Archaean Atmospheric Evolution Evidence From the Witwatersrand Gold Fields, South Africa (1)

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Earth-Science Reviews

Archaean atmospheric evolution: evidence from the

Witwatersrand gold fields, South Africa

Hartwig E. Frimmel*

Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa

Received 6 July 2004; accepted 12 October 2004

Abstract

The Witwatersrand gold fields in South Africa, the world’s largest gold-producing province, play a pivotal role in the

reconstruction of the Archaean atmosphere and hydrosphere. Past uncertainties on the genetic model for the gold caused

confusion in the debate on Archaean palaeoenvironmental conditions. The majority of Witwatersrand gold occurs together with

pyrite, uraninite and locally bitumen, on degradational surfaces of fluvial conglomerates that were laid down between 2.90 and

2.84 Ga in the Central Rand Basin. Although most of the gold appears as a precipitate within, or associated with, post-

depositional hydrothermal phases and along microfractures, available microtextural, mineralogical, geochemical and isotopic

data all indicate that this hydrothermal gold, analogous to some pyrite and uraninite, was derived from the local mobilisation of

detrital particles. Some of the key pieces of evidence are a significant correlation of the gold, pyrite and uraninite with other

heavy minerals as well as sedimentary lithofacies, local preservation of in-situ gold micronuggets and abundant rounded forms

of pyrite and uraninite, compositional heterogeneity on a microscale of the gold as well as the rounded pyrite and uraninte, and

radiometric age data that indicate an age of the gold, pyrite and uraninite that is older than the maximum age of deposition for

the host sediment. None of these observations/data is compatible with any of the suggested hydrothermal models, in which

auriferous fluids were introduced from an external source into the host rock succession after sediment deposition. In contrast,

those arguments, used in favour of hydrothermal models, emphasise the microtextural position of most of the gold, which

highlights the undisputed hydrothermal nature of that gold in its present position, but does not explain its ultimate source.

Furthermore, the macro-scale setting of the stratiform ore deposits is in stark contrast to any known type of epigenetic,

hydrothermal gold deposit. Consequently, the best-fit genetic model involves post-depositional textural and mineralogical

modification of original fluvial placer deposits.

On the Kaapvaal Craton, Witwatersrand-type mineralisation is recorded over an extended period of time from 3074 to 2642

Ma. Rounded pyrite is common in the coarser grained fractions of the siliciclastic basin fill. A lack of sulphur isotope

fractionation and typical magmatic d34S values support its detrital origin. Together with rounded uraninite, which is particularly

abundant in the older beds, it provides important constraints on the redox potential of the Meso- to Neoarchaean (3.1–2.6 Ga)

atmosphere and hydrosphere. In combination with eukaryotic steroids documented from the Pilbara Craton, Australia, the

0012-8252/$ - s

doi:10.1016/j.ea

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70 (2005) 1–46

ee front matter D 2004 Elsevier B.V. All rights reserved.

rscirev.2004.10.003

ess: Institute of Mineralogy, University of Wqrzburg, Am Hubland, D-97074 Wqrzburg, Germany. Tel.: +49 931 888 5420;

8 4620.

esses: [email protected], [email protected].

H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–462

ambient Neoarchaean oxygen fugacity is calculated as having been approximately 10�3, in equilibrium with a relatively acidic

hydrosphere (pH=6). This is in agreement with the preservation of mass-independent S isotope fractionation, which provides

independent support for an anoxic atmosphere and which has so far been recorded predominantly from sediments older than 2.3

Ga. An acidic meteoric palaeoenvironment is supported by intense chemical weathering below erosional unconformity surfaces

in the Witwatersrand Basin. In contrast to the pyrite-bearing fluvial and near-shore shallow marine deposits, marine shale

deposits contain magnetite. This supports the postulated reducing environment but also highlights total sulphur concentrations

in the ancient ocean that were orders of magnitude lower than in modern ocean water.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Witwatersrand; Palaeoplacer deposits; Anoxic Archaean atmosphere; Gold genesis

Fig. 1. Contrasting models for the evolution of atmospheric

chemistry: (A) Models involving a reducing Archaean atmosphere

pO2 evolution curve proposed by (a) Kasting (1987, 2001) and (b)

Rye and Holland (2000), remaining curves after Pavlov et al

(2001a) and Kasting (2001); (B) model involving an oxidising

atmosphere (Ohmoto, 2004).

1. Introduction

The basic conditions for the most important aspects

of modern life on Earth have been shaped as early as

in Precambrian times. These include most likely the

origin of life and definitely the evolution from the

origin of the eukaryote cell to hard bodied animals,

the growth of continental crust, the oxygenation of the

atmosphere as well as the formation of mineral

deposits without which modern civilisation would be

unthinkable. All of these aspects appear to be

interrelated but the various feed-back mechanisms

that controlled the relationships between life’s evolu-

tion, plate tectonic processes, palaeoclimate and

distribution of metals between hydrosphere and geo-

sphere throughout the critical periods of the Precam-

brian remain highly speculative. In fact, the

determination of the above relationships can be

regarded as a fundamental problem in earth sciences.

One of the more important questions concerns the

timing and cause of the rise of atmospheric O2.

Photosynthesis was the principal process by which

free O2 was produced, though some of it may also be

ascribed to hydrogen escape to space after CH4

photolysis (Catling et al., 2001). Consequently, the

evolution of the Archaean (3.8–2.5 Ga) to Proterozoic

(2.5–0.54 Ga) atmosphere must have been strongly

influenced by organisms. Production and decomposi-

tion of organic matter, in combination with the

volcanic degassing of the planet, also influenced the

atmospheric CO2 and CH4 concentrations through a

carbon cycle that involves chemical weathering and

formation of silicates and carbonates. As both CO2

and CH4 are greenhouse gases, the evolution of life

and that of Archaean to Eoproterozoic (2.5–2.05 Ga)

palaeoclimate are intimately interlinked.

General agreement seems to exist on substantial O2

levels from around 2.3 Ga onwards, but the preceding

atmospheric evolution is a matter of intense debate

(Fig. 1). Two competing schools of thought try to

explain the environments and mechanisms that con-

trolled the emergence of life: one assumes that life

originated under a reducing atmosphere (O2b1 ppm)

and that the formation of an oxic atmosphere triggered

the emergence of eukarya (Kasting and Siefert, 2002;

Knoll, 1992; Rye and Holland, 2000), whereas the

other proposes the emergence of oxygenic photo-

;

.

H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 3

synthetic organisms as early as 4 Ga with essentially

constant atmospheric O2 levels (N10%) since then

(Lasaga and Ohmoto, 2002). Towe (1990) argued for

a low (0.2–0.4%) but stable O2 concentration in the

Archaean that was maintained by aerobic respiration.

This is in contrast to anaerobic carbon cycles that are

implicit in any model of a reducing Archaean

atmosphere.

The debate extends to the question of Archaean

palaeoclimate. As the Sun’s luminosity is estimated to

have been about 30% less than today (Newman and

Rood, 1977), Earth’s hydrosphere should have been

completely frozen. No geological support exists for

such a scenario, however, and elevated greenhouse

gas concentrations have therefore been proposed to

offset this so-called Faint Young Sun Paradox (Sagan

and Mullen, 1972). A dense Archaean atmosphere is

thought to have consisted largely of CO2, and the

stronger greenhouse gases H2O and CH4. For the time

around 3.0 Ga, atmospheric CO2, CH4 and O2

concentrations of approximately 3000, 1000 and

b0.1 ppm, respectively, have been suggested by those

advocating a reduced Archaean atmosphere (Kasting,

2001; Kasting and Siefert, 2002; Rye and Holland,

2000), which implies that CH4 was the dominant

greenhouse gas. In contrast, some workers (Lasaga

and Ohmoto, 2002; Ohmoto et al., 1999; Ohmoto,

2004) argue for CO2 having been the principal

greenhouse gas at that time at a level of about 10%.

As the continents tied up vast amounts of carbon in

the form of organic matter, hydrocarbons, and

carbonaceous as well as carbonate rocks, their growth

must have led to the removal of considerable amounts

of carbon from active circulation. In addition to burial

of biogenic carbon in sediments, the consumption of

CO2 by biota and the oxidation of CH4 by biogenic O2

must have caused a further decrease in the concen-

tration of greenhouse gases throughout the Neo-

archaean (2.8–2.5 Ga) and Proterozoic. This

decrease was further exacerbated by a decrease in

radiogenic heat production, volcanic activity, and thus

a decrease in the emission rate of potential greenhouse

gases. The consequential cooling of the Earth’s

surface triggered repeated global glaciations. Owing

to their old age, the preservation potential of Archaean

glacial deposits is relatively small, but glaciogenic

deposits have been reported from Mesoarchaean (3.2–

2.8 Ga) sedimentary successions (Young et al., 1998).

They are more common in Proterozoic successions, in

which major glaciations are recorded around 2.35 Ga

and again repeatedly throughout the Neoproterozoic

(1.0–0.54 Ga). Their close stratigraphic association

with marine carbonate successions is suggestive of

severe climate fluctuations, which highlight that no

equilibrium had been achieved between solar lumi-

nosity, greenhouse gas concentrations and global

climate.

It was by no means only carbon that was removed

from active circulation in the biosphere during the

Precambrian. Apart from gravitation towards the

planet’s core during the very early stages of Earth’s

history, most heavy metals were removed from

biogeochemical circulation by both the growth of

continental crust and mineralisation processes. Indi-

cations of the latter exist in the form of huge

sedimentary Fe, Cu, Pb–Zn, Mn and possibly also U

and Au deposits that are characteristic of the Archaean

and Proterozoic Aeons (Lambert et al., 1992). To use

Fe as an example, the distribution of iron formations,

which provide the best direct evidence of Fe-depletion

of the ocean water, appears particularly intriguing.

Their occurrence in Archaean and Eoproterozoic

strata has been explained by lower atmospheric O2

concentrations at those times that made possible a

stratified ocean with Fe-rich bottom waters (Trendall,

2002). According to this model, a steady increase in

ocean oxygenation towards the end of the Eoproter-

ozoic, as living organisms and photosynthesis became

more abundant and effective, would have lowered the

interface between Fe-rich anoxic bottom waters and

oxygenated waters down to effectively seafloor level

and thereby prohibited any further deposition of iron

formations. Global ice ages seem to be particularly

favourable for the deposition of iron formations when

a largely, or possibly completely, ice-covered Earth

would have shut down ocean water circulation and led

to the development of anoxic, Fe-rich ocean bottom

waters (Kirschvink, 1992; Klein and Beukes, 1993).

Although a relationship between iron formation,

palaeoclimate, palaeolatitude and volcanism is indi-

cated, the nature of this relationship remains elusive

and is the subject of on-going debate in the current

literature.

It appears clear from the above that considerable

uncertainty exists regarding the palaeoenvironmental

conditions prior to 2.2 Ga. One of the best natural


laboratories to study these conditions is the Kaapvaal

Craton in South Africa. It hosts, similar to the Pilbara

Craton in Australia, an undisturbed sedimentary cover

(Transvaal Supergroup) that spans almost the entire

Eoproterozoic Era. This sedimentary succession con-

tains prime examples of banded iron formations and

glaciogenic diamictite deposits, as well as a number of

paleosols that have been studied extensively by

numerous researchers and provided most useful

constraints on the atmospheric evolution in the

Eoproterozoic. One of the more recent results that

emanated from these studies is the recognition that

already at 2.2 Ga weathering was comparable with

modern tropical laterites under an oxic atmosphere

(Beukes et al., 2002).

The Kaapvaal Craton is also host to some of the

world’s finest examples of Palaeoarchaean (3.6–3.2

Ga) sedimentary successions. These occur in the 3.5

to 3.2 Ga Barberton Supergroup of the Barberton

Greenstone Belt as the Fig Tree and Moodies Groups.

There the rocks contain evidence of Fe-rich hydro-

thermal vents forming iron formations on the seafloor

(de Ronde et al., 1994), extensive silicification that

signals ocean water saturated in SiO2, as well as

remnants of early stromatolites (Byerly et al., 1986)

and microfossils (Walsh, 1992).

Whereas these earliest sedimentary rocks provide

minimum constraints on the age of the oldest life

forms, and the younger Eoproterozoic successions

provide minimum constraints on the timing of an oxic

atmosphere covering the planet, the time in between,

i.e. the Meso- and Neoarchaean Eras, holds the key

for the principal question of atmospheric evolution

and its bearing on the further evolution of life

following its emergence sometime in the early

Archaean. The best available geological record to

study the palaeoenvironmental conditions during that

crucial time span is the siliciclastic sediment fill of the

Mesoarchaean Witwatersrand Basin on the Kaapvaal

Craton. It represents the world’s best preserved

Archaean sedimentary succession and it contains

redox-sensitive minerals that might hold the key for

our understanding of the Mesoarchaean atmosphere

and hydrosphere.

Iron oxides are conspicuously lacking in fluvial

sediments of the Witwatersrand, with the principal Fe-

bearing phase being pyrite. The coarse-grained

fraction of these pyrite-rich, fluvial deposits (con-

glomerate) hosts the world’s largest known accumu-

lation of gold, but also represents its largest unmined

inferred uranium resource. More than 49,400 metric

tonnes (t) of gold have been produced from these

conglomerate beds (reefs) between 1886 and 2003,

amounting to almost 40% of all the gold ever mined

during recorded history (Frimmel and Minter, 2002;

Sanders et al., 1994). South Africa is still the world’s

number one gold producer with a share of approx-

imately 16% and, according to the South African

Chamber of Mines, the remaining reserves in the

Witwatersrand Basin, estimated around 38,000 t

(Frimmel and Minter, 2002), amount to 46% of

known world reserves. Between 1952 and 1975, as

much as 1.5 t of U3O8 were produced from

Witwatesrand conglomerates at an average grade of

271 ppm (Frimmel et al., in press). Although poorly

exposed, the Witwatersrand is one of the best-

documented basins of its kind in the world thanks to

more than 100 years of underground mining and

exploration. Yet, in spite of the enormous economic

significance of the Witwatersrand gold deposits, the

genesis of these deposits is still a matter of con-

troversy. Comparing the extensive older with the more

recent literature (for a review of the older and more

recent literature see Pretorius, 1975 and Frimmel and

Minter, 2002, respectively), it appears fair to say that

the debate around this controversy has not lost

anything of its intensity since it had been described

by Davidson (1965) as bthe most disputed issue in the

history of economic geologyQ.Depending on the preferred genetic model for the

Witwatersrand gold, the pyrite and uraninite, both of

which occur predominantly as rounded particles, are

interpreted either as sedimentary heavy sands or as

hydrothermal precipitates. In the latter case, the

abundant rounded pyrite is explained as pseudomor-

phic replacement of detrital Fe–Ti oxides, Fe-pisolite,

ferricrete, banded iron formation and Fe-rich shale

(Phillips and Law, 2000) and/or product of post-

depositional dissolution and re-precipitation mecha-

nisms (Barnicoat et al., 1997; Phillips and Myers,

1989). The different genetic models have equally

different ramifications for the inferred redox state of

the atmosphere at the time of sediment deposition as

they imply either Fe-sulphides or Fe-oxides as having

been stable under the Archaean atmosphere. One of

the major goals of this paper is therefore, to review


and assess the various genetic models that have been

proposed for the Witwatersrand gold deposits, and by

implication also for the associated pyrite and uranin-

ite. It will be shown that the best-fit genetic model is

that of modified palaeoplacer deposits. The signifi-

cance thereof for the reconstruction of the Meso- to

Neoarchaean atmosphere will then be discussed.

The siliciclastic successions of the Witwatersrand

are characterised by an abundance of erosional

unconformity surfaces. Theoretically, geochemical

studies across these unconformities should make it

possible to gain insight into the extent of chemical

weathering at the time. Similarly as with the debate

around the genesis of the gold and the associated

pyrite and uraninite, a difference in opinion exists

with regard to the cause of widespread acidic

alteration of the siliciclastic mineral assemblages that

has been reported from throughout the Witwatersrand

Basin (Barnicoat et al., 1997; Phillips and Law, 1994).

Weathering under an aggressive, CO2 and/or CH4-

rich, acidic atmosphere should lead to considerable

chemical change in the weathered rock (loss of

alkalies and alkaline earths). Unfortunately, such a

chemical change would be difficult to distinguish

from acid leaching by post-depositional magmatic or

metamorphic fluids. Thus the question arises whether

systematic chemical changes observed across the

unconformities reflect paleosols or post-depositional

hydrothermal infiltration. Those workers who prefer a

hydrothermal model for the gold, pyrite and uraninite,

postulate basin-wide H+-metasomatism after sediment

deposition and link this with the formation of the gold,

pyrite and uraninite during hydrothermal infiltration

(Barnicoat et al., 1997). Others concluded from

geochemical and mineralogical studies of profiles

across stratigraphic units that metamorphism was

essentially isochemical, except for potassium (Sutton

et al., 1990). A further aim of this contribution is,

therefore, to assess this apparent discrepancy by

providing alteration profiles across various siliciclas-

tic Witwatersrand units and discussing the signifi-

cance of trends in the calculated chemical index of

alteration.

Finally, the distribution of redox-sensitive and pH-

sensitive minerals, such as Fe-sulphides versus Fe-

oxides, uraninite verus brannerite or feldspars versus

pyrophyllite will be examined across the various

stratigraphic units of the Witwatersrand Basin. Trends

in the distribution of these minerals along stratigraphic

directions as well as between different lithofacies will

be discussed in terms of their implications for the

palaeoenvironmental conditions. In summary, the

overall goal of this paper is to provide: (1) an up-to-

date review of the evolution of the Witwatersrand

Basin fill; (2) a best-fit genetic model for the world’s

largest gold province and thus also of for the redox-

sensitive monitor phases pyrite and uraninite; and (3)

constraints on the likely Mesoarchaean to Eoproter-

ozoic atmospheric and hydrospheric conditions based

on the spatial and temporal distribution of redox-

sensitive minerals.

2. Geological setting

The Witwatersrand Basin occupies a central

position on the Archaean Kaapvaal Craton (Fig. 2).

The causes of its development and early evolution are

linked with tectonic processes in and around this

craton, whose history is subdivided into two main

periods (de Wit et al., 1992). The first period (3.64–

3.08 Ga) saw the initial formation of the continental

lithosphere of the craton, with a major pulse of

accretion around 3.2 Ga. The second period (3.08–

2.64 Ga) was dominated by the development of

intracontinental basins and likely subduction-related

magmatism along the edge of the Kaapvaal Craton.

By the end of the Archaean Eon, the Kaapvaal Craton

had amalgamated with the Zimbabwe Craton along

the Limpopo Belt (Fig. 2).

2.1. Pre-Witwatersrand basement

The oldest known crustal fragment of the Kaapvaal

Craton is the ~3.64 Ga old Ancient Gneiss Complex

in Swaziland (Kroner and Tegtmeyer, 1994). Products

of Palaeoarchaean crust formation are well preserved

within the Barberton greenstone belt, where domi-

nantly basic–ultrabasic magmatism between 3.49 and

3.42 Ga followed tonalite emplacement between 3.55

and 3.52 Ga (de Ronde and de Wit, 1994). The mafic

to ultramafic rocks have been interpreted as remnants

of ocean-like lithosphere (de Wit et al., 1987). The

upper clastic part of the greenstone belt is interpreted

to comprise synorogenic deposits (Moodies Group)

related to 3.2 Ga accretion. A change from compres-

Fig. 2. Distribution of the main Archaean stratigraphic units of the Kaapvaal Craton. The Witwatersrand Basin fill comprises the West Rand and

Central Rand Groups; also shown is the outline of the three crustal blocks that are believed to have amalgamated by 2.8 Ga to form a single

craton (modified from Schmitz et al., 2004).


sional to transtensional tectonic activity around 3.1 Ga

was accompanied by widespread orogenic gold

mineralisation along late shear zones. This inversion

in the overall stress field marked the beginning of

intracontinental basin formation that eventually led to

the formation of the Witwatersrand Basin.

2.2. Witwatersrand basin development

The first stage of sediment deposition in what

eventually became the Witwatersrand Basin is recog-

nised in the Dominion Group. This group comprises

an up to 2250 m thick bimodal volcanic succession

with a thin basal siliciclastic unit for which a

continental rift basin of unknown extent is assumed.

Maximum and minimum age constraints are given by

U–Pb zircon ages of 3086F3 (the youngest age of

pre-Dominion basement: Robb et al., 1992) and

3074F6 Ma (the age of volcanism: Armstrong et

al., 1991), respectively. The basal siliciclastic unit

includes a conglomerate bed with abundant uraninite

and pyrite but relatively low gold content (Dominion

Reef). Palaeocurrent data consistently point to a

source area to the north or northeast (Frimmel and

Minter, 2002). Thus, a continuation of the original

basin to the south of the present distribution of

Dominion Group rocks is likely (Fig. 3).

The development of the proper Witwatersrand

Basin followed with a hiatus of almost 100 million

years. This great time gap implies that the tectonic

regime for the Dominion Basin is unrelated to that of

the subsequent Witwatersrand Basin. No agreement

Fig. 3. Simplified surface and subsurface geological map of the Witwatersrand Basin, also showing the distribution of Archaean granitoid

domes, the location of the gold fields, major faults and palaeocurrent directions of reefs in the Central Rand Group (from Frimmel and Minter,

2002).


exists on the detailed lithostratigraphic correlation

between the various gold fields across the Witwaters-

rand Basin, but a broad subdivision of the basin fill

into the West Rand and Central Rand Groups, both

constituting the Witwatersrand Supergroup, has long

been recognised (Fig. 4).

2.2.1. West Rand Group

The metasedimentary rocks of the West Rand

Group (Fig. 3) rest with angular unconformity above

the Dominion Group volcanic rocks. The group

attains a maximum thickness of 5150 m in the

Klerksdorp gold field and thins to the northeast. No

information is available from the Welkom gold field.

In the most distal section south of the Vredefort

Dome, about 2000 m of West Rand Group rocks have

been intersected in a borehole (Stevens and Preston,

1999). Sediment input throughout the West Rand

Group was consistently from the north or northeast

(Frimmel and Minter, 2002). Uranium–Pb data from

detrital zircon grains provide a maximum age of

2985F14 Ma for West Rand Group sedimentation

(Kositcin and Krapez, 2004). A minimum age for

most of the group is given by the Crown Formation

lava, the only volcanic unit within the succession, for

which a U–Pb single zircon age of 2914F8 Ma has

been obtained (Armstrong et al., 1991).

Three subgroups are distinguished based on vary-

ing shale/sandstone ratios and basin-wide disconform-

ities (Fig. 4). Within the basal Hospital Hill Subgroup,

the shale/sandstone ratio decreases up-section, with

the sandstone being predominantly quartz arenite,

Fig. 4. Generalised stratigraphic column for the Witwatersrand Supergroup (from Frimmel and Minter, 2002); also shown are the stratigraphic

positions of the main auriferous conglomerate beds (reefs) and their relative significance as gold producers (insert); average gold grade typical of

mined reefs from Frimmel and Minter (2002); SHRIMP U–Pb or Pb–Pb age data from (1) zircon (Armstrong et al., 1991), (2) zircon and

xenotime (Kositcin and Krapez, 2004) and (3) authigenic xenotime (Kositcin et al., 2003); Chemical Index of Alteration from Gartz and

Frimmel (1999), Sutton et al. (1990) and H.E. Frimmel (unpubl. data).


which is interpreted as subtidal deposits (Eriksson et

al., 1981). Evidence of tidal deposits has been

reported from several units. Thick–thin pairs of

siltstone–shale couplets within the upper Coronation

Formation (Fig. 4) are a particularly good example

(Eriksson and Simpson, 2004). Higher up in the

succession, feldspathic sandstone and quartz wacke

become more abundant (Law et al., 1990). The


sediment record of the West Rand Group reflects

fluctuations between distal fluvio-deltaic and shore-

face to offshore environments, ascribed by some

workers to eustatic sea level changes (Stanistreet

and McCarthy, 1991). Indirect evidence of eustatic sea

level changes comes from the presence of two

diamictite beds within the Government Subgroup,

which are possibly correlatives of the oldest known

glacial deposit on Earth in the Pongola Supergroup

(Young et al., 1998).

Of particular significance is also the presence of

several magnetic shale beds, because they contain

magnetite as principal Fe-phase and not Fe-sulphides

as in the coarser grained units. These shale beds have

been described from the northern and eastern parts of

the basin, as well as from the Klerksdorp gold field

(Tankard et al., 1982), where they occur at a number

of stratigraphic levels throughout the West Rand

Group (Fig. 4). Geophysical maps showing aeromag-

netic anomalies (Corner and Wilshire, 1989) that are

caused by the strong response of up-turned magnetite-

rich shale beds along the basin margin indicate a

basin-wide distribution of this lithotype.

2.2.2. Central Rand Group

The Central Rand Group lies unconformably above

the West Rand Group and reaches a maximum

thickness of 2880 m near the centre of the basin.

Similar to the West Rand Group, a series of cycles can

be distinguished, each of which comprises fluvially

dominated coarse-grained siliciclastic metasedimen-

tary rocks above an erosion surface, but in contrast to

the West Rand Group, shale (mudstone) plays a

subordinate role. Fluvio-deltaic processes dominated

sediment deposition, but tidal reworking has been

suggested for the interfaces between fluvial and

shallow marine systems (Els, 1998). A particularly

fine example of siltstone-shale couplets, that have

been interpreted as dominant and subordinate semi-

diurnal tidal currents, respectively (Eriksson and

Simpson, 2004), occur in the uppermost Randfontein

Formation (Fig. 4). During deposition of the Central

Rand Group, the palaeoslope direction changed from

a consistent dip to the south or southwest to east and

northeast along the western and southwestern margins

and to the southeast and south along the northwestern

and northern margins (Minter and Loen, 1991). Two

minor, locally developed, amygdaloidal basalt units

(Bird lava) in the Krugersdorp Formation provide the

only evidence of volcanism in the Central Rand

Group.

The Central Rand Group is subdivided into the

lower Johannesburg and the upper Turffontein Sub-

groups. The maximum age of deposition for the

sediments of the group is constrained by the youngest

age obtained on detrital zircon from the bottom of the

Johannesburg Subgroup (i.e. 2902F13 Ma: Kositcin

and Krapez, 2004). The youngest detrital zircon age

of 2872F6 Ma from the Krugersdorp Formation sets

the best available constraint for the age of the top of

this subgroup, whereas 2849F18 Ma based on detrital

zircon, or 2840F3 Ma based on detrital xenotime,

represents the maximum age for the top of the Central

Rand Group (Kositcin and Krapez, 2004). A mini-

mum age constraint on deposition of the Central Rand

Group is 2780F3 Ma, which is the oldest age

obtained on any authigenic mineral (xenotime) to

date (Kositcin et al., 2003).

2.2.3. Tectonic setting of the Witwatersrand Basin

General agreement seems to exist on the lower

part of the West Rand Group having been deposited

in a passive continental margin setting, facing an

open ocean to the south. Continental rift-related

rhyolite at the bottom of the Pongola Supergroup

(Nsuze Formation, Fig. 1), considered a correlative

of the Witwatersrand Supergroup (Beukes and

Cairncross, 1991), has an age of 2985F11 Ma

(Hegner et al., 1994), equivalent to that of the lower

West Rand Group. The inferred passive margin

setting is therefore explained by post-rift thermal

subsidence.

A change from upward-deepening to upward-

shallowing at the Hospital Hill/Government Subgroup

boundary has been used to suggest a change from a

passive margin to a foreland basin setting (Coward et

al., 1995). Others prefer a passive margin setting for

the entire group (de Wit et al., 1992). In a recent

SHRIMP detrital zircon study, a general decrease in

the complexity of zircon provenance age spectra is

recorded up-section through the West Rand Group

(Kositcin and Krapez, 2004). This is interpreted as

reflecting an increasing maturity of the hinterland as is

expected for an evolving passive margin, thus

supporting a thermal subsidence origin for the entire

West Rand Group.


The variation in palaeoslope directions in the lower

Central Rand Group, together with the increase in

continental to marine sediment ratio up-section,

indicates a change to a progressively shrinking

continental basin. This is well reflected by the detrital

zircon age spectra, which increase in complexity up-

section through the Central Rand Group. A progres-

sively greater variety of source rocks must have been

eroded through Central Rand Group times, reflecting

tectonic loading and thus a foreland basin setting for

this group. The same study also suggested that

granitoids were contemporaneously unroofed. Such

a provenance is more consistent with a back-arc fold-

thrust belt than with a foreland fold-thrust belt, which

led Kositcin and Krapez (2004) to postulate a retroarc

basin setting for the Central Rand Group.

With considerably more and better age data

available, more reliable integrated sedimentation rates

can be calculated. Interestingly, these indicate a

minimum of 63 m/my for the West Rand Group,

considerably higher than 18 m/my calculated for the

Central Rand Group. Such a difference seems to be

inconsistent with a simple retroarc foreland basin and

would point to a significant additional strike-slip

component, as suggested by Maynard and Klein

(1995). Stanistreet and McCarthy (1991) have already

emphasised both sinistral and dextral strike-slip along

the northern and western margins, respectively, and

postulated a southeast-directed tectonic escape basin.

However, some of the faults referred to by these

authors are post-Witwatersrand in age. Alternatively,

the lower integrated sedimentation rate calculated for

the Central Rand Group might simply reflect a higher

degree of sediment re-working, and thus less accom-

modation space, and not necessarily lower actual

sedimentation rates.

By analogy with the Pongola Supergroup, a major

folding event in upper Central Rand Group times

between 2837F5 and 2824F6 Ma, defined by pre-

and post-folding felsic intrusive rocks, has been

suggested (Gutzmer et al., 1999). Along the western

basin margin, the angles of unconformity increase up-

section in the Central Rand Group, which reflects

progressive uplift of the hinterland to the west

(Frimmel and Minter, 2002). From the above tectonic

synthesis, it becomes apparent that the generally used,

deeply entrenched term bWitwatersrand BasinQ can be

misleading, because it embraces at least two different

basin types, with a major sequence boundary between

the West Rand and Central Rand Groups. Therefore

the Witwatersrand Basin is best described as a

successor basin that contains erosional remnants of

tectonically stacked sediments originally deposited in

at least two fundamentally different tectonic settings.

The two mafic volcanic units within the Witwa-

tersrand Supergroup remain somewhat enigmatic, as

neither passive margins nor foreland basins are

particularly favourable settings for such volcanism.

Basaltic volcanism on passive margins can be caused

by reactivation of fundamental faults as a far-field

response to changes in the rate or direction of plate

tectonic processes or plume magmatism. By implica-

tion the Crown Formation lavas might record a

change from thermal to reactivated subsidence. In

contrast, the Bird volcanic interval in the Upper

Central Rand Group could be impactogenic, implying

a change from a subduction- to collision-related,

retroarc foreland domain (Stanistreet and McCarthy,

1991).

Based on the palaeoslope directions, the main

tectonic domains in the hinterland controlling the style

of sedimentation in the Witwatersrand Basin are to the

north and west, although palaeoslope directions in the

Evander gold field are more locally controlled. Along

the northern margin of the Mesoarchean Kaapvaal

Craton, east- to northeast-trending greenstone belts

(Murchison, Pietersberg and Giyani greenstone belts,

Fig. 2), surrounded by granitoids, yielded single

zircon age data that overlap with the age of

Witwatersrand sediment deposition (Brandl et al.,

1996; Kroner et al., 2000; Poujol, 2001; Poujol and

Robb, 1999; Poujol et al., 1996). All of these belts

seem to have a 3.2–3.3 Ga basement. In the

Murchison granitoid–greenstone terrain, metamor-

phosed granite, tonalite and rhyolite are dated between

3.02 and 3.09 Ga, overlapping with deposition of the

Dominion Group rocks. Felsic volcanic rocks and

granite from the same terrain, dated between 2971 and

2966 Ma, might reflect crustal thinning related to

early Witwatersrand rifting. Younger felsic intrusions

in the Giyani and Pietersberg granitoid–greenstone

terrains are dated at 2874 Ma, whereas, in the

Murchison terrain, granitoid emplacement is indicated

around 2901 and 2820 Ma. Shear-zone hosted gold

deposits occur in the mafic to ultramafic rocks of

these greenstone belts. Only the Pietersberg Belt hosts


a sequence of predominantly coarse-grained silici-

clastic metasedimentary rocks with minor palaeo-

placer gold occurrences (Uitkyk Formation) that are

possibly correlatives of the Witwatersrand Supergroup

(de Wit et al., 1992).

A series of intrusive events has also been identified

in the Amalia–Kraaipan granitoid–greenstone terrain

along the western margin of the craton (Anhaeusser

and Walraven, 1999; Poujol et al., 2002; Robb et al.,

1992; Schmitz et al., 2004). A minimum age for the

supracrustal successions is given by a U–Pb zircon

age of 3033F1 Ma obtained for a mafic intrusive

body. Ages between 2932 and 2926 Ma for meta-

morphic and anatectic zircon date both the time of

accretion of a ca. 2930 Ma volcanic arc and

continental collision between the so-called Kimberley

and Witwatersrand crustal blocks (Schmitz et al.,

2004). Deep seismic reflection profiles through the

west-central Kaapvaal Craton indicate a westward-

dipping subduction (de Wit and Tinker, 2004).

Subsequent crustal thickening and uplift led to

variable exhumation of, and decompression melting

in, the Kimberley block, with resulting high-level

granitoids as young as 2727 Ma, while the Witwa-

tersrand block underwent subsidence with sedimenta-

tion in the Central Rand Basin.

The age of the youngest granitoids in the Kraaipan

Belt (around 2790 Ma) is similar to the age of the

Gaborone Suite rapakivi-type granite and related

volcanic rocks in southern Botswana (2783 Ma;

Grobler and Walraven, 1993; Moore et al., 1993).

Considering the inferred retroarc foreland setting and

above age constraints for the Central Rand Group, the

younger granitoids in the hinterland to the north,

northwest and west might reflect the corresponding

magmatic arc, possibly related to a southward-dipping

subduction zone. The ocean that was closing at that

time probably separated the amalgamated Witwaters-

rand–Kimberley block from the Pietersburg block

(Fig. 2). Such a combination of foreland basin, related

to westward subduction, and retroarc basin, related to

the closure of an ocean further north, would help

explain the apparent inconsistency between integrated

sedimentation rates and a simple retroarc basin

mentioned above. The Limpopo orogeny, often used

as an explanation for the inferred foreland position of

the Central Rand Basin in the past (e.g. Burke et al.,

1986), took place more than 100 million years after

the Witwatersrand Supergroup rocks had been laid

down. Apart from great ambiguity about the existence

of a Himalayan-style Limpopo orogeny, the available

age data rule out any relationship between tectonic

events that shaped the Limpopo Belt and the

Witwatersrand Basin.

2.3. Post-Witwatersrand evolution of the Kaapvaal

Craton

Mild deformation in the form of block faulting and

subordinate folding and thrusting along the western

and northern basin margin, and low-grade regional

metamorphism of the Witwatersrand Basin fill, are

testimony to post-depositional alteration related to a

series of tectono-thermal events initiated during, and

succeeding, Witwatersrand sediment deposition.

Stratigraphically above the Witwatersrand Super-

group follows the Ventersdorp Supergroup, which

attains a thickness of more than 3600 m in places. In

the northern parts of the basin, the two supergroups

are separated by a regional angular unconformity that

is overlain by a thin fluvial conglomerate (Venterspost

Conglomerate Formation). In places, the conglomer-

ates are highly auriferous and form important ore

bodies (Ventersdorp Contact Reef) that are similar to

other Witwatersrand reefs. Further south, in the

Welkom gold field, the Ventersdorp Supergroup rests

paraconformably above the Witwatersrand Super-

group and the Ventersdorp Contact Reef is not

developed (Minter et al., 1986). This is explained by

a lack of re-working of the older, auriferous Witwa-

tersrand sediments along that contact there.

The thin siliciclastic sequence of the Venterspost

Conglomerate Formation is conformably overlain by a

thick succession of flood basalt (Klipriviersberg

Group), though localised erosional scours exist and

have been explained by contemporaneous eruption of

channelised lava flows over unconsolidated sediment

(Hall, 1997). A U–Pb zircon age of 2714F8 Ma is so

far the best constraint on the timing of eruption

(Armstrong et al., 1991), which implies a hiatus of at

least 50 million years for the underlying unconform-

ity. This is supported by the youngest detrital

xenotime age obtained on the Ventersdorp Contact

Reef, i.e. 2729F19 Ma (Kositcin and Krapez, 2004).

The Klipriviersberg Group basalt layers are over-

lain by siliciclastic and volcanic, predominantly felsic


and minor mafic, deposits of the Platberg Group. The

siliciclastic sediments reflect alluvial fan deposits that

prograded into lacustrine environments. A U–Pb

zircon age of 2709F4 Ma for Platberg Group basalt

(Armstrong et al., 1991) shows that the bimodal

volcanic activity recorded in the Ventersdorp Super-

group was short-lived. Deposition in a continental rift

that evolved into an Atlantic-type continental margin

is inferred from seismic profiles that indicate consid-

erable thickening of the Ventersdorp Supergroup

towards the west of the craton (Tinker et al., 2002).

A link between Ventersdorp rifting and collision

between the Kaapvaal and Zimbabwe Cratons, as

suggested previously (Burke et al., 1985), is unlikely.

Granulite facies metamorphism associated with south-

ward thrusting of the Southern Marginal Zone of the

Limpopo Belt on to the Kaapvaal Craton is dated at

2691F7 Ma Ma (Kreissig et al., 2001) and syn-

tectonic granites from the Central Zone of that belt

range in age from 2664 to 2572 Ma (McCourt and

Armstrong, 1998), later than onset of Ventersdorp

flood basalt extrusion. However, the age of the

Klipriviersberg Group overlaps with an earlier north-

ward thrusting phase, dated at 2729F19 Ma (Passer-

aub et al., 1999), that affected greenstone belts along

the northern Kaapvaal Craton. The Klipriviersberg

extension could thus be explained by southward-

directed subduction beneath the Kaapvaal Craton

prior to continental collision. Alternatively, the

ultimate cause of Ventersdorp rifting might not reside

in crustal plate tectonics but as a result of a mantle

plume (Hatton, 1995).

A second thrust event in the Witwatersrand Basin,

post-Platberg and pre-Transvaal in age (Roering,

1990), is most likely a distant expression of the

collisional tectonic processes in the Southern Mar-

ginal Zone of the Limpopo Belt. At that stage, lower

greenschist facies metamorphic conditions were

attained for the first time in the Witwatersrand rock

column, at least along the northern margin of the basin

(Coetzee et al., 1995). This phase of compression and

uplift was followed by regional peneplanation. The

corresponding unconformity represents a major

sequence boundary, separating the Ventersdorp from

the overlying Transvaal Supergroup.

Of significance to the debate on the genesis of the

Witwatersrand gold and of Archaean atmospheric

evolution is the presence of a thin, but laterally

extensive, basal conglomerate and sandstone unit

(Black Reef Quartzite Formation) that rests on the

post-Ventersdorp erosion surface. This basal silici-

clastic unit contains a conglomerate (Black Reef) that

is almost indistinguishable from the auriferous,

uraniferous and pyrite-rich Witwatersrand reefs,

except for a lower metamorphic grade. A syn-

Witwatersrand or syn-Ventersdorp age for the Black

Reef, as suggested by Phillips et al. (1997) is

untenable, because of field relationships (Els et al.,

1995) and geochronological data. The Black Reef

Quartzite Formation reflects fluvial sedimentation in

channels on a locally deeply incised palaeosurface,

braided fluvial and braid delta deposits that grade into

shallow marine deposits. A locally transitional and

conformable upper contact with the overlying pre-

dominantly chemical, marine sedimentary successions

of the Chuniespoort Group is unequivocal evidence of

an early Transvaal Supergroup age for this genetically

important auriferous formation (SACS, 1980). An

indirect age constraint for the Black Reef Quartzite

Formation exists in the form of a Pb–Pb single zircon

age of 2642F2 Ma (Walraven and Martini, 1995),

obtained for a lava in a stratigraphic correlative

(Vryberg Formation). In agreement with the above

timing of tectonic activity in the Limpopo Belt, the

sediment sources were located to the north (Barton

and Hallbauer, 1996; Els et al., 1995).

The overlying Chuniespoort Group starts with a

transgressive black shale, followed by thick platform

carbonates (Malmani Subgroup), banded iron forma-

tion (Penge Formation) and, erosive into the older

strata and reflecting a eustatic sea level fall, lacustrine

deposits including a glaciogenic diamictite (Duitsch-

land Formation). An upper age constraint is given by a

Pb–Pb single zircon age of 2550F3 Ma for a tuff layer

in the lowermost formation of the Malmani Subgroup

(Walraven and Martini, 1995), whereas U–Pb single

zircon ages of 2432F31, 2465F7 and 2480F6 Ma

reported for the stratigraphic equivalent of the Penge

Formation (Nelson et al., 1999; Walraven and Martini,

1995) set some constraints on the minimum age of the

group. Subsidence rates calculated from the thickness

of the group are in good agreement with thermal

subsidence due to lithospheric cooling after the

Ventersdorp thermal anomaly (Tinker et al., 2002).

A further major sequence boundary separates the

Chuniespoort Group from the overlying Pretoria


Group. These two groups, though unified as Transvaal

Supergroup in the literature, have little in common.

The boundary is marked by a strongly weathered

erosional unconformity reflecting a hiatus of ~80

million years, above which lies a volcano-sedimentary

succession with alluvial fan and fan delta deposits

with minor marine influence. This hiatus is likely to

reflect a glacio-eustatic sea level fall, because the

unconformity is overlain by glaciogenic diamictite.

Alluvial to lacustrine deposits grade into shallow to

deep marine deposits before a further erosional

unconformity at the top of the Timeball Hill For-

mation marks another sequence boundary. The uncon-

formity above that formation contains the paleosol

examples that have been used to constrain the rise in

atmospheric O2 during the Eoproterozoic (Beukes

et al., 2002).

Deposition of the Pretoria Group was followed

after another hiatus by the contemporaneous extrusion

of felsic volcanic rocks (Rooiberg Group), and

emplacement of the mafic to ultramafic 2059F1 Ma

(Buick et al., 2001) Rustenberg Suite and 2054F2 Ma

(Walraven and Hattingh, 1993) granitic Lebowa Suite,

both of the Bushveld Igneous Complex, to the north

of the Witwatersrand Basin. An elevated crustal

geotherm related to large-scale magmatic underplating

and intraplating also probably resulted in the intrusion

of alkali granite and mafic plutons within the

Witwatersrand Basin, some dated at 2078F12 Ma

(Moser, 1997). The last major disturbance experi-

enced by the Witwatersrand strata is reflected by a 30-

km-wide domal structure, the Vredefort Dome (Fig.

3), that could well represent the oldest (2023F2 Ma;

Kamo et al., 1996; Moser, 1997) and largest (250 to

280 km diameter; Henkel and Reimold, 1998) known

terrestrial impact structure.

2.4. Post-depositional alteration of the Witwatersrand

Supergroup

The multitude of tectono-thermal events that

affected the Kaapvaal Craton after the deposition of

the Witwatersrand Supergroup sediments is reflected

by recrystallisation, formation of metamorphic min-

eral assemblages, and likely also metasomatic reac-

tions, thus masking geochemical and mineralogical

characteristics of the sedimentary protolith. A series of

detailed petrological, geochronological and fluid

inclusion studies made it possible to distinguish

between several stages of post-depositional alteration

throughout the basin.

Initial fluid–rock interaction most probably

involved leaching by basin-wide penetration of

meteoric water shortly after sediment deposition. A

prime Phanerozoic analogue of this process has been

documented for the Parana Basin of Brazil (Franca et

al., 2003) and a similar scenario is likely to have

affected the Witwatesrand Basin fill (Phillips et al.,

1990), when uplift of at least one basin margin

provided a steep hydraulic gradient for groundwater to

flow towards the basin centre. A decrease in pH with

increasing burial is predicted from the maturation of

organic material and the release of H+ from dehy-

dration reactions. This, in turn, could have led to the

dissolution of detrital feldspars, thus generating a

secondary porosity for further diagenetic fluid flow.

Following diagenesis, dated between 2776 and

2780 Ma (Kositcin and Krapez, 2004), a first stage of

lower greenschist facies metamorphism was attained

along the northern basin margin, coeval with high-

grade metamorphism and tectonism in the Southern

Marginal Zone of the Limpopo Belt. In most parts of

the basin, regional peak metamorphic conditions of

300 to 350 8C at 3 kbar were likely achieved during

deposition of the Pretoria Group (Frimmel and Minter,

2002). Only around the Vredefort Dome were

metamorphic grades up to amphibolite facies reached

(Gibson and Wallmach, 1995). There the peak of

metamorphism is ascribed to the emplacement of the

Bushveld Igneous Complex, which is supported by an

anomalously high geothermal gradient (Frimmel,

1997).

Several stages of enhanced fluid circulation

through the Witwatersrand Basin are recognised.

These range from diagenetic basin dewatering to a

number of hydrothermal fluid infiltration events.

Hydrothermal rutile, zircon and xenotime age data

are considered most robust, whereas Rb–Sr, U–Pb and

Pb–Pb ages of various hydrothermal precipitates have

only limited geochronological meaning (Zartman and

Frimmel, 1999). The most reliable of these age data

(Armstrong et al., 1995; Kositcin and Krapez, 2004;

Robb et al., 1990) cluster around 2720 Ma (Ven-

tersdorp extension), 2580 Ma (early Transvaal thermal

subsidence), possibly around 2200 Ma (Pretoria

extension), and 2060 Ma (Bushveld event). A further


hydrothermal infiltration event has been ascribed to

the Vredefort impact, based on cross-cutting relation-

ships with dated pseudotachylyte, fluid inclusion and

gold chemistry data (Frimmel et al., 1999). Although

there is some evidence for post-2.0 Ga hydrothermal

alteration from K–Ar age spectra on very fine-grained

white mica (Zhao et al., 1999), none of these younger

events, which are far field effects of various stages of

accretion along the craton margin, are of great

significance for the post-depositional alteration his-

tory of the Witwatersrand sediments.

3. Sedimentological and mineralogical

characteristics of the orebodies

3.1. Host lithofacies

The Witwatersrand gold orebodies typically occur

in conglomerate beds (reefs) on unconformities within

major stratigraphic sequences (Pretorius, 1981). Gold

has been mined from at least 30 such reefs, the most

important of which are shown in Fig. 4, with those

from the Central Rand Group accounting for 90% of

total production. In addition, the base of the younger

Ventersdorp and Transvaal Supergroups also host, in

places, economic to subeconomic, orebodies of

comparable lithology and mineralogy. The orebodies

comprise various fluvial lithofacies that range from

clast-supported oligomictic conglomerate to loosely

packed conglomerate, pebbly arenite, or just pebble

lag surfaces associated with trough cross-bedded

quartz arenite. In exceptional circumstances, ore is

associated with immature debris flow lithofacies.

These include black argillite in the Beatrix Reef

(Minter et al., 1988), where overlying playa facies

covered eroded older ores that occurred as eroded

subcrops, and risers between terraces of the Venters-

dorp Contact Reef (Henning et al., 1994), where

Fig. 5. (A) Oligomictic pebbly quartz arenite (reef), Vaal Reef, Stilfontein

unconformity and 3 cm above the base on a bedding plane defining the to

ventifact; scale bar=1 cm. (B) same as A, but highlighting the position of

pyrite pebble lags between quartz pebble lags, Basal Reef, Welkom gold

Ventersdorp Contact Reef, Tau Lekoa mine, Klerksdorp gold field, showi

contact with Klipriviersberg Group mafic lava flow is erosive into originally

rounded, oolitic pyrite (arrow) within pyrite pebble lag shown in C (white r

cross beds in pebbly quartz arenite at the base of the Basal Reef, Welkom g

(see Minter et al., 1993); scale bar=1 cm.

underlying immature sedimentary rocks were eroded

and incorporated into the alluvium during entrench-

ment.

The Witwatersrand reef rocks are generally

described as quartz pebble conglomerates. Exceptions

exist, however, such as the Steyn Reef in the Welkom

gold field, which is polymictic with as much as 30

vol.% yellow quartz porphyry in both the pebble and

sand fractions. The Kimberley Reef in the Evander

gold field contains as much as 50% chert pebbles

(Tweedie, 1986). The average pebble assemblage is

85 vol.% vein quartz, 12 vol.% chert, 2 vol.% quartz

porphyry, and 1 vol.% metamorphic clasts (Frimmel

and Minter, 2002).

The orebodies range in thickness from decimetres

to a few metres and are confined between a basal

degradation surface, usually an angular unconformity,

and an upper planar bedding surface that separates it

from overlying quartz wacke or siltstone (Fig. 5A,B;

Minter, 1991). They have a lenslike geometry and

define fluvial bar-and-channel bedforms that display

unimodal palaeocurrent directions (Smith and Minter,

1980). Multichannel sequences of conglomerate and

quartz arenite, deposited by repeated flood- and

waning-stage flows, make up the thicker orebodies

(Fig. 5D).

Depositional environments that have been recon-

structed from the geometry of the basal degradation

surface, coarseness of the sediment, and the distribu-

tion of lithofacies, range from proximal alluvial fan

(e.g. EA Reefs in the Eldorado Formation), to terraced

fluvial deposits (e.g. Ventersdorp Contact Reef;

Henning et al., 1994) to braid plains (Composite

Reef; Tucker, 1980), to braid deltas that merge with

shoreline environments. In some places, aeolian

deflation has been suggested to account for the planar

top of the orebodies (Minter, 1999), whereas in others,

shoreline encroachment and associated wave action,

followed by burial beneath fine-grained sediment is

mine, Klerksdorp gold field; note the veneer of bitumen on the basal

p of pebbly layer; cross-bedding (so) in hanging wall, and an in-situ

fine- to coarse-grained rounded pyrite. (C) Upward fining imbricate

field; scale bar=1 cm. (D) Underground exposure of upper half of

ng multichannel sequences of conglomerate and quartz arenite; top

unconsolidated conglomerate; scale bar=50 cm. (E) Truncated, well

ectangle), scale bar=0.5 cm. (F) Concentration of gold particles along

old field. Three foresets converge with the bottomset toward the right



indicated as in the Basal Reef (Buck, 1983) and

Carbon Leader Reef (Buck and Minter, 1985).

3.2. Textural, chemical and isotopic features of the

main ore components

The bulk of gold, uraninite and pyrite, whose

origin is debated, occur together with undoubted

allogenic detrital minerals, such as mechanically

abraded zircon and chromite grains, and more rarely

PGE-minerals and diamond (for a complete list of

minerals identified to date in Witwatesrand reefs see

Phillips and Law, 2000), on degradation surfaces

marked by pebble lags or the base of clast-supported

conglomerate beds, cross-bedded foresets, bottomsets,

and coset boundaries (Fig. 5F). In thick (1–2 m)

conglomerate units, representing multichannel

sequences, the allogenic minerals are concentrated

along the basal degradation surface of each graded

bed. The highest concentrations of allogenic minerals

are found above unconformity surfaces, which reflect

a direct relationship to the amount of erosion. Cut-

and-fill structures, such a trough cross beds, contain

greater allogenic mineral concentrations than aggrada-

tional structures, such as planar cross beds. Experi-

ments (James and Minter, 1999) have confirmed that

the dominant concentration mechanism was by

selective entrainment of coarser, less dense particles

during bedload transport (Slingerland and Smith,

1986).

Average gold grades of mined reefs range from 3 to

25 g/t (Fig. 4). Case studies on the element

distribution within reefs, such as the Kimberley Reef

(Rasmussen and Fesq, 1973), the Steyn Reef (Frim-

mel and Minter, 2002), the Vaal Reef (Fox, 2002), and

the Ventersdorp Contact Reef (H.E. Frimmel, unpubl.

data) illustrate a relatively good positive correlation

between U and Au, but only poor correlation between

Au and Zr, and Cr—elements that are controlled by

detrital phases (zircon, chromite) that were not

susceptible to hydrothermal mobilisation. All data

sets show gold enrichment being linked with Zr

enrichment, but samples rich in Zr typically do not

contain elevated Au contents. A detailed study of

different lithofacies in a channel sidebar of the Steyn

Reef, Welkom gold field, where 13 aggradation events

can be separated (Frimmel and Minter, 2002) showed

that degradation surfaces were preferentially mineral-

ised with an average of 38 ppm Au, 410 ppm Zr, 1750

ppm U, and 300 ppm Cr. Higher concentrations of all

these elements were recorded in the coarser sediment

fraction, reflecting higher flow velocities. Similar

observations were also made quantitatively on other

reefs, such as the Leader Reef (Smith and Minter,

1980) and the Ventersdorp Contact Reef (H.E.

Frimmel, unpubl. data).

The Witwatersrand reefs are not only extremely

rich in gold but also are one of the world’s largest

uranium depositories. Between 1952 and 1975, as

much as 1.5�106 t U3O8 (Frimmel et al., in press) was

produced at an average grade of 271 ppm (Camisani-

Calzolari et al., 1984). One of the richest reefs was the

Monarch Reef, a small-pebble distal placer, mined at a

mean grade of 2860 ppm U3O8. The principal U-

minerals are uraninite, brannerite and leucoxene, with

a systematic decrease in uraninite/brannerite ratio up-

section. For example, in the Welkom gold field, the

uraninite/brannerite ratio is 8.7 in the Steyn Reef,

whereas in the younger Beatrix Reef it is zero (Minter

et al., 1988). Similarly, all the U in the Black Reef,

from which a specimen with as much as 3350 ppm

U3O8 has been reported (Bourret, 1975), occurs as

brannerite.

There is a broad systematic trend in the U/Au ratio

up-section. In contrast to the Central Rand Group

reefs, the Dominion Reef is highly uraniferous but

does not contain significant amounts of gold. Within a

given stratigraphic unit, both Au and U concentrations

decrease from the basin margin towards its centre, but

at different rates. A systematic increase in the U/Au

ratio down the paleoslope from 10�3 to 10 was noted

in the Welkom gold field, with uraninite being

enriched in the more distal facies, probably as a result

of hydraulic mineral sorting (Minter et al., 1986). This

is also illustrated by the coarse-grained Main Reef,

which has an average pebble size of 37 mm and a U/

Au ratio of less than 25 as opposed to the relatively

finer grained, more distal Monarch Reef, whose

average pebble size is 16 mm and U/Au ratio is 128

(Vennemann et al., 1995).

It has been noted in many studies that Witwaters-

rand gold is intimately associated with carbonaceous

matter, which occurs as stratiform seams and as

spherical, glassy globules. Both forms have been

recognised as metamorphosed solidified hydrocar-

bons, i.e. pyrobitumen (Gray et al., 1998). The


bcarbon seamsQ occur preferentially in deposits

reflecting distal environments, but are conspicuously

absent from proximal, high-energy deposits. Hydro-

carbon derivation from original algal mats is sug-

gested by the distribution of the bcarbon seamsQ onpalaeosurfaces, on sedimentary accumulation surfaces

and on trough cross-beds and ripple surfaces (Buck,

1983; Minter, 1981). In those reefs that contain such

carbonaceous matter, such as the Carbon Leader,

Basal and Vaal Reefs, the highest Au and U

concentrations are located in the zones that are

particularly rich in pyrobitumen. Nagy (1993) esti-

mated that about 40% of all mined Witwatersrand

gold was hosted by such bitumen seams. On a micro-

scale, the pyrobitumen-filled microfractures typically

contain some gold. Oil-migration, and thus by

implication gold transport, has been suggested both

prior (England et al., 2002a) and during (Jolley et al.,

2004) fracturing. It should be noted, however, that

many highly auriferous reefs contain little or no

carbonaceous matter; a number of economic reefs,

such as the Main and Kimberley reefs in the Central

and West Rand gold fields do not contain any

noteworthy amounts of bitumen.

The bitumen is interpreted to have formed by the

polymerisation and crosslinking of liquid hydrocar-

bons around irradiating grains, predominantly uranin-

ite, in the host sedimentary rock (Schidlowski, 1981).

Oil-bearing fluid inclusions provide direct evidence of

oil migration through the Witwatersrand sedimentary

rocks (Drennan et al., 1999; England et al., 2002a).

Bitumen derivation from a variety of biomass in a

reducing environment, with subsequent short-range

hydrothermal mobilisation is indicated by organo-

geochemical, bulk and molecular C isotopic studies

(Spangenberg and Frimmel, 2001).

The phases associated with the gold, which are

most important for the topic of this paper are pyrite

and uraninite. The textural, mineral chemical and

isotopic characteristics of these phases will, therefore,

be discussed in somewhat greater detail below.

3.2.1. Pyrite

Pyrite is the most common heavy mineral in all of

the fluvial deposits of the Witwatersrand. Only in

marine sedimentary rocks, such as shales in the West

Rand Group, are Fe-oxides (predominantly magnetite)

found instead of pyrite (Frimmel, 1996). Locally, such

as in parts of the Ventersdorp Contact Reef, Klerks-

dorp gold field, pyrrhotite is the stable Fe-sulphide

instead of pyrite. There the distribution of the two Fe-

sulphides seems to be controlled by locally variable

oxygen fugacity of post-depositional fluids. Gener-

ally, the pyrite occurs in a number of different textural

forms (England et al., 2002b; Hallbauer, 1986;

Ramdohr, 1958) that are grouped into (1) rounded,

compact, (2) rounded, porous, and (3) euhedral.

The rounded compact variety is by far the most

abundant form of pyrite in all reefs (Fig. 5C) except

for the Ventersdorp Contact Reef, where the pyrite

grains are predominantly euhedral. However, etching

of these euhedral grains reveals that most have one or

more rounded cores, with the euhedral outline being

an artifact of secondary, authigenic/hydrothermal

overgrowth around pre-existing rounded, compact

pyrite cores (Fig. 6I; England et al., 2002b; Frimmel

and Minter, 2002). Evidence of mechanical abrasion

of the rounded grains is given by truncation of

oscillatory growth zonation, defined by variable As

contents at grain boundaries (McLean and Fleet,

1989). A crystallographic study (Fleet, 1998)

revealed that some of these rounded grains are single

crystals and not the polycrystalline or twinned

crystals one would expect if they were pseudomorphs

after bblack sandsQ as suggested by Phillips and

Myers (1989) and Phillips and Law (2000). A further

argument against basin-wide sulphidation of bblacksandsQ is that the latter typically comprise titanomag-

netite. Pyritisation of such a precursor characteristi-

cally leads to intergrowths of minute rutile needles.

Such rutile-bearing pyrite pseudomorphs are the

exception and not the rule in the Witwatersrand reefs

(Ramdohr, 1958). The rounded pyrite grains from the

Witwatersrand are, however, devoid of Ti. Associated

with this form of pyrite are similarly rounded,

compact arsenopyrite and cobaltite particles (Saager

and Oberthur, 1984; England et al., 2002b). Genet-

ically significant mineral inclusions in rounded pyrite

are feldspar (Fig. 6F), calcite, corundum and spes-

sartine, which are absent in the metamorphic mineral

assemblage of the metasedimentary host rocks.

Examples of several centimetre-thick, almost mono-

mineralic, fining-upward pyrite beds, displaying

imbrication with the same orientation as intercalated

quartz pebble lags, occur within the Basal Reef of the

Welkom gold field (Fig. 5C).



The porous pyrite displays a variety of internal

textures, ranging from laminated aggregates, and

rounded concretions to oolitic-colloform and dendritic

forms. Many concretionary and colloform varieties

are fragmented, broken and have their internal

structures truncated, from which mechanical transport

is inferred (Fig. 5E).

Post-sedimentary, euhedral to subhedral pyrite

occurs preferentially adjacent to zones of hydro-

thermal alteration, such as veins and faults. This form

is typically associated with other authigenic/hydro-

thermal sulphides (chalcopyrite, cobaltite–gersdorf-

fite, pyrrhotite, galena and arsenopyrite) and

pyrobitumen (Gartz and Frimmel, 1999). Euhedral

pyrite overgrowths are common and, in places, are

contiguous with pyrite that fills fracture or pore

spaces.

A laser-ablation sulphur isotope study (England et

al., 2002b) revealed that the rounded pyrite forms

have a wide range in d34S values (�5.0x to +6.7x),

not only at the mine and stope-face scale but even at

the sample scale over less than 1.5 cm2. In a previous

SHRIMP study (Eldrige et al., 1993), large hetero-

geneities in d34S (�7x to +32x) were noted in

single pyrite grains and between different morpho-

logical types. Such heterogeneity is difficult to

reconcile with precipitation from a geochemically

homogeneous hydrothermal fluid, and more likely

reflects variation in pyrite from the eroded source rock

and/or microbial sulphate reduction in the depositio-

nal environment. The heterogeneity in rounded pyrite

is in contrast with the narrow range in d34S values

obtained for authigenic/euhedral pyrite (�0.5x to

+2.5x), which is also distinguished from rounded

pyrite by higher Ni and As contents as well as gold

inclusions. Of particular significance are two textur-

ally adjacent but isotopically contrasting ooid-like

Fig. 6. Photomicrographs illustrating morphological and textural features o

light, scale bars=0.2 mm, except for scanning electron microscope image

particles occurring together on a mm-scale. (B) Gold micro-nugget (Au) w

shown in A. (C) Spheroidal gold micro-nugget (Au I) with secondary,

secondary gold filling a fracture within a detrital zircon grain that pierce

overfolded rims next to rounded pyrite; all of the above from the same han

Sericite pseudomorphs after K-feldspar (Fsp–Psm) as inclusions within rou

Rand gold field; matrix silicates are chloritoid (Ctd) and muscovite (M

fragment and quartz pebble, Ventersdorp Contact Reef, Klerksdorp gold f

II), B-Reef, Free State Geduld mine, Welkom gold field. (I) Euhedral,

Ventersdorp Contact Reef, Klerksdorp gold field.

pyrite grains from the Ventersdorp Contact Reef,

which display a strong isotopic zonation but of

opposite signs (England et al., 2002b). The d34S

ratios in one of the two grains increase systematically

from �4.1x in the core to +1.4x in the rim, whereas

those in the other grain decrease from �0.8x in the

core to �4.5x in the rim. The former has been

interpreted by these authors as indicative of sulphi-

dation of an original sulphate grain. The latter grain

reflects a different provenance and highlights that the

isotopic differences must be pre-depositional and

cannot be due to fluctuations in redox potential of a

hydrothermal, sulphidising fluid that mixed with local

meteoric formation water during diagenesis as pro-

posed by Phillips and Law (2000).

Attempts to date the various forms of pyrite using

the U–Th–Pb isotope systems (Barton and Hallbauer,

1996; Poujol et al., 1999; Zartman and Frimmel,

1999) were met with mixed success. Authigenic and

hydrothermal pyrite is typically enriched in urano-

genic Pb, whereas the rounded forms have a less

radiogenic isotopic signature. However, absolute Pb–

Pb ages need to be viewed with caution as the 238U

and 235U decay schemes were likely decoupled,

presumably by the selective diffusion of 222Rn from

uraninite and its subsequent capture in hydrothermal

precipitates, leading to erroneous ages (Zartman and

Frimmel, 1999). More reliable are Re–Os data

obtained on rounded, compact pyrite from the Vaal

Reef, which yielded an age of 2.99F0.11 Ga (Kirk et

al., 2001). This is older than the time of sediment

deposition, which provides a strong argument for the

detrital origin of much of the pyrite.

3.2.2. Uraninite and leucoxene

Most of the uraninite particles are well rounded

and enclosed, or partially replaced, by bitumen, thus

f Witwatersrand gold orebodies (combined transmitted and reflected

s C and D: 0.1 mm): (A) Contrasting morphological types of gold

ith overfolded rims next to rounded pyrite (Py) from same sample as

well crystallised gold overgrowth (Au II). (D) Same as in C, but

s a rounded gold grain. (E) In-situ gold micro-nugget (Au I) with

d specimen of Basal Reef, Welkom gold field, shown in Fig. 5B. (F)

nded pyrite from the Kimberley Reef, South Roodepoort mine, West

us). (G) Hydrothermal gold with chlorite in matrix between lithic

ield. (H) Gold inclusions within secondary, hydrothermal pyrite (Py

secondary pyrite overgrowths around older, rounded cores (Py I),


explaining the good correlation between bitumen

seams and U content (Minter, 1978). Considering

the relationship between bitumen formation and

locally available radioactivity, at least some of the

uraninite must be older than the hydrocarbons, and

hence detrital or early diagenetic. Secondary, post-

depositional uraninite and other U-bearing minerals,

mainly brannerite and uraniferous leucoxene, are

considered to be products of partial mobilisation of

the earlier, rounded uraninite (England et al.,

2001a). A detrital origin of the rounded uraninite

grains is suggested by their mineral chemistry and

U–Pb geochronology. Even on a thin section scale,

adjacent grains show a great variation in Th/U ratio,

which reflects provenance from a variety of source

rocks. Many uraninite grains are rich in Th (average

3.9 wt.%), which is inconsistent with a low-temper-

ature hydrothermal origin but indicative of granitic

to pegmatitic sources (Feather and Glathaar, 1987;

Grandstaff, 1981). Direct dating of uraninite grains

from the Dominion Reef yielded an U–Pb age of

3050F50 Ma (Rundle and Snelling, 1977). This age

overlaps with that of Dominion Group sedimenta-

tion but is distinctly older than the Witwatersrand

sediments.

Another important U-mineral in the Witwatersrand

assemblages is brannerite with a composition close to

UTi2O6. It is characteristically of secondary origin,

derived from the oxidation of uraninite in the presence

of rutile, which, in turn, is an alteration product

derived from original detrital ilmenite and minor

titanomagnetite particles. Form relics of rounded

ilmenite occur concentrated on all scour surfaces

throughout the Witwatersrand Supergroup. Most of it

is, however, altered to leucoxene that is in many cases

highly uraniferous. Typical TiO2 concentrations in the

siliciclastic metasedimentary rocks are between 0.1

and 1 wt.%. The very fine-grained nature of the

rounded leucoxene pseudomorphs after ilmenite

points towards them being the result from weathering.

This is in contrast to the presence of distinct, euhedral

to subhedral rutile and brannerite grains, both of

which are related to the hydrothermal oxidation of

original ilmenite and uraninite. Direct dating of such

rutile from the West Rand Group (Robb et al., 1990)

confirms such an age relationship as it yielded a U–Pb

age of 2578F34 Ma, which is markedly younger than

the age of sediment deposition.

3.2.3. Gold

It has been noted by many workers that the

Witwatersrand gold appears late in the paragenetic

sequence (Ramdohr, 1958), with most of it occurring

in textural association with bitumen, hydrothermal/

metamorphic chlorite (Gartz and Frimmel, 1999) or

pyrophyllite (Barnicoat et al., 1997), along micro-

fractures (Jolley et al., 2004), and as inclusion within

euhedral, secondary pyrite (Fig. 6H). In contrast to the

latter, rounded pyrite is typically devoid of gold

inclusions. This generation of gold displays either

euhedral crystals or dendritic or otherwise irregularly

shaped habit. In a few instances, however, gold

particles have a completely different morphology. In

contrast to the above, they display rounded, spher-

oidal, disc-like and toroidal forms. Of particular

importance is that both morphological types, the

rounded to torroidal and the dendritic to euhedral,

secondary gold, can occur together on a micro-scale

(Fig. 6A,C,D), i.e. within millimetres in the same thin

section (Minter et al., 1993). This provides strong

evidence for a polyphase gold entrapment history,

with the rounded particles derived by mechanical

(fluvial) transport and secondary gold by precipitation

from a hydrothermal fluid.

Considerable compositional variability with

respect of Au:Ag:Hg ratios in gold particles exists

between reefs, within a given reef, and in some cases

even on a micro-scale within a given thin section

(Frimmel and Gartz, 1997; Reid et al., 1988).

Individual gold particles are, however, homogeneous,

which is readily explained by the diffusion rates of Ag

and Hg through gold at the temperatures to which the

Witwatersrand rocks were subjected during burial and

metamorphism (Frimmel et al., 1993). Only gold

particles in quartz veins that have been ascribed to the

Vredefort impact, based on field relationships with

impact-related pseudotachylite, have internal compo-

sitional variability. This is readily explained by the

short duration of that event that did not permit

diffusional homogenisation (Frimmel and Gartz,

1997).

First attempts to date the gold directly by the Re–

Os method yielded ages that are older than that of

sediment deposition (Kirk et al., 2002). Four gold

samples from the Vaal Reef define an isochron that

corresponds to an age of 3016F110 Ma and, when

combined with rounded pyrite from the same hand-


sample, define an age of 3033F21 Ma. Exceedingly

high Os concentrations, of as much as 4.16 ppm,

reported for gold from the Vaal Reef (Kirk et al.,

2001), are possibly an artifact of contamination by

minute inclusions of platinum-group element miner-

als. However, a series of subsequent analyses of gold

from the Vaal Reef (Kirk et al., 2002) and from the

Basal Reef, both of toroidal micro-nuggets and

secondary, hydrothermal gold crystals (Frimmel et

al., in press) show consistent Re concentrations

between 4 and 37 ppb and Os concentrations between

2 and 15 ppb. These values are one to four orders of

magnitude greater than those for younger gold

deposits as well as for average continental crust (Kirk

et al., 2002).

3.3. Sediment provenance

Derivation of the pebbles from mainly Archaean

granite and pegmatite (55%) as well as mesothermal

quartz veins and marine chert (45%) is indicated by

oxygen isotope data (Vennemann et al., 1992, 1995).

A granitic and/or pegmatitic source is further indi-

cated by rare grains of detrital cassiterite, molybdenite

and columbite (Feather and Koen, 1975) and by bright

cathodoluminescence of many detrital quartz grains

(Gartz, 1996). Furthermore, the concentrations of

granitophile elements, such as Zr, Ta, Th, and rare

earth elements, show a very good correlation (rN0.9)

with each other in the conglomerates.

By comparison with Archaean greenstone terrains

in the Kaapvaal Craton, which typically contain, apart

from the dominant mafic to ultramafic volcanic rocks,

also highly siliceous chemical sedimentary rocks and

felsic volcanic rocks, an equivalent to such terrains is

inferred as source region from the abundance of chert

and locally quartz porphyry pebbles in the Witwa-

tersrand conglomerates. A mafic to ultramafic com-

ponent in the source area, as expected for an Archaean

granitoid–greenstone terrane, is indicated by the

abundance of detrital chromite and subordinate

platinum group elements (PGE)-bearing minerals in

the conglomerates, but also by elevated Cr, Co and Ni

concentrations in all Witwatersrand shale units

(Wronkiewicz and Condie, 1987). The ratios between

the different PGE is surprisingly consistent through-

out the Witwatersrand gold mines with (Os+Ir)/

(Os+Ir+Pt+Ru) around 0.7, but significantly different

from younger deposits, such as the Rustenberg

Layered Suite, dunite and kimberlite, for which this

ratio is around 0.1 (De Waal, 1982). This difference

may reflect a high maturity of the placer (Cousins,

1973). However, the consistency in both the PGE

mineralogy and the PGE ratios along strike and down-

slope, and the relative proximal position of the

Witwatersrand placer deposits, may be an indication

of a specific source area characteristic, i.e. that of a

chondritic to subchondritic mantle, as recently pro-

posed for PGE alloys in the Evander gold field

(Malitch and Merkle, 2004). Furthermore, the prior

existence of a relatively stable cratonic block is

implied from the rare presence of diamond in some

reefs (Feather and Koen, 1975; Ramdohr, 1958),

probably related to the presence of kimberlite pipes in

the source area.

Detrital zircon age spectra (Kositcin and Krapez,

2004) indicate the following ages of significant felsic

rocks in the source area: 3310–3300, 3090–3060,

2990–2980, 2950–2940, and 2920–2910 Ma. The

zircon provenance age spectrum for the Central Rand

Group is considerably more complex and spans a

wider range (3450 to 2870 Ma) than that for the West

Rand Group (3300 to 2960 Ma). This confirms the

inferred tectonic setting of a passive margin for the

West Rand Group, with sediment supply from fewer

sources and no tectonic rejuvenation, and that of a

foreland basin for the Central Rand Group, with

increasingly more varied source rocks, continuous

tectonic rejuvenation and erosion to older stratigraphic

levels. Corresponding counterparts for all of the

observed detrital zircon age modes are known from

the surroundings of the Witwatersrand Basin (Frim-

mel et al., in press). Detrital xenotime ages ranging

from 2840 to 2813 Ma are not represented among the

detrital zircon ages, most likely because they were

derived from high-U granitoids, whose zircon grains

would have been metamict (Kositcin and Krapez,

2004). High-U granite and pegmatite bodies of

comparable age are known from the southern Murch-

ison Belt (Poujol, 2001; Poujol and Robb, 1999), near

the Giyani Belt (Kroner et al., 2000) and the

Barberton Belt (Meyer et al., 1994).

Almost 3 billion years of erosion would make any

of the currently exposed tectonic units improbable

source areas for the Witwatersrand sediments. A

nearly complete overlap of detrital zircon and


xenotime age spectra with ages from Palaeo- to

Neoarchaean granitoid–greenstone terranes surround-

ing the Witwatersrand Basin is, however, evident and

source areas that correspond to at least some of these

terranes are therefore possible. Ages that correspond

to the oldest detrital zircons from the Central Rand

Group are reported only from the Barberton Belt.

Equivalents to the detrital zircon age mode of 3310–

3300 Ma (West and Central Rand Groups) are known

from the Giyani and Barberton Belts. The minor

detrital zircon age modes between 3210 and 3090 Ma

reflect various granitoid bodies that form the basement

to the Witwatesrand Basin. Most of the detrital zircon

grains in both the West Rand and Central Rand

Groups represent the age modes between 3060 and

3080 Ma, which correspond to the time of felsic

volcanism in the Dominion rift. Comparable ages are

also known from felsic volcanic rocks and granitoids

in the Murchison Belt (Poujol and Robb, 1999; Poujol

et al., 1996). All younger detrital zircon grains could

have been sourced, based on sediment transport

directions and age correlations, from higher crustal

level equivalents of the Mesoarchaean Amalia–Kraai-

pan, Murchison and Giyani granitoid–greenstone

terranes.

4. Neoarchaean weathering

One of the pillars, on which recent hydrothermal

models rest, is the apparently large-scale, acidic

hydrothermal alteration of the Witwatersrand Basin

fill (Barnicoat et al., 1997). As weathering under an

acidic Archaean atmosphere would lead to similar

bulk rock chemical changes as acid leaching by post-

depositional fluids, the question arises whether

systematic chemical changes observed in the silici-

clastic successions across unconformities reflect pale-

osol horizons or whether they are related entirely to

post-depositional fluid–rock interaction.

Most of the Witwatersrand siliciclastic rocks did

not achieve thermodynamic equilibrium during post-

depositional alteration, including burial and regional

metamorphism. This is indicated by the widespread

survival of detrital clasts that are now embedded in a

metamorphic matrix. Detrital quartz is distinguished

from secondary quartz by displaying highly variable

cathodoluminescence (Gartz and Frimmel, 1999),

variable degrees of strain and different fluid inclusion

populations (Frimmel et al., 1993). Similarly, detrital

white mica can be distinguished from metamorphic

mica both on textural and compositional grounds

(Frimmel et al., 1993; Sutton et al., 1990). The most

common metamorphic silicates in these rocks are

muscovite, pyrophyllite, chlorite, sudoite and chlor-

itoid. In the coarser grained rocks these are randomly

orientated, whereas in the argillitic units they define a

slaty cleavage. Based on the silicate equilibrium

assemblages, the peak metamorphic temperatures

achieved throughout the Witwatersrand Basin range

between 300 and 400 8C (Phillips and Law, 1994),

except for the area around the Vredefort dome, where

up to amphibolite facies conditions are recorded in the

upturned, lowermost parts of the basin fill, i.e. the

West Rand Group (Gibson and Wallmach, 1995).

Away from the Vredefort dome, the presence of

kyanite and the mineral assemblage chlorite+sudoi-

te+muscovite (and/or pyrophyllite) in the middle

Central Rand Group constrain peak metamorphic

conditions at approximately 300 8C and 3 kbar

(Frimmel, 1997). Sericite pseudomorphs after anda-

lusite at the bottom of the Transvaal Supergroup and

in the Ventersdorp Supergoup (McCarthy et al., 1986)

reflect a lower pressure, i.e. an expected lower

overburden for the higher stratigraphic levels.

Metamorphic biotite is rare but has been described

from the bottom of the Central Rand Group (Phillips

et al., 1988) and also from the West Rand Group,

together with stilpnomelane, relics of K-feldspar and

chlorite (Frimmel, 1994). Feldspar is extremely rare in

the Central Rand Group (Fig 6F), but few units,

particularly within the West Rand Group contain as

much as 30 vol.% detrital K-feldspar that can be

related to a granitic source due to its perthitic textures

(Fuller, 1958). Albite is very rare in the metasedi-

mentary rocks but has been reported as metamorphic

phase from an argillite in the uppermost Johannesburg

Subgroup (Booysens Formation; Frimmel, 1994).

Pyrophyllite is widespread and becomes particu-

larly important towards the basin margin. Cross-

cutting relationships between pyrophyllite distribution

and stratigraphic boundaries have led Barnicoat et al.

(1997) to postulate basin-wide acidic hydrothermal

alteration. These workers reported intense pyrophyl-

litic alteration within mineralised conglomerate hori-

zons, as well as for more than 500 m above and below


surrounding quartzite. A postulated correlation of gold

distribution with this alteration (Phillips and Law,

1994) that postdates the hydrothermal/metamorphic

phyllosilicates forms the fundamental argument for a

hydrothermal origin of the gold.

Considering the multi-stage tectono-metamorphic

history of the Witwatersrand Basin, it appears unlikely

that all the post-depositional alteration phenomena in

the basin are ascribable to a single hydrothermal

infiltration event. Although there is little doubt that

H+-metasomatism has caused the formation of some of

the pyrophyllite at the expense of mica, as evidenced

by cross-cutting pyrophyllite veinlets, it might be

difficult, in other places, to distinguish such hydro-

thermal acid leaching from the effects of weathering

under acidic conditions. The latter can be expected to

occur along erosional unconformities, where kaolinite

would have been the starting material for pyrophyllite

formation during prograde metamorphism.

The findings of Barnicoat et al. (1997) are at

variance with those of Sutton et al. (1990) who

found, by studying the compositional and minera-

logical changes across stratigraphic units, a strong

stratigraphic control on the chemistry and mineral-

ogy of Witwatersrand arenites and concluded that,

except for potassium, metamorphism was essentially

isochemical.

4.1. Geochemical alteration profiles across

stratigraphic units

The shape of geochemical alteration profiles across

individual unconformities can be used to test whether

observed alteration patterns are related to palaeo-

weathering or to post-depositional metasomatism. In

the former case, a systematic change in bulk rock

chemistry is expected towards the top of the footwall,

but not in the hanging wall. In the latter case, both

footwall and hanging wall should show an alteration

halo around the unconformity, which is presumed to be

the principal post-depositional fluid pathway, with

dispersion causing alteration to comparable extent both

above and below that pathway. To this effect, the

chemical index of alteration (Nesbitt and Young,

1982),

CIA ¼ ½Al2O3=ðAl2O3 þ CaO4þ Na2Oþ K2OÞ�

�100

in which the oxides are expressed as molar propor-

tions and CaO* is CaO in silicates, as opposed to

carbonates and phosphates, was applied to analyses of

siliciclastic rocks across a number of reefs throughout

the Central Rand Group and the upper West Rand

Group (Figs. 4 and 7). The concentration of CaO in

most samples studied is less than 0.1 wt.% but

typically above the lower limit of detection (0.004%).

Small amounts of Na (Na2O contents are typically

around 0.2 wt.%) are bound largely to white mica as

paragonite intergrowths (Frimmel, 1994), except for

albite-bearing shale in the West Rand Group. Con-

sequently, the CIA reflects essentially the distribution

of K-feldspar, muscovite and pyrophyllite, i.e. the

extent to which K was leached out of the rock during

lateritic weathering and/or post-depositional reaction

with an acidic fluid.

Comparison of average CIA values across the

stratigraphic units (Fig. 4) shows significantly lower

indices for the few analyses available from the West

Rand Group. Analyses for most of the West Rand

Group are lacking, but data from a likely stratigraphic

equivalent, the Mozaan Group, a ca. 5000 m thick

siliciclastic succession within the Pongola Supergroup

(Beukes and Cairncross, 1991), may serve as proxies.

There, the matrix of diamictite units and associated

mudstones have significantly lower CIA values (on

average 66) than other mudstones in the Pongola

Supergroup, and these have been interpreted as

indicative of a glacial origin for the diamictite beds

(Young et al., 1998). Most of the available data for the

West Rand Group (Sutton et al., 1990) cluster between

50 and 60, whereby 50 is typical of unweathered

material, representing the composition of fresh feld-

spar. It should be noted that these data were obtained

on arenitic samples and thus reflect sediment that has

been transported over shorter distance. They are

therefore not directly comparable with those obtained

on mudstone for which the calculated CIA values

across the Witwatersrand Supergroup are variable

(70–98; Wronkiewicz and Condie, 1987). Never-

theless, a marked difference in the chemical weath-

ering intensity between West Rand and most of the

Central Rand Group sediments seems to be indicated,

similarly as noted for the Pongola Supergroup. This is

well shown by a systematic increase in the CIA to as

much as 85 towards the sequence boundary with the

Central Rand Group.

Fig. 7. Variation in Chemical Index of Alteration (CIA) and Fe/Al ratio with distance from the Denny’s, Crystalkop (Frimmel and Minter, 2002; note that the length scale in the

original source paper is incorrect) and Ventersdorp Contact Reefs (Gartz and Frimmel, 1999) at various mines in the Klerksdorp gold field (A–G) and the Welkom gold field (H.E.

Frimmel, unpubl. data; H); data are for argillitic to arenitic siliciclastic rocks, except for the hanging wall of the Ventersdorp Contact Reef, which consists of metabasalt and thus has a

lower CIA.

H.E.Frim

mel

/Earth

-Scien

ceReview

s70(2005)1–46

24


For most of the Central Rand Goup, the CIA is

around 80, which reflects above average (Nesbitt and

Young, 1996) extent of chemical weathering. Only in

the uppermost formation do the CIA values decrease

to around 70, which is in agreement with average

chemical weathering. The internal variability in CIA

within the group, particularly across unconformities,

is, however substantial (Fig. 4). A closer inspection of

this variability, using examples from the footwall and

hanging wall of the Crystalkop and Denny’s Reefs at

Vaal Reefs mine, Klerksdorp gold field (Frimmel and

Minter, 2002) and the B-Reef at Freestate No. 3 mine,

Welkom gold field (H.E. Frimmel, unpubl. data)

reveals a gradual upward increase in CIA in the

footwall towards the reef contacts (Fig. 7). This trend

is exemplified by two profiles through the Crystalkop

Reef (Fig. 7A,B). In some areas, this trend can be

traced over more than 10 m (Fig. 7A,B), whereas in

others, a systematic increase in CIA towards the reef

was found over a distance of only a few metres below

the reef (Fig. 7C,D,E,F). In the hanging wall, this

trend is reversed with an abrupt decrease in the CIA

away from the reefs (Fig. 7A,B,C,D,E,H). Note that

the very low CIA in the hanging wall of the

Ventersdorp Contact Reef (Fig. 7G) is an artifact of

the lithology as it consists of metabasalt and not of

arenitic metasedimentary rocks as in all other cases.

Superimposed on the above relatively large-scale

trends, with the highest CIA values reaching a

maximum of 95 in the footwall near the respective

reef contact, is a decrease in CIA on a smaller scale

along the reef horizons. This is best exemplified by

two profiles through the Crystalkop Reef (Fig. 7B,D),

which show a marked drop in the CIA within a few

metres both above and below the reef. Note that this

trend affects both the footwall and the hanging wall,

whereas the former, larger-scale trend appears asym-

metrical below and above the reef horizons. Similarly,

corresponding chemical changes, both above and

below the reef, have also been observed at the

Ventersdorp Contact Reef (Fig. 7G). There an earlier

potassic alteration (Gartz and Frimmel, 1999) affected

the pyrophyllite-bearing footwall arenite as well as the

metabasaltic hanging wall, causing a decrease in CIA

over the top few metres in the footwall and an increase

in CIA over the bottom few metres in the hanging-

wall. This was followed by chloritisation only along

the reef and its immediate contact zones over a few

centi- to decimetres, which is reflected by a sharp but

very local increase in CIA.

4.2. Interpretation of trends in CIA

Trends in CIA are observed on different scales, i.e.

on the hundred metres, few metres and decimetre-

scale. An example for the former is the systematic

increase in CIA towards the West Rand/Central Rand

Group unconformity over more than 100 m. This is

too long to be explained by palaeoweathering along

that unconformity, but may point to a systematic

change in the environmental conditions (increase in

temperature and/or acidity) towards that major strati-

graphic boundary.

The upwardly increasing CIA trends on the scale of

a few metres are typical of, but more extreme than,

those of both modern and Eoproterozoic paleosols

(Nesbitt and Markovics, 1997). With one exception

(Fig. 7F), the very high CIA values are confined to the

footwall, which is not the expected result if they were

related to reef-parallel acidic hydrothermal infiltration.

The very high CIA values are, therefore, ascribed to

intense chemical weathering.

Whereas Ca and Na are removed during the initial

stages of development of a weathering profile, K is

removed only during the latest stages (Nesbitt and

Markovics, 1997; Nesbitt and Young, 1984). Con-

sequently, in Al2O3–(CaO+Na2O)–K2O (A–CN–K)

space (Fig. 8), a typical weathering trend will follow a

line parallel to the A–CN join until it reaches the A–K

join from where it will continue towards the A apex.

Plotting a large data set of arenitic bulk rock analyses,

predominantly from the Central Rand Group with a

few data from the West Rand Group, in that space

(Fig. 8) reveals that the bulk of the non-mineralised

arenitic rocks from various stratrigraphic levels

between individual reef (unconformity) horizons

follows a weathering trend of progressive removal

of CaO and Na2O prior to removal of K2O, with a

starting point close to the average composition of

Neoarchaean upper continental crust. A considerable

number of samples are displaced towards the K apex,

thus indicating variable and, in places, considerable

K-metasomatism. Comparison between footwall and

hanging wall analyses (from a distance of up to 1 m

below or above a given reef), shows that especially

hanging wall samples show evidence of this K-

Fig. 8. A–CN–K (Al2O3–[CaO+Na2O]–K2O) diagram showing

alteration trends in siliciclastic metasedimentary rocks from

auriferous reefs (conglomerate), immediate footwall and hanging-

wall of reefs, and from arenite of various stratigraphic positions in

the Witwatersrand Supergroup (n=226); data from Frimmel and

Minter (2002), Gartz and Frimmel (1999), Sutton et al. (1990), Q.

Hawes (unpubl. data), and H.E. Frimmel (unpubl. data); small

triangular diagram in top left corner shows general alteration

trends: 1—chemical weathering of average Neoarchaean upper

continental crust (��) from Condie (1993), 2—Ca/Na-metasomatism,

3—K-metasomatism.


enrichment, which consequently has to be inferred as

post-depositional. Such metasomatism would explain

the local negative excursion in CIA along the

Crystalkop Reef as shown in Fig. 7D. In contrast,

most footwall samples plot very close to the A apex

(very high CIA values)—a feature that is absent in the

hanging wall samples and therefore regarded as

reflecting palaeo-weathering.

Analyses from within reef beds display a wide

spread close to the A–CN line (Fig. 8). Such a trend

cannot be explained by any weathering process but

clearly illustrates the effects of post-depositional

alteration. As most of the plotted analyses come from

the Ventersdorp Contact Reef, the enrichment in Ca,

and to a lesser extent Na, reflected by the trend

towards the CN apex, can be explained by interaction

of a hydrothermal fluid with the overlying Ca- and

Na-rich metabasalt of the Klipriviersberg Group.

Alteration patterns around this reef (Fig. 7G) also

illustrate well the effects of K-metasomatism as

described in more detail by Gartz and Frimmel

(1999). Using Ti as the least mobile reference

element, a sharp increase in K is observed with

increasing proximity to the reef, which reflects

sericitisation. Only in the immediate reef environ-

ment, within less than one metre distance, the rocks

are depleted in K and enriched in Fe, reflecting

chloritisation. As the potassic alteration over several

metres and the smaller-scale ferric alteration over a

few decimetres affected both the footwall and the

hanging wall, they cannot be related to weathering on

a palaeo-surface but must be explained by reef-

parallel post-depositional hydrothermal fluid flow as

postulated on structural grounds by Jolley et al.

(1999). This type of alteration might also be respon-

sible for the halo of elevated CIA values observed

around the Denny’s Reef (Fig. 7F) and a sericitisation

similar to that in the VCR might explain the local

sharp drop in the CIA values around the Crystalkop

Reef as shown in Fig. 7D.

In summary, large-scale trends in CIA may be the

result of an overall change in climate and/or reflect

different degrees of sediment re-working. Systematic

increases in CIA to very high values in the footwall

beneath erosional unconformities on the scale of

several metres is ascribed to deep chemical weath-

ering along these palaeo-surfaces and thus provide

indirect information on the contemporaneous atmos-

pheric composition. Small-scale variations in CIA in

the immediate vicinity of reefs, typically over centi- to

decimetres, reflect dispersive metasomatism caused

by reef-parallel fluid flow.

5. Pre- or post-depositional age of the ore?

A sedimentary model for the Witwatersrand gold

deposits, originally proposed by Mellor (1916), was

first challenged by Graton (1930), who suggested a

magmatic–hydrothermal model. Since then, workers

who have studied the ore and host rocks on a

microscopic scale have emphasised a hydrothermal

model, because gold is typically late in the para-

genetic sequence as micro-fracture fills and inclusions

in secondary, clearly hydrothermal mineral grains

(Feather and Koen, 1975; Ramdohr, 1958). In con-

trast, those who have studied the rocks on a macro-

scale noted a strong sedimentological control on gold

grade, which has been used highly successfully

throughout the history of Witwatersrand exploration

and day-to-day mining, and prompted them to

advocate a sedimentary palaeoplacer model (Minter,


1978). Following the generally accepted recognition

of regional metamorphism and post-depositional fluid

flow throughout the basin (Phillips and Law, 1994), a

palaeoplacer model sensu stricto, in which all gold

particles are perceived as detrital, has been aban-

doned. Today there is general agreement that the

majority of gold particles that appear late in the

paragenetic sequence are indeed hydrothermal precip-

itates. Thus, the debate has shifted to the question of

whether the source of that hydrothermal gold was

proximal, fluvially deposited, detrital gold within the

conglomerate beds (modified palaeoplacer model), or

external to the host rocks (hydrothermal model).

These two possibilities represent end-member models,

which are, each with variations, the focus of current

debate:

(1) In the modified palaeoplacer model, transport

of detrital gold particles into the host sediments is

assumed to have taken place by fluvial processes

with subsequent short-range mobilisation of the

gold by infiltrating hydrothermal fluids and/or

degradation of in situ hydrocarbon or hydrous

phases. Gold mobilisation and recrystallisation

induced by hydrothermal fluid infiltration has been

ascribed to the emplacement of the Ventersdorp

Supergroup lavas (Pretorius, 1991), to burial meta-

morphism in lower Transvaal Supergroup times and

the Vredefort impact event (Frimmel et al., 1999),

as well as to Pretoria Group deposition and the

emplacement of the Bushveld Igneous Complex

(Robb et al., 1997).

(2) Hydrothermal models explain the presence of

gold as the result of post-depositional hydrothermal

fluids from an external source. The presence of gold

in the conglomeratic host rocks is inferred as

consequence of long-range, basin-wide fluid flow

combined with chemical and structural controls

(Phillips and Law, 2000). One hydrothermal model

infers a separate origin for gold, uraninite and

hydrocarbons (Phillips and Law, 2000), whereas

another seeks to explain all of them as cogenetic

(Barnicoat et al., 1997). The infiltration of the

inferred auriferous hydrothermal fluids has been

linked with several different events that range from

Ventersdorp volcanism (Phillips et al., 1997),

regional metamorphism (Phillips and Myers, 1989),

or the emplacement of the Bushveld Igneous Com-

plex (Stevens et al., 1997).

All of the currently debated genetic models

include the presence of gold derived from the

movement of hydrothermal fluids, but they differ

principally in the inferred distances of hydrothermal

gold transport and the composition of that gold-

transporting fluid. The composition of the gold-

bearing fluid has been suggested to be similar to that

inferred for Archaean orogenic gold deposits (Phil-

lips and Law, 2000), in which gold is assumed to

have been transported as bisulphide complex in an

H2O–CO2 dominated, relatively reducing, low-sul-

phur, low salinity fluid. In that model, all rounded

pyrite is assumed to be the product of post-deposi-

tional hydrothermal sulphidation of black sands that

existed originally of various Fe–Ti oxides and Fe-

oxyhydroxide pisolites (Phillips and Myers, 1989).

This contrasts with models that suggest gold transport

in highly acidic, oxidising fluids (Barnicoat et al.,

1997), and models which prefer gold transport as a

hydroxy-complex (Gray et al., 1998). Reduction of

aqueous gold species to elemental gold by interaction

with pre-existing bitumen plays an important role in

all hydrothermal models and is used to explain the

strong association of gold and bitumen and the

textural position of gold grains within microfractures

in bitumen.

A major argument in favour of an external,

hydrothermal source of gold is based on mineral and

chemical zonation patterns at the deposit to hand-

specimen scale as well as elemental correlations (Fox,

2002). The recorded zonation patterns clearly dem-

onstrate hydrothermal fluid–rock interaction, but do

not contribute to solving the question of the gold

source. Hydrothermal versus detrital element correla-

tions are more crucial in this regard. Excellent

correlations between Au, U and Ag, a poor but

positive correlation of Au with Cr and (Co+Ni), and

only a very weak correlation of Au with Zr have been

noted in several studies on the Kimberley Reef

(Rasmussen and Fesq, 1973), the Vaal Reef (Fox,

2002), the Steyn Reef (Frimmel and Minter, 2002),

and the B- as well as the Ventersdorp Contact Reef

(H.E. Frimmel, unpubl. data).

The good correlation between Au and Ag is

expected, because gold is the principal sink for Ag,

with Au and Ag being alloyed in variable proportions

with each other and Hg (Frimmel and Gartz, 1997;

Reid et al., 1988; Utter, 1979). Similarly, the positive


correlation with Co and Ni is readily explained by

precipitation of cobaltite–gersdorffite together with

hydrothermal gold particles (Fox, 2002; Frimmel et

al., 1993) and Co as well as Ni enrichment in co-

existing pyrite. The good correlation between Au and

U as well as Cr can be ascribed to a spatial association

with heavy sands that contained detrital uraninite and

chromite particles. As the formation of bitumen is

genetically and spatially related to detrital uraninite,

hydrothermal gold precipitation by a locally available

reductant would have taken place preferentially near,

or within, fractures of uraninite, thus further exacer-

bating the positive Au–U correlation.

The poor correlation between Au and Zr has been

used as evidence against a detrital origin of the gold,

because the distribution of Zr is controlled by zircon,

almost all of which is detrital (Fox, 2002). Apart from

gold and zircon being derived from different source

rock types, the poor correlation could be a function of

differences in the hydraulic behaviour during repeated

sediment re-working (Smith and Minter, 1980). It is

noteworthy that many Zr-rich samples are devoid of

Au, simply indicating a gold-poor source area,

whereas only very few samples show elevated Au

but low Zr contents and these typically contain

auriferous veinlets indicative of local gold mobilisa-

tion. If the gold had been introduced into the host

sedimentary rocks by hydrothermal fluids, both

zircon-rich and zircon-poor domains should have

been affected to a similar degree. This is not the case.

A less than perfect correlation between Au and other

elements concentrated in detrital minerals, such as Zr,

is likely the result of dispersion of the gold by short-

range hydrothermal mobilisation.

Other arguments for a hydrothermal origin include

the observation that the bulk of gold grains are located

within or near micro-fractures, filled with bitumen and

uraninite/brannerite, which post-date early, bedding-

parallel pyrite–pyrrhotite–quartz filled fractures that

contain no gold (Fox, 2002). While this observation

demonstrates the undisputed hydrothermal nature of

the bulk of the gold particles, it does not clarify the

distance of hydrothermal gold transport. The same

applies to effectively all other arguments that have

been brought forward in favour of a hydrothermal

model (Table 1). In contrast, a number of observations

and analytical data can hardly be explained other than

in terms of a modified palaeoplacer model.

Some of the most important lines of evidence for a

modified placer model are the observation of gold

micro-nuggets that are spatially associated with

secondary, locally remobilised, hydrothermal gold.

Although these spheroidal to torroidal micro-nuggets

are very rare and have so far only been found in

samples from the Basal Reef, Vaal Reef, B-Reef and

Crystalkop Reef (Frimmel and Minter, 2002; Minter et

al., 1993), their existence gives a clear clue as to the

primary process of gold enrichment in the Witwaters-

rand sediments. Strong support for such a sedimentary

gold enrichment process also comes from the Re–Os

isotope data, which indicate an age for the gold that is

clearly older than that of host sediment deposition, but

the same as for rounded pyrite (3033F21 Ma; Kirk et

al., 2001, 2002). Notwithstanding local evidence of

sulphidation of Fe–Ti oxides (Ramdohr, 1958), chert

and iron formation pebbles (Hallbauer, 1986; Hirdes

and Saager, 1983), the Re–Os data imply a detrital

origin of the most abundant, rounded form of pyrite.

All of the above examples of sulphidation can be

related to the same fluids that caused the formation of

secondary pyrite at various stages throughout the

complex post-depositional alteration history of the

Witwatersrand sedimentary rocks. The noted textural

association of gold as inclusions within secondary

pyrite and the lack of gold inclusions in rounded pyrite

further indicate that hydrothermal gold transport

cannot be linked with a postulated sulphidation event

that supposedly formed all the pyrite. A similar

argument can be applied to uraninite. The U–Pb age

of 3050F50 Ma obtained for uraninite from the

Dominion Reef (Rundle and Snelling, 1977) is older

than the age of Witwatersrand sediment deposition

and, though subject to a considerable uncertainty,

effectively rules out a post-depositional introduction of

significant amounts of U into the basin fill.

In particular, the interpretation of the Re–Os data

appears to be very robust as the only alternative to the

interpretation given above would be that of isotope

mixing between two or more different ad-hoc end

members. Any mixing model would be difficult to

reconcile with the excellent agreement between the

Re–Os isochron ages of gold and rounded pyrite from

various localities, the high precision of the isochron

ages, and the initial Os isotopic compositions that are

identical within error to the Os isotopic composition

of the mantle at 3 Ga.

Table 1

Main arguments for a hydrothermal and modified palaeoplacer model, respectively, for the Witwatersrand gold

Hydrothermal model Modified palaeoplacer model

Gold is late in paragenetic sequence (Barnicoat et al., 1997; Feather

and Koen, 1975)

Rare co-existence of rounded gold micro-nuggets with secondary,

hydrothermal gold on mm-scale (Minter et al., 1993)

Gold is associated with acid metamorphic alteration (Phillips and

Myers, 1989; Barnicoat et al., 1997)

Composition of fluid inclusions in auriferous hydrothermal quartz

indicate neutral to basic pH (Frimmel et al., 1999)

Basin-wide distribution of pyrophyllite related to large-scale

H+-metasomatism (Barnicoat et al., 1997)

Increase in chemical index of alteration in footwall towards reef

related to chemical weathering under acid atmosphere (Sutton et al.,

1990; Frimmel and Minter, 2002)

Abundant rounded pyrite and uraninite particles associated with the

gold ore are of post-depositional hydrothermal origin (Barnicoat

et al., 1997)

Isotopic data for rounded sulphides and uraninite yield ages older

than time of sedimentation (Rundle and Snelling, 1977; Kirk et al.,

2001)

Rounded pyrite derived from basin-wide sulphidation of dblacksandsT (Phillips and Myers, 1989)

Pyrite morphology, cyrstallography, and truncated growth zonation

patterns indicate detrital nature of rounded grains (McLean and Fleet,

1989; Fleet, 1998; England et al., 2002b)

Strong correlation between gold and hydrothermal pyrobitumen

(Nagy, 1993; Gray et al., 1998)

Derivation of hydrothermal pyrobitumen from local intrinsic oils,

based on bulk and molecular y13C data (Spangenberg and Frimmel,

2001) and fluid inclusion studies (England et al., 2001a)

Local variability in gold composition within a given reef reflects

differences in source areas (Frimmel and Gartz, 1997)

Conglomerate beds were preferred channels for infiltration of

gold-bearing hydrothermal fluids (Barnicoat et al., 1997)

Sedimentological control on gold distribution (Minter, 1978);

Negative correlation between authigenic/hydrothermal xenotime and

ore bodies (Kositcin et al., 2003)

Gold was introduced into the Witwatersrand Basin after sediment

deposition (Phillips and Myers, 1989; Barnicoat et al., 1997)

Re–Os ages of the gold are older than time of sedimentation (Kirk

et al., 2002)

Hydrothermal introduction of gold into the basin during peak

metamorphism, coeval with 2.06 Ga Bushveld event (Phillips and

Law, 1994)

Lack of significant secondary permeability after N600 my of burial,

diagenesis and low-grade metamorphism (Frimmel, 1997)

Hydrothermal introduction of gold into the Witwatersrand basin

during global 2.7–2.6 Ga gold-forming thermal event, coeval with

Ventersdorp Supergroup volcanism (Phillips et al., 1997)

Sedimentary reworking of Witwatersrand gold ore in

late-Ventersdorp diamictite was followed by Witwatersrand-style

gold mineralization in the post-Ventersdorp Black Reef

Lack of suitable source area for placer gold (Phillips and Myers,

1989)

Calculated background value for eroded source area corresponds to

mean Au content of Archaean granitoid–greenstone crust (Loen,

1992)


If the gold were brought into the Witwatersrand

Basin during post-depositional fluid infiltration, large

fluid/rock ratios would be expected. The auriferous

fluids would have had to flow preferentially along the

conglomerate beds in order to explain the apparent

sedimentological control on the basin-wide ore dis-

tribution. Although some evidence exists for bedding-

parallel fluid flow (Jolley et al., 1999), mass balance

calculations (Gartz and Frimmel, 1999) point to rather

limited external fluid infiltration into the reefs. Only if

all the pyrophyllite is explained by post-depositional

H+-metasomatism (Barnicoat et al., 1997), can a case

be made for large-scale fluid infiltration. Bearing in

mind that the loss of alkalies on the scale of tens of

metres can be attributed to palaeoweathering, as

outline above, it is more likely that a large proportion

of pyrophyllite in the Witwatersrand metasedimentary

rocks is derived from the prograde metamorphism of a

kaolinite-bearing protolith, thus revoking the neces-

sity for significant post-depositional fluid infiltration.

However, even if all the pyrophyllite were hydro-

thermal–metasomatic, it has been shown from fluid

inclusion analyses that in those cases studied, hydro-

thermally mobilised gold was transported by a fluid of

a composition that is incompatible with the stability of

pyrophyllite (Frimmel et al., 1999). Only limited

interaction between potentially auriferous fluids and

the host conglomerate beds is also supported by

studies on the chemistry and age distribution of

xenotime, which is a particularly useful monitor for

hydrothermal infiltration (England et al., 2001b;

Kositcin et al., 2003). These studies revealed that


diagenetic xenotime is relatively abundant in the reefs

compared to their surroundings, but hydrothermal

xenotime is present in much lesser amounts within the

reefs. This implies that the reefs were more permeable

during diagenesis but less permeable during post-

diagenetic hydrothermal events, which is in disagree-

ment with proposed hydrothermal models.

Apart from the textural, geochemical and isotopic

evidence, there is also important geological evidence

that speaks against a hydrothermal model not only

for the Witwatersrand gold, but also for the

associated pyrite and uraninite. The Witwatersrand

ore bodies must have formed prior to deposition of

the Platberg Group sediments as the latter contain

pebbles of mineralised Witwatersrand reef material

(Phillips et al., 1997). Yet, Witwatersrand-style

mineralisation is evident in the Black Reef at the

base of the Transvaal Supergroup that is some 70

million years younger than the Platberg Group.

Formation of a placer deposit under similar environ-

mental conditions repeatedly through time is to be

expected, but the same style of hydrothermal metal

introduction into the basin at different stages of basin

development in different tectonic settings is unlikely.

If there had been a second major hydrothermal ore-

forming event in the central Kaapvaal Craton,

including the conglomerates at the base of the

Transvaal Supergroup, it would have affected the

Witwatersrand rocks after they had undergone dia-

genesis and first low-grade metamorphism. Very

little permeability would have remained at that stage

in the Witwatersrand strata, and basin-wide, large-

scale, predominantly bedding-parallel fluid flow, as

postulated by the various hydrothermal models,

would have been effectively impossible.

Problems with the hydrothermal models extend

from the micro- to the macroscale. On a tectonic

scale, the Witwatersrand deposits have often been

compared with orogenic gold deposits by those who

favour a hydrothermal model. This is mainly for the

similarity in the paragenetic sequence and in the

common gold–pyrite–hydrocarbon association (Phil-

lips and Myers, 1989). There are, however, a number

of significant differences between the two styles of

deposit (Frimmel et al., in press; Groves et al., 2003).

The rounded, sub-spherical morphology of most of

the Witwatersrand pyrite and its highly variably

geochemical and S isotopic composition (England

et al., 2002b) are in contrast to the typically subhedral

to euhedral, compositionally restricted pyrite found in

orogenic deposits. Wallrock alteration, inferred to

have taken place over several hundreds of metres

across stratigraphic boundaries throughout the Wit-

watersrand Basin (Barnicoat et al., 1997), is orders of

magnitude more extensive than known from any

orogenic gold deposit. The latter are characterised

by the abundance of auriferous quartz veins, but the

basin that hosts the by far greatest known concen-

tration of Au is characteristically devoid of a

plentitude of such veins. A foreland/retroarc basin

setting is indicated for the bulk of Witwatersrand

deposits. This is analogous with modern placer gold

deposits but in stark contrast to orogenic gold

deposits, which typically occur in near-arc or arc

settings. Finally, the geometry of the Witwatersrand

orebodies (gently dipping decimetre to metre thick,

laterally extensive sheets) is unlike the overall shape

of most known epigenetic or orogenic deposit.

5.1. Best-fit genetic model

A modified palaeoplacer models accounts best for

all the available data and observations. Clastic sedi-

ments, first laid down in the 3074 Ma Dominion rift

graben, were largely derived from felsic sources,

which led to enrichment in detrital uraninite but only

low gold contents. During subsequent West Rand

Group sedimentation (2985–2914 Ma), progressively

more mafic rocks from Mesoarchaen greenstone belts

in the hinterland were eroded, but only during Central

Rand Group times (2902–2780 Ma) were the high

levels of gold concentrations in the placer sediments

reached, for which the Witwatersrand has become so

famous. These extraordinary gold concentrations are

explained by a combination of factors that range from

fertile source regions, to tectonic setting and to

palaeoenvironmental conditions. Initially, the Central

Rand Basin took a foreland position relative to the

overriding Kimberley block in the west, with the

intervening Amalia–Kraaipan greenstone belt repre-

senting an obducted slice of former oceanic crust. A

subsequent change in the continental stress field led to

the accretion of the Murchison greenstone belt to the

north, with the Central Rand Basin taking a retroarc

position (Frimmel et al., in press). Gold and chromite,

predominantly from the surrounding greenstone belts,


together with uraninite and zircon from the associated

granitoids, were brought into this basin by fluvial

transport and concentrated by mineral sorting. Aeolian

transport and re-sedimentation on a number of ero-

sional unconformities led to further up-grading of

gold along the basal degradation surfaces. Sedimen-

tary re-working in braided stream systems and

effective wind sorting were particularly vigorous

because of intense acid weathering and a lack of

vegetation and widespread organisms.

The complex post-Witwatersrand tectono-thermal

evolution of the central Kaapvaal Craton affected the

gold to a variable extent. The bulk of fluid flow

through the Witwatersrand Basin must have taken

place due to diagenetic dewatering. Different parts of

the basin experienced further fluid flow to various

degrees as consequence of the following events: (1)

low-grade burial metamorphism; (2) the syn-Venters-

dorp thermal anomaly, including syn-Platberg rifting;

(3) dynamic metamorphism in a compressional stress

field during thrusting of the Southern Marginal Zone

of the Limpopo Belt on to the Kaapvaal Craton; (4)

lower Transvaal thermal subsidence-induced exten-

sion; (5) the thermal anomaly related to the Bushveld

magmatic event; and (6) pervasive fracturing due to

the Vredefort impact. Some of these events caused the

mobilisation of the gold, together with other detrital

phases, such as pyrite and uraninite. Hydrocarbons,

derived from oil/bitumen, played an important role in

the re-precipitation of the mobilised gold by acting as

reductants. The distances over which gold, pyrite and

uraninite were mobilised, in general were in the order

of millimetres to centimetres. Locally, fracture-con-

trolled fluid flow allowed transport of Au over longer

distances. Most of the hydrocarbons, and thus by

implication gold, was first mobilised during dia-

genesis, but meteoric waters that percolated through

Vredefort impact-related secondary interconnected

(micro-) fracture space were also capable of trans-

porting hydrocarbons, sulphides and gold.

6. Neoarchaean palaeoenvironment

6.1. Neoarchaean atmosphere

It has long been recognised that attempts to

constrain the Archaean atmospheric composition

hinge essentially on four lines of evidence: (1)

presence of detrital uraninite; (2) presence of detrital

pyrite; (3) composition of detrital gold particles; and

(4) the presence and composition of paleosols (Hol-

land, 1984). Since then the debate has shifted from a

previously favoured palaeoplacer model to various

hydrothermal models for the Witwatersrand gold, thus

invalidating the reliability of the above pieces of

evidence (Holland, 1994). The recognition that the

best-fit model for the Witwatersrand deposits is that of

a modified palaeoplacer, reaffirms however the use-

fulness of the above pieces of evidence for the debate

on the Archaean atmospheric composition.

The significance of the Witwatersrand in this

regard is obvious, considering that all four lines of

evidence can be tested there. However, other Meso-

archaean to Eoproterozoic deposits that bear strong

similarities with those of the Witwatersrand outside

the Kaapvaal Craton should not be ignored and can

contribute useful additional information. These

include the 2.13 to 2.10 Ga Tarkwaian System

(Ghana), the 2.09 to 1.88 Ga Jacobina and the poorly

dated 2.8 to 2.2 Ga Moeda deposits (Brazil), as well as

the 1.90 Ga Roraima Supergroup (northern South

America). Similar styles of mineralisation with

abundant detrital pyrite and uraninite, but a conspic-

uous lack of gold, are known from the 2.9 to 2.6 Ga

Bababudan Group (India) and the 2.45 Ga lower Elliot

Lake Group (Huronian Supergoup) in Canada. All of

these deposits have in common that a case for a

palaeoplacer origin can be made (Frimmel et al., in

press), and that the fluvial siliciclastic host sediments

were laid down in foreland/retroarc basins. An

Archaean to Palaeoproterozoic greenstone terrain is

suggested as the most likely source area for all of the

above gold palaeoplacer deposits. In addition, detrital

pyrite, uraninite, and signficantly also siderite, have

been reported from effectively unmetamorphosed

fluvial siliciclastic sedimentary rocks from the Pilbara

Craton (Rasmussen and Buick, 1999).

The first piece of evidence to be assessed is the

presence of detrital uraninite. Based on thermody-

namic grounds U4+-minerals, such as uraninite, are

not stable under modern atmospheric conditions, as

they would oxidise rapidly (Fig. 9). The dependence

of the uraninite stability on oxygen fugacity is almost

independent of pH and the fugacities of other critical

species, such as CO2 and CH4. Consequently, these

Fig. 9. Oxygen fugacity versus temperature diagram showing the

conditions at which oxidation of pyrite to goethite (solid lines) takes

place at variable pH, as well as that of uraninite to a dissolved

oxyhyroxide (dashed line); calculated using PHREEQC (Parkhurst

and Appelo, 1999).


reduced U-minerals are unlikely to survive mechan-

ical transport in the fluvial environment under an

oxidising atmosphere (e.g. Holland, 1994). Localised

modern occurrences of detrital uraninite, such as in

sand from the Indus river, are confined to domains

that have experienced virtually no chemical weath-

ering and can, therefore, not be used as analogues for

the Witwatersrand occurrences (Maynard et al., 1991),

for which intense chemical weathering is indicated.

Detrital uraninite is abundant from the oldest

siliciclastic rocks, the Dominion Reef, throughout

the entire Witwatersrand Supergroup, to the base of

the Ventersdorp Supergroup. Its occurrence thus spans

a period from 3074 to 2714 Ma. The uraninite/

brannerite ratio varies greatly between reefs. Branner-

ite is not important in the Dominion Reef. A system-

atic decrease in the uraninite/brannerite ratio towards

younger stratigraphic levels has been noted in the

Welkom gold field, where this ratio changes from 8.7

in the Steyn Reef to zero in the Beatrix Reef (Minter

et al., 1988). In the Klerksdorp gold field, the younger

Ventersdorp Contact Reef does contain uraninite, but

its textural relationships (typically as inclusions within

bitumen) do not permit to distinguish between a

detrital and a hydrothermal derivation. No uraninite is

known from the 2642 Ma Black Reef, in which all U

occurs as brannerite. This trend towards lower

uraninite/brannerite ratios up-section might be inter-

preted as reflecting repeated sediment re-working

under an atmosphere that became slightly more

oxidising in the course of the Neoarchaean Aera. As

the brannerite is secondary, such an interpretation is,

however, not imperative and this trend might equally

reflect a more oxidising hydrothermal fluid composi-

tion at shallower crustal levels. The latter explanation

is preferred, because of the abundance of detrital

uraninite in the Elliot Lake Group, which is signifi-

cantly younger than the Black Reef.

The second piece of evidence concerns rounded

pyrite, whose detrital nature has been established. The

stability of pyrite requires even lower oxygen fugacity

than that of uraninite (Fig. 9) thus supporting a

reducing environment. Kinetic limitations to the

solubility of pyrite are unlikely to explain the

preservation of pyrite in the fluvial to fluvio-deltaic

sedimentary rocks of the Witwatersrand. In particular,

along those unconformities that reflect repeated re-

working of the underlying sediment, pyrite must have

been exposed to the meteoric environment over

sufficient lengths of time to equilibrate with its

immediate environment.

The rounded pyrite type includes a variety of

textural forms, ranging from compact rounded to

ooid-like particles, all of which have formed prior to

sediment deposition. However, as the mentioned

petrographic and S isotope study (England et al.,

2002b) revealed, some varieties represent pseudo-

morphs after other minerals, including sulphates, with

replacement having taken place before erosion of the

source rocks. It must be emphasised that rounded,

detrital pyrite is not restricted to some small, localised

occurrences, but is a characteristic feature of all fluvial

deposits that range in time from the Dominion Group

to the bottom of the Transvaal Supergroup (3074–

2642 Ma), and in space over several hundred square

kilometres. The common occurrence of detrital pyrite

in fluvial sediments is not a peculiarity of the

Kaapvaal Craton, but also typical of all other known

Neoarchaean to early Palaeoproterozoic fluvial depos-

its, such as those of the Pilbara Craton in Australia,

the Elliot Lake Group in Canada, the Bababudan

Group in India and the Moeda deposits of Brazil, all

of which are older than 2.2 Ga.

The third piece of evidence, the composition of

detrital gold particles is more problematic and less

conclusive. Modern placer gold tends to be depleted

in Ag, owing to an increased mobility of Ag relative

to Au in oxidising waters (Morrison et al., 1991).

Modern placer gold typically shows compositional


zonation with an increase in Au/Ag ratio towards the

rim of an individual particle (Groen et al., 1990). Such

a zonation pattern is absent in all of the thousands of

studied gold particles from the Witwatersrand. This is

readily explained by diffusional homogenisation with

respect to Ag (and Hg) in gold particles at the

temperatures to which the Witwatersrand gold was

exposed (Frimmel et al., 1993). In spite of this intra-

grain homogenisation, inter-grain homogenisation

was not achieved as evidenced by considerable

differences in gold composition between reefs (Utter,

1979), between different domains within a given reef

(Frimmel and Gartz, 1997) and even between indi-

vidual grains on a hand-specimen scale (Reid et al.,

1988). These differences are most likely a function of

provenance and reflect gold sources of variable

composition.

Mobilisation of detrital gold particles during post-

depositional hydrothermal alteration may have led to

some modification of the composition, because of

different mobility of Au- and Ag-bearing dissolved

species. Consequently, individual gold particles that

formed as hydrothermal precipitates should not be

used as reference for deciphering contemporaneous

atmospheric oxidation potential. However, the com-

positions of preserved detrital gold particles may be

more illuminating in this regard. The only example of

well studied detrital gold particles from the Witwa-

tersrand concerns the Basal Reef in the Welkom gold

field, for which average Ag and Hg concentrations of

8.9 and 1.2 wt.%, respectively, have been obtained

(Frimmel et al., 1993). Such a composition is in very

good agreement with Archaean greenstone-hosted

gold (for compilation see Morrison et al., 1991).

The elevated Ag contents could thus be used as

argument against an oxidising environment during

fluvial transport, but it may be argued that Ag-

depleted rims that had developed were subsequently

mechanically eroded during fluvial transport. Further-

more, Au-rich rim formation in placer gold is

probably related to a combination of self-electro-

refining and cementation instead of selective leaching

of Ag and intra-grain diffusion (Groen et al., 1990), in

which case little inferences can be made regarding

atmospheric conditions.

Last but not least, paleosols can provide some of

the most reliable information on the composition of

the atmosphere at the time of surface exposure. In an

extensive review of paleosols described from the time

period of interest here (Rye and Holland, 1998), it was

concluded that all paleosols older than 2.2 Ga are

characterised by significant Fe-loss. This includes

examples from the Witwatersrand, but they are all

problematic because of their complex post-depositio-

nal alteration history. Metamorphic chloritisation or

formation of pyrophyllite at the expense of mica

would cause enrichment or depletion in Fe, respec-

tively. In many cases, it is not clear, whether an

observed loss in Fe is related to palaeo-weathering or

hydrothermal alteration. At least some of the inferred

paleosols from the Witwatersrand, previously used to

make inferences regarding Archaean atmospheric

conditions, appear to reflect hydrothermally altered

zones (Palmer, 1986), as documented, for example,

for the so-called Deelkraal paleosol below the

Ventersdorp Contact Reef (Jolley et al., 1999).

From the geochemical changes across stratigraphic

boundaries, outlined above, chemical weathering over

several metres below unconformity surfaces can be

deduced and distinguished from hydrothermal alter-

ation that took place, in many cases, over only a very

limited distance away from a given reef. Assuming

that Al behaved conservatively, the extent of Fe-

enrichment or depletion may be illustrated by the total

Fe/Al ratio. A sympathetic relationship between Fe/Al

ratio and CIA can reflect either hydrothermal chlori-

tisation, as exemplified by the Ventersdorp Contact

Reef (Fig. 7G) or a stratigraphic difference in

sediment grain size and thus clay content of the

protolith, as is the case in the footwall of the B-Reef

(Fig. 7H). The former is typically found over very

short distances (Fig. 7F,G), whereas the latter is found

over longer distances across stratigraphic boundaries

(Fig. 7H). In contrast, an antipathetic relationship may

be explained by weathering under an acidic, reducing

atmosphere. Indeed, some Witwatersrand reefs show a

distinct depletion in Fe in their immediate footwall

that is not linked to a corresponding decrease in CIA

(Fig. 7B,C,D,F). Even in profiles in which no distinct

trend is recognisable, the overall Fe/Al ratios tend to

be very low, i.e. less than 0.2. This compares well

with numerous examples of Meso- to Neoarchaean

paleosols from other areas, for which Fe depletion has

been described (Rye and Holland, 1998).

The exact acidity of the contemporaneous atmos-

phere is difficult to constrain, but an upper limit on pH


can be placed from the lack of detrital feldspar in

otherwise relatively immature siliciclastic sediment.

The principal chemical weathering product was

probably kaolinite, which subsequently gave rise to

the abundant metamorphic pyrophyllite in the succes-

sion. In a number of samples, detrital white mica can

still be recognised (Frimmel et al., 1993; Sutton et al.,

1990). A pH close to the boundary between the

kaolinite and muscovite (illite) stability fields, calcu-

lated as 6.2 at a temperature of 25 8C (for aK+=0.1), is

therefore likely. Such acid weathering agrees well with

the very high CIA values obtained for most footwall

sections beneath unconformities, where they exceed

80 (for comparison, the CIA of modern tropical stream

sediments is around 75, Maynard et al., 1991).

Further evidence for a reducing sedimentary

environment during upper Witwatersrand times comes

from the bulk and molecular isotopic composition of

the indigenous kerogen component in bitumen (Span-

genberg and Frimmel, 2001). The C isotopic compo-

sition and distribution of n-alkane in stratiform

bcarbon seamsQ within reefs point to considerable

input from autochthonous algal-bacterial lipids. It may

be argued that the spatial association between these

hydrocarbons and all the other evidence provided

above for a reducing Mesoarchaean atmosphere may

render this evidence inconclusive. The argument

could be that localised areas that were covered by

terrestrial biomass, provided islands in which reduc-

ing conditions prevailed under an overall oxidising

atmosphere. However, the distribution of detrital

pyrite and uraninite, and even gold, is not at all

restricted to domains rich in hydrocarbons. The bulk

of siliciclastic fluvial to fluvio-deltaic sedimentary

rocks of the Witwatersrand do not contain bcarbonseamsQ (they are restricted to localised lithofacies on

unconformities) but contain abundant rounded pyrite

as well as elevated uraninite concentrations.

A reducing Meso- to Neoarchaean atmosphere,

inferred here from multiple aspects of the nature of the

Witwatesrand placer deposits, is also in agreement

with totally different types of data from mass-

independent fractionation of S isotopes. This phenom-

enon, which refers to the deviation from the mass-

dependent relationship between S isotopes typical of

most processes in aqueous solution or solid phase

(d33Si0.515d34S, d36Si1.91d34S), has so far been

reported from Archean to Eoproterozoic sedimentary

sulphate and sulphide minerals, volcanic beds in ice

cores and modern sulphate aerosols (Farquhar and

Wing, 2003). A large scatter in d33S, which does not

obey mass-dependent relationships between S iso-

topes, is evident in samples older than 2.45 Ga, with a

transition for the loss of this peculiar isotopic

behaviour spnning from 2.45 to 2.09 Ga (Farquhar

et al., 2000). The phenomenon, which has since been

verified with data from a number of older cratons

(Farquhar and Wing, 2003; Mojzsis et al., 2003), is

explained by photochemical reactions, such as SO2

photolysis. As SO2 and SO photolysis are caused by

intense ultraviolet radiation, the existence of such

photolytic reactions in the Archaean atmosphere

implies a lack of ozone and oxygen, which are the

principal atmospheric components that absorb ultra-

violet radiation. According to the photochemical

model of Pavlov and Kasting (2002), the preservation

of the observed mass-independent S isotope fractio-

nation is only possible in an atmosphere with O2

concentrations less than 10�5 times the present

atmospheric level. Furthermore, the preservation of

the mass-independent S isotopic signatures is only

possible in the absence of large-scale, homogenising,

mass-dependent bacterial S processing in a marine,

sulphate-rich reservoir (Farquhar et al., 2000). Thus,

the mass-independent S isotope fractionation that

characterises Archaean to Eoproterozoic sediments

provides a very strong, independent argument for an

anoxic atmosphere at those times.

Reducing atmospheric conditions must have pre-

vailed until at least 2.64 Ga as indicated by the pyrite-

rich Black Reef at the bottom of the Transvaal

Supergroup and probably lasted until at least 2.45

Ga, taking into consideration the uraninite-bearing

conglomerates of the Elliot Lake Group. A minimum

age for a reducing atmosphere is given by the 2.2 Ga

Hekpoort paleosol, for which lateritic weathering with

Fe-enrichment in the top zone has been documented

(Beukes et al., 2002)—in agreement with paleosols of

similar age in other areas, for which weathering under

a highly oxidising atmosphere is indicated (Holland

and Rye, 1997).

6.2. Neoarchaean hydrosphere

Having established that the Archaean atmosphere

was most likely reducing, one might intuitively jump


to the conclusion that the contemporaneous hydro-

sphere was equally reduced. Such a conclusion seems

not to be justified, however, because syngenetic barite

deposits as old as Palaeoarchaean, with good exam-

ples known from the 3.5–3.2 Ga Barberton Super-

group in the Barberton greenstone belt (for a

compilation of all known occurrences see Huston

and Logan, 2004), clearly evidence that sulphur was

present in its oxidised state in the form of sulphate as

dominant hydrous S-species at least in some parts of

the Archaean ocean. The main reductants supplied to

the Archaean hydrosphere by hydrothermal discharge

as well as metamorphic and weathering fluxes involve

Fe2+, H2, CO, H2S, and SO2. In this context it is

particularly interesting to scrutinise the distribution of

Fe-oxides, principally magnetite, Fe-sulphides, essen-

tially pyrite, and sulphates in the various sedimentary

environments. The Witwatersrand rock record pro-

vides some pivotal information to this effect that can

be summarised as follows: (1) Pyrite is stable in all

fluvial to fluvio-deltaic deposits, even in those that

experienced extensive sedimentary re-working and

thus prolonged exposure to the meteoric environment;

(2) various textural forms of evidently detrital pyrite

show a complex S isotopic composition; and (3)

magnetite is present instead of pyrite in most marine

shale deposits (Fig. 4).

The significance of the first criterion has already

been discussed in the previous section and it suffices

to state here that the above constraints on the

Neoarchaean atmosphere are equally applicable to

the meteoric hydrosphere and even the oceanic top

waters. Sulphur isotopic composition has been used

repeatedly to constrain Archaean O2 distribution (for

review of available data see Strauss, 2003). Palae-

oarchaean barite, at least some of which is supposedly

sedimentary, has a fairly uniform S isotopic compo-

sition (d34S=+2.7x to +8.7x) that matches the

composition of Palaeoarchaean pyrite from the Bar-

berton Supergroup (d34S=�3.1x to 8.8x), which, in

turn, corresponds to that of magmatic sulphur. Larger

variations in, and deviation from 0x, of d34S is

typical of younger sedimentary sulphides and is

usually ascribed to enhanced S isotope fractionation

between seawater sulphate and reduced sulphide,

accomplished by biological S recycling. However,

some authors pointed out relatively large variations in

d34S already in Archaean sediments (e.g. Ohmoto et

al., 1993; Shen et al., 2001), which prompted them to

postulate microbial sulphate-reduction to have taken

place as early as in Palaeoarchaean times. The S

isotopic composition of the rounded, evidently detrital

pyrite from the Witwatersrand corresponds to the

limited range typical of Palaeoarchaean pyrite. This

does not seem to support a model of extensive

microbial sulphate reduction, but in the absence of a

good control on the difference between the isotopic

composition of the S source (marine sulphate) and the

product (pyrite), no definitive conclusions on the role

of microbial sulphate reduction in the Archaean

hydrosphere can be drawn.

The third of the above criteria highlights an

apparent stratification with regard to deep ocean

waters and near surface or freshwater environments

during the Neoarchaean. Although the magnetite in its

current textural position is clearly metamorphic, it is

implausible to derive it from the oxidation of original

sulphide, bearing in mind the comparatively over-

whelming proportion of sulphide in the entire strati-

graphic column. Any post-depositional fluid that

percolated through the Witwatersrand Basin fill is

more likely to have been enriched in S-species from

the partial dissolution of the abundant pyrite in the

stratigraphic succession than to have been capable of

selectively oxidising a hypothetical primary sulphide

in the intercalated shale beds. This is clearly evidenced

by the common observation of secondary pyrite

overgrowths at many stratigraphic levels. Furthermore,

the shale beds were most likely those with the lowest

permeability and thus least likely to have been affected

by chemical change due to fluid circulation across

stratigraphic boundaries. Derivation of the magnetite

in these marine shale deposits from an oxide precursor

is therefore assumed (Frimmel, 1996).

As illustrated by Huston and Logan (2004), Fe

solubility in the system Fe–S–O is highest and lowest

in the magnetite and pyrite stability fields, respec-

tively (Fig. 10). Magnetite is only stable under

reduced conditions (ASO4/AH2Sb10�2.5) and very

low total sulphur concentrations, whereas pyrite can

be stable even under oxidising conditions (ASO4/

AH2Sb108), provided the total sulphur concentration

is high, i.e. close to that of modern seawater. In the

presence of Ba, the pyrite stability field decreases to

reducing conditions (ASO4bAH2S) as barite precip-

itates under relatively oxidising conditions (ASO4/

Fig. 10. Phase relationships and Fe solubilities in the system Fe–

Ba–S–O as a function of redox potential (shown as ASO4/AH2S)

and total sulphur concentration normalised to modern ocean water

composition at a temperature of 25 8C and pH=7.8 (from Huston

and Logan, 2004).


AH2SN10�2). In contrast to barite, Ca-sulphates are

highly soluble and a considerable degree of evapo-

ration is needed to precipitate gypsum. No occurrence

of sedimentary sulphate is known to date from the

Witwatersrand. However, based on S isotopic evi-

dence, England et al. (2002b) concluded that some of

the rare ooid-like pyrite grains found in the Venters-

dorp Contact Reef represent the product of microbial

sulphate reduction. This might reflect sulphate in the

source area, such as barite that has been described

from the Palaeoarchaean Barberton greenstone belt

(see compilation by Huston and Logan, 2004).

The apparent lack of sulphates in the Witwaters-

rand rock record may be explained in various ways.

From a thermodynamic standpoint it should reflect a

decrease in redox potential and/or total S concen-

tration, or a lack of Ba. The latter possibility is not

considered further as the principal Ba source is

hydrothermal discharge, which is unlikely to have

been shut down over several hundred million years

during Witwatersrand sediment deposition. Whether

the lack of sulphate in the Witwatersrand rocks in

particular, and in Meso- to Neoarchaean rocks in

general, reflects a combined increase in oceanic Fe

and decrease in oceanic sulphate as well as total S

concentrations relative to the Palaeoarchaean ocean,

remains contentious and dependent on the interpreta-

tion of Palaeoarchaean sulphate deposits. An abrupt

change in ocean water chemistry from a stratified

ocean with sulphate-bearing top waters to a sulphate-

free, Fe-rich ocean has been suggested to have

occurred around 3.2 Ga (Huston and Logan, 2004)

and has been ascribed to heavy bombardment of

Earth’s surface by meteorites (Glikson, 2001). Palae-

oarchaean sulphate precipitation may well have been

restricted to isolated oases of marine evaporative

ponds (Pavlov and Kasting, 2002; Shen et al., 2001)

and not representative of the world ocean. The lack of

comparable deposits in the Witwatersrand succession

could be merely a function of a lack of suitable

environments for such oases in the Witwatersrand

Basin at that time, and of preservation as evaporite

deposits would be prone to erosion especially in a

tectonically active foreland/retroarc setting.

The presence of magnetite in many of the

Witwatersrand shales mirrors a global surge in iron

formation occurrences during the Meso- to Neo-

archaean (Trendall, 2002), which reflects a reduced

ocean, in which high concentrations of Fe2+ were

possible in the bottom waters. The total S concen-

tration in that ocean must have been less than 10�5

that of modern ocean water (Fig. 10). In contrast, the

prevalence of pyrite in fluvial to shallow marine

Witwatersrand deposits points to relatively higher

total S concentration and/or higher O2 levels (Fig. 10)

in the oceanic top waters and the meteoric realm.

One of the magnetite-rich shale beds in the West

Rand Group rests above diamictite deposits (Corona-

tion Formation, Fig. 4). A causal link between iron

formation and global ice age, as suggested for the

younger Neoproterozoic iron formations, which are

viewed as result of isolation of the oceans from the

atmosphere by global, or near-global, ice cover

(Kirschvink, 1992; Klein and Beukes, 1993), may

therefore be applicable also to the magnetite-rich shale

beds of the Witwatersrand. According to that model,

melting of the ice cover would have triggered Fe-

oxide precipitation following hydrothermal Fe-enrich-

ment during glaciation. It should be noted, however,

that a number of magnetite shale beds in the West

Rand Group are not associated with diamictite.

Furthermore, the above model is not universally

accepted, not even for the Neoproterozoic deposits

(for critical assessment of existing evidence see

Young, 2004). Consequently, glaciation might have

further facilitated the deposition of iron formation

(magnetite shale) in the Meso- to Neoarchaean, but

was probably not the primary cause of it.


6.3. Quantifying Neoarchaean oxygen levels

Apart from the mentioned preservation of mass-

independent S isotope fractionation in pre-2.3 Ga

sediments, the two main constraints on the O2

concentration in the palaeo-atmosphere between 3.0

and 2.6 Ga are based on the detrital mineralogy of

placer deposits and on biochemical data on early

microfossils. As pyrite requires even lower oxygen

fugacity to be stable than uraninite, it is the

preferred phase for setting an upper limit on ancient

O2 levels. A lower limit is given by lipid biomarker

data from sedimentary rocks of the Pilbara Craton

in Western Australia. They provide evidence of

oxygenic photosynthesis at least as early as 2.7 Ga

(Brocks et al., 1999), and a lower limit of at least

1% of present atmospheric O2 is set by the

presence of steranes, found in these rocks, because

eukaryotic steroids require free oxygen (Jahnke and

Klein, 1983). Although some workers have sug-

gested the presence of aerobic bacteria already in

Palaeoarchaean times, the evidence for that is

problematic and a question of debate (Canfield

and Raiswell, 1999).

Fig. 11. Rate of oxidation of Fe2+ to Fe3+ in a solution of modern seawat

oxygen fugacity ( fO2), pH and temperature (T): (A) pH of modern seawat

modern atmospheric O2 pressure corresponds to fO2=10�0.67; calculated u

In order to quantify the redox conditions that are

required to explain the presence of pyrite in the

given environments, the physico-chemical conditions

for, and the rate of, oxidation of divalent to trivalent

iron were calculated (Fig. 11). Considering that

pyrite was also stable in Meso- to Neoarchaean

shallow marine environments, a modern seawater

composition (pH=8.22) was used for the initial

calculations. Under these conditions, oxidation of

Fe2+ to Fe3+ would be effectively instantaneous at

fO2 of 10�3, very slow at fO2 of 10�7 and

impossible at fO2 of less than 10�8 at a temperature

of 25 8C, with two orders of magnitude lower fO2

required for comparable reaction rates at a temper-

ature of 50 8C (Fig. 11A). This result is incompatible

with the above biochemical limit. As indicated by

the geochemical data, the atmosphere at the time of

interest was likely acidic. A lower pH of the

contemporaneous seawater would be more in line

with the combined evidence from detrital pyrite and

eukaryotic steroids. At a pH of 6.0, which would

correspond to the inferred silicate weathering to

kaolinite and the partial survival of detrital musco-

vite, oxidation of Fe2+ to Fe3+ would be very rapid

er composition as a function of time, in dependence of atmospheric

er, (B) pH=6; left panel—T=25 8C, right panel—T=50 8C; note thatsing data generated from PHREEQC (Parkhurst and Appelo, 1999).


at fO2 of 10�0.67—the modern atmospheric O2

level—and extremely slow at fO2 of 10�3—the

biochemical limit (Fig. 11B). Consequently, the

biochemical fO2 limit of 1% present atmospheric

level cannot have been exceeded significantly.

7. Conclusions

The Witwatersrand gold fields in South Africa,

which are the world’s largest gold producing

province, hold important keys for understanding

Archaean atmospheric and hydrospheric evolution.

Crucial to the debate around atmospheric O2 levels

at that time is the genesis of redox-sensitive

minerals that are associated with the gold.

Although most of the gold appears as a precipitate

within, or associated with, post-depositional hydro-

thermal phases and along microfractures, available

microtextural, mineralogical, geochemical and iso-

topic data, as well as the macro-scale stratiform

distribution of the ore bodies and its strong

sedimentological control, all indicate that this

hydrothermal gold, analogous to the associated

pyrite and uraninite, was derived from the local

mobilisation of detrital particles. Some of the key

pieces of evidence are a significant correlation

between gold and other heavy minerals as well as

sedimentary lithofacies, local preservation of in-situ

micro-nuggets with well preserved delicate textures

indicative of aeolian abrasion, compositional heter-

ogeneity on a microscale, and radiometric age data

that indicate an age of the gold, pyrite and

uraninite (3.03 Ga) that is older than the maximum

sedimentation age for the host sediment (2.90 Ga).

None of these observations/data is compatible with

a hydrothermal model, in which auriferous, poten-

tially sulphidising fluids were introduced from an

external source into the host rocks after sediment

deposition. In contrast, those arguments, used in

favour of hydrothermal models, emphasise the

microtextural position of most of the gold, which

highlights the undisputed hydrothermal nature of

much of the gold in its present position, but does

not explain the ultimate source of that gold.

Similarly, microtextural features, S isotopic

heterogeneity within and between grains, as well

as direct dating by the Re–Os method indicate that

the by far most abundant morphological variety of

pyrite in the Witwatersrand deposits, i.e. rounded

pyrite, is detrital. The same applies to rounded

uraninite, which is responsible for the Witwaters-

rand, together with the siliciclastic deposits of the

2.45 Ga Elliot Lake Group, representing the

world’s largest inferred U resource. Mineral chem-

ical characteristics and particularly variability

between grains, together with direct dating by the

U–Pb method, provide independent evidence of its

detrital nature. As with the gold, both pyrite and

uraninite were mobilised during post-depositional

fluid–rock interaction to variable degree, whereby

partial to complete oxidation to brannerite affected

the detrital uraninite, especially at higher strati-

graphic levels.

The currently available data point to mechanical

transport of gold, pyrite and uraninite from eroded

source regions that bear similarities to granitoid–

greenstone terranes currently exposed to the north

and west of the Witwatersrand Basin. The propor-

tion between felsic and mafic/ultramafic source

rocks varies strongly between and within reefs,

and this explains the only poor correlation between

Au and elements that are concentrated in detrital

minerals. Sediment deposition took place initially in

a continental rift (Dominion Group), followed by a

passive margin (West Rand Group), and then in a

foreland setting relative to the collision between the

Witwatersrand and Kimberley crustal blocks, with a

subsequent change to a retroarc position in con-

sequence of oceanic basin closure to the north of

the craton (Central Rand Group). Most of the gold

accumulated in the foreland/retroarc setting as that

setting favoured particularly intense sediment-

reworking on a series of erosional unconformities.

The recognition that rounded pyrite and uranin-

ite, both of which are characteristic of 3.07 to 2.64

Ga fluvial to fluvio-deltaic siliciclastic sediments on

the Kaapvaal Craton are not the products of post-

depositional hydrothermal fluid infiltration, as sug-

gested repeatedly in the past (e.g. Barnicoat et al.,

1997; Phillips and Law, 2000; Phillips and Myers,

1989), but detrital components of numerous placer

deposits re-affirms their pivotal importance for

constraining the evolution of O2 concentrations in

the Archaean atmosphere and hydrosphere. The lack

of Fe-oxides/hydroxides and the survival of pyrite


and uraninite during fluvial transport strongly

support a reducing Meso- to Neoarchaean atmos-

phere. In contrast to the fluvial to near-shore coarser

grained deposits, marine shale deposits contain

magnetite instead of pyrite. Comparably reducing

conditions are therefore inferred also for the oceanic

bottom waters, but with orders of magnitude lower

total sulphur concentrations than in modern ocean

water. Thus the observed distribution of pyrite and

magnetite in the Witwatersrand strata are in support

of the hypothesis, largely based on sedimentological

and S isotopic evidence, that Fe2+ was the principal

oceanic redox buffer prior to 2.4 Ga, whereas after

1.8 Ga, following an intervening transition period,

sulphate took over that role.

During the Meso- to Neoarchaean Aeras, total S

levels in the oceans must have been extremely low.

Sulphate-reducing bacteria have been inferred from as

early as Palaeoarchaean times (Shen et al., 2001). Any

microbial sulphate reduction, together with hydro-

thermal inorganic sulphate reduction, would have

removed sulphate rapidly to precipitate Fe-sulphides

from Fe-rich ocean waters. Consequently, the oceans

during Witwatersrand times would have been essen-

tially free of sulphate. Furthermore, the lack of

oxidative weathering of terrestrial pyrite would have

prevented the supply of sulphate to the oceans at a rate

that was greater than sulphate removal by Fe-sulphide

precipitation.

Most other evidence used for constraining

Archaean atmospheric O2 concentrations, such as

mass-dependent S isotope fractionation between

Archaean sulphate and sulphide, chemical compo-

sition of gold grains, and geochemical character-

istics of inferred paleosols, are not as conclusive as

the abundant occurrence of detrital pyrite and

uraninite in sediments that were laid down over

extensive areas on several cratons. Independent

support for a reducing Archaean atmosphere comes

from mass-independent S isotope fractionation. An

acid atmosphere, inferred from geochemical data,

was probably in equilibrium with a correspondingly

acid ocean. Kinetic calculations of the oxidation

from Fe2+ to Fe3+ show that a pH of 6 is required

in order to explain both the survival of detrital

pyrite and the presence of eukaryotic steroids in

Neoarchaean sediments. The calculated fO2 of 10�3

is in good agreement with the atmospheric evolu-

tion suggested previously by Kasting (1987, 2001;

Fig. 1A: curve a).

In summary, the available data endorse an acidic

hydrosphere beneath a reducing atmosphere during

the Archaean and early Palaeoproterozoic. Such an

acid environment requires elevated concentrations of

greenhouse gases, predominantly CO2. In addition,

the postulated anoxic atmosphere would have

favoured methanogenic bacteria that could contribute

to elevated atmospheric CH4 concentrations, in

agreement with the available C isotope record for

the Archaean (Pavlov et al., 2001b). Thus the

Archaean atmosphere was likely to be enriched in

effective greenhouse gases that would have effi-

ciently offset the lower solar luminosity in the early

history of Earth as suggested by Walker et al.

(1983). Such a palaeoclimate model explains the

mineralogy of ancient placer deposits. Variably

modified palaeoplacer deposits are known from the

Mesoarchaean to the Palaeoproterozoic from a

number of cratons, but only those older than about

2.4 Ga contain detrital pyrite and uraninite, whereas

in the younger deposits, detrital sulphides and

uraninite are conspicuously lacking and Fe-oxides

occur instead.

Acknowledgements

The author is indebted to Lawrie Minter who

has never hesitated in sharing his enormous

experience on the sedimentology and economic

geology of the Witwatersrand and related deposits

and who has given unrestricted access to his

valuable sample collection, some of which has

historic value as it contains material from mined

out areas that are not accessible anymore. The work

presented is based on numerous visits to under-

ground mines and core yards, which would have

been impossible without the cooperation of a

number of mining companies. Of particular impor-

tance for the conclusions presented in this paper

was the logistic support granted by Anglogold and

its staff. J.B. Maynard is thanked for a constructive

review of the manuscript. Parts of the work were

funded through grants from the South African

National Research Foundation and the University

of Cape Town.


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Hartwig E. Frimmel, PhD, is Professor at

the University of Wqrzburg, Germany,

where he heads the Institute of Mineralogy.

He is also Professor at the Department of

Geological Sciences, University of Cape

Town, where for the past 15 years he has

worked inter alia on the genesis of the

Witwatersrand gold deposits. Other major

research interests include the relationship

between plate tectonics, palaeo-climate and

syn-sedimentary ore-forming processes,

with particular focus on the Neoproterozoic Aera, and the geo-

dynamic evolution of Precambrian supercontinents, with regional

emphasis on Antarctica, Africa and South America. Since 1998 he

has been the leader of the Earth Science subprogramme within the

South African National Antarctic Programme. He has served on

several editorial boards, supervised numerous post-graduate stu-

dents and has over 80 publications in international journals and

books to his credit.

Documents

Archaean Atmospheric Evolution Evidence From the Witwatersrand Gold Fields, South Africa (1)