Sulfur sources of buckshot pyrite in the Auriferous
Conglomerates of the Mesoarchean Witwatersrand and
Ventersdorp Supergroups, South Africa
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Citation
Guy, B. M., S. Ono, J. Gutzmer, Y. Lin, and N. J. Beukes. “Sulfur
Sources of Sedimentary ‘buckshot’ Pyrite in the Auriferous
Conglomerates of the Mesoarchean Witwatersrand and
Ventersdorp Supergroups, Kaapvaal Craton, South Africa.” Miner
Deposita 49, no. 6 (April 9, 2014): 751–775.
As Published
http://dx.doi.org/10.1007/s00126-014-0518-3
Publisher
Springer-Verlag
Version
Author's final manuscript
Accessed
Wed May 25 22:46:33 EDT 2016
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http://hdl.handle.net/1721.1/98891
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Creative Commons Attribution-Noncommercial-Share Alike
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Sulfur Sources of Buckshot Pyrite in the Auriferous Conglomerates of the Mesoarchean Witwatersrand
and Ventersdorp Supergroups, South Africa
B. M. Guya,*, S. Onob, J. Gutzmera,c, Y. Linb,1, A. Bauerb and N. J. Beukesa
a
Department of Geology, University of Johannesburg, Auckland Park, 2006, South Africa
email: nbeukes@uj.ac.za
b
Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, 77
Massachusetts Avenue, Cambridge, Massachusetts, 02139, USA
email: sono@mit.edu, annbauer@mit.edu
c
Helmholtz-Institute Freiberg for Resource Technology, Halsbrücker Str. 34, 09599 Freiberg, Germany.
email: j.gutzmer@hzdr.de
*
Present address: Mineral Services, SGS South Africa, 58 Melville Street, Booysens, Johannesburg, 2091, South
Africa.
email: bradley.guy@sgs.com
1
Present address: School of Earth and Space Sciences, University of Science and Technology of China, Hefei,
230026, P. R. China
email: yinglin@ustc.edu.cn
Abstract
Large rounded pyrite grains (>1 mm), commonly referred to as ‘buckshot’ pyrites, are a characteristic feature of
the auriferous conglomerates (reefs) in the Witwatersrand and Ventersdorp supergroups. Detailed petrographic
analyses of the reefs indicated that the vast majority of the buckshot pyrites are of reworked sedimentary origin,
i.e., that the pyrite originally formed in the sedimentary environment during sedimentation and diagenesis.
Forty-five of these reworked sedimentary pyrite grains from the Main, Vaal, Basal, Kalkoenkrans, Beatrix and
Ventersdorp Contact reefs were analyzed for their multiple sulfur isotope compositions (δ34S, Δ33S and Δ36S) to
determine the source of the pyrite sulfur. In addition, four epigenetic pyrite samples (pyrite formed after
sedimentation and lithification) from the Middelvlei and the Ventersdorp Contact reefs were measured for
comparison. The δ34S, Δ33S and Δ36S values of all 45 reworked sedimentary pyrite grains indicate clear
signatures of mass-dependent and mass-independent fractionation and range from −6.8 to +13.8‰, −1.7 to
+1.7‰ and −3.9 to +0.9‰, respectively. In contrast, the four epigenetic pyrite samples display a very limited
range of δ34S, Δ33S and Δ36S values (+0.7 to +4.0‰, −0.3 to +0.03‰ and −0.3 to +0.1‰, respectively).
Despite the clear signatures of mass-independent sulfur isotope fractionation, very few data points plot along
the primary Archean photochemical array suggesting a weak photolytic control over the data set. Instead, other
factors command a greater degree of influence such as pyrite paragenesis, the prevailing depositional
environment and non-photolytic sulfur sources. In relation to pyrite paragenesis, reworked syngenetic
sedimentary pyrite grains (pyrite originally precipitated along the sediment-water interface) are characterized by
negative δ34S and Δ33S values, suggesting open system conditions with respect to sulfate supply and the
presence of microbial sulfate reducers. On the contrary, most reworked diagenetic sedimentary pyrite grains
(pyrite originally precipitated below the sediment-water interface) show positive δ34S and negative Δ33S values,
suggesting closed system conditions. Negligible Δ33S anomalies from epigenetic pyrite suggest that the sulfur
was sourced from a mass-dependent or isotopically homogenized metamorphic/hydrothermal sulfur reservoir.
Contrasting sulfur isotope compositions were also observed from different depositional environments, namely
fluvial conglomerates and marine-modified fluvial conglomerates. The bulk of the pyrite grains from fluvial
conglomerates are characterized by a wide range of δ34S values (−6.2 to +4.8‰) and small Δ33S values (±0.3‰).
This signature likely represents a crustal sulfate reservoir derived from either volcanic degassing or from
weathering of sulfide minerals in the hinterland. Reworked sedimentary pyrite grains from marine-modified
fluvial conglomerates share similar isotope compositions, but also produce a positive Δ33S/δ34S array that
overlaps with the composition of Archean barite, suggesting the introduction of marine sulfur. These results
demonstrate multiple sources of sulfur, including atmospheric, crustal and marine sulfur reservoirs. The
prevalence of the mass-dependent crustal sulfur isotope signature suggests that sulfate concentrations were
probably much higher in terrestrial settings in comparison to marine environments, which were sulfate-deficient.
However, the optimum conditions for forming sedimentary pyrite with a crustal sulfur isotope signature were
probably not during fluvial progradation but rather during the early phases of flooding of low angle
unconformities, i.e., during retrogradational fluvial deposition, coupled in some cases with marine
transgressions, immediately following inflection points of maximum rate of relative sea level fall.
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Introduction
Pyrite in the gold-bearing conglomerates of the upper Witwatersrand Supergroup has traditionally been
classified into three groups, namely ‘compact rounded’, ‘porous rounded’ and hydrothermal (Ramdohr 1958;
Saager 1970; Utter 1978; Hirdes and Saager 1983; England et al. 2002). The use of descriptive terminology for
the first two groups is primarily a reflection of genetic uncertainty, i.e., whether the rounded pyrite grains
represent clastic components of the conglomerate or whether they represent post-depositional pseudomorphs
after rounded Fe-oxide minerals (Barnicoat et al. 1997; Law and Phillips 2005). Despite this uncertainty, most
authors regard ‘compact rounded’ pyrite and ‘porous rounded’ pyrite as fragments of pyrite that were derived
from the erosion of granitic-greenstone and sedimentary sources, respectively (Hallbauer 1986; England et al.
2002).
Reworked pyrite grains of a sedimentary origin have been documented by most Witwatersrand researchers,
i.e., the so-called ‘buckshot’ pyrites, which are generally large rounded pyrite grains (>1 mm in size). However,
the bulk of the literature has focused on compact rounded pyrite (<1 mm in size) and hydrothermal pyrite
because of their significance in gold exploration in the Witwatersrand Supergroup (McLean and Fleet 1989;
Phillips and Myers 1989; Myers et al. 1993; Barnicoat et al. 1997; Fleet 1998; Kirk et al. 2001; England et al.
2002; Frimmel et al. 2005). Consequently, reworked sedimentary pyrite grains are routinely overlooked, despite
the fact that they comprise the bulk of total pyrite in several auriferous conglomerates of the Witwatersrand
Supergroup (Ramdohr 1958; Utter 1978; Tucker 1980; Hirdes and Saager 1983; Hallbauer 1986; Hutchinson
and Viljoen 1987; England et al. 2002) and that they can contain substantial contents of solid-solution gold
and/or nano-particle gold (Utter 1978; Koglin et al. 2010; Large et al. 2013). Considering the scientific and
economic importance of this world class Au-U deposit, it is somewhat surprising that virtually no one has
questioned how the substantial quantities of sedimentary pyrite could have formed if the Archean atmosphere
was highly depleted in oxygen (Farquhar et al. 2000), implying that sulfide weathering would have been
negligible and that both continental runoff and ocean water would have essentially been sulfate-free (Habicht et
al. 2002; Frimmel 2005).
Previous authors have reported relatively small δ34S variations from Witwatersrand pyrite (−0.5 to +6.5‰)
(Hoefs et al. 1968; Hattori et al. 1983; Palmer 1986; Strauss and Beukes 1991; Zhao 1998). The small range of
δ34Spy values from barren and mineralized horizons, and between different paragenetic associations, has been
used to suggest a magmatic source for the sulfur. However, recent small-scale analyses of the ‘porous rounded’
pyrite (laser ablation and micro-drilling) have revealed a much wider range of δ34Spy values (−17 to +16‰),
indicating that a portion of Witwatersrand pyrite might have formed through microbial sulfate reduction
(England et al. 2002; Hofmann et al. 2009; Large et al. 2013). This information, supplemented with multiple
sulfur isotope data (e.g., Hofmann et al. 2009), provides further evidence for a sedimentary origin as first
suggested by the petrographic studies of Ramdohr (1958), Saager (1970), Utter (1978), Dimroth (1979), Hirdes
and Saager (1983) and Hallbauer (1986).
In this paper, the results of a combined petrographic and multiple sulfur isotope investigation (δ 34S, Δ33S and
36
Δ S) are presented in order to constrain the source of the sulfur for reworked sedimentary pyrite grains of the
upper Witwatersrand and lower Ventersdorp supergroups (~2.9 to 2.7 Ga). In the first part of the investigation,
pyrite is classified according to textural and paragenetic constraints, where it will be argued that the vast
majority of the buckshot pyrite grains were derived from a sedimentary source. In the second part of the
investigation, it is shown that the bulk of the sulfur isotope compositions are controlled by terrestrial geological
processes with only minor atmospheric input. Finally, a model for sedimentary pyrite formation in the upper
Witwatersrand and lower Ventersdorp supergroups is presented.
Geological setting
The Witwatersrand Supergroup represents a ~7 km thick siliciclastic sequence that was deposited between 2.985
and 2.780 Ga over the central parts of the Archean Kaapvaal Craton (Fig. 1a) (Kositcin et al. 2003; Kositcin and
Krapez 2004; SACS 2006). The succession unconformably overlies a granitoid-greenstone basement (3.48 to
3.03 Ga) and the volcano-sedimentary rocks of the 3.09 to 3.07 Ga Dominion Group (Armstrong et al. 1991;
Robb et al. 1992; Poujol et al. 2003). The Witwatersrand Supergroup has experienced only low-grade
greenschist facies regional metamorphism, although certain areas have experienced higher grades (e.g., the
northwestern parts of the basin, the overturned collar of the Vredefort Dome and the northeastern part of the
Evander gold field) (Gibson and Wallmach 1995; Frimmel 1997; Phillips and Law 2000; Law and Phillips
2005). Stratigraphically, the Witwatersrand Supergroup is divided into the lower West Rand Group (~5 km
thick) and the upper Central Rand Group (~2 km thick) (Fig. 1b).
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Fig. 1.
The West Rand Group is composed of arenite and shale, with minor banded iron formation (BIF),
conglomerate, diamictite and lava. Marine shelf depofacies are prominent in the lower parts of the group,
whereas fluvial braidplain and distal marine shelf depofacies characterize the middle to upper parts (Watchorn
1981; Tankard et al. 1982; Beukes 1995). Most authors agree that the lower Hospital Hill Subgroup was
deposited in a passive margin setting, with an open ocean to the southeast. However, there is little consensus
regarding the tectonic setting of the middle and upper parts of the West Rand Group. Whilst some authors have
argued for a passive margin setting for the entire West Rand Group (Kositcin and Krapez 2004; Koglin et al.
2010), others have suggested that the change to a foreland basin setting occurred at the junction between the
Hospital Hill and Government subgroups (Beukes 1995; Coward et al. 1995; Schmitz et al. 2004). Alternatively,
Catuneanu (2001) proposed a foreland basin setting for the entire Witwatersrand Supergroup, where the West
Rand and Central Rand groups represent the underfilled and overfilled stages, respectively. The maximum age
of the West Rand Group is 2985 ± 14 Ma (detrital zircon grains, Kositcin and Krapez 2004) and the minimum
age for most of the group is 2914 ± 8 Ma (the age of the lavas in the Crown Formation, Armstrong et al. 1991).
In contrast to the West Rand Group, the Central Rand Group contains a greater proportion of arenaceous and
rudaceous lithotypes, with minor argillaceous and volcanic units such as the Booysens Formation and the Bird
Lava, respectively. The Central Rand Group is dominated by fluvial braidplain sedimentary rocks that were
deposited in a restricted retroarc foreland basin setting (Catuneanu 2001; Kositcin and Krapez 2004; Schmitz et
al. 2004; Frimmel et al. 2005). Fluvial sedimentation was controlled by deep-seated thrust faults along the basin
margin, which led to synsedimentary folding and progressive uplift, source area rejuvenation, erosion and
recycling of older sediments (Myers et al. 1990; McCarthy 1992; Coward et al. 1995). Marine incursions are
recorded at various times during deposition of the Central Rand Group in the form of laterally continuous
mudstones (Booysens Formation) and transgressive shallow marine quartz arenites. A number of prominent
gold-bearing conglomerates are overlain by transgressive shallow marine quartz arenites (e.g., the Main, Carbon
Leader, Vaal, Basal and Denny’s reefs; Fig. 2). Marine modification of these fluvial conglomerates, via
winnowing and wave reworking, likely played a crucial role in upgrading their gold and heavy mineral content
(Minter 1972; Verrezen 1987; Watchorn and O’Brien 1991; Beukes and Nelson 1995; Catuneanu and Biddulph
2001). The maximum age of the Central Rand Group is 2902 ± 13 Ma (detrital zircon grains, Kositcin and
Krapez 2004) and the minimum age is 2780 ± 3 Ma (authigenic xenotime, Kositcin et al. 2003). However, based
on stratigraphic correlations between the Mozaan Group and the Witwatersrand Supergroup, and indications that
the two successions were deposited in the same basin, i.e., the Greater Witwatersrand basin (Beukes and
Cairncross 1991), it is possible that Central Rand Group sedimentation came to an end around 2837 Ma
(Gutzmer et al. 1999).
The auriferous Ventersdorp Contact Reef (VCR) overlies the Witwatersrand Supergroup with an erosional
contact. Deposition probably took place around 2780 Ma based on the erosional event that preceded the
correlative Derdepoort, Dennilton and Kanye-Lobatse volcanic successions (Beukes and Nelson 1995; Hartzer
2000). The reef forms part of the Venterspost Formation and represents terraced fluvial and debris-flow deposits
(Krapez 1985; Frimmel et al. 2005).
Methods and samples
Forty-four pyrite-bearing conglomerate samples of the Central Rand Group and Ventersdorp Contact Reef were
collected from fresh drill-core and underground mine workings from the East Rand, Central Rand, West Rand,
Carletonville, Klerksdorp and Welkom gold fields (Fig. 1b). Sampled reefs include the North, Carbon Leader,
Main Reef Leader, Main, Middelvlei, Basal, Vaal, Saaiplaas, Kalkoenkrans, Kimberley, Beatrix, Kloof, Elsburg
and the Ventersdorp Contact reefs (Fig. 2). From these samples, polished blocks were prepared and examined
using reflected light and scanning electron microscopy (SEM). Minerals were identified in the SEM by energydispersive spectroscopy (EDS). Petrographic studies indicated that seven conglomerates contained reworked
sedimentary pyrite grains suitable for micro-drilling (i.e., grains larger than 500 μm) (Fig. 2):
Fluvial conglomerates
a.
b.
c.
d.
Beatrix Reef, Welkom gold field (Minter et al. 1988; Genis 1990)
Kalkoenkrans Reef, Welkom gold field (H Brouwer 2013, pers. comm.)
Middelvlei Reef, Carletonville gold field (Myers et al. 1993)
Ventersdorp Contact Reef, Carletonville gold field (Krapez 1985)
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Marine-modified fluvial conglomerates
e.
f.
g.
Vaal Reef, Klerksdorp gold field (Catuneanu and Biddulph 2001)
Basal Reef, Welkom gold field (Frimmel et al. 2005)
Main Reef, Carletonville gold field (Beukes and Nelson 1995)
From these seven reefs, 45 reworked sedimentary pyrite grains and four epigenetic pyrite samples were
selected for multiple sulfur isotope analyses (33S/32S, 34S/32S and 36S/32S). Sulfur was extracted by conventional
Cr reduction and Ag2S precipitation (Canfield et al. 1986). Ag2S was then heated to 300°C, in the presence of
F2, to form SF6. The SF6 (3 to 6 μmol) was purified by a gas chromatography (GC) system similar to that
described in Ono et al. (2006), and introduced to an isotope ratio mass spectrometer using a dual-inlet mode for
measurements of masses 127, 128, 129, and 131. The mass spectrometry measurement consists of three runs,
with each run consisting of eight cycles of reference versus sample injection. Reproducibility for complete
analysis, from fluorination, GC purification, and isotope ratio analysis are 0.1, 0.2 and 0.4‰ (1σ) for δ33S, δ34S
and δ36S, respectively, and 0.01 and 0.1‰ (2σ) for Δ33S and Δ36S, respectively. Sulfur isotope values are
reported relative to the Vienna Canyon Diablo Troilite Standard (V-CDT) by defining IAEA S-1 to be: δ33S =
−0.055‰, δ34S = −0.300‰ and δ36S = −1.14‰.
Fig. 2.
Classification of pyrite
The auriferous conglomerates contain a wide variety of paragenetic and morphological types of pyrite, which
are typically concentrated on footwall contacts along with other detrital heavy minerals and gold (Ramdohr
1958; Saager 1970; Utter 1978; Hirdes 1979; Hirdes and Saager 1983; Hallbauer 1986; Fleet 1998; England et
al. 2002). In this contribution, pyrite is classified according to its position in the paragenetic sequence (e.g.,
Hallbauer 1986), where each paragenetic association is further subdivided according to textural constraints. For
ease of reference, the genetic classification system employed by Guy et al. (2010) for pyrite in the West Rand
Group is utilized in the current contribution (Table 1). Application of this classification to the conglomerates of
the Central Rand Group and Ventersdorp Supergroup required two minor adjustments, namely:
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1) the inclusion of a syngenetic (synsedimentary) pyrite association (SYN) and
2) the allocation of certain diagenetic pyrite textures (e.g., banded and bladed pyrite), which were not
observed in the West Rand Group, into the nodular and concretionary category (DIA-5).
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It should be noted that the syngenetic and diagenetic pyrite grains in the auriferous conglomerates occur as
reworked fragments of earlier sedimentary pyrite aggregates, and not as components formed in situ as reported
for the West Rand Group (Guy et al. 2010). Furthermore, many of the delicate pyrite types that were observed in
the shales of the West Rand Group were not observed in the conglomerates of the Central Rand Group and
lower Ventersdorp Supergroup. These include finely disseminated pyrite (DIA-3), pyrite aggregates (DIA-4),
dendritic skeletal pyrite (DIA-7), pyrite laths (DIA-8), phyllosilicate pyrite (DIA-9) and pyrite whiskers (DIA10). Chert pebbles with bands of finely disseminated pyrite, as documented by Hirdes and Saager (1983), are
not included in this classification system since they were not observed during the petrographic examinations.
However, this particular type of pyrite will be revisited in the discussion.
Pyrite descriptions
In the descriptions that follow, each type of pyrite will be discussed, with specific reference to pyrite texture,
stratigraphic occurrence and other salient aspects, including mechanisms of formation and a review of previous
literature (where applicable).
Detrital pyrite
The detrital pyrite association (DET) pertains to pyrite that was initially eroded and transported from an external
source area, and then incorporated into the auriferous conglomerates. There are two types of detrital pyrite:
compact rounded pyrite (DET-1) and pyrite in composite lithoclasts (DET-2) (Table 1).
Table 1.
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Compact rounded pyrite (DET-1) represents inclusion-free, non-porous euhedral pyrite crystals (usually as
single crystals; Fleet 1998) that have been rounded to various degrees (Fig. 3a). The average grain size ranges
from 0.1 to 1.5 mm in diameter, although grains as small as 10 μm and as large as 5 mm have been documented
(Feather and Koen 1975; Utter 1978). The majority of compact rounded pyrite grains identified in this study
were less than 0.5 mm in diameter. Detrital grains of compact pyrite and cobaltite are common in the Main,
Main Reef Leader, Carbon Leader and Elsburg reefs. Conversely, detrital grains of compact pyrite are less
common in hydrothermally altered reefs (e.g., the Middelvlei Reef) and in reefs dominated by reworked
sedimentary pyrite grains (e.g., the Beatrix, Kalkoenkrans and Ventersdorp Contact reefs). Most authors agree
that compact rounded pyrite was derived from an Archean granite-greenstone hinterland (Utter 1978; Hirdes and
Saager 1983). However, recent chemical and Pb-Pb isotopic studies have shown that compact rounded pyrite
could also represent rounded fragments of diagenetic and/or metamorphic pyrite (Guy et al. 2010; Large et al.
2013). Thus, compact rounded pyrite could have been derived from a variety of geological source areas. It is
also acknowledged that an unknown percentage of compact rounded pyrite grains may represent pseudomorphs
after detrital iron-oxide/iron-titanium-oxide minerals (Ramdohr 1958; Saager 1970) but also rutile, chlorite,
shale, chert, quartz and other lithic fragments (Myers et al. 1993; Zhou et al. 1995). Note that compact rounded
pyrite grains were not analyzed for their multiple sulfur isotope compositions since they were not the focus of
this investigation (due to small grain sizes and uncertainties relating to provenance).
The second type of detrital pyrite relates to pyrite situated in lithoclasts (DET-2). The designation of this
type of pyrite into the detrital paragenetic association is appropriate since the host rock or mineral particle from
which it came was lithified, eroded and transported to its current position (Hallbauer 1986; Fig. 7A of England
et al. 2002). The term is generally applied when the provenance of the pyrite-bearing lithoclast is equivocal.
Reworked syngenetic pyrite
The reworked syngenetic pyrite association (SYN) refers to pyrite that originally formed in the water column or
at the sediment-water interface (i.e., synsedimentary pyrite). This contrasts with diagenetic pyrite that formed
below the sediment-water interface. There are two types of reworked syngenetic pyrite present in the auriferous
conglomerates, namely oolitic pyrite (SYN-1) and fine-grained micro-porous pyrite (SYN-2) (Table 1).
Oolitic pyrite, or concentrically laminated pyrite (SYN-1), are spherical to near-spherical pyrite grains (up to
7 mm in size, Hallbauer 1986) that contain fine concentric pyrite laminae ranging between 5 and 50 µm in size.
Some examples may possess a core of radiating pyrite crystals (Fig. 3b, c; Figs. 5A and 15A of England et al.
2002). Oolitic pyrite has been documented from the Vaal, Kimberley, Elsburg, Composite, Ventersdorp Contact
and Black reefs (Ramdohr 1958; Utter 1978; Hirdes 1979; Tucker 1980; Hallbauer 1986; England et al. 2002).
In this study, oolitic pyrite grains were observed in the Carbon Leader, Middelvlei, Elsburg and Ventersdorp
Contact reefs. Petrographic studies reveal the presence of galena, chalcopyrite and gold inclusions within the
cortical laminae of the pyrite oolites (Fig. 3d). A replacement origin of precursor carbonate or ironstone ooids
has been suggested by previous authors (e.g., England et al. 2002). However, these explanations are unlikely
considering that carbonate detritus has not been reported from the auriferous conglomerates and that
phyllosilicate minerals occur within the pyrite laminae (Tucker 1980; Agangi et al. submitted; this study).
Furthermore, variable δ34S signatures across the cortical envelope would argue against post-depositional
sulfidation (England et al. 2002; Schieber and Riciputi 2005). Schieber and Riciputi (2005) envisaged that
oolitic pyrite formed via the accretion of iron sulfide laminae under shallow burial conditions, which were
subject to periods of intermittent reworking, grain abrasion and reburial.
Only one example of fine-grained, micro-porous pyrite (SYN-2) was observed, namely in the Ventersdorp
Contact Reef (Fig. 3e, f). The micro-porous pyrite occurs in a laminated pyrite grain. Individual laminae are
generally less than 500 µm thick. The pyrite is very fine-grained and contains numerous microscopic pores
(~0.5 µm in size), as well as inclusions of galena. These features, together with the absence of siliciclastic
mineral inclusions, suggest that pyrite nucleation was rapid and that precipitation occurred from highly
supersaturated fluids (Scott et al. 2009). The most interesting feature, however, is the occurrence of microspherical structures (Fig. 3f). These micro-spherical structures, which are composed of quartz, are orientated
parallel to the pyrite laminae and exhibit a uniform size range between 10 to 20 µm. Although speculative, they
may represent pseudomorphs after gaseous fenestrae associated with microbes or decaying organic matter.
Reworked diagenetic pyrite
Diagenetic pyrite (DIA) refers to pyrite that formed below the sediment surface after sedimentation but before
lithification and metamorphism. Six major types of reworked diagenetic pyrite were observed, namely
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framboidal pyrite (DIA-1), small euhedral pyrite crystals (DIA-2), pyrite nodules and concretions (DIA-5),
euhedral pyrite crystals (DIA-6), pyrite overgrowths (DIA-11) and replacive pyrite (DIA-12) (Table 1). Of
these, pyrite nodules and concretions are the most abundant.
Framboidal pyrite (DIA-1) was observed in samples from the Vaal, Beatrix and Kalkoenkrans reefs (Fig.
3g). The individual pyrite framboids are relatively large (10 to 20 µm) and together form sizeable aggregates (1
to 2 mm in size). Thin compact overgrowths were commonly observed around the framboids. Putative atoll
structures, which are thought to reflect the diagenetic transformation of framboidal pyrite (England and Ostwald
1993), were observed in the Kalkoenkrans Reef (Fig. 3h). An early diagenetic origin for the framboidal pyrite is
inferred by the large diameter of the pyrite framboids (~10 µm), as syngenetic framboids that form in euxinic
water columns are generally much smaller (<5 µm) (Wilkin et al. 1996; Wilkin et al. 1997). However, the
uniform size range of the framboids, their close packing and lack of clastic mineral inclusions may point
towards an earlier stage of formation (i.e., syngenetic). Pyrite framboids in carbonaceous chert have also been
reported by Hirdes (1979) and Hirdes and Saager (1983) in the Evander gold field. It is interesting to speculate
on the source area of the above-mentioned pyrite-chert pebbles, since a greenstone belt provenance has been
inferred for much of the detritus in the Kimberley Reef (e.g., a high proportion of BIF pebbles, chromite grains
and platinum group minerals; Hirdes and Saager 1983).
Small (<30 µm) euhedral pyrite crystals (DIA-2) were observed in reworked and rounded carbonaceous
lithoclasts in the Basal, Vaal, Elsburg and Ventersdorp Contact reefs (Fig. 3i). In situ analogues of this type of
pyrite were also observed in carbonaceous shale units from the West Rand Group (Guy et al. 2010) and
generally form during the early stages of diagenesis (Taylor and Macquakor 2000).
Pyrite nodules and concretions (DIA-5) constitute the bulk of the reworked and rounded diagenetic pyrite in
the auriferous conglomerates (Saager 1970). Volumetrically, pyrite nodules are subordinate to pyrite
concretions. Pyrite nodules have been interpreted to form during the earliest stages of diagenesis in shallow
burial depths where the sediment is characterized by a high degree of plasticity, high fluid content and an
isotropic stress field (Selles-Martinez 1996). The defining criterion for pyrite nodules is the general lack of
siliciclastic mineral inclusions. Three broad subtypes of pyrite nodules are recognized: a) dendritic, b) bladed,
radial and c) banded.
Dendritic pyrite nodules are characterized by cone-in-cone or branching structures, and are common in the
Basal (Fig. 3j, k) and Steyn reefs (Fig. 6A of England et al. 2002). Similar in situ dendritic pyrite nodules have
been reported from the West Rand Group (Guy et al. 2010) and from Phanerozoic black shale successions
(Woodland 1975; Carstens 1985; Carstens 1986; Ruppert et al. 2005).
Bladed and radial pyrite nodules were observed in the Kalkoenkrans, Basal and Ventersdorp Contact reefs
(Fig. 3b, l, m), but have also been reported from the Kimberley, Steyn, Composite and Elsburg reefs (Tucker
1980; Hirdes and Saager 1983; Hallbauer and von Gehlen 1983; England et al. 2002; Reimold et al. 2004).
Although pyrite can develop extremely elongated habits (Ohfuji and Rickard 2005), previous authors have
suggested that the bladed and radial forms of pyrite were initially derived from marcasite that had converted to
pyrite during greenschist facies metamorphism (Hallbauer and von Gehlen 1983; Hallbauer 1986; England et al.
2002; Falconer 2003; Falconer et al. 2006). For example, the radial pyrite of Hirdes and Saager (1983, Plates
10E and F) displays a striking resemblance to the radial marcasite of Falconer et al. (2006, Fig. 6D). In this
study, marcasite ancestry is supported by the observation of a conspicuous texture that resembles the
recrystallized marcasite of Falconer et al. (2006) – compare Fig. 3n and Fig. 54 of Ramdohr (1958) with Fig. 6F
of Falconer et al. (2006). Interestingly, abiotic dissolution or microbial tunneling is documented from one of
these samples (Fig. 3o). Alternative explanations for the bladed and radial textures include derivation from
massive sulfide deposits and replacement of precursor evaporitic gypsum (England et al. 2002). However, the
lack of massive sulfide detritus (e.g., detrital base metal sulfides) and evaporitic gypsum in the Witwatersrand
Supergroup, coupled with the recognition of sedimentary equivalents, supports a sedimentary provenance.
Banded pyrite nodules are characterized by various textures associated with growth banding, i.e., the
banding does not reflect pyritized sedimentary layers (see replacive pyrite, DIA-12). Banded pyrite nodules
have been recorded from the Kimberley, Steyn and Black reefs (Ramdohr 1958; Hirdes and Saager 1983;
England et al. 2002). The ‘chevron-textured’ and ‘vein-fill’ pyrite of England et al. (2002) probably represents
two varieties of banded pyrite (Fig. 6E of Falconer et al. 2006). In this study, only a small number of banded
pyrite nodules were identified from the Kalkoenkrans, Basal and Ventersdorp Contact reefs (Fig. 4a) and only
one example of colloform pyrite was observed (Fig. 4b). Interestingly, Ketzer et al. (2003) observed similar
colloform pyrites in a paralic sandstone bed overlying coal layers. Hirdes and Saager (1983) and England et al.
(2002) suggested that banded pyrite was derived from sulfidic-banded iron formations and volcanogenic
massive sulfides, respectively. However, we prefer an early diagenetic origin due to the lack of massive sulfide
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detritus in the auriferous conglomerates and the recognition of comparable early diagenetic pyrite textures in
shale, sandstone and conglomerate of different ages (Frizzo et al. 1991; Ketzer et al. 2003; Falconer et al. 2006).
Pyrite concretions, in contrast to pyrite nodules, incorporate the surrounding host rock material by filling in
the pore spaces between siliciclastic grains (Selles-Martinez 1996). Concretions that form during early
diagenesis are characterized by few point contacts between quartz grains, whereas those concretions that form
during the latter stages of diagenesis, replace the siliciclastic framework (Selles-Martinez 1996). In this study,
pyrite concretions were observed in a range of sedimentary rock types, which include sandstone, siltstone and
mudstone. Pyrite concretions that formed in sandstone were observed in the Main, Kimberley and Ventersdorp
Contact reefs (Fig. 4c). Previous authors have documented this type of pyrite in the proximal facies of several
Witwatersrand and Ventersdorp reefs (e.g., quartz-pyrite aggregates of Ramdohr 1958; Hirdes and Saager 1983;
Hallbauer 1986). Although Ramdohr (1958) advocated a replacement origin (i.e., of magnetite and hematite
quartzite), comparable pyrite and marcasite concretions in recently deposited sandstone and conglomerate
suggest a sedimentary origin (Ketzer et al. 2003; Falconer 2003). The bulk of the pyrite concretions in the
auriferous conglomerates are characterized by silt-sized quartz grain inclusions and not clay- or sand-sized
quartz grain inclusions (Fig. 4d, e). Mineral inclusions such as quartz, chlorite, muscovite, feldspar, rutile and
chromite infer a sedimentary protolith. Note that concretions are characterized by mineral inclusions and not
round/elliptical pores. (Hence, the frequent usage of the term ‘porous pyrite’ for concretionary pyrite is a
misnomer.) Interestingly, polygonal syneresis cracks were observed in one concretion from the Vaal Reef. The
cracks subsequently became filled with compact pyrite (Fig. 4f). Pyrite concretions with negligible rounding
reflect a local source, but not an in situ origin as suggested by England et al. (2002), since the mineralogy and
small size of the inclusions (fine-grained phyllosilicate minerals) is not consistent with their formation in quartzpebble conglomerates (Fig. 4e and Table 2).
Diagenetic euhedral pyrite crystals (DIA-6) are relatively uncommon in the auriferous conglomerates of the
Central Rand Group and Ventersdorp Supergroup. However, it is possible that many of the euhedral pyrite
crystals have been rounded to form of compact rounded pyrite grains (e.g., Guy et al. 2010). Only one example
was observed from the Carbon Leader Reef as part of a larger reworked composite grain (Fig. 4g; compare with
an in situ example in supplementary Fig. S1j).
Diagenetic pyrite overgrowths (DIA-11) are also uncommon in the auriferous conglomerates of the Central
Rand Group and Ventersdorp Supergroup. Nevertheless, they were observed as thin coatings (~10 µm) around
framboidal pyrite and small euhedral pyrite crystals in the Basal and Kalkoenkrans reefs (Fig. 3g). Diagenetic
pyrite overgrowths are generally much smaller than epigenetic overgrowths and occur in sediments that show
little evidence of metamorphic/hydrothermal fluid flow.
Replacive pyrite (DIA-12) was observed in a small number of sedimentary lithoclasts in the Kalkoenkrans
and Beatrix reefs (Fig. 4h, i, j). There are two types, namely variably-pyritized black shale clasts and
‘mineralized bacteria’. Interestingly, Tucker (1980) found that the former type was common in samples with
high gold grades. Historically, the term ‘mineralized bacteria’ has been applied to pyrite morphologies that
exhibit small microspherical structures (10-45 µm in size). These structures were interpreted to represent
pyritized bacterial colonies or the remains of primitive life forms (Ramdohr 1958; Schidlowski 1965; Saager
1970). However, following the work of numerous authors (e.g., Scott et al. 2009 and references therein), it has
become apparent that a large proportion of the ‘mineralized bacteria’ textures represent framboidal pyrite or
structures derived from framboidal pyrite (England and Ostwald 1993). Nevertheless, not all of these textures
represent pyrite framboids or atoll structures. In one sample from the Beatrix Reef, small 5 to 20 µm-sized
elliptical microstructures and wrinkled pyrite laminae were observed (Fig. 4h, i; Figs. 31 and 32 of Saager
1970). These features are consistent with the morphology of bacterial debris (Fig. 3A of Lott et al. 1999) and
may represent the pyritized remains of ancient organic material (Bubela and Cloud 1983; Schieber 2002). SEMEDS analyses of the minerals inside the elliptical structures indicate the presence of quartz, galena, chalcopyrite,
Fe-Mg-Mn carbonates, pyrophyllite and muscovite.
Epigenetic pyrite
The epigenetic pyrite association (EPI) consists of pyrite that formed after sedimentation and lithification,
during regional metamorphism and local hydrothermal events. There are five types of epigenetic pyrite:
euhedral crystals (EPI-1), overgrowths (EPI-2), veinlets (EPI-3), replacive pyrite (EPI-4) and pyrrhotite and
base metal sulfides (EPI-5) (Table 1). Although epigenetic pyrite is not the primary focus of the investigation,
some important observations are noted below.
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Euhedral pyrite and pyrite overgrowths (EPI-1 and EPI-2) (both porous and compact) comprise the largest
proportion of epigenetic pyrite in the auriferous quartz-pebble conglomerates (Fig. 4k, l). The crystals are
generally coarse-grained and can reach several millimeters in size (Feather and Koen 1975). They commonly
occur in aggregates, and overgrow compact rounded pyrite grains and reworked sedimentary pyrite grains. They
are particularly abundant in reefs that contain an appreciable amount of reworked diagenetic pyrite (e.g., the
Basal and Ventersdorp Contact reefs). Hirdes and Saager (1983) and Hallbauer and von Gehlen (1983) observed
that epigenetic euhedral pyrite crystals in the Kimberley and Ventersdorp Contact reefs were commonly
associated with local heat sources, such as dykes and lavas of the Ventersdorp Supergroup. Similar occurrences
were reported from the Promise Formation in the West Rand Group, where epigenetic euhedral pyrite crystals
were closely associated with diabase sills and early diagenetic pyrite (Guy et al. 2010; see also Wagner and
Boyce 2006). Rounded fragments of putative epigenetic porous pyrite were observed in the Beatrix,
Kalkoenkrans, Elsburg, North and Vaal reefs (Fig. 4k), but also in the Basal Reef (Fig. 47 of Saager 1970) and
in the Coronation Formation, West Rand Group (Fig. 7I of Guy et al. 2010). In situ analogues of this type of
pyrite have been reported from a number of metamorphosed sedimentary successions (Large et al. 2009; Guy et
al. 2010; Thomas et al. 2011).
Pyrite veinlets (EPI-3) are relatively uncommon in comparison to euhedral pyrite crystals and pyrite
overgrowths. The vast majority of veinlet-associated pyrite occurs in micro-fractures within quartz and chert
pebbles (Fig. 4m). However, one sample consisting of brittle and granular pyrite was collected from a major
auriferous and carbonaceous vein from the Klerksdorp gold field (Cloete 2009).
Only a small number of replacement textures (EPI-4) were observed during this study (e.g., Reimer and
Mossman 1990; Fleet 1998). These include pyrite replacing quartz grains, shale fragments and Fe-Ti oxide
minerals from the Basal, Kalkoenkrans and Ventersdorp Contact reefs (Fig. 4n). In each case, however, the
replacement was partial and the precursor mineral was recognizable. Previous researchers have documented
pyrite and pyrrhotite replacing hematite, magnetite, titaniferous magnetite, ilmenite, rutile, phyllosilicates, chert,
quartz pebbles and other lithic rock fragments (Ramdohr 1958; Saager 1970; Hirdes and Saager 1983; Phillips et
al. 1987; Myers et al. 1993; Phillips and Dong 1994; Zhou et al. 1995; Barnicoat et al. 1997; Phillips and Law
2000). However, some of these examples probably reflect textural varieties of diagenetic pyrite (e.g., the quartzpyrite aggregates of Ramdohr 1958).
Epigenetic pyrrhotite and base metal sulfides (EPI-5) were rarely observed in this study (Fig. 4o). However,
the scarcity of pyrrhotite and base metal sulfides probably reflects the sampling bias towards unaltered reef
sections that contained well-preserved sedimentary pyrite grains. Zhou et al. (1995) reported that pyrrhotite was
common in the northern parts of the basin (West Rand and Carletonville gold fields) but less common in the
East Rand gold field and absent in the Evander and Welkom gold fields. In addition, Zhou et al. (1995) also
documented pyrrhotite in areas that were unaffected by contact metamorphism and suggested a regional
metamorphic influence. However, a close relationship between pyrrhotite and carbonaceous lithologies in the
Witwatersrand Supergroup (Tweedie 1968; Guy 2012) suggests that the pyrite to pyrrhotite transformation may
also be linked to the presence of organic carbon and is not only dependent on regional metamorphism (Ferry
1981; Thomas et al. 2011).
Fig. 3.
Fig. 4.
Multiple sulfur isotope data
The δ34S, Δ33S and Δ36S values of the entire sample set range from −6.8 to +13.8‰, −1.7 to +1.7‰ and −3.9 to
+0.9‰, respectively (Figs. 5, 6 and Table 2). Sulfur isotope compositions vary according to paragenetic
association, pyrite type and depositional environment.
Variation according to paragenetic association
Reworked syngenetic pyrite grains display negative δ34S values (−6.2 to −1.5‰) and small negative Δ33S
anomalies (average −0.3‰) (Fig. 5a). In contrast, the majority of reworked diagenetic pyrite grains display more
positive δ34S values (most between ~0 and +5‰) with variable Δ33S anomalies (−1.7 to +1.7‰) (Fig. 5a). Not
all diagenetic samples exhibit positive δ34S values, as a few samples are characterized by negative δ34S values
(down to −6.8‰). In addition, reworked diagenetic pyrite grains display much larger Δ36S offsets in comparison
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to reworked syngenetic pyrite grains (relative to the Archean Δ36S/Δ33S array of −1 ± 0.2). Epigenetic pyrites,
which include both in situ samples and putative reworked fragments (Fig. 4k), are characterized by highly
attenuated δ34S, Δ33S, and Δ36S values (average +2.5, −0.1 and −0.1‰, respectively) (Figs. 5a and 6a).
Variation according to pyrite type
Connections between sulfur isotope composition and textural pyrite type are most pronounced for oolitic pyrite
(SYN-1), fine-grained, micro-porous pyrite (SYN-2), dendritic pyrite nodules (DIA-5), pyrite concretions with
sand-sized quartz inclusions (DIA-5) and reworked epigenetic porous pyrite aggregates (EPI-1) (Fig. 5b).
However, many other diagenetic pyrite types, such as small euhedral crystals (DIA-2) and concretions with finegrained mineral inclusions (DIA-5), lack appreciable systematic isotope variation (Fig. 5b).
Oolitic pyrite (SYN-1) yields negative δ34S excursions (down to −6.2‰) and small negative Δ33S anomalies
(average −0.4‰) (Fig. 6b, d, i). The single fine-grained, micro-porous pyrite sample (SYN-2) from the VCR
exhibits negative δ34S and negligible Δ33S values (−5.2 and −0.07‰, respectively) (Fig. 6e). Dendritic pyrite
nodules (DIA-5) are marked by small negative δ34S values (average −1.3‰) but relatively large negative Δ33S
anomalies (average −1.4‰). Pyrite concretions with sand-sized quartz grain inclusions (DIA-5) display the most
variable δ34S and Δ33S compositions (−6.6 to +13.8‰ and −0.6 to +1.7‰, respectively) (Fig. 6h). Dendritic
pyrite nodules and pyrite concretions with sand-sized quartz grain inclusions also exhibit some of the largest
offsets in Δ36S (~3‰). Reworked porous pyrite aggregates (EPI-1) display limited variations in δ34S and Δ33S,
and average +3.2 and −0.02‰, respectively.
Variation according to depositional environment
Excluding two outliers, reworked sedimentary pyrite grains from the fluvial-dominated Beatrix, Kalkoenkrans
and Ventersdorp Contact reefs display low magnitude Δ33S (±0.3‰) and variable δ34S values (−6.2 to +4.8‰)
(Fig. 5c). Conversely, pyrite grains from the marine-reworked conglomerates of the Main, Vaal and Basal reefs
are characterized by a wider range of Δ33S anomalies (−1.7 to +0.9‰) and a Δ33S/δ34S array that intercepts the
Δ33S and δ34S axes at ~−1.5 and +12‰, respectively (Fig. 5c).
Fig. 5.
Fig. 6.
Table 2.
Discussion
Detailed petrographic analysis of the buckshot pyrite from the gold-bearing conglomerates of the upper
Witwatersrand and lower Ventersdorp supergroups suggest that the majority of the large pyrite grains were
derived from a sedimentary source rock. In the discussion that follows, we expand on this premise and use the
multiple sulfur isotope data to provide further constraints on their origin.
Sedimentary provenance for buckshot pyrite
A number of source terranes have been proposed to account for the abundance and startling morphological
variety of the large rounded pyrite grains in the auriferous conglomerates of the Witwatersrand and Ventersdorp
supergroups. These include volcanogenic massive sulfide deposits, carbonate platforms, evaporites (Hutchinson
and Viljoen 1987; England et al. 2002) and sulfidic-banded iron formations (Hallbauer and von Gehlen 1983;
Hirdes and Saager 1983; Hutchinson and Viljoen 1987; Koglin et al. 2010). However, the requisite bulk rock
detritus from these terranes is conspicuously absent and evidence for pre-burial sulfidation is lacking, e.g.,
detrital clasts of partially pyritized carbonates were not observed (Myers et al. 1993). Similarly, it is also
unlikely that the large quantity of pyrite in the auriferous conglomerates was sourced from Paleoarchean granitegreenstone belts, as these typically contain little pyrite (Köppel and Saager 1974; Phillips and Powell 2011).
Rather, it may be suggested that the majority of the buckshot pyrite was derived from an intrabasinal
sedimentary setting, i.e., that they are of endoclastic origin and were not transported over very long distances
(Dimroth 1979; Hallbauer 1986; Large et al. 2013). In support of this interpretation, virtually every single one of
the pyrite morphologies observed in the Witwatersrand conglomerates possesses an in situ modern or ancient
textural equivalent from a sedimentary setting (Fig. 7). Additional lines of evidence for an intrabasinal
sedimentary source includes immature grain shapes, fragile internal textures, siliciclastic mineral inclusions,
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presence of shale clasts and trace element compositions (Minter 1976; Dimroth 1979; Hallbauer and von Gehlen
1983; Hirdes and Saager 1983; Hallbauer 1986; Phillips and Law 2000; Agangi et al. submitted; Large et al.
2013). Possible source rocks of this affinity include sediments of the stratigraphically lower West Rand Group
and contemporaneous Central Rand Group.
Note that at least two types of rounded pyrite were probably derived from an extrabasinal source area. These
include rounded fragments of epigenetic porous euhedral pyrite and pyrite-bearing chert clasts. The former
probably reflects a metamorphic provenance and the latter, a greenstone belt provenance (Hirdes and Saager
1983; Phillips and Dong 1994).
Fig. 7.
Geological context of multiple sulfur isotope signatures
Paragenetic association
Links between pyrite paragenesis and multiple sulfur isotope compositions are most pronounced for sedimentary
rocks of Archean age – but less so for post-Archean sedimentary rocks that were deposited in more sulfate-rich
settings (Kamber and Whitehouse 2007; Partridge et al. 2008; Ono et al. 2009b; Farquhar et al. 2013). This
statement holds true for the reworked sedimentary pyrite grains in the Mesoarchean Witwatersrand and
Ventersdorp supergroups. More specifically, reworked syngenetic pyrite grains are characterized by sulfur
isotope values that plot to the left of the Archean Reference Array (ARA), whereas most of the sulfur isotope
values from reworked diagenetic pyrite grains plot to the right of the ARA (Fig. 5a). Ono et al. (2009a)
attributed this variation to reflect the openness of the system, where samples that plot to the left of the ARA
represent open system conditions (i.e., abundant sulfate supply) and those that plot to the right of the ARA
represent closed system conditions below the sediment-water interface (i.e., restricted sulfate supply).
Interestingly, very little data plots along the ARA in the Δ33S-δ34S plot (Fig. 5a), suggesting either microbial
reworking of the primary atmospheric signal (Johnston, 2011), changes in atmospheric photochemistry
(Farquhar et al. 2007, Domagal-Goldman et al. 2008), existence of non-photolytic sulfur reservoirs (Guy et al.
2012; Roerdink et al. 2012) or a combination thereof. Similarly, the bulk of the multiple sulfur isotope data in
the Δ36S-Δ33S plots below the ARA (Fig. 5a). Whereas syngenetic pyrite displays minor offsets, some diagenetic
pyrites are characterized by impressive Δ36S offsets of ~3‰. Deviation from this array can be explained by
overprinting of mass-dependent isotope effects (Δ36S/Δ33S of −6.8; Ueno et al. 2008; Ono et al. 2009b) and/or
isotope self-shielding due to increased volcanic SO2 loading (Δ36S/Δ33S of −4.6; Ono et al. 2013). Despite these
offsets, the Δ36S/Δ33S ratio of syngenetic pyrite (−1.0) is consistent with the main Archean S-MIF reaction (Fig.
5a) (Kaufman et al. 2007; Ono et al. 2009b; Ono et al. 2013).
Pyrite type
Five morphological types of pyrite reveal unique sulfur isotopic compositions (Fig. 5b). Syngenetic oolitic
pyrite (SYN-1) from the Ventersdorp Contact Reef yields relatively large negative δ34S values (−1.5 to −6.2‰)
and weakly negative Δ33S values (~−0.4‰) (Fig. 6b, d, i). In addition to relatively large negative δ34S values,
England et al. (2002) observed core-to-rim δ34S gradients in oolitic pyrite from the same reef (see their Fig. 15).
Similar δ34S gradients have been documented in modern oolitic pyrite laminae (Fig. 10 of Schieber and Riciputi
2005) and point to changes in pore water sulfate concentrations; a feature probably caused by microbial sulfate
reduction in an environment subject to short-lived high-energy pulses and sedimentary reworking. In support of
a microbial origin, Schieber and Riciputi (2005) found thin sheet-like structures on the cortices of modern pyrite
ooids that resembled organic biofilms. They suggested that the biofilms might have facilitated in the
precipitation of sulfide minerals by providing an organic substrate for sulfate-reducing bacteria and for acting as
a growth template.
Although only one example of syngenetic fine-grained, micro-porous pyrite (SYN-2) was observed, the δ34S
value of −5.2‰ is similar to that of oolitic pyrite (SYN-1) (Fig. 5b). Indeed, the unique textural characteristics
of the sample are suggestive of elevated sulfur concentrations that may have enabled microbial discrimination of
sulfur isotopes.
In contrast to the oolitic and micro-porous pyrite varieties, diagenetic dendritic pyrite nodules from the Basal
Reef (DIA-5) display small δ34S values (~−0.4‰) and relatively large negative Δ33S values (~−1.6‰), which
plot near the ARA (Fig. 5b). Dendritic nodules from the West Rand Group are also characterized by large
negative Δ33S values that plot near the ARA (−3.1 to −4.0‰, Guy et al. 2012). This implies that these
morphological types of pyrite are adept at acquiring large Δ33S fractionations with minimal mass-dependent
reworking (Ono et al. 2003; Ono et al. 2009a; Johnston 2011). Although the reasons for this relationship are
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unclear, it may be related to rapid disequilibrium precipitation (as indicated by the dendritic growth textures) in
a closed system (e.g., Carstens 1985; Carstens 1986; Staude et al. 2007).
Diagenetic pyrite concretions (DIA-5) with sand-sized quartz grain inclusions reveal a wide range in sulfur
isotope compositions, both in relation to δ34S and Δ33S (−6.6 to +13.8‰ and −0.6 to +1.7‰, respectively) (Fig.
5b). These results contrast with the limited fractionation range of pyrite concretions with fine-grained mineral
inclusions. The variability of the sulfur isotope signatures could reflect the porous and permeable nature of the
host sandstone, i.e., high hydraulic conductivities and open system diagenetic conditions.
Three examples of reworked epigenetic porous pyrite aggregates (EPI-1) from the Beatrix and Kalkoenkrans
reefs yield remarkably uniform δ34S and Δ33S compositions (average +3.2 and −0.02‰, respectively) (Fig. 5b).
Based on the internal texture of the pyrite, the presence of silicate mineral inclusions and mass-dependent Δ33S
values, it could be speculated that these porous aggregates were derived from a source area consisting of
metamorphosed sedimentary rocks.
Depositional environment
Links between depositional facies and sulfur isotope compositions have contributed to a better understanding of
sulfur cycles in a broader depositional context (Kamber and Whitehouse 2007; Ono et al. 2009a; Guy et al.
2012). Similarly, in the auriferous conglomerates of the Witwatersrand and Ventersdorp supergroups, there are
tangible differences between the sulfur isotope ratios of pyrite from fluvial conglomerates and marine-modified
fluvial conglomerates (Fig. 5c).
Reworked sedimentary pyrite grains collected from fluvial conglomerates (i.e., the Kalkoenkrans, Beatrix
and Ventersdorp Contact reefs) generally yield small Δ33S values (±0.3‰) and variable δ34S values (−6.2 and
+4.8‰) (Fig. 5c). With respect to δ34S-Δ33S systematics, two major trends were observed. The first trend is
characterized by relatively large negative δ34S values (down to −6.2‰) and small negative Δ33S anomalies
(~−0.4‰). Such values are indicative of microbial sulfate reduction (Ono et al. 2003; Farquhar et al. 2013) and
probably reflect elevated concentrations of sulfate in abandoned fluvial channels, stagnant pools and ephemeral
ponds (Fig. 8a). The second trend is characterized by negligible Δ33S anomalies and δ34S values between ~0 and
+5‰. This range of δ34S values is typical for crustal (riverine) sulfate today and may signify reservoir effects
associated with microbial sulfate reduction (Turchyn et al. 2013) (solid line ‘MSR’, Fig. 8a). Alternatively, it
may represent mixing between two sulfur reservoirs, i.e., a juvenile volcanogenic reservoir and a putative crustal
mass-dependent reservoir (dashed line ‘V’ to ‘C’, Fig. 8a). Nevertheless, the δ34S composition of this reservoir
is comparable to the inferred mass-dependent reservoir observed from fluvial deposits in the middle and upper
West Rand Group (‘C’ in Fig. 8b) (Guy et al. 2012). This reservoir was initially interpreted to represent
continentally-derived sulfate due to the close association of dampened Δ33S anomalies with early diagenetic
pyrite, fluvial braidplain deposits and a foreland basin tectonic setting (Guy et al. 2012). Sulfate was probably
supplied to the sedimentary environment via oxidation of sulfide minerals (Reinhard et al. 2009; Lefticariu et al.
2010; Wacey et al. 2011; Stueeken et al. 2012) and/or degassing from subaerial volcanoes (Kump and Barley
2007; Gaillard et al. 2011; Binder and Keppler 2011; Philippot et al. 2012). Evidence for a terrestrial sulfur flux
can be found in other parts of the Kaapvaal and Zimbabwe cratons during the Mesoarchean and Neoarchean
eras. Reworked pyrite grains of sedimentary origin have been documented from the ~3.1 Ga Dominion Reef
(Rantzsch et al. 2011), the ~3.0 Ga Mozaan Contact Reef (Saager et al. 1986; Hofmann et al. 2009), the ~2.9 Ga
conglomerates of the Uitkyk Formation (Saager and Muff 1986; de Wit et al. 1993; Kröner et al. 2000), the ~2.7
Ga Ventersdorp Contact Reef (Hallbauer 1986; England et al. 2002), the ~2.7 Ga Belingwe greenstone belt
(Grassineau et al. 2001; Grassineau et al. 2002; Hofmann et al. 2009) and the ~2.6 Ga Black Reef (Barton and
Hallbauer 1996; Hofmann et al. 2009). In contrast, reworked sedimentary pyrite grains are scarce/absent in the
chert-dominated conglomerates of the ~3.5 to 3.2 Ga Barberton greenstone belt (Hessler and Lowe 2006; de Wit
et al. 2011). A similar narrative applies to the temporal distribution of chert and chert-pyrite clasts in Archean
conglomerates on the Kaapvaal Craton (Figs. 3.11C and 4.5D in Guy 2012; Plate 12 of Hirdes and Saager 1983;
Fig. 1 of Phillips and Dong 1994).
Fig. 8.
Reworked sedimentary pyrite grains from marine transgressive conglomerates (i.e., the Main, Vaal and
Basal reefs) typically yield larger Δ33S values in comparison to pyrite grains from fluvial conglomerates (Fig.
5c). With respect to δ34S-Δ33S systematics, three trends were observed. The first trend relates to the small Δ33S
anomalies that overlap with the sulfur isotope compositions of reworked sedimentary pyrite from fluvial
conglomerates. This trend probably reflects the incorporation of pyrite from underlying fluvial channels, i.e.,
pyrite with a crustal signature (Fig. 5c). The second trend is characterized by a positive Δ33S/δ34S array that
intercepts the Δ33S and δ34S axes at ~−1.5 and +12‰, respectively (dashed line ‘A’ to ‘M’, Fig. 8a). These endmembers have been observed in previous studies and were interpreted to represent two sulfate reservoirs,
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namely a mass-independent reservoir derived from photochemical processes (‘A’ in Fig. 8a) (Farquhar et al.
2000, Ono et al. 2003; Ueno et al. 2008) and a mass-dependent reservoir derived from seafloor hydrothermal
systems (‘M’ in Fig. 8a) (Bao et al. 2007; Ueno et al. 2008). The occurrence of a similar mass-dependent sulfur
reservoir in the marine-dominated parts of the West Rand Group, Witwatersrand Supergroup (Fig. 8b; Guy et al.
2012) and the close compositional fit with Archean barite (Fig. 8c; Bao et al. 2007; Ueno et al. 2008; Roerdink
et al. 2012) suggests that this array represents a marine sulfate signature (‘M’) (Fig. 8a). If valid, the
preservation of this array is consistent with low levels of sulfate in the Archean ocean (Habicht et al. 2002; Ono
et al. 2009a; Jamieson et al. 2012). The third trend is related to the data points that plot to the left of the ARA in
the second quadrant (Fig. 8a). Previous studies have interpreted such data as evidence for S0 disproportionation
(Philippot et al. 2007) and oxidation of photochemical sulfur (Kaufman et al. 2007). Alternatively, the data
points may also point to mixing between polysulphides derived from an elemental sulfur precursor (Δ33S>0) and
hydrogen sulphide derived from microbial sulfate reducers (δ34S<0 and Δ33S<0) (e.g., Farquhar et al. 2013).
Conceptual models for sedimentary pyrite formation
Establishing the original depositional setting and source of buckshot pyrite in the auriferous conglomerates
has proven problematic since the pyrite segregations have been dislodged from their enclosing sediment. Indeed,
published accounts of in situ sedimentary pyrite in the auriferous conglomerates are exceedingly scarce (see
Tucker 1980 for two notable exceptions). One explanation could be that the sedimentary pyrite grains were
derived from the erosion and reworking of pyrite-bearing strata from the stratigraphically lower West Rand
Group. This could have happened during deposition of the Central Rand Group in a shrinking foreland basin
with strata of the West Rand Group being exposed along the tectonic margins of the basin. However, such an
explanation may only be partly applicable since there is a clear difference in sulfur isotope populations between
fluvial and marine-modified conglomerates (Fig. 5c). If the sedimentary buckshot pyrites were exclusively
derived from erosion of older strata, then the fluvial and marine-modified conglomerates should have displayed
similar sulfur isotope patterns. However, because we see only partial overlap of the data (Fig. 5c), it can be
concluded that a major proportion of the buckshot pyrites were formed and reworked during deposition of the
two aforementioned conglomerate facies, i.e., they are syndepositional and strictly endoclastic in origin.
Reworked sedimentary pyrite grains with a crustal sulfur isotope signature are interpreted to have formed in
a braided fluvial depositional environment, i.e., on the floor of abandoned fluvial channels, stagnant pools and
interchannel flats (Tucker 1980; Hallbauer and von Gehlen 1983). In this continental setting, Tucker (1980)
suggested that sedimentary pyrite would have formed during periods of quiescence in abandoned channels under
low-energy flow conditions. However, during short-lived high-energy pulses, the pyrite would be ripped up and
re-deposited. Reworked sedimentary pyrite grains with a marine sulfur isotope signature are interpreted to have
formed in sedimentary rocks associated with marine transgressions.
There are, however, many caveats to this model. For instance, it is important to realize that all major
auriferous conglomerates in the Witwatersrand Supergroup overlie laterally extensive and very flat erosional
unconformities. The conglomerates thus represent retrogradational braided fluvial complexes, sometimes in
combination with marine transgressions. It is envisaged that conditions for sedimentary pyrite formation were
most favorable when accommodation space in the basin started to increase following degradation and the
development of the erosional unconformities (Fig. 9). It is suggested that under such conditions, sedimentary
pyrite (carrying the crustal sulfur isotope signature) formed in flooded abandoned fluvial channels and on
interchannel flood plains, whereas sedimentary pyrite (carrying the marine sulfur isotope signature) formed in
paralic marine embayments (Tucker 1980; Hallbauer and von Gehlen 1983; Saager and Muff 1986; Eriksson et
al. 1998). As relative sea levels increased, these deposits were reworked by wave action to form the marinemodified conglomerates. (Fig. 9). This succession of events would help to explain the sheet-like architecture of
the major auriferous conglomerates, the pyrite-rich matrix and the deposition of thin quartz arenites overlying
many of the auriferous conglomerates (Verrezen, 1987; Catuneanu and Biddulph 2001). Note that the formation
of sedimentary pyrite extended into deeper water settings, as evidenced from the Booysens Formation in the
Evander gold field, where pyrite layers in carbonaceous shale can be traced for over 23 km (e.g., the upper and
lower pyrite markers) (Tweedie 1968; Bailey et al. 1990). The pyrite layers are individually composed of pyrite
seams, nodules and concretions that exhibit concentric and radiating growths of pyrite (Tweedie 1968; personal
observation). Erosion of such strata, via fluvial or marine processes, would also help to explain the textural
variations of pyrite in the auriferous conglomerates (see supplementary Fig. S1 for examples of pyrite textures
in black shale) and perhaps provide a source for some of the gold (see Fig. 7.3 of Guy 2012 for an example of a
gold-bearing diagenetic pyrite nodule in West Rand Group shale).
In contrast to the retrogradational fluvial conglomerates, it is informative to note that the thick successions of
progradational braided fluvial deposits and associated discontinuous lenticular conglomerates formed during
periods of decreasing accommodation space in the basin, prior to the development of low angle erosional
unconformities (Fig. 9). These progradational fluvial deposits are typically characterized by dispersed
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concentrations of buckshot pyrite (e.g., Catuneanu and Biddulph 2001). This observation supports the
conclusion that sedimentary pyrite formed preferentially during periods of increased rates of relative sealevel
rise and drowning of peneplained land surfaces. However, this is not to say that sedimentary pyrite did form not
in progradational fluvial sediments. During degradation and unconformity development, such dispersed
sedimentary pyrite accumulations may have become reworked and concentrated in heavy mineral lags or
pediment deposits along unconformity surfaces.
Fig. 9.
Conclusion
Petrographic and multiple sulfur isotope analyses were carried out on buckshot pyrite grains from auriferous
conglomerates in the Witwatersrand and Ventersdorp supergroups to determine the origin of the pyrite sulfur.
Textural and geochemical constraints indicate that the majority of the buckshot pyrite grains were derived from
a sedimentary provenance and that the wide range of pyrite morphologies are characteristic of sedimentary
pyrite, and not multiple source areas (Dimroth 1979). Forty-five of these reworked sedimentary pyrite grains
were analyzed for their multiple sulfur isotope compositions. The δ34S, Δ33S and Δ36S values range from −6.8 to
+13.8‰, −1.7 to +1.7‰ and −3.9 to +0.9‰, respectively. Relationships between δ34S and Δ33S point to the
presence of three sulfur reservoirs: 1) mass-independent photochemical SO42- and S0 (data points on the ARA),
2) mass-dependent marine SO42- (positive Δ33S/δ34S array in the fourth quadrant) and 3) mass-dependent crustal
SO42- (δ34S = ~0 to +5‰, Δ33S = ~0‰). Of these three, the latter is the most prominent reservoir and implies a
significant contribution of continentally-derived sulfur for most of the pyrite grains in the auriferous quartz
pebble conglomerates. The emergence of a continentally-derived crustal sulfur reservoir may be linked to the
colonization of land by sulfide-oxidizing microbes (Wacey et al. 2011; Stueeken et al. 2012; Turchyn et al.
2013) and/or the advent of modern plate tectonic processes and oxidized volcanic gases (Chaussidon et al. 1989;
Ohmoto and Goldhaber 1997; Kump and Barley 2007; Smithies et al. 2007; Frimmel 2008; Binder and Keppler
2011; Gaillard et al. 2011; Schneiderhan et al. 2011; Philippot et al. 2012; Roerdink et al. 2012). In contrast, the
preservation of a marine sulfur isotope signature (Fig. 8a) is consistent with low levels of sulfate in the Archean
ocean (Habicht et al. 2002; Jamieson et al. 2012) since elevated concentrations of sulfate would have erased the
photochemical signature – producing a Δ33S trend resembling the modern sulfur cycle (Bekker et al. 2004; Guo
et al. 2009; Johnston 2011). The conclusion that most of the large rounded pyrite grains in the auriferous
conglomerates are of sedimentary and endoclastic origin has important ramifications regarding the cycling of
sulfur during the Archean, but also perhaps for the metallogenesis of gold in the conglomerates. More detailed
sulfur isotope studies, in the context of depositional systems tracts, are required to test the model presented here.
Acknowledgments
This work is supported by a research grant to NJB by the National Research Foundation, Pretoria (GUN
2053184) and a NASA Exobiology grant, NNX07AU12G to SO. Danie Thompson and Dawie Strydom of
Taung Gold, Louis Cloete of AngloGold Ashanti and particularly Hans Brouwer of Gold Fields (now Sibanye
Gold) are gratefully acknowledged for providing access to drill-core and underground mine samples. This
manuscript benefited greatly from discussions with Donna Falconer and Eva Stueeken. The editor and two
anonymous reviewers are thanked for their constructive comments on an earlier version of the manuscript.
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Fig. 2. Simplified stratigraphic column of the upper Witwatersrand Supergroup and sampled reefs. Column 1:
Groups. Column 2: Subgroups. Column 3: Formations. Column 4: Genetic successions after Beukes and
Nelson (1995). Column 5: Lithology. Column 6: Member/bed/reef. Sampled reefs for sulfur isotope analyses in
bold. Column 7: SHRIMP U-Pb or Pb-Pb ages after Frimmel et al. (2005), where (1) zircon (Armstrong et al.
1991), (2) zircon and xenotime (Kositcin and Krapez 2004), (3) authigenic xenotime (Kositcin et al. 2003) and
(4) zircon (Gutzmer et al. 1999).
Table 1. Paragenetic and morphological classification of pyrite in the auriferous conglomerates of the
Witwatersrand and Ventersdorp supergroups.
Table 2. Multiple sulfur isotope data of reworked sedimentary pyrite grains from the Witwatersrand and
Ventersdorp supergroups.
Fig. 1. a Geographic positions of the Witwatersrand and Pongola supergroups on the Archean Kaapvaal Craton.
b Geological map of the Witwatersrand Supergroup, illustrating major gold fields, paleoslopes, major faults and
the lower West Rand and upper Central Rand groups.
Fig. 3. SEM-BSE and reflected light (RF) photomicrographs of detrital pyrite and reworked syngenetic and
diagenetic pyrite. Detrital pyrite. a Compact rounded pyrite grains (DET-1) (see arrows). Reworked
syngenetic pyrite. b Oolitic or concentrically laminated pyrite (SYN-1) with a radiating pyrite core. c Oolitic
pyrite (SYN-1) with an abraded rim. d Oolitic pyrite laminae (SYN-1). Bright mineral is galena. e, f Finegrained, micro-porous pyrite (SYN-2). Note scale in f. Reworked diagenetic pyrite. g Framboidal pyrite (DIA1) with compact pyrite overgrowths (DIA-11). h Putative atoll structures after framboidal pyrite (DIA-1). i
Small euhedral crystals (DIA-2) in a shale clast. j Dendritic pyrite nodule (DIA-5) (cone-in-cone). k Branching
dendritic pyrite nodule (DIA-5). l Bladed, radiating pyrite nodule (DIA-5). m Bladed pyrite rim on concretion
(DIA-5). n Pyrite pseudomorphs after marcasite spheroids? (DIA-5). o Abiotic dissolution/microbial tunneling.
Refer to Table 2 for numbers 1 to 7 (top left corner of selected photomicrographs).
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Fig. 4. SEM-BSE and reflected light (RF) photomicrographs of reworked diagenetic pyrite and epigenetic
pyrite. Reworked diagenetic pyrite. a Banded pyrite nodule (DIA-5). b Colloform banded pyrite nodule (DIA5). Note truncation. c Quartz-pyrite concretion (DIA-5). Note compact cement and sand-sized quartz grains. d
Rounded pyrite concretion (DIA-5). e Fine-grained pyrite concretion (DIA-5). Note lack of rounding. Vertical
line is polishing scratch. f Polygonal syneresis cracks in a rounded pyrite concretion (DIA-5). g Zoned euhedral
pyrite crystals. h, i ‘Mineralized bacteria’ (DIA-12). Note wavy contortions of pyrite laminae. j Pyritized shale
clast (DIA-12). Epigenetic pyrite. k Detrital fragment of porous euhedral pyrite (EPI-1, 2). l Compact euhedral
pyrite crystals (EPI-1) and overgrowths (EPI-2). m Pyrite veinlet in quartz pebble (EPI-3). n Partial pyrite
replacement of a quartz grain (EPI-4). o Pyrrhotite enveloping a compact rounded pyrite grain (EPI-5). Refer to
Table 2 for numbers 8 to 13 (top left corner of selected photomicrographs).
Fig. 5. Plots of Δ33S vs. δ34S (‰) and Δ36S vs. Δ33S (‰) for a paragenetic association, b pyrite type and c
depositional environment.
Fig. 6. Hand specimen (4.7 cm in length) from the Ventersdorp Contact Reef. a Pyrite veinlet (EPI-3). b Oolitic
pyrite or concentrically laminated pyrite (SYN-1). c Fine-grained pyrite concretion (DIA-5). d Oolitic pyrite or
concentrically laminated pyrite (SYN-1). e Fine-grained, micro-porous pyrite (SYN-2). f Pyrite concretion
(DIA-5). g Putative replacive pyrite (DIA-12/EPI-4). h Quartz-pyrite concretion. i Oolitic pyrite (SYN-1) with
core of radiating pyrite. Values in permil (‰). Scale bar corresponds to 500 µm.
Fig. 7. Schematic representation of the pyrite types identified in the auriferous conglomerates (left) and
examples of their in situ textural equivalents (right). References: (1) Guy et al. (2010). (2) Black shale specimen
from the Transvaal Supergroup. (3) Carstens (1985, 1986). (4) Schieber and Riciputi (2005). (5) Ketzer et al.
(2003). (6) Falconer et al. (2006). (7) Frizzo et al. (1991). (8) Woodland (1975) (9) Wiese and Fyfe (1986). (10)
Large et al. (2009). (11) Ruppert et al. (2005).
Fig. 8. Plots of Δ33S vs. δ34S (‰) for a conglomerates of the Central Rand Group and b non-conglomeratic
sedimentary rocks from the West Rand Group (Guy et al. 2012). Dashed lines indicate reservoir mixing,
whereas solid lines indicate microbial processing (Johnston 2011). c Comparative Δ33S-δ34S data sets of barite
(Bao et al. 2007; Ueno et al. 2008; Roerdink et al. 2012) and reworked sedimentary pyrite from the
Witwatersrand and Ventersdorp supergroups (Hofmann et al. 2009).
Fig. 9. Conceptual models for sedimentary pyrite formation in the upper Witwatersrand and lower Ventersdorp
supergroups. a Highstand systems tract characterized by high accommodation space, rapid sedimentation and
progradation of braided fluvial systems. Dispersed sedimentary pyrite may have formed in the fluvial and
marine depositional environments. b Late highstand systems tract characterized by increasing degradation and
sedimentary bypassing. Sedimentary pyrite with a crustal signature may have formed in abandoned channels and
ephemeral ponds. c These sedimentary pyrite grains would have been concentrated into heavy mineral lags
during the development of low angle unconformities (e.g., Beatrix, Kalkoenkrans and Ventersdorp Contact
reefs). d Sedimentary pyrite formation (both with crustal and marine signatures) would have followed the
development of the unconformities due to very flat land surfaces and flooding of river channels and paralic
marine embayments. e These pyrites were then reworked into marine-modified conglomerates during marine
transgressions (e.g., Main, Basal and Vaal reefs).
Fig. S1. Variations in pyrite texture in black shale (Transvaal Supergroup). a Radiating pyrite spheroids (DIA5). Note framboid core (DIA-1). b Putative atoll structures (DIA-1). c Spheroidal pyrite nodules (DIA-5).
Compare with Fig. 3n. d Bladed, radiating pyrite (DIA-5). e Banded pyrite (DIA-5). f Fine-grained radiating
pyrite (DIA-5). g, h Pyrite concretion (DIA-5). Note differential pyritization. i, j Euhedral pyrite crystals (DIA6). Note zoning in j and compare zoning with Fig. 4g.
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