Ore Petrology, Chemistry, and Timing of Electrum ... - CiteSeerX

Econoraic Geology Vol 93, 1998, pp. 271-291

Ore Petrology, Chemistry, and Timing of Electrum in the Archcan Hypozonal Transvaal Lode Gold Deposit, Western Australia

STEFFEN G. HAGEMANN?

Department of Geology and Geophysics, Centre for Teaching and Research in Strategic Mineral Deposits, University of Western Australia, Nedlands, Western Australia 6907, Australia

PHILIP E. B•tOWN,

Department of Geology and Geophysics, University of Wisconsin, Madison, Wisconsin 53706

JOHN RIDLEY, *

Key Centre for Strategic Mineral Deposits, Department of Geology and Geophysics, University of Western Australia, Nedlands, YVestern Australia 6009, Australia

PETER STERN, AND JOHN FOURNELLE Department of Geology and Geophysics, University of Wisconsin, Madison, Wisconsin 53706

Abstract

The Transvaal lode gold deposit is hosted in ultramarie schists and graphite-bearing pelites that are metamorphosed to lower and mid-amphibolite facies assemblages. The orebody is strueturally controlled by the Transvaal shear zone which overprinted and reactivated the penetrative regional fabric. The main tabular orebody is formed of a massive quartz sulfide vein and related splays and adjacent hydrothermally altered wall rock. Prograde alteration in the ultramarie rocks consists of a proximal diopside-aetinolite-quartz zone which grades outward into a distal amphibole (hørnblende-aetinølite)-biøtite-plagiøelase-quartz schist. In pelitie metasedimentary rocks, a proximal biotite-muscovite-quartz-graphite zone grades outward into a distal cordieritc-quartz-biotite-graphite schist. Retrograde ehlorite, tale, and zoisite replaced prograde hydrothermal alteration minerals and exhibit generally postkinematic fabrics. The dominant sulfide assemblages in the hydrothermal alteration zones at Transvaal are pyrrhotite-loellingite, pyrrhotite-arsenopyrite 1, and loellingite- arsenopyrite 1. Other assemblages include ehaleopyrite-pyrrhotite, pyrite-loellingite, and pyrrhotite-eubanite. These assemblages are interpreted to consist of a mixture of equilibrium assemblages that relate to a prograde hydrothermal alteration event and disequilibrium assemblages that formed during cooling and reequilibration of the rocks. Arsenopyrite 1 is interpreted to have formed both during the prograde alteration event and during high-temperature retrogressive replacement of loellingite. Average compositions of arsenopyrite i in equilibrium with loellingite and pyrrhotite indicate palcotemperatures of about 510 ø _+ 20øC, about 40øC lower than estimates of the peak metamorphic temperatures. Low and high Fe pyrrhotite and arsenopyrite 2, for instance, would not be in equilibrium with other sulfides and are interpreted to be retrograde.

Both electrum and gold occur in the Transvaal gold deposits but electrum is the more abundant gold- bearing phase and contains between 24 and 56 wt percent silver. The compositions of electrum grains correspond to gold finenesses of about 431 to 732 with silver the dominant impurity. Electrum occurs mostly at the interface between arsenopyrite 1-1oellingite, or as inclusions in loellingite, pyrrhotite, and arsenopyrite 1. More rarely, it occurs as as fracture fill within arsenopyrite 1. The differing silver contents of electrum can be explained by differing fluid-rock reactions. Low Ag electrum (about 24 wt % Ag) occurs where the lode is entirely hosted by ultramarie schists, whereas high Ag electrum is at the ultramarie sehist-graphitie-pelite contact. It is proposed that fluid interaction with reducing pelites produced Ag-rieh electrum through destabili- zation of gold bisulfide and silver chloride complexes, whereas fluid that reacted with more oxidized ultramarie rocks formed Au-rieh electrum as silver chloride complexes were not destabilized.

The main phase of hydrothermal alteration, which included the crystallization of prograde silicates, pyrrhotite, and loellingite, was synteetonie with respect to the Transvaal shear zone and at approximately peak metamorphic conditions. Electrum in equilibrium with pyrrhotite and/or loellingite was thus deposited at synpeak metamorphic conditions. However, the majority of electrum grains coexist with arsenopyrite i and composite loellingite-arsenopyrite i grains and are interpreted to have formed from submicroscopic gold released from loellingite during high-temperature retrogressive replacement by arsenopyrite. These are also considered to be part of the main gold-bearing hydrothermal event.

Retrograde low-temperature, postmetamorphic silicate and sulfide minerals such as ehlorite, muscovite, pyrite, and cubanitc are evidence for processes that occurred after the main, gold-depositing hydrothermal

• Corresponding author: email, shageman@geøl'uwa'edu'au * Present address: ETH Zurich, Zurich, Switzerland.

0361-0128/98/1970/271-2156.00 271

272 HAGEMANN ET AL.

activi• took place. Renewed flux of hydrothermal fluids during the protracted metamorphic history of the terrane could have produced the low-temperature retrogression; the occurrence of low-temperature sulfides is likely the result of in situ reequilibration during cooling of the rocks.

Introduction

MOST case studies on Archcan lode gold deposits have been concerned with classic mesozonal deposits in the cratons of Western Australia, Canada, Brazil, southern Africa, and India (Gebre-Mariam et al., 1995). In many of these deposits gold occurs as free gold (>700 fineness) in medium-temperature (250ø-450øC) quartz carbonate veins and breccias and/or as inclusions (< 100/a) in pyrite and/or arsenopyrite. Gold deposition is syntectonic and postpeak metamorphic relative to the regional tectonic history. However, there is still considerable uncertainty as to the fluid source responsible for gold mineralization. Recent studies on high P-T, hypozonal Archcan lode gold deposits in Western Australia (Frasers and Griffins Find, Barnicoat, 1989; Marvel Loch, Mueller et al., 1991; Coolgar- die, Knight et al., 1993; Norseman, McCuaig et al., 1993; Southern Cross, Bloem et al., 1994) and Canada (Eastmain River, Couture and Guha, 1990; Hemlo, Kuhn et al., 1994) show that gold mineralization is associated with high-temperature (450ø-650øC) metasomatic silicate (e.g., Ca amphibole, diopsidic pyroxene, and garnet) and sulfide assemblages (e.g., pyrrhotite, loellingite, and arsenopyrite). Gold in hypozonal deposits generally occurs also as high fineness (>700) free gold in quartz veins, in intensely altered wall rock, or as inclusions within loellingite. The relative timing of the introduction and possible remobilization of gold in hypozonal deposits with respect to the regional tectonic history is in many cases regarded as equivocal. Several hypotheses have been invoked: (1) prepeak metamorphic introduction with various degrees of remobilization (e.g., Big Bell, Phillips and de Nooy, 1988; Hemlo, Kuhn et al., 1994; Kolar schist belt, Hamilton and Hodgson, 1986), (2) synpeak metamorphic introduction (e.g., Griffins Find: Barnicoat et al., 1991; Red Lake, Andrews et al., 1986), and (3) postpeak metamorphic introduction (e.g., Marvel Loch, Mueller et al., 1991; Detour Lake mine, Marmont, 1986; Renco, Tabcart, 1987; Hemlo, Pan and Fleet, 1992).

This study of the Transvaal lode gold deposit in the South- ern Cross greenstone belt of Western Australia employs detailed ore petrology and X-ray mapping techniques to decipher exact locations and chemistry of gold and associated sulfide minerals. A major difference to the case studies listed above is the predominant occurrence of electrum rather than gold as the major gold-bearing phase. Furthermore, the close association of electrum with high-temperature (510øC) composite arsenopyrite-loellingite grains, coupled with the lack of evidence of pervasive gold remobilization, indicates that at Transvaal gold mineralization was related to high-temperature metasomatic assemblages and was introduced peak to slightly postpeak with respect to regional metamorphism. This study presents further evidence that significant Archcan lode gold deposits formed in high-temperature (>450øC), deep, hypozonal el'uStal levels (> 10-km depth) and that, in the case of the Transvaal deposit, electrum was deposited broadly contemporaneously with the peak metamorphic event of the terrane.

Regional Geologic Setting The Southern Cross greenstone belt is located in the

Southern Cross tectonic province of the Archcan Yilgarn block in Western Australia (Fig. 1). The rocks likely formed as a platform-style Arch can basin in which layered sequences of tholeiitic-komatiitic volcanic rocks and minor felsic volca-

nic rocks were deposited (Gee et al., 1981; Groves and Bart, 1984). Domal uplift of granitic intrusions and simultaneous regional deformation squeezed these sequences into a series of narrow north-northwest-trending greenstone belts (e.g., Bloem et al., 1994). The stratigraphy of the Southern Cross greenstone belt is poorly exposed due to extensive laterite cover but may be subdivided into an up to 3,000-m- thick lower greenstone sequence consisting mainly of metakomati- ires and metabasalts, intercalated with thin layers of graphitic quartz-biotite schist and quartz-grunerite banded iron-formation, and lens-shaped metagabbro sills; and an upper gray- wacke sequence (> 1,500 m thick), composed mainly ofpelitic metasedimentary rocks with a thin zone of carbonaceous metasedimentary rocks and sulfidic cherts developed at its base (Gee, 1982; Keats, 1991; Mueller, 1991; McQueen, 1992). The lower sequence is host to most of the major hypozonal lode gold deposits (e.g., Marvel Loch, Mueller, 1991; Frasers, Barnicoat et al., 1991), however the volcanic massive sulfide-style Great Victoria gold deposit (McQueen, 1992) and the Transvaal deposit are sited at or just above the contact between the two sequences. The greenstones are intruded by granitoid plutons. Metamorphic isograds are subparallel to the granitoid-greenstone belt contacts with high- to medium- grade rocks (granulite to amphibolite facies) near the margins of the belt, and lower grades, down to upper greenschist facies, in the central parts (Ahmat, 1986; Mueller, 1988; Dals- traet al., in press). The large-scale structure of the belt is characterized by regional upright folds with variably plunging axes and by a northwest-striking and steeply southwest- and northeast-dipping system of transcurrent ductile shear zones, that follow the southwestern side of the Ghooli dome (Fig. 1). One of these ductile shear zones, the Transvaal shear zone, controls the Transvaal gold deposit.

Methodology Forty-five polished sections and 40 doubly polished thin

sections of ore samples were examined using reflected light microscopy. Selected samples of sulfides and gold were examined by electron microprobe, using a Cameca SX51 at the University of Wisconsin-Madison. Analytical conditions for sulfide and gold analysis were 25 or 20 keV accelerating volt- age, with 20 nA of beam (Faraday cup) current, using a fLxed beam. The following standards were used: for arsenopyrite and loellingite, ASP200 (natural arsenopyrite; Kretschmar and Scott, 1976); for pyrite, a natural pyrite standard (Elba); for pyrrhotite, a troilite standard from the Staunton meteorite (American Museum of Natural History); for chalcopyrite, a chalcopyrite standard provided by G. Czamanske; for gold- electrum, an electrum (Au60Ag40) standard provided by G. Czamanske. Backscattered electron images and X-ray maps

ARCHEAN HYPOZONAL TRANSVAAL LODE GOLD DEPOSIT, W. AUSTRALIA 273

,•--• Granitoid Pelitic Adamellite [• Rocks Metasedimentary • Mafic-ultramafic Metavolcanic Rocks

.;:i..':• Metamorphosed Felsic Proterozoic Gneiss Volcanic Rocks • ..- Dolerite Dike

FIG. 1. Generalized geologic map of the Southern Cross greenstone belt (modified after Gee, 1982) with location of major high-temperature (hypozonal) gold deposits, including the Transvaal lode gold deposit.

were also collected using the SX51 electron microprobe. The backscattered electron images mode was used to decipher cryptic compositional zoning in arsenopyrite and pyrrhotite and to detect minute particles of gold. X-ray maps were collected for defining compositional variations over areas ranging from 256 x 256 •m, to 1,024 x 1,024 •m. Three microprobe analyses of electrum from the Transvaal deposit were obtained by Dalstra (1995) and used in this study.

Geologic Setting of the Transvaal Gold Deposit The Transvaal gold deposit is located toward the center of

the Southern Cross greenstone belt, south of the town of Southern Cross (Fig. 1) near an ultramafic schist-pelite contact. Structures and fabrics of two phases of deformation are preserved in the mine area. A penetrative regional S2 foliation is here essentially axial planar to a tight to isoclinal asymmetric fold (local F,•) of the contact. The foliation is overprinted, reactivated, and intensified within D3 ductile shear zones, which are up to a few meters wide, can be traced over a few hundred meters strike length, display only minor (<10 m) displacement of earlier structures, and as they are a few degrees oblique to the S.• fabric, intersect the moderately steep, northerly plunging large-scale fold of the ultramafic schist- pelite contact. Based on detailed mineral stability and geo- thermobarometric analyses, Dalstra (1995) proposed lower to mid-amphibolite peak metamorphic conditions for the Transvaal area, with evolving P-T from low pressure-high temperature (3.0-4.5 kbars at 550øC) to very low pressure-

high temperature (<2.9 kbars at 550øC) conditions. Peak metamorphic minerals outside the immediate mine area com- prise in ultramafic rocks, hornblende, cummingtonite, actinolite _ tremolite, gruneritc, chlorite, plagioclase, talc, _ carbonate, ilmenite, magnetite, Cr spinel, chromite, and in recta- sedimentary rocks, andalusite, cordieritc, almandine, biotite, quartz, plagioclase, limehire, titanitc, magnetite, graphite, _ chlorite, and _ muscovite (Dalstra et al., in press). Microtex- tures indicate that the peak of the earlier phase of slightly higher pressure metamorphism was late- to post-D,• and that the lower pressure phase was syn- to post-Ds (Dalstra, 1995).

The Transvaal deposit is hosted by the most prominent of the Ds shear zones in the area, the Transvaal shear zone. A single major tabular and well-defined ore shoot, formed of a massive quartz sulfide vein, with some splays, and immediately adjacent altered wall rock, between 0.5 and 3 m in width, about 150 m long, trending north-northwest, and plunging moderately steep to the south, has been mined to about a 200- m depth at the deposit. The plane of the host shear zone and lode is on average 5 ø to 10 ø oblique to both the penetrative fabric and the ultramafic-pelite contact, however, it is deflected so as to follow the ultramafic-pelite contact over several tens of meters strike length, especially in the upper levels of the mine. The shear zone transgresses progressively across the fold form and the structural succession to deeper levels. Three smaller subparallel lodes occur to the east (Fig. 2A).

The lode is characterized by pinch and swell structures, and alteration is restricted to within a few meters of the ore

274 HAGEMANN ET AL.

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FIG. 2. Level plans and long section of the Transvaal lode gold deposit. A. Surface projection from Dalstra et al. (in press). B. Level 171 (40 m RL; modified after James, 1995). c. Long section (9,800 m E; modified after Denig, pers. commun.). Mine development shown as at end of 1994.

zone. The term "ore" in this paper is defined as a mixture of sulfide and oxide minerals that can be observed macroscopi- cally. Regardless of host rock, the lode zones are characterized by at least six distinct ore textures. These are, in order of decreasing abundance: (1) foliation-parallel zones that occur in both quartz-rich veins and wall rocks, (2) massive wallrock ores that immediately border veins, (3) irregular blebs of sulfides within the wall rocks and veins, (4) ore in zones of microbrecciation in the veins, mostly at the contact between wall-rock clasts and the quartz matrix, (5) finely disseminated ore in wall rocks and veins, and (6) irregular monomineralic fracture fill. All six modes carry economic grades, but gold is typically most abundant in types I and 2.

In view of the Da, near-peak metamorphic timing of formation of the host shear zone, it is interpreted that hydrothermal alteration and mineralization cannot have been prepeak metamorphic but must have been either broadly at the peak of metamorphism or postdated it. There is no evidence that the shear zone followed an earlier zone of alteration.

Hydrothermal Alteration, Textures, and Timing with Respect to Deformation and Metamorphism

General characteristics

Hydrothermal alteration envelopes around the lodes are typically limited to a few meters, extending across the lami-


A.

Ultramafic Distal Proximal rocks Least Wallrock!Vein -Wallrock!Veins

altered i lets j SILICATES

Amphibole i i Actinolite/Trem

Grunerite ; i Cummingt. onitc Hornblenoe

Pyroxenes • i _ Diopside ... Garnet r i Plagioclase

Chlorite ....... . Chlorite (R) ....... • i Biotite ...... •r.USCovite (.R) .......

muscovite (R) Talc I . .i Pumpellyite (R) .......

Zircon . CARBONATES ..................... , ..... OXIDES i Magnetite .............. .. Titanite ....... I. .... i Chromite .............. • ........ . ...... Ilmenite ....... I I Cr-spinel .......

i i PHOSPHATES Apatite '- ....

SUEFIDES i i Pyrrhotite :' ' ' ' Loellingite I I Arsenopyrite I ....... '- - - - Arsenopyrite 2* i i Pyrite* . ............. Chalcopyrite* i Gersdorfite • ....

Niccolite i i ..... Galena .. .....

Cubanite* i ELEM/ALLOYS

Gold i" .... i Electrum .. .....

a.

Metased. Distal Proximel rocks Least Wallrock!Vein - Wallrock!Veins

altered i lets i SILICATES

Andalusite i ...... .i Cordierite

Garnet (almandine) ...... j i Plagioclase i i ...... Quartz Chlorite ....... I I

Chlorite (R) .............. . ....... . Biotite Muscovite ...... Tourmaline ....... i ' ' i Zoisite (R)

CARBONATES OXIDES i i Magnetite ............. , Titanite ...... •' .... i Ilmenite .......

SUEFIDES i i

Pyrrhotite ' ."." .': .' .' •' "' I Loellingite . Arsenopyrite 1 -- - Arsenopyrite 2* r i Pyrite* - ..... •- - - - [- ..... Chalcopyrite* Cubanite* i i ......

ELEM/ALLOYS

Graphite ';--- • ..... Electrum : ,- ....

* = possibly result of re-equilibration during cooling of the rocks

(R) = retrograde mineral with respect to main phase of hydrothermal alteration and gold mineralization

F•G. 3. Schematic mineral zonation profile through an orebody hosted in (A) ultramarie, and (B) metasedimentary rocks. Thick and thin lines indicate major and minor mineral occurrence, respectively. Dashed lines indicate trace occurrence.

nated veins, and hosting ductile shear zones into the adjacent wall rocks. Based on mapping of ore sections in the underground mine, drill core logging and petrographie studies from two continuous diamond cores (TDD 103 and TDD 95, see location--intersection with long section--in Fig. 2) and on selected samples from different underground levels, prograde and retrograde alteration minerals (Fig. 3) were distin- guished. The prograde minerals form the main, synkinematic fabric in the D3 shear zone in both ultramarie and metasedi- menta W host rocks and are restricted to the shear zone environment. Retrograde minerals are low-temperature alteration minerals that mostly replace prograde minerals and almost always exhibit postkinematic textures. They are common within and outside the Transvaal shear zone in volcano-sedi-

menta W lithologies throughout the greenstone belt (Dalstra, 1995).

A zoned distribution of prograde assemblages occurs in ultramarie rocks (Fig. 3A), typically with a proximal diopside- actinolite-quartz zone which grades outward into a distal amphibole (hornblende-actinolite)-biotite-plagioclase-quartz schist. The transition from the amphibole-chlorite-plagioclase-talc ultramarie host rock to the distal alteration zone is

extremely gradational and is marked by the onset of quartz- plagioclase veinlets and the appearance of quartz and biotite. Similar alteration zoning in high-temperature, hypozonal gold deposits has been described in the Pilbara eraton (Mount York district, Neumayr et al. 1993b, 1995) and the Yilgarn eraton (Southern Cross and Coolgardie greenstone belts,

Mueller and Groves, 1991; Knight et al. 1993, respectively) of Western Australia.

In pelitie metasedimentary rocks (Fig. 3B) a proximal biotite-muscovite-quartz-graphite zone grades outward into a distal eordierite-quartz-biotite-graphite schist which grades into unaltered andalusite-eordierite-almandine-biotite-plagioelase-quartz __ graphite metapelite host rock. The contact between the distal alteration zone and unaltered rocks is gradational and in many eases not dearly definable because of heterogeneities in the pelitie schists. The distal zone is characterized by more abundant biotite and quartz-biotite-carbonate veinlets, the presence of disseminated pyrrhotite and loellingite, and less andalusite and garnet than in unaltered rocks. Similar descriptions of alteration zoning in metasedimentary rocks of high-temperature, hypozonal lode gold deposits are rare with, perhaps, the Crixas deposit in the S•o Francisco eraton of Brazil (Thomson, 1986; Thomson and Fyfe 1990), another well-studied example. Mineralogieal changes associated with wall-rock alteration and eraplacement of veins in both ultramarie and pelitie metasedimentary host rocks are presented in Figure 3.

Alteration mineralogy and textures Distal alteration in ultramarie rocks is characterized by nar-

row intervals (< 1 m) and irregular clusters or flakes of biotite or amphibole, with interstitial quartz and plagioclase (Fig. 4A, B). Biotite is typically intergrown with metamorphic and hydrothermal amphibole (cummingtonite; Fig. 4A). Locally,

276 HAGEMANN ET AL.

F]c. 4. Characteristic alteration minerals and textures at the Transvaal lode gold deposits. All photomicrographs are about 4 mm across. A. Photomicrograph (TV 103-80A) of distal alteration in the ultramarie host rock in transmitted, cross- polarized light. Quartz (Qtz) vein is bordered by bands of cummingtonite (Cum) and biotite (Bio). Polysynthetically twinned plagioclase (Plag) grains occur interstitial within cummingtonite and biotite bands. B. Photomicrograph (TV 103-81A) of typical distal alteration in the ultramarie host rock in transmitted, cross-polarized light. Prograde actinolite (Act) crystals in contact with plagioclase (Plag) and talc (Tc), the latter two partially replaced by retrograde muscovite (Mu). C. Photomicro- graph (TV 103-81A) of distal alteration in ultramarie host rock in transmitted, cross-polarized light. Retrograde chlorite +__ zoisite (Chl ___ Zoi) replace large euhedral plagioclase crystal. D. Photomicrograph (TV 103-85A) of proximal alteration in


biotite forms narrow (<3 mm) monomineralic bands parallel to the shear zone fabric. Retrograde alteration minerals in- elude ehlorite _+ zoisite, which locally replaced plagioelase (Fig. 4C), and muscovite that replaced plagioelase and tale (Fig. 4B). Finely disseminated sulfides occur in the distal wall rocks (see below). Metamorphic amphibole, ehlorite, plagioelase, and tale are largely preserved but gradually diminish with proximity to the ore zone. Quartz-plagioelase + calcite veinlets erosscut the rocks irregularly and in places contain apatite and titanitc. Proximal alteration in the ultramarie rocks is characterized by an increase in the abundance of laminated quartz-plagioelase and quartz-diopside veins (up to 5 em; Fig. 4D) that alternate with zones of biotite and/or hydrothermal aetinolite (Fig. 4D). The veins include interspersed biotite and aetinolite, that locally overprinted diopside, and in places irregular patches of retrograde Cr-bearing mica or muscovite. Most sulfide minerals and gold are associated with these veins (see below). Monomineralie diopside veins (up to 5 em in width) are observed locally in the proximal zones. Biotite- and aetinolite-rieh zones (Fig. 4D) contain interstitial carbonate and ehlorite and, locally, euhedral zircons. Away from the biotite- and aetinolite-rieh zones, the wall rock contains euhedral garnets coexisting with dispersed diopside and quartz. Retrograde ehlorite, pumpellyite, and zoisite occur only locally and replace biotite, garnet, and diopside.

Distal alteration in metasedimentary rocks is characterized by an increase in quartz, muscovite, biotite, and irregular dots or fine veins (up to i mm) of graphite. Tourmaline and magnetite occur as euhedral minerals finely distributed throughout the wall rock. Retrograde ehlorite has locally replaced biotite (Fig. 4E). Metamorphic minerals are ubiquitous and include mostly porphyroblasts of cordierire and large, up to centimeter size, andalusite with garnet inclusions (Fig. 4E) and interspersed quartz. Metasedimentary rocks in proximal alteration zones are dominated by alternating biotite-muscovite-graphite bands and quartz-rich veins (up to 2 em). The former contains locally interspersed quartz and euhedral plagioelase grains (0.5 mm), whereas the latter generally contains flakes of biotite and muscovite, most of the sulfide minerals, and gold (see below). Quartz displays a range of grain sizes, from large euhedral (up to 2 mm) to very fine anhedral mosaiclike (0.01 mm) grains, suggesting local reerystallization. Wall rocks contain zones of biotite-muscovite-graphite that coexist with garnet (Fig. 4F) and interspersed quartz, and locally, euhedral to anhedral cordieritc. Garnet generally is less abundant toward the ore zone. Retro- gressive ehlorite and zoisite replace biotite and muscovite (Fig. 4F ).

Relative timing of prograde hydrothermal alteration with respect to deformation and metamorphism

Hydrothermal alteration minerals occur only within the Da ductile shear zone and its immediate wall rocks. In altered

ultramarie rocks, alteration minerals such as biotite or amphibole define the well-developed S.• synkinematic fabric. Lo- cally, however, actinolite occurs both parallel to the foliation and randomly overprinting it, indicating continued growth after shear zone deformation. In proximal alteration zones laminated quartz-plagioclase and quartz-diopside veins and adjacent alternating monomineralic zones of biotite and/or actinolite also define the Sa fabric (Fig. 4D). Locally, postkinematic, large (0.5 ram) randomly oriented hydrothermal plagioclase grains crosscut the foliation, indicating that plagioclase growth outlasted Ds deformation.

In altered pelitic metasedimentary rocks biotite, muscovite, and graphite define the Sa synkinematic fabric (Fig. 4F), whereas cordieritc, andalusite, and garnet are both synkinematic and postkinematic with respect to the Sa shear zone fabric (see also Dalstra et al., in press). Locally, postkinematic euhedral to subhedral garnets (up to 1.5 ram) overprint the foliation, suggesting that growth of this mineral continued after Da shear zone deformation (Fig. 4 E, F).

In summary, these microstructural observations indicate that (1) alteration minerals which formed during the Da shear zone eraplacement define a strong synkinematic Sa fabric, (2) the growth of metamorphic minerals such as cordieritc, andalusite, and garnet occurred both prior to Da shear zone deformation, as part of the regional metamorphic event, and syn- and post-Da shear zone deformation; this is consistent with the growth of metamorphic minerals outlasting the deformation event, and (3) alteration minerals, such as large euhedral plagioclase crystals in ultramafic rocks, locally formed post-Da shear zone eraplacement. We interpret these petrographic-microstructural observations as evidence that crystallization of the prograde alteration minerals at Transvaal was broadly contemporaneous with regional peak metamor- phisre, deformation, and Da shear zone eraplacement.

Sulfide Petrology and Chemistry The most abundant sulfide minerals at the Transvaal gold

deposit are pyrrhotite, arsenopyrite, and loellingite. Although there is clearly a greater abundance of these sulfides in the ore zones than in the outer alteration zone, there is no evidence for a lateral zonation of sulfide assemblages. The sulfide assemblage is characterized by an overall low sulfidation state.

The most abundant sulfide mineral, pyrrhotite, generally forms massive aggregates and backscattered electron images reveal internal exsolution textures, characterized by variations in iron and sulfur contents (Fig. 5A). Pyrrhotite commonly occurs in textural equilibrium with arsenopyrite, rarely with loellingite, and locally displays exsolution and anhedral intergrowth textures with cubanitc (Fig. 5B). A common equilibrium texture in the proximal alteration zone is anhedral pyrrhotite rimming composite arsenopyrite-loellingite grains. A test with a handheld magnet on Transvaal samples suggests that the majority of pyrrhotite is monoclinic. Electron micro-

the ultramarie host rock in transmitted, cross-polarized light. Actinolite crystals in contact with diopside-quartz vein (Di and Q, respectively). E. Photomicrograph (TV 103-89) of distal alteration in metasedimentary host rocks in transmitted, partially cross-polarized light. Large, euhedra] andalusite (And) porphyroblast rimmed by quartz and overprinted by biotite which is partially replaced by retrograde chlorite. Andalusite porphyroblast contains garnet (Gar) inclusions of variable sizes. F. Photomicrograph (TV 103-88) of proxima] alteration in metasedimentary host rock in transmitted, cross-polarized light. Subhedral garnet within groundmass of biotite and muscovite which are partially replaced by retrograde chlorite.

278 HAGEMANN ET AL.

Otz

ß

FIG. 5. Backscattered electron microscopy images of characteristic sulfide minerals, zoning, and textures at Transvaal. A. Euhedral arsenopyrite 2 (Apy) crystal located within chalcopyrite. Note the anhedral intergrowth of pyrrhotite (Po) and chalcopyrite (Cpy) and the internal exsolution texture of pyrrhotite. The light and dark shades of gray correspond to low and high Fe content, respectively. Sample TV 69A2. B. Cubanitc (Cub) exsolution in pyrrhotite and anhedral intergrowth texture of cubanitc and pyrrhotite. Sample TV 69A2. C. Euhedral arsenopyrite 2 with inclusions of pyrrhotite. Sample TV 122B2. D. Euhedral, arsenopyrite 2 with well-developed minimum of 4 growth zones. The As content (wt %) of growth zone (GZ) I = 35.03, GZ 2 = 33.28, GZ 3 = 32.28, and GZ 4 = 31.38. Sample TV 124.4.

probe analyses of pyrrhotite (Table 1) from the Transvaal proximal alteration zone revealed a range of composition which could be grouped, with respect to Fe (at. %) content, into a major population that contains 46.9 to 48.2 at. percent Fe, (mean 47.4 +__ 0.3 at. % Fe; I •r, n - 16) and that corre- lates to monoclinic pyrrhotite of nominal composition Fe?Ss. Two outTiers of Fe-rich pyrrhotites, which correlate to hexagonal pyrrhotite, contain 49.5 and 49.8 at. percent Fe (Ta- ble 1).

The second most abundant sulfide mineral, arsenopyrite, occurs in two generations. The first is synkinematic (with respect to the Transvaal shear zone) euhedral rhombshaped crystals and massive anhedral to subhedral mineral aggregates that are commonly intergrown with pyrrhotite. This type of arsenopyrite is described as arsenopyrite I throughout the text, figures, and tables. The synkinematic occurrence is evi- denced by the alignment of arsenopyrite I grains within the

shear zone foliation and veinlets that are part of the shear zone fabric. Many euhedral to anhedral arsenopyrite I grains are in textural equilibrium with anhedral loellingite or occur as composite arsenopyrite-loellingite grains. Microprobe analyses reveal that most arsenopyrite 1 grains are homoge- neous with respect to As (Fig. 6A) and S content and define a single population with 33.9 to 36.9 at. percent As. The trace element content of arsenopyrites is generally low; slighfiy elevated Co and Ni values were only locally observed (Ta- ble 1).

The second generation of arsenopyrite forms postkinematic, mostly small (•0.05 mm long) euhedral crystals which overprint the ductile (foliation) fabric or occur in late-stage fractures and veins. These texturally late-stage arsenopyrite 2 grains are rare and generally characterized by multiple, small (•0.02 mm) pyrrhotite and cubanitc inclusions (Fig. 5C). Significanfiy, arsenopyrite 2 is characterized by a lack of


equilibrium textures with other sulfides. Microprobe analyses indicate that arsenopyrite 2 is mostly strongly zoned with respect to As (Fig. 6B) and S, with differences in As content of up to 4 at. percent in individual crystals (Fig. 5D), and has significantly lower As contents than arsenopyrite 1. In addition, microprobe analyses reveal that arsenopyrite 2 contains locally up to 4.3 wt percent Cu.

Loellingite is ubiquitous in gold-rich zones and occurs as either composite grains with synkinematie arsenopyrite 1 or, rarely, as discrete anhedral grains coexisting with pyrrhotite. The grain boundaries between anhedral to subhedral composite loellingite and arsenopyrite 1 grains vary from straight to embayed. The ubiquitous rimming of loellingite by arsenopyrite 1 (the reverse is never detected) and the locally observed embayment textures suggest that loellingite was progressively replaced by the more sulfur rich arsenopyrite 1. Microprobe analyses indicate that some loellingites incorporate up to 10 wt percent nickel in their lattice (Table 1) and thus form part of the loellingite-rammelsbergite solid solution series (Ramdohr, 1980). The sulfur content in loellingite is relatively constant and ranges from 1.3 to 2.5 wt percent S.

Chaleopyrite is rarely developed as euhedral or anhedral mineral grains (Fig. 5A) in equilibrium with pyrrhotite. Euhe- dral ehaleopyrite overprints locally the alteration silicate fabric and is spatially associated with arsenopyrite 2, although there is no clear evidence of textural equilibrium between the two. Nieeolite and gersdorffite occur as accessory sulfide minerals, mostly coexisting with loellingite and arsenopyrite 1. Other minor accessory minerals include galena, pentlan- ditc, and pyrite.

Arsenopyrite geothermometry and sulfur activity

Experiments on the stability of arsenopyrite in the system Fe-As-S have been used to calibrate a geothermometer by using appropriate buffer assemblages (e.g., arsenopyrite-pyrite and -pyrrhotite + loellingite) and arsenopyrite compositions (Clark, 1960; Kretschmar and Scott, 1976; Sharp et al., 1985). The experimental work by Kretschmar and Scott (1976) yielded a pseudobinary T-X section relating atomic percent arsenic in arsenopyrite to temperature in different assemblages. Sharp et al. (1985) presented a logfs.z-T diagram with isopleths of arsenopyrite compositions superimposed on Barton's (1969) logfs.z-T plot of arsenopyrite stability. These isopleths are vertical in the pyrite and loellingite field (Fig. 7). No pressure correction of the deduced temperature is required due to the simultaneous shift of the equilibria and arsenopyrite isopleths (Sharp et al., 1985).

Arsenopyrite 1 coexists with loellingite and/or pyrrhotite and therefore allows an estimation of mineralization temperatures using the geothermometer (Kretschmar and Scott, 1976; Sharp et al., 1985). The lack of buffer assemblages in equilibrium with arsenopyrite 2 precludes its use as a geother- toometer. All arsenopyrite 1 grains were tested for compositional zoning and core-rim pairs that showed > 1 at. percent As variation were rejected for geothermometry (Fig. 6). In addition, combined minor element contents of the arsenopyrite used for temperature calculations are <1 wt percent (Table 1) as recommended by Kretschmar and Scott (1976). Unzoned arsenopyrite 1 does not depart markedly from the

stoiehiometrie ratio Fe/(As + S) = 0.5 and direct application of the arsenopyrite geothermometer is, therefore, possible.

Twenty-four analyses (from eight ore samples) of arsenopyrite 1 grains in contact with loellingite and pyrrhotite show a range and mean of 34.1 to 36.0 and 35.1 _ 0.5 at. percent arsenic, respectively (1 a, n = 24). This corresponds to a range of palcotemperatures of 480 ø to 560øC with a mean of 510 ø _+ 20øC (1 a, n = 24; Fig. 7) using the Sharp et al. (1985) version of the geothermometer. An about 100øC lower mean temperature is obtained if the temperature-atomic percent As in the arsenopyrite diagram (Kretsehmar and Scott, 1976) is applied. In this paper we are using the higher temperature estimate because it is in general agreement with independent mineralization temperatures (garnet-biotite geothermometry) obtained by Dalstra (1995) and Dalstra et al. (in press), and textural observations presented in this con- tribution (see above) and Dalstra et al. (in press) indicate that crystallization of the prograde alteration minerals, including arsenopyrite 1, at Transvaal was broadly contemporaneous with that of regional peak metamorphic minerals. Since the closure temperature for reequilibration of arsenopyrite is between 500 ø and 650øC (Sharp et al., 1985), it is possible that the mean arsenopyrite composition is not representative of equilibration at peak conditions. In this respect, it is noted that the highest As contents indicate temperatures of around 500øC, using the atomic percent As-T diagram of Kretschmar and Scott (1976). For the equilibrium assemblage arsenopyrite 1-1oellingite-pyrrhotite, the sulfur fugacity can be determined from the arsenic content of the arsenopyrite 1 and is approximated to be 10 -6'3 to 10 -s 5 bars with a mean of 10 -7'5 _+ 0.7 bars (1 a, n = 24; Fig. 7).

Pyrrhotite and sulfur activity Using the electrum tarnish method, Toulmin and Barton

(1964) determined that the Fe content of hexagonal pyrrho- tire is a function of changing sulfur fugacity and temperature in the pyrite-pyrrhotite system. Pressure effects on the composition of pyrrhotite are small (Toulmin and Barton, 1964) and, therefore, no corrections for pressure are necessary. However, pyrrhotite is known to easily reequilibrate during cooling (Barton and Skinner, 1979) and, thus, the analyses might not record the original composition during crystallization. Assuming a mineralization temperature of 510øC, based on arsenop rite eothermomet rrhotite records a ran e Y 52 g, -0 ry' py g offs2 from 10- ' to 10-" bars (1 a, n = 7; Fig. 7). Assuming the same mineralization temperature, the two Fe-rich pyrrho-

118 13 8

tite analyses show substantially lowerfs2 of 10- ' and 10- ß bars (Fig. 7).

Interpretation of sulfide paragenesis The sulfide minerals in the hydrothermal alteration zones

at Transvaal are interpreted to consist of a mixture of equilibrium assemblages that relate to the prograde alteration event and disequilibrium assemblages that formed during cooling and reequilibration of the rocks. The dominant prograde sulfide minerals loellingite and arsenopyrite 1 are almost always spatially associated with gold-bearing lodes and specifically with the mineral electrum (see below). Rarely, coexisting chalcopyrite-arsenopyrite 1 is observed in the proximal alteration zone. Equilibrium textures such as the intergrowth of

280 HAGEMANN ET AL.

• .


+1 +1 +1 +1 +1 +l +l

+1 +1 +1 +1 +1

+1 +1 +1 • +1 +1 +1 +1 +1

+1 +1 +1 +1 +1 +1 +1 +1

• .

c•

+1 +1 +1 +l +1 +1 +1 +1

9,82 HAGEMANN ET AL.

Ai Unzoned Arsenopyrltes

al Zoned Arsenopyrltes

31 32 33 34 35 36 37 31 32 33 34 35 36 37 Atomic % As Atomic % As

FIG. 6. Diagrams displaying the range of atomic percent As in (A) unzoned and (B) zoned arsenopyrites of type 1 and 2. Points on a tie line represent analyses from a single grain of arsenopyrite.

arsenopyrite and pyrrhotite with plagioclase, biotite, or hornblende indicate that sulfide and silicate minerals formed eon-

temporaneously (el. Knight et al., 1993; MeCuaig et al., 1993; Neumayr et. al., 1995).

Based on the textural relationships outlined above, we pro- pose that the main phase of metasomatism which included the deposition of prograde silicates, pyrrhotite, and loellingite occurred synteetonieally with respect to the Transvaal D3 shear zone at peak metamorphic conditions (el. Dalstra et al., in press). The most pertinent textural and mineral chemistry criteria, defining the timing of the sulfide assemblage at Transvaal are:

1. The synkinematie sulfarsenide assemblage occurs in textural equilibrium with (prograde) synkinematie metasomatie minerals of the proximal alteration zone such as biotite, diopside, garnet, plagioelase, and quartz. Evidence for this in- eludes intergrowth textures of pyrrhotite, amphibole, and diopside in the proximal alteration zone, similar grain size of sulfides and silicate minerals, sulfide and silicate textures are part of the shear zone fabric (i.e., aligned subparallel to the foliation fabric), and metasomatie diopside locally encloses sulfides and vice versa.

2. Geothermometry of early-stage, synkinematie arsenopyrite 1 reveals temperatures of about 510 ø + 20øC, which are broadly comparable, albeit slightly lower (40øC), than the peak metamorphic temperature conditions (550 ø + 25øC) obtained by Dalstra et al. (in press).

Of particular significance is the common presence of the reaction assemblage arsenopyrite-loellingite-pyrrhotite and the typical rimming of single and multiple Ioellingite cores by arsenopyrite 1, which in turn is rimmed by pyrrhotite. Importantly, the reverse is not observed. This texture, combined with the embayed grain boundaries, clearly indicates that arsenopyrite 1 partially replaced Ioellingite. The same assemblage and similar reaction textures has also been recorded at a number of other hypozonal lode gold deposits in Western Australia (Marvel Loch, Mueller, 1991; Frasers and Griffins Find, Barnieoat et al., 1991; Mount York, Neumayr

et al., 1993b; Norseman, McCuaig et al., 1993), as well as in tin-tungsten deposits of the eastern Andes in Bolivia (Kelly and Turneaure, 1970) and in historic gold deposits in the Oberpfalzer Wald, Germany (Grundmann et al., 1985). Dal- stra et al. (in press) and Neumayr et al. (1993a) discussed three possible processes that could have resulted in the p?- rhotite-loellingite-arsenopyrite assemblage: (1) simultaneous crystallization of pyrrhotite and loellingite and a subsequent retrograde solid-solid reaction to produce arsenopyrite, (2) crystallization of arsenopyrite and a subsequent prograde solid-solid reaction to loellingite and pyrrhotite, and/or (3) changing fluid chemistry. For example wall-rock sulfidation and a resulting decrease in the H2S content of the fluid could potentially explain the paragenetic relationships and rimming of loellingite by arsenopyrite. Alternatively, crystallization of Ioellingite from a fluid with high As activity could have been followed by subsequent crystallization of arsenopyrite and pyrrhotite from fluids with lower As activity. These authors favor (1) based on textural evidence such as the rimming of Ioellingite by arsenopyrite, resoption textures of Ioellingite, and the occurrence of pyrrhotite and loellingite as syntectonic crystals but not arsenopyrite.

The replacement of loellingite by arsenopyrite i at Transvaal is also interpreted to be the result of a solid-solid retrogressive reaction with decreasing temperatures, in which

1000/T(K-1) -2 1.7 1.6 1.5 1.4 1.3 1.2 1.1 1.0

• 47•- -3• IPyrrhotite atsbility • range / .,.'•

-$

'• •/ ••1 • i•. •rrhotlte •

"•/" •. '1 49.5 I

•00 4oo s00 •00 •00

T•p•r•t•r• (o•)

Fro. 7. Sulfur aetM•-temperature projection of the stabili• field of arsenopy•te 1. Thin solid contours are atomic percent arsenic. Modified from Kretsehmer and Scott (197B) and Sha• et g. (1985). Also sho• is the logfs, stabili• range (ve•ieal hatched bar at the Y •s) and temperature range of arsenopy•tes 1 (ve•ieal hatched bar at the bottom of the X •s) for the Transvaal lode gold deposit. Supe•mposed are isopleths of pyrrhotite e•ressed in atomic percent Fe (dashed lines) as a fune•on of S2 aetMty and temperature after Ba•on and S•nner (1979). Unee•n• in p•rhotite isopleth is approximately z 0.35 in log f s,. •lso sho• are log f s, stabili• ranges for p•rhotite and Fe-•eh p•rhotite based on microprobe data (exam- ples sho• in Table 1).

ARCHEAN HYPOZONAL TRANSVAAL LODE GOLD DEPOSIT, IV. AUSTRALIA 283

loellingite reacted with pyrrhotite to form arsenopyrite i and a more iron rich pyrrhotite (el. Neumayr et al., 1993a):

FeAs.2 + (1/x)FeSl• + xl = 2FeAsS + ([1 - x]/x)FeS, (1)

or approximately for the pyrrhotite composition at Transvaal:

FeAs.2 + 10FeS•.• = 2FeAsS + 9FeS. (2)

The occurrence of iron-rich pyrrhotite (about 2 at. % more Fe, see Table 1) within the proximal alteration zone at Transvaal may be a product of this reaction. The slightly retrograde, postpeak metamorphic nature of this process is supported by the somewhat lower arsenopyrite 1 temperatures of 510 ø +_ 20øC when compared to the metamorphic temperature conditions of about 550øC (Dalstra et al., in press).

The coexistence of ehaleopyrite-pyrrhotite, pyrrhotite-pyrite, pyrite-loellingite, pyrrhotite-eubanite, and low and high Fe pyrrhotite cannot have been stable assemblages at near- peak metamorphic conditions. Chaleopyrite and pyrrhotite are generally intergrown and in apparent textural equilibrium with high-temperature silicate phases, such as amphibole, diopside, and garnet. However, experiments on sulfide stabilities (Craig and Scott, 1976) suggest that chaleopyrite is not stable in coexistence with pyrrhotite above 350øC. Pyrrhotite and euhedral pyrite are locally observed in apparent equilibrium, but investigations on metamorphism of pyrite and pyritie ores led Craig and Vokes (1993) to argue that during retrograde cooling and reerystallization of ores, pyrrhotite releases sulfur that could be absorbed by pyrite to initiate retrograde growth of euhedral pyrite crystals. Furthermore, pyrite-loellingite assemblages should not be stable based on experimental investigations in the Fe-As-S system (Kret- schmar and Scott, 1976). Cubanitc that only occurs together with pyrrhotite is not stable at the prevailing mineralization temperatures of 510 ø +_ 20øC at Transvaal and, therefore, is interpreted to have formed during the cooling and reequilibration of the altered rocks below about 210øC (Cabri, 1973; Craig and Scott, 1976). Pyrrhotite in the Fe-S system occurs over a wide range of T-XFe space (Craig and Scott, 1976) and is characterized by its relative ease of reequilibration down temperature. The rare high Fe pyrrhotite may represent primary, high-temperature hexagonal pyrrhotite. Arsenopyrite 2 is, based on textural and compositional data, interpreted to be retrograde in nature.

Occurrence, Petrology, and Chemistry of Gold Electrum and gold have been observed at the Transvaal

gold deposit, but electrum (Ramdohr, 1980) is by far the most abundant gold-bearing phase and contains between 24 and 57 wt percent Ag (Fig. 8). Gold has only been observed in one section (two grains) and contains a minor amount (5 wt %) of Ag and, thus, is defined as gold sensu stricto. The compositions of electrum grains correspond to gold finenesses of 431 to 732, whereas gold sensu stricto has a fineness of 947. In all cases, Ag is the dominant impurity (Table 2). Metallurgical data indicate that about 15 percent of the gold at the deposit is genuinely refractory (James, 1995). By analogy with secondary ion mass spectrometry analyses on gold and loellingite from other deposits in the Yilgarn and Pilbara

blocks (Neumayr et al., 1993a), this gold is likely to be distributed as invisible Au in loellingite. Dalstra (1995) detected trace gold within loellingite by electron microprobe analysis.

Electrum in the Transvaal deposit is located in the proximal alteration zones of both ultramarie and metasedimentary rocks. It is observed as grains of various sizes coexisting with the following sulfides regardless of host-rock type (Fig. 9): (1) as grains (<100 pro) at the contact between arsenopyrite 1 and loellingite (Fig. 9A, B), (2) as small inclusions (<50 pro) in loellingite (Fig. 9A), (3) as grains (<600 pm) coexisting with pyrrhotite, arsenopyrite 1, and loellingite-rammels- bergitc (Fig. 9C, D), and (4) as fracture fill within arsenopyrite 1 (Fig. 9B). One ease of small (<50 pm) grains of free gold in a quartz vein is documented.

Electrum-loellingite-pyrrhotite-arsenopyrite 1 mineralization

Electrum is most abundantly observed at the arsenopyrite 1-1oellingite interface and within loellingite. Due to its similar reflectance to arsenopyrite, the exact extent of loellingite has been determined by X-ray mapping. In addition, these X- ray maps have been used to determine the precise areas of electrum occurrence (Fig. 10). Figure 10 depicts that in this setting, electrum is located at the arsenopyrite 1-1oellingite interface and within loellingite where there is textural evidence of rimming and replacement of loellingite by arsenopyrite 1. In these settings, it is absent in arsenopyrite and pyrrhotite.

A similar siting of gold has been described at other deposits in Western Australia by Barnieoat et al. (1991) and Neumayr et al. (1993a). Neumayr et al. (1993a) in an innovative study on the distribution and timing of gold in the amphibolite and lower granulite-hosted Mount York and Griffins Find deposits, respectively, used detailed secondary ion mass spee- trometry to show that the majority of gold (apart from free gold) occurs as invisible Au in loellingite, whereas arsenopyrite is essentially gold free. They conclude that the lack of gold in arsenopyrite is probably due to the inability of the arsenopyrite structure to accommodate the Au at high, amphibolite-granulite facies temperatures. Dalstra et al. (in press) confirm the earlier conclusions of Barnicoat et al. (1991) and Neumayr et al. (1993a) that, at high temperatures, the formation of native gold at loellingite-arsenopyrite grain boundaries resulted from the inability of arsenopyrite to incorporate Au in its structure, and that a retrogressive reaction of loellingite to arsenopyrite as discussed above was responsible for the formation of native gold but not for the initial precipitation of Au in the loellingite structure.

This study supports the interpretations of Dalstra et al. (in press) but in addition shows electrum of variable composition as the main gold-bearing phases. Electrum mostly coexists with synkinematic composite grains of loellingite-arsenopyrite 1 and more rarely with discrete loellingite, arsenopyrite 1, pyrrhotite, and gersdorffite. Electrum that is in equilibrium with pyrrhotite and/or loellingite was likely deposited metaso- matically at slightly postpeak conditions. Electrum in this case was liberated from the loellingite lattice during the replacement of loellingite by arsenopyrite 1 and preferentially rede- posited at loellingite-arsenopyrite 1 grain boundaries.

As loellingite and arsenopyrite 1 are interpreted to be part

284 HAGEMANN ET AL.

Low Ag (<28 wt% Ag) Electrum

Sample 69a2: eleopo-a•l

Sample69A1:•1

Sample 71B: elseaWl

Gold

[] [] [] High Ag (>35 wt% Ag) [] Electrum

• [] Sample 71A [] [] Sample 71B []

[] Sample 71B []

• [] Sample71 B [] [] [] Sample 121595 []

[] Sample'IV 103 [][]l•J I•[][] [] (elec-qtz) [][][] I•Jl• I•1 [] [] [][][]K'I [] EI[]r•Et []

I ......... I ......... I ......... I ......... I ......... I ......... I

0 includes 58 values 10 (int =1) 20 30 40 50 60

wt. % Ag FIG. 8. Histogram displaying the silver content (wt %) of electrum and gold grains in different samples. Also shown

are the different electrum, sulfide, and quartz assemblages. Analyses from sample 121595 were obtained by Dalstra (1995). Abbreviations: apy = arsenopyrite, elec= electrum, po = pyrrhotite, qtz = quartz.

of the prograde alteration or formed during high-temperature retrogression, the presence of gold within loellingite, and together with arsenopyrite 1, implies that it was introduced broadly synchronous with the peak of metamorphism or under slightly retrograde metamorphic conditions (cf. Barnicoat et al., 1991).

State of gold--some speculations The formation of free electrum at loellingite-arsenopyrite

interfaces at Transvaal likely resulted from the inability of arsenopyrite 1 to incorporate electrum in its lattice at these high temperatures (cf. Neumayr et al., 1993a; Dalstra et al. in press). Moeller and Kersten (1994) show with galvanic cell arrangements that As-rich arsenopyrite can represent a p- type conductor (electron acceptor) and that gold is electro- chemically accumulated on arsenopyrite, whereas loellingite (n-type conductor--electron donor) is anodieally dissolved. However, Mumin et al. (1994) demonstrated via secondary ion mass spectrometry analyses for the mesothermal Bogosu- Prestea gold system (Ashanti gold belt) in Ghana that invisible Au can occur in arsenian pyrite and arsenopyrite. Based on the Au distribution patterns, they concluded that primary Au was apparently precipitated in solid solution with pyrite and arsenopyrite. The question whether invisible gold in loellingite or arsenopyrite at Transvaal is present as particulate (col- loidal-size) submicroscopic gold and/or geochemically bound in the structure of loellingite and arsenopyrite remains open and awaits further detailed secondary ion mass spectrometry, high-resolution transmission electron microscopy (Cabri et al., 1989), and Moessbauer analyses (Wagner et al., 1986; Cathelineau et al., 1989) on the gold-bearing minerals at Transvaal.

Thermodynamic constraints of electrum deposition at Transvaal

Recent studies (e.g., Morrison et al., 1991; Huston et al., 1992; Gammons and Williams-Jones, 1995) have emphasized the significance of Au-Ag alloys (e.g., electrum) in ore-form- ing systems and have established a geochemical data base for the system. Due to the lack of thermodynamic data at higher temperatures (>400øC), quantitative modeling of electrum paragenesis at Transvaal awaits further experiments on the hydrothermal geochemistry (e.g., solubilities) of electrum at high temperatures, however, the available data may be used to interpret qualitatively the processes of electrum transport, deposition, and parageneses at the Transvaal deposit. Al- though electrum grains along loellingite-arsenopyrite interfaces are secondary, and the Au in them was likely deposited initially as invisible Au in loellingite, electrum in other sites is interpreted to have precipitated directly from the hydrothermal fluid. In any one sample, electrum has the same composition, regardless of mineralogical setting. The deposition of Au and Ag is thus discussed with respect to the stability of electrum.

At low temperatures (<350øC) one of the most important phases in the Au-Ag-S-C1-H20 system is argentitc (AgsS), even though the formation of auriferous argentitc is restricted to conditions of relatively highf s2 (Gammons and Williams- Jones, 1995). In the presence of AgsS the composition of the coexisting Au-Ag alloy is constrained by the following reaction:

Ag(auoy) + 1/4S2(gas)= 1/2AgsS(solid). (3)

This reaction forms the basis of the electrum tarnish method


TABLE 2. Electrum-Sulfide + Quartz Assemblages, Au/Ag Content, and Fineness of Electrum

Ag Au Sample no. Assemblage (wt %) (wt %) Fineness Figure Comments

Electrum (high Ag)

71B-3-1 apy 39.18 61.26 610 anhedral grain within apy 71B-3-2 apy + lo 37.72 61.68 621 euh grain in contact w lo + apy 71B-l.1 apy + lo 38.24 60.98 615 6A grain at Io/apy contact 71B-1.2 apy + Io 38.86 61.60 613 6A grain at lo/apy contact 71B-1.3 apy + Io 38.97 61.79 613 6A grain at Io/apy contact 71B-2.1 apy + lo 39.92 60.22 601 6A grain at lo/apy contact 71B-2.2 apy + Io 40.35 60.00 598 6A grain at lo/apy contact 71B-3.1 apy + lo 39.70 60.54 604 6A grain at Io/apy contact 71B-3.2 apy + Io 40.15 60.50 601 6A grain at lo/apy contact 71B-4.1 apy + po + qtz 38.30 61.29 615 6A grain w apy + po surf by qtz 71B-4.2 apy + po + qtz 38.96 61.39 612 6A grain w apy + po surf by qtz 71A-1-1.1 qtz (apy + Io + po) 38.33 61.62 617 6C several <25/•m grains w qtz 71A-1-1.2 qtz (apy + lo + po) 38.52 61.98 617 6C several <25/•m grains w qtz 71A( 1 ) qtz 40.41 58.43 591 71A(4) qtz 39.23 61.21 609

Electrum (low Ag)

69A2-1-1 apy + po 26.47 74.18 737 6B mass grain contact w apy + po 69A2-1-2.1 apy + po 25.44 75.28 747 6B mass grain contact w apy + po 69A2-1-2.2 apy + po 25.85 74.22 742 6B mass grain contact w apy + po 69A2-1-3.1 apy + po (lo) 26.95 73.73 732 6B surr apy and cont w po 69A2-1-3.2 apy + po (1o) 26.84 73.72 733 6B surr apy and cont w po 69A2-1-3.3 apy + po (1o) 26.72 73.72 734 6B surr apy and cont w po 69A2-1-3.4 apy + po (lo) 26.47 74.03 737 6B surr apy and cont w po 69A2-2-4.1 apy + po (lo) 26.59 73.27 734 6B surr apy and cont w po 69A2-2-4.2 apy + po (lo) 26.76 73.83 734 6B surf apy and cont w po 69A2-2-4.3 apy + po (1o) 26.92 73.39 732 6B surr apy and cont w po 69A2-2-4.4 apy + po (lo) 26.65 73.62 734 6B surr apy and cont w po 69A2-2-4.5 apy + po (1o) 26.97 73.77 732 6B surr apy and cont w po 6932-2-4.6 apy + po (1o) 26.42 73.22 735 6B surr apy and cont w po 69A2-2-4.7 apy + po (lo) 26.74 73.79 734 6B surr apy and cont w po 69A2-1-4.1 apy 26.35 74.43 739 6B fracture in apy 69A2-2-5.2 qtz (po) 25.92 73.99 741 6B euh pseudore. apy in qtz vein 69A2-2-5.1 qtz (po) 25.89 74.13 741 6B euh pseudore. apy in qtz vein 69a2-1-1.1 po + apy 25.18 75.48 750 grain surf by po + apy 69a2-1-1.2 po + apy 25.04 75.35 751 grain surf by po + apy 69a2-1-1.3 po + apy 24.94 74.29 749 grain surr by po + apy 69a2-2-2.1 qtz + po 27.00 73.22 731 grain surr by po + qtz 69a2-2-2.2 qtz + po 27.02 73.45 731 grain surr by po + qtz 69a2-2-3.1 apy + po + qtz 26.74 73.93 734 grain surr by qtz + po-apy 69a2-2-3.2 apy + po + qtz 26.31 73.64 737 grain surr by qtz + po-apy 69a2-2-4 apy + po 25.83 72.66 738 surf by apy + po within qtz mat 69A1-2-1.1 apy + po 26.53 73.10 734 within po and <20/•m apy 69A1-2-1.2 apy + po 26.56 73.28 734 within po and <20/•m apy 69A1-2-1.3 apy + po 26.49 74.45 738 within po and <20/•m apy 69A1-2-1.1 po + qtz 26.45 73.28 735 part surr by qtz within po 69A1-2-1.2 po + qtz 26.08 74.45 741 part surr by qtz within po 69A1-1.1 apy + lo (po) 27.08 73.39 730 69A1-1.1 apy + lo (po) 26.90 73.23 731 69A1-1.1 apy + lo (po) 26.86 73.66 733 69A1-1.1 apy + Io (po) 26.63 73.58 734 69A1-1.1 apy + Io (po) 26.51 73.50 735 71B-2-1 apy 25.10 71.07 739 incl in apy

Gold

103-1-1.1 qtz 5.04 94.16 949 grains within qtz vein 103-1-1.2 qtz 5.09 94.69 949 grains within qtz vein 103-1-1.3 qtz 5.04 94.09 949 grains within qtz vein 103-1-1.4 qtz 5.26 94.83 947 grains within qtz vein

Analyses are from 23 grains of electrum from 5 different samples, all samples occur in the proximal alteration zones (possibly except 103) within ultramarie host rocks The following electrum-sulfide + quartz assemblages are observed: (1) apy, (2) po, (3) qtz, (4) lo + po, (5) apy + lo, (6) apy-po-lo, (7) apy + po + qtz, (8) po + qtz, (9) apy

+ Io + po-qtz, and (10) Au gersdorffite: for the latter assemblage no microprobe analyses are available due to the small size of the gold grain; high Ag electrum is uniquely characterized by assemblages 2, 5, 6, and 8, whereas low Ag electrum is uniquely characterized by assemblage 9; both high and low Ag electrum are observed in assemblages 1, 4, and 7; all arsenopyrite is of type 1. Abbreviations: apy = arsenopyrite, cont = contact, eu = euhedral, incl = inclusion, Io = loellingite, mat = matrix, po = pyrrhotite, pseudore. = pseudomorph, qtz = quartz, surr = surrounded, w = with

286 HAGEMANN ET AL.

c

Otz

Otz •;•,

F]c. 9. Backscattered electron microscopy images of equilibrium assemblages of electrum. A. Electrum-loellingite- arsenopyrite 1-pyrrhotite assemblage (Elec, Loe, Apy, Po, respectively). Sample TV 71B. Note that this arsenopyrite, as in all other images of this figure, is of type 1. B. Electrum partially surrounding arsenopyrite I and loellingite. Sample TV 69A2. C. Electrum in contact with quartz (Qtz), arsenopyrite I and pyrrhotite but not with chalcopyrite (Cpy). Sample TV 7lB. D. Electrum grain sandwiched between loellingite and loellingite-rammelsbergite (L-R). Sample TV 103-86B.

which Barton and Toulmin (1964) used to determine the fugacity of sulfur in the argentitc-electrum system. In Figure 11 low and high Ag electrum from the Transvaal gold deposit are shown in fs2-T space. Due to the lack of argentitc at Transvaal the determined sulfur fugacities from electrum represent maximum values only. At mineralization temperatures of 510 ø ___ 20øC, obtained by arsenopyrite geothermom- et , the deduced maximumfs2 for low and hi A electrum ß ry -31 -40 -40 --53 gh . g is 10 ' to 10 - and 10 ' to 10 ' bars, respectively. These fs• values are higher than the fs• values determined by the arsenic content of arsenopyrite 1.

The sulfide contents of the hydrothermal fluid at Transvaal are constrained by the above equilibria and the sulfide assemblages discussed earlier. There are as yet no fluid inclusion data to constrain salinites of the fluid, but it is assumed that they were in the same range as those recorded in nearby deposits that are similar except for their host-rock setting, and hence low (2-8 equiv wt % NaC1, e.g., Bloem and Brown 1991). For these constraints on fluid composition, Ridley et al. (1996) have argued that Au bisulfide complexes would

dominate over Au chloride complexes at 500 ø to 550øC, by two to three orders of magnitude. There is less data for Ag complexes, however, assuming that the most likely Au and Ag complexes in solution are AuCI•, Au(HS)•, AgCI•, and Ag(HS)•, the relative importance of chloride and sulfide spe- cies for Au and for Ag transport can be estimated by consider- ing the exchange reaction:

AuCI• + Ag(HS)• = AgCl• + Au(HS)•. (4)

From the thermodynamic data for the dissolution reactions of gold and silver given by Gammons and wfiliams-Jones (1995), log K (4) - 5.85 at 300øC and can be estimated as between 3 and 5 for the higher temperatures of interest here. Given this value of K (4), the ratio of Au bisulfide to Au chloride complexes should therefore be at least three orders of magnitude higher than the ratio of Ag bisulfide to Ag chloride complexes, whatever the chemistry of the solution. It is, therefore, considered likely that Ag was dominantly transported as a chloride complex, or that Ag bisulfide and Ag chloride complexes were of simfiar concentration.

ARCHEAY HYPOZONAL TRANSVAAL LODE GOLD DEPOSIT, IV. AUSTRALIA 287

IA

200 un •uxr BbE 25 k¾ "0 nA 200 ;m Auxl S ka 25 kV 20 nA

c D

200 un •uxr ½.- ka .'$ k¾ 20 nA 200 un Auxr Au ka 25 k¾ 20 a a,

F•c. 10. A. Color-enhanced backscattered electron image of a typical electrum-loellingite-arsenopyrite 1-pyrrhotite assemblage (sample 25' 69A2; same as Fig. 9A). Color-enhanced backscattered electron image of the same assemblage, showing relative concentration and distribution of sulfur (B), arsenic (C), and gold (D).

At Transvaal there is a very variable composition of electrum through the deposit but a generally constant composition within an), sample (Fig. 8), even in samples in which individual grains of gold have contrasting mineralogical settings. If Au were in solution dominantly as bisulfide complexes and Ag as chloride complexes, differing fluid wall-rock reactions are a potential explanation of this variable electrum

composition in the deposit. It is noted that the occurrence of low Ag electrum is farthest away from the graphitic pelite host rock, whereas the high Ag electrum is generally where the lode is along the ultramafic-pelitic contact. Given the thermodynamic constraints outlined above, and the fact that the graphite-bearing pelite is a much more reducing rock than the ultramafic rocks, it appears feasible that the Au-

288 HAGEMANN ET AL.

-14 300 400 500 600 700 800 900

Temperature (øC)

FIG. 11. Modified diagram from Barton and Touhnin (1964) showing the Nag range of electrum at the Transvaal lode gold deposit at the temperature of mineralization (510 ø _+ 20øC deduced from arsenopyrite geothermometry; see Fig. 7). The two boxes along the y axis indicate the range of maximum logfs2 obtained for low and high Ag electrum (10 -3 ]-10 -40 bars) and high Ag electrum (10-4ø-10 -s3 bars).

bearing hydrothermal fluid which reacted with reducing pel- ires produced Ag-rich electrum while fluid that reacted with more oxidized ultramafic rocks formed Au-rich electrum. A

dramatic change infs• orfo_• of the ore fluid during the evolution of the palcohydrothermal system could equally explain the different Ag/Au ratio. However, no mineral or thermodynamic evidence of a significant change with time of the ore fluid composition has been observed and the latter explanation could also not easily account for the spatial distribution of low and high Ag electrum (Figs. 2 and 8).

Summary

In summary, minor amounts of electrum have been observed coexisting with pyrrhotite and loellingite and are interpreted as the product of prograde, peak metamorphic and syntectonic processes. The majority of electrum, however, is detected coexisting with arsenopyrite 1, or composite arsenopyrite 1-1oellingite grains and interpreted to have formed during early retrograde, slightly postpeak metamorphic processes, most likely by release of submicroscopic gold from loellingite. However, textural evidence for synkinematic electrum-arsenopyrite 1 and electrum-arsenopyrite 1-1oellingite assemblages indicate that the introduction of the Ag and Au occurred during the main, prograde hydrothermal mineralization event.

Postmetamorphic Sulfide and Silicate Minerals Petrographic and microprobe analyses of the metasomatic

silicate and sulfide minerals at the Transvaal gold deposit have produced evidence for retrograde, postmetamorphic formation of many sulfide and metasomatic silicate minerals. This evidence includes postkinematic overprinting and replacement textures, growth of low-temperature silicate minerals, and thermal stability limits of some observed sulfide assemblages.

Postkinematic textures are documented by the occurrence of arsenopyrite 2 that overprints the existing sulfide minerals and fabric (Fig. 5A). Their c axes are oriented randomly, commonly oblique to the existing shear zone fabric. Other typical retrograde replacement textures are chlorite pseudo- morphs after biotite (Fig. 4E), chlorite that is oblique to the shear zone fabric (Fig. 4F), and chlorite _+ zoisite and muscovite replacement of plagioclase (Fig. 4B, C). Low-temperature silicate minerals such as pumpellyite fill some voids and fractures. Pyrrhotite displays a wide range in X•e (Fig. 7) and, given the ease with which pyrrhotite reequilibrates (Craig and Scott, 1976), these compositional variations are interpreted to be the effect of retrograde low-temperature reequilibration. Cubanitc has a thermal stability limit of 210øC (Cabri et al., 1973) that is incompatible with the deduced mineralization temperature of 510øC at Transvaal. Co- existing cubanitc and pyrrhotite thus provides evidence for reequilibration at lower temperatures. The retrogression and reequilibration of silicate and sulfide minerals is not associated with a penetrative deformation fabric and hence is clearly postkinematic with respect to the shear zone fabric at Transvaal, but small fractures and veinlets are locally filled with carbonate, chlorite, zoisite, and muscovite.

The formation of retrograde silicate and sulfide minerals within the distal and proximal alteration zones may be the result of one or several of the following events or processes: (1) deposition during the waning stages (minor pulses) of the main, gold-depositing phase of hydrothermal activity but after most of the ductile deformation took place, (2) an unrelated, later stage hydrothermal event, possibly associated with renewed tectonic activity (i.e., uplift and/or igneous activity), and (3) recrystallization and remobilization of existing mineral assemblages during the cooling of the rocks. Hydrothermal histories involving pulses of fluid of significantly different chemistry have recently been shown by Vennemann et al. (1993) for an epithermal gold deposit in the Dominican Re- public and by Hagemann (1993) for the epizonal Wiluna lode gold deposit in Western Australia. By analogy to these studies, a late-stage pulse of hydrothermal fluid at Transvaal could have caused the deposition of the arsenopyrite 2. However, fluid flow and shear zone deformation are generally considered to occur contemporaneously (Sibson, 1990), i.e., high fluid pressures are considered the main driving force for the deformation along the shear zone, hence arsenopyrite that crystallized from a waning pulse of fluid should also be texturally synkinematic to the shear zone, except if late movement was along a narrow zone within the shear zone. In the case of the retrograde silicate minerals, the substantial temperature difference between the peak metasomatic minerals at 510øC (e.g., arsenopyrite 1, diopside) and the retrograde, postkinematic silicate minerals (e.g., muscovite, pumpellyite) which are most likely to have formed at less than 350øC (cf. Mueller et al., 1991) also suggests that they were not the result of a single hydrothermal event. More detailed studies (e.g., fluid inclusions and stable isotopes) on silicate and sulfides are necessary to further elucidate the possibility of multiple pulses of hydrothermal activity.

An unrelated, later stage hydrothermal event associated with renewed tectonic activity would be consistent with re- sults from a metamorphic study of the Transvaal area (Dalstra

ARCHEAN HYPOZONAL TRANSVAAL LODE GOLD DEPOSIT, W. AUSTRALIA 9,89

et al., in press) that proposes a rather protracted metamorphic history for the greenstone belt. Using Ar-Ar dating, Napier (1993) concluded that the terrane spent 50 m.y. at temperatures above 300øC following the peak of metamorphism. Con- tinued metamorphic activity would likely be accompanied by nonseismic fluid flow, and thus deposition of minerals, after the main gold-depositing hydrothermal event ceased. The retrogression to lower temperature minerals, such as zoisite, ehlorite, or pumpellyite, has also been observed outside the immediate mine area (Dalstra et al. in press) and is believed to be related to hydrothermal activity at a greenstone belt scale, probably during uplift with possible incursion of surface waters. This process could explain the formation of low-temperature silicate minerals. However, arsenopyrite 2 occurs without coexisting low-temperature sulfide and silicate minerals and has not been observed outside the shear zone. There-

fore the greenstone belt wide retrogression did not involve significant arsenic and sulfur metasomatism. However, renewed hydrothermal fluid activity that affected the Transvaal shear zone could well have been associated with remobiliza-

tion of arsenic and sulfur and thus explain the restriction of arsenopyrite 2 to the shear zone.

Reerystallization and remobilization of existing mineral assemblages during cooling is favored for the formation of par- tieular sulfides such as cubanitc, pyrite, ehaleopyrite, and low Fe pyrrhotite. Experimental data by Craig and Scott (1976) and Barnes (1979) and studies on metamorphosed strata- bound base metal sulfide deposits (Vokes and Craig, 1993) support the possibility that reerystallization of these sulfide minerals can occur during cooling. The lack of equilibrium assemblages between cubanire, low Fe pyrrhotite, and retrograde low-temperature silicates and arsenopyrite 2, however, restricts the application of this process to the formation of cubanitc, pyrite, ehaleopyrite, and low Fe pyrrhotite only.

In summary, the formation of retrograde, postmetamorphic minerals is evidence for processes that occurred after the main, gold-bearing hydrothermal activity. A likely mechanism that could have produced the retrogression of the silicate

minerals is the influx of late hydrothermal fluids during the protracted metamorphic history of the terrane. The occur- renee of low-temperature sulfides is likely the result of in situ reequilibration during cooling. The cause of formation of arsenopyrite 2 is equivocal but could be the result of a late fluid flux at the end of the main gold-bearing hydrothermal event or related to a later stage influx of hydrothermal fluids.

Conclusions

This study is the first description of electrum as the major gold-bearing phase in an Arehean lode gold deposit in West- ern Australia and possibly, besides the epizonal Ross mine in Ontario (Akande, 1985), the first worldwide (cf. Morrison et al., 1991). Even more remarkably, electrum in the Transvaal deposit occurs in a high-temperature, hypozonal gold deposit rather than in the extensively documented low-temperature epithermal environment (e.g., Japan and Korea; Shelton et al., 1990). The low fineness of 431 to 732 for electrum (Ag the main impurity) in the Transvaal deposit sets it apart from classic mesozonal and hypozonal Archcan lode gold deposits, where gold fineness generally ranges from 780 to 960 (Fig. 12; Ridley et al., 1995).

Ore petrography, mineral chemistry, and X-ray mapping of ore samples reveal that electrum coexists with loellingite, arsenopyrite, pyrrhotite, gersdorffite, and quartz. The most commonly observed equilibrium texture, however, is of electrum with composite loellingite-arsenopyrite 1 grains which, in many cases, are rimmed by pyrrhotite. The replacement ofloellingite by arsenopyrite 1 is explained by a high-temperature retrograde reaction which was proposed for other hypozonal lode gold deposits in the Yilgarn block and Pilbara where loellingite replacement by arsenopyrite is observed (cf. Barnicoat et al., 1991; Neumayr et al., 1993b).

Combined textural and mineral chemistry analyses of metasomatic silicates, sulfides, and electrum suggest that the main phase of metasomatism occurred syntectonically with respect to the Transvaal shear zone and was broadly contemporaneous with the amphibolite facies metamorphic event that af-

Gold Fineness

Deposit 4o0 soo •00 700 I I i

Zakanaka

Transvaal ß

Hopes Hill Corlnthia

Mystery Hill Bayleys Kings Cross Lindsays

Three Mile Hill Tindais

Wiluna lode-gold

FIG. 12. Range and mean of gold fineness at the Transvaal lode gold deposit compared to the epizonal Wiluna lode gold deposit and other high4emperature (hypozonal) gold deposits in the Yilgarn block of Western Australia. Modified after Ridley et al. (1995).

290 HAGEMANN ET AL.

feeted the greenstone belt. In detail, however, only a small portion of electrum (i.e., electrum that is in textural equilibrium with pyrrhotite and loellingite without arsenopyrite) was directly precipitated from the hydrothermal fluid and thus crystallized during peak metamorphism. The majority of electrum grains (i.e., coexisting with composite loellingite-arsenopyrite 1 grains and with arsenopyrite 1 only) were formed through the release of Au and Ag that had been earlier con- tained within loellingite. The replacement of loellingite, however, occurred under temperatures about 40øC lower than the regional peak metamorphism, thus this process is considered to be part of the main (prograde) gold-bearing hydrothermal event. The retrogressive replacement of loellingite has been documented in other high-temperature, hypozonal lode gold deposits in the Yilgarn and Pilbara eratons (Barni- coat et al., 1991; MeCuaig et al., 1993; Neumayr et al., 1993b)

Acknowledgments

We would like to thank D.I. Groves, H. Dalstra, E. Bloem, and C. Gammons for helpful discussions and suggestions. S.G.H. and P.E.B. acknowledge National Science Foundation grants INT 90-15198, EAR 93-05245, EAR-9508257, and EAR-94-06683. S.G.H. also acknowledges the Deutsche Forsehungsgemeinsehaft Habilitandenstipendium 2327/1-1, 1-2. We also thank two Economic Geology reviewers for their insightful corrections and very helpful suggestions which greatly improved the final version of the mansueript.

October 1, 1996; January 13, 1998

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