets that may
have extended to sea-level at low latitudes (e.g. Evans
et al., 1997; Williams and Schmidt, 1997; Schmidt and
Williams, 1999; Sohl et al., 1999; Kempf et al., 2000).
A comparative analysis of the two glaciogenic intervals reveals striking similarities, in particular both are
coincident with supercontinent rifting (Young, 1988),
and differences, most notably the unusual presence of
carbonates and iron formations in the younger, but their
general absence in the older.
Neoproterozoic glacial diamictites deposited during the rifting and breakup of the Neoproterozoic
supercontinent known as Rodinia are commonly enriched in iron oxides with local accumulations of
thick iron formations in basinal facies, and are commonly overlain by enigmatic “cap” carbonates (e.g.
Kaufman et al., 1991; Kennedy, 1996; Hoffman and
Schrag, 2002; Young, 2002). Geochronologic and
chemostratigraphic constraints argue for at least three
discrete ice ages during the Neoproterozoic (Kaufman
et al., 1997; Bowring et al., 2003) which has been
considered as “snowball Earth” events (Kirschvink,
1992; Hoffman et al., 1998). Initiation of these glacial
episodes has been attributed to various processes, including: (1) CO2 sequestration due to high rates of
primary productivity and subsequent organic carbon
burial (Kaufman et al., 1997), (2) intense weathering of the high-standing supercontinent during rifting
and breakup (Hoffman and Schrag, 2000; Goddéris
et al., 2003; Donnadieu et al., 2004), (3) oxidation
of a pre-glacial atmosphere dominated by methane
(Schrag et al., 2002; Halverson et al., 2002), or (4)
changes in ocean circulation associated with the low
latitudinal position of Rodinia (Smith and Pickering,
2003).
Early Paleoproterozoic ice ages (≤2.45 to >2.22 Ga)
are believed to be associated with assembly and rifting of a Late Archean supercontinent (Kenorland of
Williams et al., 1991), which was likely positioned
in low latitudes (Evans et al., 1997). Although up to
three discrete episodes of glaciation are recorded, there
are no true iron formations in this interval and very
few carbonates. Based on paleomagnetic, chemostratigraphic, and lithostratigraphic data, at least one of these
ice ages has been interpreted as an older snowball
Earth event potentially related to the transition from
an anoxic methane-rich atmosphere to an oxygenated
atmosphere with high CO2 levels (Evans et al., 1997;
Kirschvink et al., 2000; Pavlov et al., 2000; Bekker
et al., 2001). Lacking thick, well-preserved carbonates, little is known of temporal variations in carbon
and strontium isotope composition of seawater during the early Paleoproterozoic glacial epoch. However,
chemostratigraphic studies of later Paleoproterozoic
carbonates from Fennoscandia and elsewhere defined
a remarkable positive carbon isotope excursion with an
age between ca. 2.22 and 2.10 Ga (Karhu, 1993; Karhu
and Holland, 1996; Bekker et al., 2003).
The early Paleoproterozoic Huronian and Snowy
Pass supergroups in southern Ontario and Wyoming
preserve lithologic and geochemical (chemical index
of alteration of siliciclastic units) evidence of three
separate glaciations with intervening periods of warm
climate (Fig. 1; Young, 1991). In both successions,
carbonates occur only above the middle glacial diamictite. Veizer et al. (1992) studied post-glacial carbonates of the Espanola Formation, Huronian Supergroup,
Ontario and interpreted their negative ␦13 C values as
reflecting a lacustrine environment of deposition. In
contrast, carbonates above the glacial diamictite of the
lower Duitschland Formation, which is sandwiched
between two glacial deposits of the early Paleoproterozoic Pretoria Group, South Africa have negative
carbon isotope values interpreted to reflect seawater
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169
Fig. 1. Correlation of the Snowy Pass Supergroup, Medicine Bow Mountains, WY, USA, the Huronian Supergroup, ON, Canada, and the
Transvaal Supergroup, South Africa based on tectonomagmatic event at ca. 2.48–2.45 Ga, glacial diamictites, carbonates with negative carbon
isotope values (based on this paper) and mature Al- and hematite-rich quartzites overlying the middle and upper glacial diamictites respectively,
and chemostratigraphy of post-glacial carbonates (Bekker et al., 2001, 2003). References to age constraints: 1: U–Pb TIMS baddeleyite age
(Andrews et al., 1986); 2: U–Pb TIMS zircon age for volcanic rocks (Krogh et al., 1984); 3: U–Pb TIMS zircon and baddeleyite ages (Premo
and Van Schmus, 1989; Snyder et al., 1995; Cox et al., 2000); 4: U–Pb TIMS detrital zircon age (Premo and Van Schmus, 1989); 5: Pb–Pb
whole rock isochron age (Cornell et al., 1996); 6: Re–Os isochron age of early diagenetic pyrite (Hannah et al., 2004); 7: U–Pb SHRIMP zircon
age of ash beds (Pickard, 2003); 8: U–Pb SHRIMP zircon age of ash beds (Martin et al., 1998) (Bekker et al., 1996; Buick et al., 1998; Master
et al., 1993; Schidlowski et al., 1975; Swart, 1999).
composition (Bekker et al., 2001). The carbonates of
the lower Duitschland Formation are possibly correlative with carbonates immediately above the middle
glacial diamictite in the North American successions
based on lithostratigraphic and chemostratigraphic data
(Bekker et al., 2001, 2003), but geochronologic constraints are lacking. Notably, the upper Duitschland
Formation provides the only example of a positive carbon isotope excursion in Paleoproterozoic pre-glacial
carbonates.
Worldwide, the end of the Paleoproterozoic glacial
epoch is marked by mature, Al-rich quartzites suggesting climate change to greenhouse conditions favoring strong chemical weathering (Young, 1973;
Karlstrom et al., 1983; Marmo, 1992; Ojakangas,
1997; Bekker et al., 2001), quite likely under a high
CO2 atmosphere. The ca. 2.22–2.10 Ga carbon isotope excursion (Karhu and Holland, 1996) also follows
the ice ages, but the temporal relationship between
the beginning of the carbon isotope excursion and
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the onset of extreme weathering is poorly constrained.
Whereas glacial deposits are known throughout
Earth’s long history, only Proterozoic glacial diamictites appear to be immediately overlain by unique and
isotopically anomalous carbonates deposited during
post-glacial transgression. Neoproterozoic cap carbonates have been additionally defined by their unusual
sedimentary structures (Corsetti and Kaufman, 2003),
which may have resulted from anomalously high carbonate alkalinity in the ocean driving rapid carbonate precipitation. With the exception of the unusual
sedimentary structures, the known post-glacial Paleoproterozoic carbonates may also be accepted as “cap
carbonates”, and are considered so in this paper, which
describes sedimentology and chemostratigraphy of the
carbonates in the early Paleoproterozoic Snowy Pass
Supergroup of Wyoming and Huronian Supergroup of
Ontario. Results of this study are presented to (1) evaluate the environment of deposition of the carbonate
rocks; (2) provide a record of temporal carbon isotope
variations during the Paleoproterozoic glacial epoch
and its relationship to climatic changes, which may provide constraints on global correlations, and (3) suggest
a model for deposition of the Paleoproterozoic postglacial carbonates, which are linked to changes in the
redox state of the atmosphere.
2. Regional setting and stratigraphy
2.1. Huronian Supergroup
The best exposed and studied early Paleoproterozoic successions in North America are the thick, predominantly siliciclastic packages preserved along the
southern margin of the Superior Craton. The Huronian Supergroup outcrops along the north shore of
Lake Huron, Ontario, Canada and in a series of large
outliers east to Cobalt and north to Timmins (Young,
1991; Fig. 2). The lower part of the Huronian Supergroup was deposited in a rift setting possibly connected
with an ocean to the east, whereas the upper part starting with the Gowganda Formation is interpreted to
represent a passive margin succession dominated by
siliciclastic sediments (Zolnai et al., 1984; Young et
al., 2001). Local disconformities and unconformities
at the base of the Serpent and Gowganda formations
(Young and Church, 1966) as well as abundant clastic dikes, breccias, and syn-depositional faults contemporaneous with deposition of the Espanola Formation
provide evidence for tectonic reactivation (Eisbacher,
1970; Young, 1972), likely related to extension accompanying rift-drift transition. Further evidence of
extension and related episodic large-scale gravitational
displacement includes ‘zig-zag’ or diffuse boundaries
between units, chaotic folding and incorporation of
large fragments of sediments into adjacent units (Long
et al., 1999). The Huronian Supergroup is subdivided
by unconformities into four groups; the upper three
are climatically controlled cycles with basal glacial diamictites overlain by deltaic shale or carbonate and
thick fluvial sandstones (Young et al., 2001). The only
thick and extensive carbonate unit of the Huronian
Supergroup, the Espanola Formation, occurs directly
above the middle glacial diamictite (Bruce Formation)
and is overlain by fluvial sandstones of the Serpent Formation.
The basal Huronian Supergroup rests unconformably, with locally preserved paleosol, on basement rocks of the Archean Superior Province (Prasad
and Roscoe, 1996). Basal sediments above the reduced
paleosol contain conglomerates with detrital pyrite
and uraninite, suggesting low atmospheric pO2 (Livingstone Creek and Matinenda formations; Fig. 1).
Interlayered mafic volcanic rocks and intrusive contacts with the Murray and Creighton granites constrain the age of the basal Huronian to 2.49–2.45 Ga
(Krogh et al., 1984, 1996; Smith and Heaman, 1999),
whereas the whole Huronian Supergroup is cut by the
2217.5 ± 1.6 Ma Nipissing sills and dikes (Andrews et
al., 1986). Red sandstone and shale in the upper part of
the Gowganda Formation and the oxidized Ville Marie
paleosol below the red beds of the Lorrain Formation
provide the first evidence for an oxidizing atmosphere
(Wood, 1979; Rainbird et al., 1990; Panahi et al., 2000;
Young et al., 2001). Pseudomorphs after anhydrite in
the Gordon Lake Formation (Chandler, 1988) are consistent with an intensification of the oxidative part of
the S cycle. Al-rich quartzites of the Lorrain Formation
above the glacial diamictite of the Gowganda Formation imply rigorous chemical weathering and therefore
rapid climate amelioration after the glaciation (Young,
1973; Long et al., 1999). Facies change dramatically
across the latitudinal Murray Fault Zone (M.T.Z.; see
Fig. 2), which separates shallow water sequences with
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
171
Fig. 2. Location of Huronian rocks and studied sections of the Espanola Formation (modified from Fedo et al., 1997). Insets show position of
the Huronian Basin in North America and along the north shore of the Lake Huron on the southern margin of the Archean Superior Province.
Arrows within the Huronian outcrop represent generalized paleocurrent directions for the Huronian Supergroup, M.F.Z. is Murray Fault Zone.
subgreenschist to lower greenschist facies of metamorphism (Card, 1978) and open folds to the north from
deep-water, thicker sequences with higher metamorphic grade and tight upright folds to the south. Paleocurrent indicators in the lower Huronian sandstones
along with facies analysis and thickness changes suggest a paleoslope towards the southeast (Fralick and
Miall, 1989; Rousell and Long, 1998) locally complicated by basement highs (Robertson, 1976, 1977).
2.2. Snowy Pass Supergroup, Wyoming
Paleoproterozoic sedimentary and subordinate volcanic rocks are well exposed along the southeastern
margin of the Wyoming Craton in the Sierra Madre
and Medicine Bow Mountains, Wyoming (see inset in
Fig. 3). These sequences were deformed during the
1.78–1.74 Ga Medicine Bow Orogeny when the passive margin of the Wyoming Craton collided with island
arcs that accreted to the southern margin of Laurentia
(Karlstrom and Houston, 1984; Chamberlain, 1998).
The early Paleoproterozoic sequences of the Medicine
Bow Mountains experienced greenschist facies grade
of metamorphism reaching biotite isograde and are
preserved in broad open folds (Houston et al., 1981).
The maximum age of the supergroup is constrained by
U–Pb TIMS dating of detrital zircons from the Magnolia Formation in the Sierra Madre to <2460 Ma (Premo
and Van Schmus, 1989). The Snowy Pass Supergroup
either unconformably overlies or is in structural contact
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with Late Archean successions. The supergroup is subdivided by rotated thrusts into the Deep Lake Group,
and the lower and upper parts of the Libby Creek Group
(Figs. 1 and 3; Houston et al., 1992). Both the Deep
Lake Group and the lower Libby Creek Group contain
glaciogenic sediments.
The Deep Lake Group and the lower Libby Creek
Group were deposited in rift and passive margin settings, whereas the upper Libby Creek Group is considered to reflect a transition from passive margin
to foredeep (Karlstrom et al., 1983) or dissection
of a mature passive margin (Bekker and Eriksson,
2003). The supergroup was intruded by mafic sills
and dikes that range in composition from tholeiitic
to calc-alkaline. Tholeiitic gabbros that intruded prefolded Deep Lake Group are likely related to rifting (Karlstrom et al., 1981) and were dated in the
Sierra Madre and Laramie Mountains at 2.09–2.01 Ga
(U–Pb TIMS zircon and baddeleyite ages, Premo and
Van Schmus, 1989; Snyder et al., 1995; Cox et al.,
Fig. 3. (A and B) Detailed maps of two areas in the Medicine Bow Mountains (after Houston and Karlstrom, 1992) with sampled localities of
Vagner carbonates shown. Smaller insets on Fig. 3A show the location of Wyoming in North America and Sierra Madre and Medicine Bow
Mountains with exposed early Paleoproterozoic sedimentary successions north of the Cheyenne Belt (C.B.) in Wyoming. Larger inset on Fig. 3A
(after Houston and Graff, 1995) shows location of Fig. 3A and B in the northern part of the Medicine Bow Mountains.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Fig. 3. (A and B) (Continued ).
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2000). Pegmatitic metagabbro intruding the Cascade
Quartzite in the Sierra Madre provides a minimum age
for the Deep Lake Group in the Medicine Bow Mountains of ca. 2.09 Ga (U–Pb TIMS zircon age, Premo
and Van Schmus, 1989).
The Snowy Pass Supergroup commences with
fluvial rift deposits, including quartz-pebble conglomerates with detrital pyrite and uraninite (Magnolia
Formation) and contains three glacial units sandwiched
between quartzites, metapelites, and a thin carbonate
unit that overlies the middle glacial horizon (Vagner
Formation). A transition from intracratonic rift to open
marine settings coincides with deposition of the Vagner
Formation (Karlstrom et al., 1983). The youngest
glacial diamictite is overlain by thick, mature, Alrich and hematite-bearing quartzites (Medicine Peak
Quartzite).
The Paleoproterozoic succession in the Sierra
Madre is strongly deformed and dismembered by a late
cataclastic event (Houston, 1993), but, based primarily on lithostratigraphic similarities, it is considered to
be correlative with the Snowy Pass Supergroup in the
Medicine Bow Mountains (Houston et al., 1992). Possible correlative horizons in these successions and in the
Huronian Supergroup include a ca. 2.45 Ga siliciclastic rift sequence with basal conglomerates containing
detrital uraninite and pyrite, three glacial diamictites,
carbonates above the middle glacial diamictite, Alrich hematite-bearing mature quartzites, and carbonate
units above the upper glacial diamictite with similar
13 C-enrichment (Fig. 1; Bekker et al., 2003).
3. Analytical methods
The carbonate-dominated sections were measured
and fresh carbonate samples lacking shear, foliation, or
fracture were collected from carbonate intervals. Samples collected for whole rock analysis were cleaned,
cut, and powdered; these are marked by asterisk in
Appendix A. Other samples were petrographically
characterized and least altered (lacking veins, discoloration, weathering rinds, and silicification) and
finest-grained portions of polished thick sections were
microdrilled with 1 mm diamond drills. Thin sections
contain xenotopic mosaics of anhedral, coarsely crystalline to very coarsely crystalline crystals of either
dolomite or calcite with pseudospar and veins and frac-
tures filled with quartz, very coarsely crystalline carbonate and muscovite. Whereas micrite and spar did
not survive through neomorphism and recrystallization
during diagenesis and metamorphism, gross primary
textures (e.g. laminations) are still present. Mn, Sr, and
Fe concentrations and all isotopic measurements were
completed and quantified using procedures outlined in
detail by Bekker et al. (2003). Samples with carbonate
content below 50% were not analyzed for major and
trace elements.
4. Huronian Supergroup
4.1. Espanola Formation
The post-glacial Espanola Formation outcrops
across a wide region in both the western and southern parts of the Huronian Basin. However, due to poor
exposure and erosional contacts with overlying units, a
complete section was only measured in a drill core from
Quirke Lake (proximal section; Fig. 4) and at Moose
Point in the Whitefish area (distal section; Fig. 5).
In addition, partial sections of the Espanola Formation were measured and sampled in proximal sections
(Geneva Lake, Echo Lake, and outcrops in the Quirke
Lake area; Figs. 6–8, respectively). In studied sections,
the Espanola Formation either rests unconformably on
Archean basement or comformably on the glacial diamictites of the Bruce Formation. When exposed, the
lower contact with the Bruce diamictite is sharply overlain by an up to 3 m-thick transgressive shale interval.
The thickness of the formation is highly variable over
the outcrop area ranging from as little as 15 m in proximal sections to over 565 m in distal sections. The upper
contact is either comformable and gradational with the
Serpent Formation or disconformable with the Gowganda Formation.
The Espanola Formation in the Quirke Lake area
(Fig. 2; see Appendix B for detailed description of
all studied sections) is informally subdivided – in ascending order – into limestone, siltstone, and dolostone
members (Young, 1973a). The formation thickens to
the south of the Murray Fault Zone with a notable
increase in thickness of the siltstone member. In the
southern area, the upper dolostone member is discontinuous; it is included in the siltstone member which
is overlain by a heterolithic member (Bernstein and
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175
Fig. 4. Measured section of the Espanola Formation from the Kerr-McGee Corp. drill core 150/1 collared in the Bouck Township 150. ␦13 C
values of carbonate and TOC are shown.
Young, 1990). The lower two members were deposited
in a subtidal settings whereas the heterolithic member
was tide- and storm-influenced (Bernstein and Young,
1990). Stromatolites and other evidence for biological activity are conspicuously lacking in this formation
except that crustose and columnar stromatolites were
found in loose blocks in the Denison Mine area, Quirke
Lake (Hofmann et al., 1980) and stratiform stromatolites in the Geneva Lake area (Fig. 9A).
The lower limestone member in proximal sections
includes, in addition to limestone, thin graded beds of
sandstone and siltstone with erosional bases and crossbedding, shales with convolute bedding, rare stratiform
stromatolites, intraformational flat-pebble conglomerates (Fig. 9B) that are sometimes confined to channels,
and rare meter-size slumped blocks. Synsedimentary
folding and faulting, soft sediment deformation, injection and flame structures are abundant in this member and indicate episodically high sedimentation rates
and slope instability. In addition, distal sections contain
laminations and climbing ripples, suggesting rapid accumulation.
The gradationally overlying siltstone member consists of cross-bedded and laminated argillaceous siltstone and shale with graded beds, ball-and-pillow
structures (Fig. 9C) and Bouma beds, intraformational conglomerates with limestone and shale pebbles, and thin carbonate beds and lenses. Shales at
the top of sandy beds in distal section have flaser
and wavy bedding and both symmetrical and asymmetrical wave ripples with internal cross-laminations.
Both this member and the overlying dolomite member
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Fig. 5. Measured section of the Espanola Formation on the Moose Point (see Appendix B for location) with ␦13 C variations shown. Note that
only the basal part of the heterolithic member is shown as the rest of the member is covered by water.
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Fig. 6. Measured sections of the undivided Espanola Formation at Localities 2 and 3 in the Geneva Lake area (see Appendix B for location)
with ␦13 C variations shown. Sawtooth pattern at the base of the Espanola Formation indicates unconformity. Due to poor outcrop only partial
section of the Espanola Formation was measured at Locality 2.
contain soft sediment deformation and injection structures. Synsedimentary faults overlain by disrupted beds
and 35 m thick intraformational breccia-conglomerate
with boulders up to 1 m in size are present at Moose
Point.
The uppermost dolomite member in the proximal sections consists of graded beds of carbonate,
shale, and intraclasts. Scour surfaces, “beach rosettes”,
abundant wave ripples with internal cross-laminations
(Fig. 9D), mudcracks, cross-bedding, rare fenestral
textures, and peloids indicate a shallow-water environment above wave base. The heterolithic member at Moose Point consists of upward-fining cycles
of sandstone, siltstone, shale, and thin impure carbonates. Small conglomerate channels, herringbone
cross-stratification, wave ripples, and trough and large
low-angle cross-bedding indicate a tidally influenced
shallow-water environment.
4.2. Serpent and Gowganda formations
Carbonates and calcareous siliciclastic beds have
been previously described from units that overlie the
Espanola Formation (e.g. Card et al., 1977; Long,
1976). For this study, carbonates of the Serpent Formation were sampled on highway 6 south of Loon Lake,
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Fig. 7. Measured section of the limestone member, Espanola Formation at Marble Point of Echo Lake (see Appendix B for location) with ␦13 C
variations shown.
100–200 m south of La Chance Drive (Appendix A).
Yellow carbonates form two massive and resedimented
layers with soft sediment deformation structures. The
lower layer is 1 m thick and contains in its upper part
shale and sandstone clasts and carbonate rip-ups in
sandy carbonate and sandstone matrix. Carbonate layers and lenses of the Serpent Formation were also sampled on the eastern side of the East Simpson Island
in the Bay of Islands. Carbonate lenses in laminated
shales of the Gowganda Formation were sampled on
the portage between Ishmael and Low Lakes. On close
inspection, most carbonates of both the Serpent and
Gowganda formations appear to be carbonate cements
in sandstone or post-sedimentary features likely related
to hydrothermal activity. The only known exception is
sedimentary carbonates of the Serpent Formation near
Loon Lake.
4.3. Geochemical data
All geochemical data for the Espanola Formation
and the overlying Serpent and Gowganda formations
are tabulated in Appendix A. Carbon and oxygen
isotope data for Espanola Formation limestone and
dolomite samples show a large range with consistently
negative carbon isotope values in all sampled sections
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Fig. 8. Measured section of the Espanola Formation in the Denison Mine area, Quirke Lake (see Appendix B for location). Note that due to poor
outcrop, sample heights and thickness of the members are approximate. However, relative stratigraphic position of samples is well established.
The siltstone member was not logged in this area due to limited outcrop.
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Fig. 9. Sedimentary structures in carbonate rocks of the Espanola Formation. A—stratiform stromatolites in the Geneva Lake area, coin is 3.5 cm in diameter; B—intraformational
conglomerate, Marble Point of Echo Lake, scale is 10 cm in length; C—ball-and-pillow structure in the siltstone member of the Espanola Formation, Moose Point area; D—ripple
marks with unidirectional cross-laminations and bundled upbuilding indicating wave activity, coin is 2 cm in diameter.
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Fig. 10. Sedimentary structures in the Vagner Formation. A—angular lonestones in coarse-grained subarkosic matrix, pen is 14 cm long; B—finely laminated carbonates at the base
of the section at the BE-C-95-8 locality, hammer is 70 cm long; C—flat-pebble conglomerate at the top of the section at the BE-C-95-8 locality, pencil for scale.
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varying from −6.7 to −0.8‰ V-PDB. Oxygen isotope values have a much larger range (∼10‰) than
those of carbon isotopes and are significantly depleted
in 18 O; all samples have ␦18 O values of less than
−10‰. Manganese, iron, and Sr contents in these samples also range widely. Total organic carbon (TOC)
contents are low in most samples from the Espanola
Formation, ranging from 0.03 to 1.23 mg C/g sample for the whole dataset with most samples having TOC content below 0.2 mg C/g sample. Carbon
isotope values of organic matter are highly variable.
Measured differences between organic and carbonate
carbon yield δ values of less than 26.6‰; the average for the best-preserved drill core samples is around
20‰.
Carbonates of the Serpent Formation in the Loon
Lake area have highly variable ␦13 C values ranging from −6.3 to +0.9‰ whereas strongly negative
␦18 O values show much smaller range (−21.0 to
−19.8‰; Appendix A). These samples have also high
Mn (4170–4406 ppm) and Sr (254–463 ppm) contents
and high Mn/Sr ratios (9–17.3).
5. Snowy Pass Supergroup, Wyoming
5.1. Vagner Formation
The Vagner Formation overlies the Cascade
Quartzite and is up to 800 m thick (Karlstrom et al.,
1983; Fig. 1). A three-fold subdivision is recognized,
including (1) a basal diamictite up to 300 m thick;
(2) a middle marble unit with variable thickness from
5 to 60 m; and (3) an upper quartz-rich phyllite and
fine-grained quartzite unit (Karlstrom, 1977; Karlstrom
and Houston, 1979). An unconformity at the base of
the Vagner Formation is inferred from the variable
thickness of the underlying Cascade Quartzite, the
presence of channels along the contact, and pebbles
of Cascade Quartzite in the basal Vagner Formation
(Karlstrom, 1977; Bekker, 1998). The diamictite includes angular to subangular clasts in a greenish to reddish coarse-grained subarkosic matrix (Fig. 10A). The
basal diamictite has been interpreted as a glaciomarine deposit, based on the presence of dropstones in
laminated phyllite, faint stratification in some conglomerate layers, and major element analyses of finegrained matrix from the diamictite (Sylvester, 1973;
Houston et al., 1981). The contact with the overlying carbonate unit was not observed in the field but
is considered to be conformable (Karlstrom et al.,
1983).
The marble unit consists of 1–5 cm-thick rhythmically interlayered massive, fine-grained, grey–blue
to brown carbonate and brown coarse-grained
siliciclastics-rich carbonate. Contacts between these
carbonates are transitional. These carbonates are recrystallized into a coarse to very coarse calcite mosaic.
Convolute bedding, thin laminations, normal grading,
small-scale hummocky cross-stratification, lensoidal
bedding, and flat-pebble conglomerate are restricted
to the upper part of this unit. The marble unit locally
includes basal rhythmites of sandstone–siltstone–
carbonate (e.g. BE-C-95-24 locality; Fig. 3A). Phyllite
with lonestones of granitic composition ranging from
1 to 5 cm in diameter is interlayered with carbonate
in some localities. The only relatively complete
section of carbonates (BE-C-95-8 locality; Fig. 3B)
is in close proximity to a mafic sill. This section has
a shallowing-upward trend; laminated carbonates
(Fig. 10B) with stratiform stromatolites and small
oölites occur at the base and grade upsection into massive and siliciclastics-rich carbonates with flat-pebble
conglomerate (Fig. 10C) at the top. Quartzites in the
overlying phyllite-quartzite unit contain climbing
ripples, plane and wavy bedding, and rare cross-beds
and ripples. Carbonate clasts in this unit are rarely
present; but neither warping nor piercing of layers
beneath clasts was observed.
Due to poor exposure, only partial sections of the
marble unit were measured and sampled (Fig. 11). Additional samples were collected across the whole outcrop area of the marble unit (Fig. 3A and B) from
siliciclastics-poor carbonates to test variability of stable isotope compositions. Samples from the marble
unit of the Vagner Formation consist of low-Mg calcites (Bekker, 1998; Appendix A) and have variable
contents of Mn (295–1461 ppm) and Sr (21–217 ppm),
and a wide range of Mn/Sr ratios (1.8–16). Oxygen isotope values are strongly depleted in 18 O ranging from
−20.9 to −14.5‰. Carbon isotopes are uniformly depleted in 13 C with ␦13 C values ranging from −3.8 to
−1.5‰. Analyzed samples have very low TOC content (0.05–0.3 mg C/g sample) which is highly 13 Cenriched (−19.6 to −11.0‰) resulting in small measured difference in ␦13 C values between carbonate and
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
183
Fig. 11. Two closely spaced partial sections (BW-series samples in Appendix B) of the marble unit, Vagner Formation sampled near the small
lake to the north of Dipper Lake (41◦ 22 23 N, 106◦ 20 39 W; see Fig. 3A for location) with ␦13 C variations shown.
organic phases. Two samples with relatively high Sr
contents and low Mn/Sr ratios were analyzed for Sr
isotopes; the lowest 87 Sr/86 Sr value of ∼0.7081 is significantly lower than those measured from the correlative Espanola Formation (0.71128, Veizer et al.,
1992).
5.2. Bottle Creek Formation, Snowy Pass Group,
Sierra Madre
The Bottle Creek Formation (sensu Houston et al.,
1992; previously named as Vagner Formation in Graff,
1978) is considered correlative with the Vagner to Heart
formations of the Medicine Bow Mountains based on
lithostratigraphic similarity (Fig. 1). It includes diamictite that is overlain successively by marble, quartzite,
a second diamictite, and quartzite (Houston et al.,
1992). Carbonate is confined to the western part of
the Sierra Madre (see Graff, 1978). The lower diamictite and greenish, fine-grained, finely-laminated marble, which is up to 30 m thick and interlayered with
sandstone and siltstone, are considered as glaciomarine offshore facies equivalents of the Vagner Formation in the Medicine Bow Mountains (Houston
et al., 1992). Samples were collected about 1.6 and
1.9 km to the east along the Big Sandstone Creek
from the intersection with the Rawlings Road (see
Appendix A).
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Like their Vagner Formation counterparts, oxygen
isotope values of marble samples from the Bottle Creek
Formation are also highly 18 O depleted, ranging from
−16.1 to −12.5‰ V-PDB. Most carbon isotope values
vary from −3.4 to −2.2‰ V-PDB.
6. Discussion
6.1. Diagenesis and metamorphism
It is clear from the highly recrystallized textures of
the Espanola, Vagner, and Bottle Creek formation carbonates as well as the significantly depleted 18 O values and elevated 87 Sr/86 Sr compositions (see Appendix
A; Veizer et al., 1992) that these lonesome early Paleoproterozoic carbonates have been altered by secondary processes, thereby hampering interpretation of
carbon isotope values. The low ␦18 O values, in particular, suggest equilibration with hot metamorphic or
diagenetic fluids (Veizer, 1983), which must have also
affected trace element concentrations. Similarly, the
low abundance of TOC and the decreased fractionation between inorganic and organic carbon isotope
compositions is likely related to metamorphic alteration of carbon, which can greatly alter 13 C abundances in TOC (Hayes et al., 1983). If correct, some
degree of 13 C depletion in carbonate may reflect either the addition of newly-formed phases subsequent
to organic matter degradation or metamorphic devolatization. However, only metamorphism of siliceous carbonates can result in 13 C depletion of carbonate through
Rayleigh distillation associated with decarbonation reactions (Valley, 1986; Baumgartner and Valley, 2001)
and the formation of certain metamorphic minerals
(e.g. wollastonite). The lack of these minerals in most
of our samples, which are predominantly pure carbonate, argues against a metamorphic explanation for the
low ␦13 C values of the Espanola, Vagner, and Bottle Creek carbonates. Furthermore, mass balance considerations suggest that “organic” additions to these
organic-lean carbonates should not affect the ␦13 C values by more than a per mil (cf. Pelechaty et al., 1996)
and argue for a primary origin of these negative ␦13 C
values. Lastly, the general lack of correlation between
13 C abundances in these carbonates and those of 18 O
or elemental compositions (Figs. 12 and 13; Appendix
A) support this notion.
Carbon isotope values of the least altered samples of the Espanola carbonates range from −4.0 to
−0.8‰, but stratigraphic trends are lacking except in
the Quirke Lake drill core. In this section we note a
poorly defined trend of 13 C-enrichment in carbonate
and 13 C-depletion in organic matter upsection. Carbon isotope values of Vagner carbonates also do not
show a stratigraphic trend (Fig. 11). In combination
with regional coverage provided by the rest of samples
and excluding the most altered samples, carbon isotope values of the Vagner carbonates vary from −2.6 to
−1.5‰.
6.2. Paleoproterozoic cap carbonates
Carbon isotope data for Espanola carbonates from
both distal and proximal parts of the Huronian Basin
and for the whole outcrop area of Vagner and Bottle Creek carbonates suggest that dissolved seawater
bicarbonate was moderately depleted in 13 C after the
second early Paleoproterozoic glacial event. A marine
rather than lacustrine (cf. Veizer et al., 1992) origin is
supported by: (1) the transgressive nature of the deposits; (2) the long range lithostratigraphic correlation
of these unique North American carbonate accumulations; (3) the presence of herringbone cross-bedding
recording tidal influence (Bernstein and Young, 1990);
and (4) the size of the basins as well as the inferred
tectonic setting for both units being a failed rift arm of
a triple junction extending to the ocean (Young, 1983;
Karlstrom et al., 1983). Insofar as all Late Archean and
early Paleoproterozoic carbonates record ␦13 C values
near 0‰ or higher (e.g. Bekker et al., 2003a), we relate
this Paleoproterozoic negative carbon isotope anomaly
to environmental changes in response to global glaciation. Support for this view comes from our proposed
correlation of the North American post-glacial carbonates with carbonates from the lower Duitschland Formation, South Africa (Bekker et al., 2001). The lower
Duitschland carbonate was deposited above glacial diamictite on an open-marine margin during oceanic
transgression and records very similar degrees of 13 Cdepletion in carbonates with ␦13 C values ranging from
−3.7 to 0.1‰, as well as 13 C-enrichment in organic
matter (Bekker et al., 2001). Other Paleoproterozoic
successions worldwide apparently do not contain the
record of the second glacial event and overlying cap
carbonate.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
185
Fig. 12. ␦13 C vs. ␦18 O values of carbonates of the Espanola Formation analyzed in this study; data are grouped by locality. Analyses for outcrop
and drill core samples from the Quirke Lake area are shown separately. Note that all datasets show no correlation between these two parameters.
Fig. 13. ␦13 C vs. ␦18 O scatter diagram for carbonates of the Vagner Formation. Note lack of correlation between these parameters.
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A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Our interpretation of Paleoproterozoic carbonates in
North America and South Africa as post-glacial “cap
carbonates” relies, in part, on comparison with known
Neoproterozoic occurrences. For example, both Neoproterozoic and Paleoproterozoic glaciogenic successions were deposited in similar tectonic settings. Both
glaciogenic intervals coincide with supercontinent rifting (Kenorland and Rodinia for the Paleoproterozoic
and Neoproterozoic eras, respectively) resulting in
thick packages of siliciclastic sediments including
glacial diamictites (Young, 1988; Hoffman, 1991),
and variable thicknesses of carbonate deposited during
post-glacial transgression. The post-glacial carbonates
of both intervals are typically subtidal deposits, recognized by the abundance of fine laminations, hummocky cross-bedding, flat-pebble conglomerates, and
graded event beds; many of them represent carbonate
formed on shelves that was reworked and deposited
in thin graded units further offshore (Kennedy, 1996).
The cap carbonates of the Neoproterozoic and Paleoproterozoic are also similar in their depletion of 13 C in
carbonates and 13 C-enrichment in co-existing organic
matter (e.g. Kaufman and Knoll, 1995; Kaufman et al.,
1997; Bekker et al., 2001).
In contrast, carbon isotope values of the Vagner,
Bottle Creek, and Espanola carbonates do not show
convincing stratigraphic or facies-related trends in
␦13 C values, whereas heavily-sampled thick carbonatedominated Neoproterozoic successions record rapid
changes in ␦13 C values with time (Kaufman et al.,
1991, 1997; Kennedy, 1996; Hoffman et al., 1998;
Halverson et al., 2002; Xiao et al., 2004); furthermore,
on Neoproterozoic carbonate platforms ␦13 C compositions are more negative on the slope and increase
across the shelf (Kennedy, 1996; James et al., 2001).
Locally, where the uppermost strata of Neoproterozoic carbonate platform were not eroded below the
diamictites, carbon isotope values show a trend upsection from highly positive to highly negative values (Kaufman and Knoll, 1995; Kaufman et al., 1997;
Halverson et al., 2002; McKirdy et al., 2001). These
isotopic trends are less obvious in Neoproterozoic
siliciclastics-dominated sequences where the isotopic
records are truncated. Most (but not all) of the Neoproterozoic cap carbonates are notably dolomitic at their
base and are overlain by limestones that formed during maximum transgression, or by siliciclastic rocks,
depending on sediment supply and other factors; the
dolomite-limestone boundary is marked by the presence of a thin barite interval in Mackenzie Mountains, northwest Canada (Hoffman and Schrag, 2002).
Above this limestone interval, Neoproterozoic cap carbonates become increasingly dolomitic as accommodation space filled and the basins shallowed. On the
other hand, the Paleoproterozoic cap carbonates of
North America and South Africa begin with deep-water
limestone above thin transgressive deep-water shale.
As relative sea-level fell, the limestone accumulations
were overlain by shallow-water dolomite or siliciclastic
sediments.
Furthermore, the Paleoproterozoic post-glacial carbonates lack some of the notable sedimentary structures
recognized in many Neoproterozoic examples, including: (1) giant wave ripples (Allen and Hoffman, 2005),
(2) pseudomorphs after aragonite fans, (3) finelylaminated inversely-graded micro- and macropeletal
dolomites with roll-up structures, (4) sheet-crack cements, and (5) tube-like stromatolites (e.g. Hoffman
and Schrag, 2002; Corsetti and Kaufman, 2003). Some
of these structures were related to precipitation from
seawater oversaturated with carbonate alkalinity and
Mg (James et al., 2001; Corsetti and Kaufman, 2003)
and seismicity during isostatic rebound following the
glaciation (Nogueira et al., 2003). Whereas some of
these structures may have been present in the Paleoproterozoic cap carbonates and are now obscured by
recrystallization, others were likely not developed at
all. Some of the sedimentary structures common to
Neoproterozoic cap carbonates but lacking in Paleoproterozoic cap carbonates are likely related to high rates
of inorganic carbonate precipitation in shallow-marine
conditions from the post-glacial ocean supersaturated
with carbonate alkalinity and Mg, but with low oxygen
and sulfate contents. Giant wave ripples might reflect
the higher storm and cyclone intensity in response to
higher surface temperature gradients in the aftermath
of the snowball Earth (Emanuel, 1987; Hoffman and
Schrag, 2002).
Extreme alkalinity in the Neoproterozoic postglacial ocean has been related to ocean overturn at the
end of glaciation (e.g. Kaufman et al., 1991; Grotzinger
and Knoll, 1995; Hoffman et al., 1998) and high rates
of carbonate weathering on continents (Hoffman and
Schrag, 2002). Notably, some of the unusual sedimentary structures in the Neoproterozoic cap carbonates
(i.e. pseudomorphs after aragonite fans, cement-filled
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
sheet cracks, marine cement laminae, and roll-up structures) are also common on Late Archean carbonate
platforms and have been related to high carbonate alkalinity in the Archean ocean (Simonson et al., 1993;
Sumner and Grotzinger, 2000). Their apparent absence in the Paleoproterozoic cap carbonates, suggesting lower overall oceanic carbonate saturation in
the glaciogenic interval, might reflect that either atmospheric pCO2 levels were low, thereby affecting
continental weathering, or high, causing oceanic pH
to decline and carbonate to dissolve. Studies of the
chemical index of alteration of mature siliciclastic
sediments between diamictites in the Huronian and
Snowy Pass supergroups (Sylvester, 1973; Houston
et al., 1981; Nesbitt and Young, 1982; Bekker, 1998)
suggest extreme weathering of continental protolith.
This process ultimately produces carbonate alkalinity
in the oceans through reaction with atmospheric carbon dioxide. Due to the general absence of carbonates
in underlying Paleoproterozoic successions, continent
derived alkalinity could have only been provided by
silicate weathering, which apparently led to generally undersaturated conditions in the Paleoproterozoic
post-glacial oceans and may explain the absence of
the unusual sedimentary structures in the Espanola
and equivalent cap carbonates. On the other hand, if
atmospheric CO2 was very high, silicate weathering
might have delivered ample alkalinity to seawater but
the induced drop in oceanic pH and the rise in the
carbonate compensation depth would delay carbonate
accumulation.
6.3. A depositional model for the Paleoproterozoic
cap carbonates of North America
In our depositional model, post-glacial flooding of
alkalinity-laden seawater temporally shut down siliciclastic influx leading to accumulation of carbonate on
a relatively flat siliciclastic shelf. Based on the general
lack of traction-induced sedimentary structures except
near their tops, Paleoproterozoic cap carbonates in
North America were deposited below wave base
as reworked carbonates originally inorganically
precipitated in shallower waters; this interpretation is
supported by the scarcity of stromatolites and oncolites
in both distal and proximal settings. The abundance of
carbonate breccia in both distal and proximal sections
of the Espanola Formation suggests that the carbonate
187
platform rapidly developed an oversteepened rim, thus
shedding carbonate debris into deeper environments
due to high rates of carbonate accumulation. The
buildup of the platform caused a relative sea-level
fall allowing for siliciclastic deposition to resume
in marginal marine and, later, in deltaic and fluvial
settings (Karlstrom et al., 1983). At this point,
some margins returned to ramp-like morphology,
which favored synsedimentary to early diagenetic
dolomitization of the overlying delta plain deposits
(dolomite member of the Espanola Formation), which
were overlain by prograding fluvial sandstones of the
Serpent Formation, the Huronian Supergroup (Long,
1976).
6.4. Coupled early Paleoproterozoic change of the
atmospheric redox state and glaciations
Recent geochemical models of the early Paleoproterozoic glaciation have focused on the oxidative conversion of atmospheric methane, a remarkably strong
greenhouse gas, to carbon dioxide, which is more transparent to infrared radiation and hence less efficient as
a greenhouse gas, during the rise of atmospheric oxygen (Pavlov et al., 2000). This atmospheric transition
may have resulted in surface refrigeration at a time of
lower solar luminosity (∼87% of modern), but stratigraphic and sedimentologic evidence argues for three
discrete glacial events separated by long time intervals
with greenhouse conditions. A fully integrated explanation for these ice ages then must account for three
glacial events in the early Paleoproterozoic, as well as
the unusual appearance of “cap carbonates” above only
the middle diamictite. We hypothesize that the glacial
events were driven by a misbalance between oxygen
production (through photosynthesis) and consumption
(through oxidation of reduced compounds in seawater),
and its effects on atmospheric CH4 and CO2 concentrations; with their changes directly affecting surface
temperatures.
Since methane likely acted as the dominant greenhouse gas before the rise of atmospheric oxygen, the
glacial epoch might represent climatic adjustments during the transition to an atmosphere with CO2 acting
as a major greenhouse gas. Empirical constraints on
the atmospheric CO2 and CH4 contents across the
early Paleoproterozoic glacial epoch, however, are virtually nonexistent. The single geochemical estimate for
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atmospheric CO2 content around this time is based on
the absence of siderite in paleosols older than 2.2 Ga,
which suggests an upper limit of 10−1.4 atm (∼100
PAL; Rye et al., 1995). Using the same constraint in a
one-dimensional, radiative-convective climate model,
Pavlov et al. (2000) suggested that Archean atmosphere
should have contained up to 1000 ppmv methane to
overcome significantly lower solar luminosity at that
time. There are no constraints on atmospheric CO2
content in the Paleoproterozoic after the rise of atmospheric oxygen but since the methane contents should
have decreased to less than 100 ppmv in an oxidizing atmosphere (Pavlov et al., 2003), atmospheric
pCO2 must have increased dramatically to keep the
Earth from freezing over again. Accepting that oxidation of atmospheric methane to carbon dioxide resulted in each of the three discrete glaciations requires
that atmospheric pCH4 remained high throughout the
interval.
Constraints on the atmospheric oxygen level across
the Paleoproterozoic glacial epoch may ultimately provide a test for the Pavlov et al. (2000) model. The correlation between Paleoproterozoic successions of North
America and South Africa (Fig. 1) is based on their
stratigraphic and temporal relationship with respect to
the 2.48–2.45 Ga tectonomagmatic event, three glacial
events, climatic amelioration after the last glacial event,
and >2.2–2.1 Ga carbon isotope excursion. If this correlation is accepted then we can combine constraints
for atmospheric oxygen level from these successions
(Fig. 14).
Geological and geochemical indicators of atmospheric redox state indicate that the atmosphere
remained anoxic across the 2.48–2.45 Ga tectonomagmatic event (Fig. 14). These include (1) reduced paleosols at the base of the Huronian succession; (2)
well-documented detrital grains of pyrite and uraninite in the Matinenda and Magnolia formations of the
Huronian and Snowy Pass supergroups, respectively;
(3) diagenetic sulfides with small range of ␦34 S values
and non-mass dependent fractionation of sulfur isotopes, and (4) the absence of a Ce anomaly in iron formations of the Ghaap Group, South Africa. Similarly,
units overlying the oldest glacial diamictites in both
North American successions contain radioactive and
pyritiferous conglomerates and sandstones (Roscoe,
1969; Meyn, 1970; Karlstrom and Houston, 1979a;
Meyn and Matthews, 1980, 1991; Mossman and
Harron, 1983; Long, 1987) suggesting that the atmospheric oxygen content remained low between the first
and second glacial events. It was recently confirmed
based on three S isotope study of pyrite grains from the
pyritiferous conglomerates and sandstones of the Missassagi Formation, Huronian Supergroup that at least
some of these grains are detrital in origin (Bekker et al.,
2005). This observation constrains the maximum level
of atmospheric oxygen after the first glacial event,
whereas sulfur isotope data for sulfides extracted from
shales and siltstones that sandwich the oldest Huronian glacial diamictite reveal a small range of 33 S
values from −0.06 to +0.27‰ V-CDT (Farquhar et al.,
2000; Wing et al., 2002, 2004), suggesting that immediately before the oldest Paleoproterozoic glaciation the atmosphere became more oxidizing than in the
latest Neoarchean (Farquhar et al., 2000; Ono et al.,
2003).
There are several lines of evidence suggesting significant amounts of oxygen in the atmosphere between
the second and third Paleoproterozoic glacial events including hematitic ironstone, highly negative ␦34 S values, and the lack of non-mass dependent fractionation
in S isotopes in South Africa (Bekker et al., 2004a,
2004b). Puzzlingly, the Noomut Formation, below the
glaciogenic Padlei Formation of the Hurwitz Group,
Canada that is considered correlative with the upper
glacial diamictite of the Huronian Supergroup contains variegated beds and detrital pyrite (Aspler and
Chiarenzelli, 1997). Carbonates of the upper Duitschland Formation, South Africa that were also deposited
between the second and third glacial events have highly
positive ␦13 C compositions, suggesting an excess in organic carbon burial and a release of photosyntheticallyproduced oxygen to the atmosphere (Bekker et al.,
2001). Units overlying the third and youngest Paleoproterozoic glacial diamictite contain multiple evidence
for the oxygenated atmosphere (Fig. 14).
Put together, these observations indicate that pCO2
rose dramatically during the early Paleoproterozoic
glacial epoch, but likely fluctuated across the three
glaciations. We speculate here for a stepwise increase
in atmospheric oxygen across these climatic events
that (1) commenced the glaciations through oxidation
of atmospheric methane, and (2) explains why cap
carbonate is present only above the middle of three
glacial diamictites. The lack of carbonate platforms
after the ca. 2.5 Ga and before the beginning of the
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
189
Fig. 14. Indicators of atmospheric and ocean redox state (green—reduced; red—oxidized) arranged with respect to 2.48–2.45 Ga tectonothermal
events, three glaciogenic events (triangles, question marks indicate uncertainty on their age), and carbon isotope variations. Carbon isotope
composition of the early Paleoproterozoic carbonates (grey curve) is from Karhu and Holland (1996) with additional data (dashed bold line)
from Bekker et al. (2001) and this paper. Indicators of atmospheric and ocean redox state are organized in somewhat subjective way from the
most sensitive to the oxygen level at the bottom to the least sensitive at the top. The Paleoproterozoic glacial epoch is bracketed between 2.42
and 2.22 Ga; 2.32 Ga Re–Os isochrone age of pyrites from black shales deposited before the last glacial event (Hannah et al., 2004) provide
another age constraint for these glacial events. References: detrital pyrite, siderite, and uraninite (Roscoe, 1969, 1981, 1996; Roscoe and Minter,
1993; Meyn, 1970; Meyn and Matthews, 1980, 1991; Karlstrom and Houston, 1979a; Mossman and Harron, 1983; Long, 1987; Coetzee, 2001;
Bekker et al., 2005); red/variegated beds, copper stratiform deposits (Wood, 1979; Rainbird and Donaldson, 1988; Chandler, 1989; Houston et
al., 1992; Aspler and Chiarenzelli, 1997; Dorland, 1999; Beukes et al., 2002); ␦34 S record (Cameron, 1982, 1983; Hattori et al., 1983); 33 S
record (Farquhar et al., 2000; Bekker et al., 2004a, 2004b; Wing et al., 2002, 2004); paleosols (Rainbird et al., 1990; Marmo, 1992; Rye and
Holland, 1998; Panahi et al., 2000; Beukes et al., 2002a; Yang and Holland, 2003); sulfate evaporites (Chandler, 1988); Mn deposits (Beukes,
1983; Tsikos and Moore, 1998); Ce anomaly (Alibert and McCulloch, 1993; Bau et al., 1998; Murakami et al., 2001; Maynard, 2004). (For
interpretation of the references to colour in this figure legend, the reader is referred to the web version of the article)
Paleoproterozoic glacial epoch might be related to the
supercontinent assembly. In this case, enhanced chemical weathering would have progressively drawn down
atmospheric CO2 (cf. Nance et al., 1988) eventually
attenuating alkalinity fluxes, whereas terrestrial silici-
clastic fluxes would increase precluding carbonate accumulation. The resulting methane-rich and extremely
CO2 -poor atmosphere would have been highly unstable as photosynthetically-produced oxygen started to
build up in the atmosphere. It is possible that decrease
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in hydrothermal and terrestrial flux of reduced compounds during the supercontinent assembly allowed
oxidation of the surface environments. As the cumulative greenhouse power of the atmospheric methane and
carbon dioxide fell below threshold limits, global temperatures dropped and the first glaciation ensued. No
carbonate was deposited after the first Paleoproterozoic
glaciation because of persistent carbonate undersaturation. Assuming that volcanic CO2 and biogenicallyderived methane continued to be released to the atmosphere during and after the first glaciation, surface
temperatures would have warmed and enhanced silicate weathering and continental fluxes of alkalinity. A
following increase in the photosynthetically-produced
oxygen flux resulting in drawdown of atmospheric
methane and decrease in the atmospheric CO2 levels
due to chemical weathering led to the second glaciation. In this case, carbonate alkalinity fluxes were sufficient to result in the formation of carbonate in the
aftermath of glaciation. A further reduction of the atmospheric methane conceivably associated with the positive carbon isotope excursion in seawater composition
recorded in the upper Duitschland Formation carbonates (Bekker et al., 2001) required even more elevated
atmospheric CO2 levels to overcome low albedo effect
and to warm surface temperatures in the aftermath of
the third glaciation. Carbonate deposition was apparently delayed as evidenced by the absence or scarcity
of carbonates associated with deeply weathered sediments above the ultimate Paleoproterozoic tillite until
carbonate alkalinity levels in the ocean were sufficient to counteract dissolution induced by extremely
high atmospheric CO2 levels. We therefore infer that
the atmospheric CO2 levels in the aftermath of the
second glaciation reached a level sufficient to provide enough carbonate alkalinity to seawater through
chemical weathering but not so high that carbonate preservation would be inhibited by higher overall
acidity.
Carbonate again becomes abundant in the marine
sedimentary record during the 2.22–2.10 Ga positive
carbon isotope excursion. We speculate that warming
of the oceans and enhanced biological uptake at that
time (Karhu and Holland, 1996) drew down oceanic
CO2 so that alkalinity levels in the ocean were sufficient to counteract dissolution. The concurrent rise
of pO2 associated with the ca. 2.22–2.10 Ga carbon
isotope excursion likely swept most of the remaining
methane from the atmosphere, leaving CO2 as the dominant greenhouse gas thereafter (cf. Kaufman and Xiao,
2003).
We conclude that although Neoproterozoic and Paleoproterozoic glacial epochs have many similarities,
the early Paleoproterozoic glacial epoch has certain
unique characteristics that likely reflect the transition from an anoxic methane-rich atmosphere to the
oxygenated atmosphere with carbon dioxide being a
dominant greenhouse gas. This speculative model of
stepwise or pulsed atmospheric oxidation provides
a plausible explanation for multiple early Paleoproterozoic glaciations, the general absence of carbonate
throughout the glaciogenic interval, and the singular
presence of a cap carbonate in the middle of the glacial
epoch.
Acknowledgements
This project started when A.B. was a MS student at the University of Minnesota, Duluth under
the direction of R.W. Ojakangas. D. Long, G. Bennett, G.M. Young, and P.J. Graff helped in field
and with logistics. D. Long pointed to outcrops in
the Geneva Lake area with stratiform stromatolites.
K. Chamberlain provided helpful discussions on the
tectonic setting and geochronology of the Snowy
Pass Supergroup. D. Rimstidt provided access to and
helped with AAS work. Support came from Department of Geology, University of Minnesota, Duluth; GSA and southeastern section of GSA; Colorado Scientific Society; and the graduate school and
Department of Geological Sciences, VPI and SU.
Funding for participation by AJK and isotopic analyses at the University of Maryland was provided
by NSF grants EAR-98-17348 and EAR-0126378
and NASA Exobiology grant NAG 512337. JAK
thanks the Geological Survey of Finland for permission to publish data produced in their laboratories
and A. Henttinen for assistance. AB gratefully acknowledges financial support from PRF/ACS 270953
(37194-AC2) grant to H. D. Holland, Harvard University. Constructive reviews by R. H. Rainbird and
an anonymous reviewer are gratefully acknowledged.
AB dedicates this paper to fond memories of Olga
M. Haring.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
191
Appendix A. Supplementary materials
Carbon, oxygen, and strontium isotope values and trace and major element contents of studied carbonates
Sample number
Height
above
base (m)
Espanola Formation
Geneva Lake
Locality 1
GE-4
4.2
GE-5
8.7
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
Calc. si-ne
Calc. si-ne
−5.8
−18.6
Calcite
Calcite
Calcite
Cal. sa-ne
Calcite
Calcite
Calcite
Calcite
−1.3
−0.8
−1.5
−1.3
−2.1
−1.0
−0.9
−14.8
−14.6
−14.8
−12.8
Locality 2
GE-8
GE-9
GE-11
GE-12
GE-13
GE-14
GE-16
GE-17
0.0
1.6
3.0
3.3
5.3
7.6
9.2
10.2
Locality 3
GE-22
GE-23
GE-24
GE-25
2.7
3.2
5.4
5.7
Calcite
Calcite
Calcite
Calcite
−1.0
−1.2
−1.4
−2.3
−16.6
−11.7
−14.1
−11.7
Echo Lake
EL-5
EL-6
EL-7
EL-8
EL-9
EL-10
EL-11
EL-12
EL-13
EL-14
EL-15
EL-16
EL-17
EL-18
EL-19
EL-20
EL-21
EL-22
EL-23
EL-24
EL-25
EL-26
El-27
EL-28
EL-29
EL-30
EL-31
EL-32
EL-33
EL-34
EL-34A
1.0
2.2
2.7
3.7
5.2
6.7
8.2
9.7
11.0
12.0
13.0
14.5
16.0
17.2
18.0
19.5
21.0
22.5
24.0
25.5
27.0
28.5
30.0
31.0
32.0
33.5
35.6
37.7
39.8
40.0
40.0
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−2.3
−2.0
−2.1
−2.2
−2.6
−1.9
−2.0
−2.1
−1.7
−1.6
−2.1
−2.2
−1.9
−2.1
−2.0
−1.9
−2.1
−1.8
−1.8
−2.0
−2.1
−2.2
−2.3
−2.2
−2.0
−1.9
−2.0
−1.2
−2.1
−2.0
−1.1
−14.4
−15.2
−15.3
−15.7
−16.5
−15.4
−16.0
−15.0
−15.7
−16.4
−13.9
−16.8
−16.1
−15.9
−14.6
−13.1
−14.6
−14.1
−14.2
−14.6
−14.0
−11.8
−14.8
−12.7
−12.2
−15.1
−15.7
−14.5
−14.5
−14.8
−14.1
−11.1
−12.9
−12.7
TOC,
mg C/g
sample
␦13 CTOC ,
‰ V-PDB
0.08
−11.7
δ
Mn
(ppm)
Sr
(ppm)
2053
tr.
1301
1378
1987
50
tr.
4
527.2
19.3
363
226
182
437
240
118
113
20
1.5
1.9
1.6
22.2
10396
9963
6047
8092
tr.
40
9
tr.
248.8
687.4
261
166
188
165
154
223
281
173
337
99
262
248
361
277
130
174
251
65
2.6
0.6
0.8
0.5
0.6
1.7
1.6
0.7
5.1
294
221
205
235
221
332
212
171
163
176
83
319
234
93
458
153
168
141
172
208
151
143
156
224
329
154
98
165
198
191
185
114
68
168
57
46
209
280
316
327
200
372
2.1
1.4
0.9
0.7
1.4
3.4
1.3
0.9
0.9
1.0
0.7
4.7
1.4
1.6
10.0
0.7
0.6
0.4
0.5
1.0
0.4
0.17
0.12
0.56
0.05
−21.9
19.6
0.06
−15.7
13.4
0.05
0.08
0.21
0.11
0.07
0.07
0.30
0.06
0.38
0.11
0.11
−15.5
13.4
−23.0
20.4
−20.2
−14.5
−17.8
−23.6
−19.2
−11.1
−15.8
−22.3
−17.9
18.0
12.9
16.2
21.6
17.0
9.2
13.6
20.3
16.0
−16.1
14.3
−12.8
−20.6
10.7
18.4
−12.4
10.2
−15.7
−10.3
−18.2
13.8
8.4
17.1
−12.2
−21.2
10.2
20.1
0.08
0.07
0.11
0.12
0.07
0.10
0.07
0.06
0.11
Mn/Sr
25.8
−22.2
−20.2
0.20
0.12
0.04
0.07
0.09
Fe
(ppm)
87
Sr/86 Sr
192
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Appendix A (Continued )
Sample number
EL-35
EL-35A
EL-36
El-37
El-38
EL-39
EL-40
EL-41
EL-42
EL-42A
EL-43
EL-44
EL-45
EL-46
EL-47
EL-48
EL-49
EL-50M
EL-50C
El-51
EL-52
EL-53
EL-54
El-55
EL-56
EL-57
EL-58
EL-59
EL-60
EL-61
El-62
EL-63
EL-64
EL-65
Height
above
base (m)
42.0
42.0
43.5
44.3
45.0
47.0
48.1
49.6
51.1
51.2
51.6
53.1
54.6
55.6
56.5
58.4
59.9
61.7
61.7
63.2
65.5
66.3
66.9
68.1
69.6
71.1
72.6
74.1
75.6
77.4
78.9
80.4
81.9
83.4
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
TOC,
mg C/g
sample
␦13 CTOC ,
‰ V-PDB
δ
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−2.1
−2.0
−1.6
−2.1
−1.6
−2.1
−1.9
−2.1
−2.0
−1.6
−2.1
−1.8
−1.9
−2.1
−2.2
−1.7
−2.0
−2.0
−1.7
−1.7
−1.7
−1.5
−2.1
−2.1
−1.9
−1.8
−1.5
−2.2
−2.1
−2.1
−1.7
−2.0
−2.1
−2.3
−14.4
−15.4
−14.0
−14.7
−13.5
−13.7
−14.0
−13.6
−14.9
−14.1
−14.4
−13.7
−14.3
−14.3
−14.1
−14.3
−15.8
−14.2
−14.4
−14.9
−13.5
−14.4
−13.9
−15.8
−15.7
−15.9
−14.1
−15.6
−15.8
−15.5
−15.5
−15.0
−13.9
−15.9
0.05
0.10
0.05
0.20
0.04
0.06
0.06
0.06
0.08
0.07
0.05
0.06
0.08
−14.4
12.3
−12.0
10.3
−12.5
10.9
−11.4
9.3
−8.3
6.5
−10.0
−8.7
8.3
6.7
−23.1
21.4
−8.6
6.7
−8.4
6.8
−22.4
−15.7
20.3
13.5
−12.9
10.9
0.05
0.11
0.09
0.11
0.06
0.16
0.06
0.05
0.06
0.08
0.06
0.06
0.14
0.17
0.20
0.15
0.07
0.05
Quirke Lake (Kerr-McGee Corp. Drill core 150/1, Bouck Township 150, 46◦ 27 41 , 82◦ 36 04 )
Dolostone Member
DM-1722
524.9
m. calcite
−2.1
−12.9
0.11
DM-1735
528.8
m. dolomite −1.9
−14.0
0.24
−23.9
DM-1748
532.8
m. calcite
−2.5
−16.4
0.49
−23.3
DM-1749
533.1
m. calcite
−2.6
−16.7
0.1
−19.4
DM-1755 A 534.9
m. calcite
−2.7
−16.5
0.11
DM-1765
538.0
m. dolomite −2.1
−12.3
0.23
−26.1
DM1772
540.1
m. calcite
−2.2
−15.1
0.12
−14.8
DM-1782 6 543.2
m. calcite
−2.9
−15.5
0.14
−19.1
DM-1791
545.9
m. calcite
−2.2
−13.0
0.1
DM-1796
547.4
m. calcite
−2.5
−12.7
0.14
−17.0
DM-1806
550.5
m. calcite
−2.3
−12.0
0.29
−24.7
DM-1816
553.5
m. dolomite −2.2
−13.4
−13.7
DM-1829
557.5
calc. mud-ne
0.15
−23.6
DM-1842
561.4
m. calcite
−3.0
−15.9
0.45
−25.4
DM-1852
564.5
m. dolomite −2.5
−14.0
0.36
−24.8
DM-1862
567.5
m. calcite
−3.1
−15.5
0.3
−22.2
DM-1872
570.6
m. dolomite −2.2
−12.9
0.08
−27.5
DM-1882
573.6
m. dolomite −2.7
−15.9
0.33
−21.3
DM-1892
576.7
m. calcite
−3.3
−13.9
0.26
−14.3
DM-1902
579.7
m. dolomite −2.3
−12.1
0.23
−24.3
DM-1912 4 582.8
Calcite
−3.1
−16.5
0.25
−23.2
Fe
(ppm)
Mn
(ppm)
Sr
(ppm)
Mn/Sr
196
199
170
164
181
164
197
182
241
335
335
263
333
242
326
223
144
138
0.6
0.6
0.6
0.5
0.7
0.5
0.9
1.3
1.7
155
249
161
267
161
188
169
213
258
237
188
265
196
237
211
98
222
302
0.7
1
0.6
1.4
0.7
0.9
1.7
1.0
0.9
257
264
287
234
159
210
251
268
234
227
194
143
126
182
357
290
313
299
257
257
1.3
1.8
2.3
1.3
0.4
0.7
0.8
0.9
0.9
0.9
289
226
399
226
92
140
1.3
2.5
2.8
46
67
46
54
115
tr.
tr.
75
52
41
27
19
22.8
45.4
18.3
15.2
9.2
14.5
22.4
11.5
1036
3030
843
817
1060
895
696
760
2304
2026
1139
1584
22.4
22.2
19.2
25.3
18.6
11.0
22.0
20.1
2090
2792
934
1882
1023
1575
3370
2075
112
tr.
122
tr.
tr.
tr.
48
367
18.7
22.0
20.7
16.8
23.9
12.6
16.2
10.1
44.5
49.6
42.9
82.3
7.7
70.5
5.6
87
Sr/86 Sr
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
193
Appendix A (Continued )
Sample number
DM1917
DM-1919
DM-1929
DM-1934 6
DM-1944 6
DM-1959
DM-1968
Height
above
base (m)
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
584.3
584.9
588.0
589.5
592.5
597.1
599.8
m. dolomite
m. dolomite
m. dolomite
m. calcite
m. calcite
m. dolomite
m. dolomite
−3.9
−2.5
−3.3
−3.4
−3.6
−4.0
−3.5
−17.7
−12.9
−15.5
−16.2
−15.7
−16.2
−16.1
Siltstone Member
DM1977
602.6
DM1990
606.6
DM2000
609.6
DM2010
612.6
DM2020
612.6
DM2030
618.7
DM2040
621.8
DM2050
624.8
DM2060
627.9
DM2080
634.0
DM2111
643.4
DM2127
648.3
DM2138
651.7
DM2148
654.7
DM2151 10 655.6
DM2183
665.4
DM2193
668.4
DM2203
671.5
DM2210
673.6
DM2223
677.6
DM2227
678.8
DM2240
682.8
∗
DM2250
685.8
DM2283
695.9
DM2302
701.6
DM2308 7 703.5
Calc. si-ne
Graywacke
Calc. si-ne
Calc. si-ne
Calc. si-ne
Mud-ne
Calc. si-ne
Mud-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Mud-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne
Mud-ne
Mud-ne
Calc. si-ne
Calc. si-ne
Calc. si-ne −3.0
Calc. si-ne
Calc. si-ne
Limestone Member
DM2317 4 706.2
DM2318 8 706.5
DM2323 3 708.1
DM2328
709.6
DM2332 9 710.8
DM2339
712.9
DM2341
713.5
DM2343 9 714.1
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−2.9
−3.5
−1.5
−2.2
−1.8
−2.9
−2.4
−2.7
−16.5
−18.3
−14.3
−18.8
−17.1
−19.3
−18.8
−19.7
Denison Mine, Quirke Lake
Limestone Member
QE-1A
0.0
QE-2
6.5
QE-3
7.5
QE-4
9.0
QE-5
10.5
QE-6
17.0
QE-7
22.0
QE-8
27.0
QE-9
30.0
QE-10
36.0
QE-11
42.0
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−1.4
−3.9
−2.1
−1.2
−1.0
−1.4
−0.9
−1.4
−1.7
−1.1
−1.7
−19.3
−20.1
−20.5
−18.7
−19.2
−19.2
−18.9
−18.0
−17.2
−15.5
−16.4
−18.0
TOC,
mg C/g
sample
␦13 CTOC ,
‰ V-PDB
δ
0.12
0.1
0.59
0.59
0.27
0.49
−21.0
18.6
−23.9
−24.9
−22.8
−22.6
20.5
21.3
18.8
19.2
0.20
0.21
0.25
0.19
0.17
0.41
1.23
0.11
0.63
0.21
0.10
0.15
0.13
0.06
0.14
0.12
0.14
0.12
0.18
0.07
0.09
0.07
0.11
0.35
0.23
0.15
−23.4
−26.8
−15.0
−17.6
−24.7
−12.7
−18.2
−16.1
−13.8
−17.7
−17.8
−17.0
−11.5
−15.1
−19.1
−14.1
−16.7
−17.3
−24.4
−11.7
−15.8
−17.2
−12.8
0.11
0.14
0.06
0.16
0.19
0.07
0.09
0.16
Fe
(ppm)
Mn
(ppm)
Sr
(ppm)
3086
589
399
1265
744
792
83
tr.
tr.
115
tr.
63
475
tr.
Mn/Sr
37.0
11.0
12.6
−11.9
−15.0
−11.4
−15.9
7.9
14.4
1342
106
12.6
−21.9
−29.5
20.2
26.6
2162
210
10.3
−17.3
14.6
1767.0
3168.8
1759.1
1830.6
1796.0
1346.5
1169.2
625.4
516.4
332.7
307.2
308.8
413.6
117.6
241.5
250.4
85.3
82.3
117.7
146.4
62.2
174.5
5.7
7.7
15.0
7.6
7.2
15.8
14.2
5.3
3.5
5.3
1.8
0.60
0.11
0.26
−31.6
−21.4
−21.2
30.6
20.0
20.3
0.16
0.21
−19.6
−20.2
17.9
19.1
87
Sr/86 Sr
194
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Appendix A (Continued )
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−1.5
−1.4
−1.5
−2.1
−1.4
−1.5
−1.5
−1.5
−1.3
−2.9
−2.0
−1.0
−15.2
−15.4
−15.1
−15.2
−14.5
−14.5
−14.9
−15.2
−14.8
−17.1
−15.6
−11.9
Dolostone Member
QE-24
0.0
QE-25
3.0
QE-26
5.0
QE-27
8.0
QE-29
16.5
QE-30
23.5
QE-31
24.0
QE-32
24.5
QE-33
27.1
QE-34
28.4
QE-35
30.1
QE-37
39.0
QE-38
50.4
QE-39
52.7
QE-40
57.1
QE-41
59.0
QE-42
60.0
QE-44
65.0
Esp-1-3∗
Esp-1-9∗
Esp-2-1∗
Esp-3-1∗
Esp-5-1∗
Calcite
Dolomite
Calcite
Calcite
Calcite
Calcite
Dolomite
Dolomite
Calcite
Dolomite
Calcite
Calcite
Calcite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
Dolomite
−3.0
−2.5
−2.7
−2.7
−2.0
−2.8
−2.4
−2.9
−2.3
−2.5
−2.8
−2.7
−2.3
−2.2
−2.0
−2.3
−2.5
−1.8
−2.5
−2.2
−2.1
−2.1
−2.5
−18.5
−14.3
−13.0
−13.9
−14.8
−16.4
−15.9
−18.2
−14.4
−17.8
−16.0
−15.6
−17.7
−13.5
−13.4
−16.5
−15.1
−14.9
−16.5
−16.7
−16.0
−15.7
−13.6
Moose Point
MP-4
MP-7
MP-9
MP-10
MP-11
MP-12
MP-27
MP-28
MP-29
MP-30
MP-31
MP-32
MP-33
MP-34
MP-35
MP-36
MP-38
MP-39
MP-40
Cal. si-ne
Dolomite
Calcite
Dolomite
Dolomite
Dolomite
Calcite
Calcite
Calc. si-ne
Calc. si-ne
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calc. si-ne
Calcite
Calcite
−3.2
−3.1
−3.1
−2.8
−3.7
−3.6
−4.2
−12.9
−13.5
−14.5
−14.1
−15.1
−16.4
−18.2
Sample number
QE-12
QE-12A
QE-13
QE-13A
QE-14
QE-15
QE-16
QE-17
QE-18
QE-20
QE-21
QE-22
Height
above
base (m)
46.0
50.8
52.0
52.0
53.1
54.3
55.3
56.3
57.0
80.0
79.0
77.8
2.1
5.1
7.1
8.1
9.1
10.1
140.0
148.5
155.7
159.0
170.6
178.8
182.2
191.0
194.0
198.7
204.9
207.4
212.3
−5.1
−6.7
−3.7
−3.5
−3.5
−3.9
−5.2
−3.9
−3.7
−18.1
−18.0
−17.8
−15.9
−16.8
−16.1
−17.6
−15.8
−15.4
TOC,
mg C/g
sample
0.18
0.04
0.20
0.03
0.19
␦13 CTOC ,
‰ V-PDB
δ
−18.4
16.3
−14.6
13.1
Fe
(ppm)
0.17
0.36
0.09
0.11
43100
38000
37800
34900
0.07
0.06
0.13
−16.0
−16.8
−21.4
13.6
18.3
0.46
−17.4
13.7
0.08
0.13
−17.3
−18.4
0.07
0.09
Mn
(ppm)
Sr
(ppm)
Mn/Sr
282.4
191.7
220.8
199.3
179.0
210.4
249.6
281.4
183.4
641.7
119.0
1882.7
161.0
184.6
194.3
209.6
174.7
172.1
156.8
180.8
235.8
32.4
257.1
tr.
1.8
1.0
1.1
1.0
1.0
1.2
1.6
1.6
0.8
19.8
0.5
677.3
1882.7
3566.1
2642.0
3762.1
2099.3
3819.8
1238.0
2963.2
3360.7
5171.9
661.7
1534.7
4077.2
4707.0
2294.5
2649.9
606.0
118.0
tr.
65.8
38.1
113.1
307.9
55.6
25.7
55.5
22.8
101.0
tr.
tr.
56.1
67.2
tr.
48.2
tr.
5.7
5825.8
5289.0
4685.4
5044.0
186.2
182.1
153.9
131.9
5935
5005
5882
4931
5049
2609
1852
1
10
11
19
27
233
125
8551.3
520.5
515.6
260.1
187.4
11.2
14.8
1756
2771
2927
2970
3321
3358
2756
4261
5477
27
162
214
159
189
173
81
272
237
65.3
17.1
13.7
18.6
17.5
19.5
34.1
15.6
23.1
54.2
69.3
33.3
6.8
68.7
48.2
53.4
147.1
51.2
72.7
70.1
55.0
31.3
29.0
30.4
38.2
87
Sr/86 Sr
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
195
Appendix A (Continued )
Height
above
base (m)
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
216.4
227.3
238.6
241.0
265.0
267.8
270.0
274.1
278.1
297.0
Calcite
Calc. si-ne
Calcite
Calcite
Calc. si-ne
Calcite
Calcite
Calcite
Calcite
Calc. si-ne
−3.7
−15.9
−3.8
−3.8
−3.3
−3.0
−3.0
−3.0
−3.2
−3.1
−15.9
−16.0
−15.3
−15.2
−12.6
−15.9
−15.4
−16.4
Serpent Formation
SE-1-1∗
SE-2.3∗
SE-2.4a∗
SE-2.4b∗
SE-2.5a∗
SE-2.6∗
ES-1
ES-5
SE.2.1
SE.2.2
SA-98-1-A
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−2.9
0.7
−6.3
−1.0
−0.7
0.7
−8.0
−0.1
−2.5
0.9
−0.5
−20.8
−21.0
−21.0
−20.9
−20.7
−20.3
−16.2
−7.3
−20.2
−19.8
−20.3
Gowganda Formation
IS-GO
Calcite
−8.8
−9.7
Vagner Formation
BE-C-95-1∗
BE-C-95-2∗
BE-C-95-3∗
BE-C-95-4∗
BE-C-95-7∗
BE-C-95-8∗
BE-C-95-9∗
BE-C-95-10a∗
BE-C-95-10b∗
BE-C-95-11∗
BE-C-95-12∗
BE-C-95-20∗
BE-C-95-21∗
BE-C-95-24∗
BW99.1
BW99.2
BW99.3
BW99.4
BW99.5
BW99.6
BW99.7
BW99.8
BW99.9
BW99.10
BW99.11
5
6
8
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
Calcite
−2.3
−2.5
-2
−2.5
−1.9
−1.5
−2.6
−2.3
−2.2
−1.9
−1.7
−1.7
−2.2
−2.2
−2.3
−2.1
−2.4
−2.0
−2.6
−2.1
−2.0
−2.0
−1.9
−2.1
−1.9
−2.9
−3.6
−3.8
−19.8
−19.4
−20.9
−18.2
−18.8
−19.9
−16.4
−17.4
−19.6
−19.4
−18.9
−19.0
−17.7
−16.7
−15.3
−15.7
−16.1
−15.4
−16.7
−15.9
−14.5
−16.8
−15.8
−17.9
−18.3
−18.6
−18.1
−17.2
Sample number
MP-41
MP-43
MP-46
MP-47
MP-52
MP-52A
MP-53
MP-54
MP-54A
MP-55B
TOC,
mg C/g
sample
␦13 CTOC ,
‰ V-PDB
0.25
−16.0
δ
Fe
(ppm)
0.25
504
Mn
(ppm)
Sr
(ppm)
Mn/Sr
4655
262
17.8
3580
4980
915
2122
2534
1577
1854
1677
231
188
4
113
28
106
113
8
15.5
26.4
212.4
18.7
91.5
14.9
16.4
201.9
998
2374
4170
4406
4227
35
36
463
254
305
28.7
66.9
9.0
17.3
13.8
1055
128
8.3
87
Sr/86 Sr
0.05
−17.7
15.3
0.14
0.23
0.09
−20.4
−32.8
−19.6
92
17.9 94
30.9 6306
18.1
496
443
1461
117
187
181
4.2
2.4 0.708075
8.1
0.1
0.08
−12.6
−14.8
393
520
831
630
217
187
217
187
1.8 0.715257
2.8
3.8
3.4
0.08
0.07
0.3
−18.8
−14.1
−19.1
10.3 288
12.6 390
2228
1829
17.1
11.9
16.9 748
529
295
352
453
523
527
513
571
568
468
408
480
343
548
566
145
59
61
59
48
59
95
125
161
73
52
62
21
42
36
3.7
5.0
5.8
7.7
10.9
8.9
5.4
4.6
3.5
6.4
7.8
7.8
16.3
13.0
15.7
196
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Appendix A (Continued )
Sample number
Height
above
base (m)
Mineral/
rock
␦13 C,
‰ V-PDB
␦18 O,
‰ V-PDB
Bottle Creek Formation, Sierra Madre
SE1/4, SE1/4, SE1/4, Sec. 17, T14N, R87W
BS-1
6.5
Calcite
−2.2
BS-2
5
Calcite
−2.4
BS-3
0
Calcite
−3.4
BS-4A
13.5
Calcite
−3.1
BS-4B
13.5
Calcite
−2.8
BS-4C
13.5
Calcite
−3.6
BS-5
19
Calcite
−4.8
−14.2
−13.9
−14.5
−14.1
−12.5
−15.1
−16.1
SE1/4, SE1/4, SE1/4, Sec. 16, T14N, R87W
BS-6
Calcite
−2.5
−13.4
TOC,
mg C/g
sample
␦13 CTOC ,
‰ V-PDB
δ
Fe
(ppm)
Mn
(ppm)
Sr
(ppm)
Mn/Sr
87
Sr/86 Sr
Samples marked with the asterisk were analyzed at the Geological Survey of Finland using whole rock powders, for the rest of samples C, O, and
Sr isotope analyses of carbonates were done at the University of Maryland and Mn and Sr contents were measured at VPI & SU. TOC content
was measured at the University of Maryland; carbon isotope values of organic carbon were measured at the Mountain Mass Spectrometry. M:
cement, C: clast.
Appendix B. Detailed description of studied
sections
B.1. Geneva Lake
In the Geneva Lake area, the Espanola Formation
rests unconformably on Archean basement or, locally,
conformably on conglomerates of the Bruce Formation. Its thickness ranges from less than 15 m to over
90 m with an average around 45 m (Card and Innes,
1981). The upper contact with the Serpent Formation
is conformable and gradational (Card and Innes, 1981).
The lower part of the Espanola Formation includes a
15–60 m thick limestone-dolomite member with thin
graded beds of sandstone and siltstone with carbonate
cement and mudchips, shales, and carbonates (Locations 2 and 3 in Fig. 15). Intraformational flat-pebble
conglomerates with imbricated rounded pebbles of
laminated carbonate up to 2–3 cm in diameter are locally present (Loc. 2; Fig. 6). Graded sandstones have
erosional base and cross-bedding, whereas overlying
shales have convolute bedding and other features of
soft-sediment deformation. The lower part of the section in Locality 3 is marked by a basal conglomerate overlain by white to pink carbonates with shale
seams, stratiform stromatolites (Fig. 9A) and convolute bedding (Location 3, Fig. 6). Shaly carbonate with
sandstone caps the lower part of the section. The upper
part of the Espanola Formation in this area is 15–30 m
thick and consists of laminated siltstones and sandstones with carbonate cement and thin carbonate beds
(Location 1 in Fig. 15). Parallel, cross-, and lensoidal
bedding and graded beds are common in the upper
part.
B.2. Echo Lake
The lower contact of the Espanola Formation is
not exposed in the sampled section (Fig. 16). Here,
the Gowganda Formation disconformably overlies the
limestone member of the Espanola Formation with the
higher members and the Serpent Formation missing.
Three facies are interlayered in this section (Fig. 7): (1)
intraformational conglomerate with carbonate clasts
up to 8 cm in size (Fig. 9B); (2) finely interlayered
carbonate, marl, and shale; and (3) massive carbonate with thin shale beds. Some intraformational conglomerates include thin beds of shaly and massive
carbonates. The shale-dominated facies have convolute bedding and include channelized beds of intraformational conglomerate and lenses and beds of
massive carbonates ranging in thickness from 0.5
to 60 cm. There are also meter-size slumped blocks
in these facies. Massive carbonate beds increase in
thickness and abundance upsection. This section displays evidence for synsedimentary faulting and folding and was later intricately folded forming sheath
folds.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
197
Fig. 15. Geological map of the Huronian Supergroup in the Geneva Lake area (modified from Card and Innes, 1981) with sampled localities
shown. See Fig. 2 for location of the Geneva Lake area.
B.3. Quirke Lake
Two sections of the Espanola Formation were
logged and sampled in the Quirke Lake area; one is
the Kerr-McGee drill core 150/1 from the Bouck (150)
Township and the other is along the Quirke Lake at
Denison Mine (Fig. 17). The contact with the diamictite of the Bruce Formation in the drill core is sharply
overlain by thin shale interval (not shown on Fig. 4)
succeeded by white limestones with thin shaly lenses
and beds (Fig. 4). The limestone interval contains thin
beds of limestone pebbles in a siliciclastics-rich matrix
with soft sediment deformation, injection, and flame
structures indicating high sedimentation rates and
slope instability. A thin bed with spherules up to 2 mm
in diameter filled with megaquartz with undulose radial extinction occurs in the upper part of this member.
These spherules are similar to those formed after anhydrite in Phanerozoic and Neoproterozoic successions
(Milliken, 1979). The overlying siltstone member
consists of dark argillaceous siltstone with graded beds
and intraformational conglomerates with limestone
and shale pebbles, and thin carbonate beds and lenses.
Graded beds have sharp erosional lower contact. This
member also contains soft sediment deformation and
injection structures. The upper member is dolomitic
in the field but in the drill core is identified by higher
carbonate content in silty beds within thick mudstone
198
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
Fig. 16. Geological map of the Echo Lake area (modified from Bennett, 1982). The Marble Point section was sampled in this study. See Fig. 2
for location of the Echo Lake area.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
199
Fig. 17. Geological map of the Quirke Lake area (modified from Robertson, 1968). See Fig. 2 for location of the Quirke Lake area. Sampled
section is at the shore of Quirke Lake between Quarry Bay and No. 2 Shaft and sampled drill hole 150/1 is along the southern border of
the map.
200
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
intervals containing soft sediment deformation and
injection structures. This member has graded beds of
sandy carbonate, carbonate, and shale and contains
shale and carbonate intraclasts. Sandy carbonates have
basal scour surfaces and cross-bedding whereas overlying carbonates contain rare fenestral textures and
peloids.
The limestone member in the outcrop section starts
with a 1.5–3.0 m thick shale overlain by thin-bedded
grey siltstone and limestone (Fig. 9). Limestone beds
are thinly laminated and contain soft sediment deformation structures. The overlying siltstone member
contains ball-and-pillow structures and Bouma beds.
The contact between the limestone and siltstone members ranges from gradational to erosional. Conglomerate beds with closely packed pebbles of carbonate,
rare granite, and chert occur above this contact. Intraformational breccias are common at the top of the
member. The overlying dolostone member consists of
interlayered ferruginous dolomite, calcareous siltstone,
and limestone. The member contains flat-pebble conglomerate beds that resemble “beach rosettes” suggesting deposition in oscillatory flow. Other sedimentary
structures of the dolostone member are small-scale
and sigmoidal cross-bedding, abundant wave ripples
with internal laminations (Fig. 9D), convolute bedding,
thin laminations, rare ball-and-pillow structures, and
graded beds. The contact with the underlying member
is gradational. Several meter thick upward-deepening
cycles contain wave-rippled dolomite at the base grading to massive dolomite and overlain by calcareous
siltstone. The uppermost part of the member contains
mudcracks, symmetrical and ladder ripples, and wave
ripples indicating a shallow-water depositional setting.
The contact with the overlying Serpent Formation is
conformable in this area.
Fig. 18. Geological map of Moose Point (modified from Bernstein and Young, 1990) with sampled profile shown. See Fig. 2 for location of the
Moose Point area.
A. Bekker et al. / Precambrian Research 137 (2005) 167–206
B.4. Moose Point near Whitefish Falls
The Espanola Formation in this section (Fig. 18)
shows middle to upper greenschist facies grade metamorphism (Card, 1978). The limestone member is
6–17 m thick (Bernstein and Young, 1990) and sharply
overlies diamictite of the Bruce Formation (Fig. 5).
The basal part contains laminated sandy shales overlain
by interlayered grey limestones and dolostones, silty
limestones and dolostones, and calcareous siltstones.
The following sedimentary structures indicate a lowenergy, deep-water environment with episodic highenergy fluxes due to storm events, turbidity currents,
or slope failure resulting from fluidization: laminations,
rare climbing ripples, graded beds, convolute and crossbedded layers, and synsedimentary folds. A 0.5-m thick
conglomerate present in the upper part consists of imbricated clasts up to 30 cm in size of non-calcareous
to slightly calcareous shales, calcareous siltstones, and
silty limestones.
The gradationally overlying 230–310 m thick siltstone member consists of cross-bedded and laminated
calcareous siltstone and sandy shale with ball-andpillow structures (Fig. 9C) and lenses and beds of impure carbonates. Flaser and wavy bedding and symmetrical and asymmetrical wave ripples with internal
cross-laminations are present in shales at the top of
sandy beds. The member also contains syndepositional
faults overlain by disrupted beds. The siltstone member contains at the base a 35 m thick intraformational
breccia-conglomerate with pebbles and boulders up to
1 m in size of calcareous and non-calcareous siltstone
and sandstone.
The 212–238 m thick heterolithic member consists
of upward-fining cycles of coarse-grained sandstone,
siltstone, shale, and thin impure dolostone and limestone. Sandstones have trough and large low-angle
cross-bedding, ball-and-pillow structures, herringbone
cross-stratification and climbing, symmetrical, and
asymmetrical ripples. Small channels with granite and
quartzite pebbles up to 5 cm in diameter occur in this
member. Fine-grained rocks have convolute bedding,
thin laminations, starved ripples, and slump and load
structures.
The Espanola Formation at Moose Point records
a regressive cycle with the limestone member deposited below the wave base, the siltstone member
above the wave base, and the heterolithic member
201
influenced by tidal processes (Bernstein and Young,
1990).
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