Platinum group elements in a 3. 5 Ga nickel-iron occurrence: Possible evidence of a deep mantle...

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 94, NO. B1, PAGES 795-813, JANUARY 10, 1989

Platinum Group Elements in a 3.5 Ga Nickel-Iron Occurrence' Possible Evidence of a Deep Mantle Origin

MARIAN TREDOUX, 1 MAARTEN J. DE WIT, 2' 3 RODGER J. HART, 1 RICHARD A. ARMSTRONG, 2' 4 NICHOLAS M. LINDSAY 1 AND JACQUES P. F. SELLSCHOP 1

The Bon Accord (BA) Ni-Fe deposit occurs in chemically depleted ultramafic rocks of the circa 3.5 Ga Jamestown ophiolite complex in the Barberton greenstone belt of the Kaapvaal craton, South Africa. The host rocks of the BA body are high-temperature ultramafic tectonites which probably represent the upper mantle residue from which the overlying magmatic rocks were separated during partial melting. BA is unusual both mineralogically and chemically. It consists of a rare Ni-rich assemblage: Ni-oxide (bunsenite), -spinels (trevorite, nichromite), and-silicates (e.g., liebenbergite, the Ni end-member olivine) and their altered equivalents. NiO (average = 38%) and FeO + Fe203 (average = 34%) are the major chemical constituents. The Cu and S contents are very low (both < 100 ppm), and no evidence for replacement of preexisting sulphides has been detected. BA is enriched in platinum group elements (PGE) and shows bimodal patterns on the standard PGE diagrams. Differences between the patterns are much greater for Os and Ir than for the other PGE. This observation is interpreted as being indicative of a high-temperature process (> 2000øC; lower mantle). High Ni/Fe and Ni/Co ratios (relative to C-1 chondrite) suggest that BA might have been derived from siderophile-rich material that remained in the lower mantle after inefficient core formation. A model is presented wherein such a metal-silicate heterogeneity is fractionated and oxidized during ascent through the mantle in a thermal plume which originates in the lowermost mantle (the D"layer). Some of its residue is finally incorporated as pods into the lithospheric mantle during formation of the Archean oceanic crust. Inclusion of such fractionated pods in the old subcontinental "keel" of the Kaapvaal craton might constitute a potential PGE source, which could have been tapped by subsequent magmatic activity.

INTRODUCTION

In 1920 an unusual Fe-Ni-Co body was discovered in the metaperidotites of the Barberton greenstone belt [Trevor, 1920]. The body, known as the Bon Accord (BA) deposit (Figure la), first attracted attention because of its extremely high Ni contents (38% NiO) and later because of its Ni-rich, but sulphide deficient, mineralogy (see Table 1). Walker [1923] identified trevorite (NiFe204), the nickeliferous spinel, in this body; Partridge [1944] described the rare serpentine, nepouite (Ni3Si2Os(OH)4); S. A. de Waal and L. C. Calk (as cited by de Waal [1978]) discovered the pure Ni end-member of the olivine group, liebenbergite (Ni2SiO4). De Waal [1979] reported five other new minerals, including the Co-rich and the Ni-rich spinels, cochromite and nichromite. No Cu-minerals have been reported, and sulphides are rare. Chemically, the BA deposit differs from all known Ni deposits in that it is practically devoid of Cu and S (Cu, S < 100 ppm) [de Waal, 1978] and much richer in Ni.

Renewed interest in the BA occurrence has focused on its

origin, following the suggestion by de Waal [1978, 1979] that

•Wits-Council for Scientific and Industrial Research Schonland

Research Centre for Nuclear Sciences, Johannesburg, Republic of South Africa.

2Bernard Price Institute for Geophysical Research, University of the Witwatersrand, Johannesburg, Republic of South Africa.

3Lunar and Planetary Institute, Houston, Texas. 4Now at Department of Geochemistry, University of Cape

Town, Rondebosch, Republic of South Africa.

Copyright 1989 by the American Geophysical Union.

Paper number 88JB03466. 0148-0227/89/88JB-03466505.00

BA might represent an oxidized form of a metallic Fe-Ni occurrence, a conclusion that he reached after elimination of some other possible origins. He argued that neither magmatic segregation nor simple residual (e.g., lateritic) enrichment was likely to have caused the exceptionally high Ni content. Further, de Waal [1979] dismissed a previously proposed [Anhaeusser, 1963] hydrothermal origin related to fluids derived from the nearby Nelspruit granitoid intrusion; he reported that the deposit's distinctive mineralogy and Ni mineralization predate the thermal effects of the granitoid pluton.

De Waal [1978, 1979] suggested that the Fe-Ni body could have had one of three possible forms: (1) awaruite, an alloy associated with the serpentinization of ultramafic rocks [Ramdohr, 1950; Dick, 1974]; (2) josephinite, speculated to be either a deep-mantle-derived alloy [Bird and Weathers, 1979] or a special form of awaruite [Dick, 1974]; or (3) an iron meteorite. De Waal [1978, 1979] reasoned that this last possibility was the most likely because (1) his chemical data suggested a compatibility between BA and iron meteorites; (2) the large size of the BA occurrence exceeds that of known awaruite and josephinite samples by several orders of magnitude; (3) the Ni/Fe ratio of BA (-0.9) is different from that of awaruite and josephinite (both -3.0); and (4) at that time it was believed that the BA occurrence was re-

stricted to only a single locality; this became an overriding factor in favor of the paleometeorite hypothesis.

Recently, however, it has been pointed out that B A is associated with a laterally extensive zone of Ni mineralization, mainly low-grade sulphides, which occurs in highly deformed and altered ultramafic rocks to the southeast of the BA oc-

currence [Keenan, 1986]. Ni values in the Scotia Talc Mine (Figure 2), for example, range between 0.1% and 2% over 10-m-thick zones (Eland Exploration Ltd, unpublished data,

795

796 TREDOUX ET AL.' PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE

COMP!

Fig. 2 (BON ACCORD)

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Localities of sheeted intrusives

ß DETAILED STRAIN ANALYSIS

(GAY, 1969)

I STRUCTURAL SECTION (FRIPP e_t a_l., 1980)

a

\\\

o lO I I

s

b

Fig. 1. (a) Schematic diagram of the Barberton greenstone belt, South Africa, showing the location of the BA occurrence. Also shown are localities with detailed strain analysis [Gay, 1969]. Shaded area on inset map is extent of the Kaapvaal craton. (b) A section across the northern margin of the belt, modified after Fripp et al. [1980]; dark dashed lines are thrusts; light dashed lines depict lithologic layering.

1974). This observation might not tally well with a meteorite hypothesis.

In this paper, new field observations from the area and geochemical analysis of BA and surrounding rocks are reported, with particular emphasis on trace element, platinum group element (PGE), and isotopic (Pb, Nd, Os) compositions. The data provide a new insight into the possible mechanisms for the formation of a BA-type deposit. We suggest an interpretation of BA as a siderophile-rich heterogeneity remaining in the deep mantle after a process of incomplete core formation. Such a model would have

interesting implications for the study of core-mantle segregation and the geochemistry of the lowermost mantle.

GEOLOGICAL SETtING

Regional Geology and Geochronology The BA deposit occurs within the lowermost units of the

Barberton greenstone belt (see Figures la and 2), which consist of altered ultrabasic to basic intrusives and lava flows

of the Tjakastad Subgroup [Keenan, 1986]. It has recently been shown [de Wit et al., 1987a] that most of this subgroup

TREDOUX ET AL.' PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE 797

TABLE 1. Nickel Minerals From the Bon Accord Occurence

Mineral Composition

Ni-olivine* Liebenbergite [Ni2SiOn] Ni-oxide Bunsenite [NiO] Ni-spinel Trevorite [NiFe2On] Ni-Co-Cr spineIs* Nichromite [NiCr2On]

Cochromite [CoCr204] Ni-serpentine Nepouite [Ni3Si2Os(OH)4] Ni-talc* Willemsite [ (Ni Mg)3Si40•0(O H)2] Ni-chlorite* Nimite [ (Ni Mg)6A12Si30•o(OH)8] Ni-borate* Bonaccordite [Ni2Fe BOs] Ni-carbonate Gaspeite [(Ni Mg)CO3]

After de Waal [1978, 1979]. *New minerals first discovered and described from the Bon

Accord deposit.

constitutes a cross section through Archean simatic crust. When reconstructed, an integral section through the green- stones of the Barberton belt consists of a lower (rare earth element (REE) depleted) peridotitic tectonite zone, overlain by a zone which consists of a complex array of magma chambers and conduits (in part sheeted), which in turn in-

trude and are covered by a substantial carapace of pillow lavas and thin cherts. Such a sequence can be classified as an ophiolite senso stricto according to the "Steinmann Trinity" and the Penrose Conference definition of the specific pseudostratigraphy to which ophiolites should conform. These ophiolite rocks are overlain by fine-grained sediments of the Fig Tree Group and the younger arenaceous sequence of the Moodies Group [Anhaeusser, 1963; de Wit et al., 1987b]. The rocks of the greenstone belt are totally sur- rounded by a complex terrain of granite-gneiss plutons (Figure la).

The REE and oxygen isotope profiles across the Bar- berton ophiolitic pseudostratigraphy are similar to those of Phanerozoic ophiolites [Hoffman et al., 1986; de Wit et al., 1987a]. Petrographic and geochemical data on the hydrothermal metamorphism and alteration of these greenstone rocks can also be correlated with ocean floor rocks and Pha-

nerozoic ophiolites. Geochronological studies indicate that the mafic-ultramafic rocks were both formed and metamor-

phosed (i.e., pervasively hydrated) between 3.45 and 3.49 Ga [de Wit et al., 1987a, b]. These combined data have been interpreted by de Wit et al. [1987a] in a framework of simatic

LEGEND

I--• BANDED GREY-WHITE CHERT, OFTEN MYLONITIC ..• QUARTZ- SERICITE SCHISTS, WITH OCCASIONAL CHLORITE - AMPHIBOLE LENSES

.• AMPHIBOUTES ::"..."•.....';• TALC SC.,STS. WIT. MASS,VE TALC LE.SES :::• MASS•.E TO SC.,STOSE S•RPE.*,.,*ES. PREDOMINANTLY META - DUNITE

•vl• STENTOR GRANITOIDS: PREDOMINANTLY MIGMATITES, TONAUTE - GRANODIORITE GNEISSES

""'"' -.... LATE FAULT

• DIP, STRIKE OF LAYERING 26

•8 DIP, STRIKE OF SCHISTOSITY • MINERAL STRETCHING UNEATION

• RIVER

"'"-"• _- TRACK

BON ACCORD

HORNBLENDE

vv

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MINERALIZATION ....

KAAP RIVER

0 100 200 300 400 500

METERS

Fig. 2. Geological map around the BA occurrence, modified after Keenan [1986]. Note the zone of nickel-sulphide mineralization east of BA and the hematite-magnetite-chromite occurrence (locality H, center of map).

798 TREDOUX ET AL..' PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE

BON ACCORD Talc-carbonate schist

• / Tectonic contact Tecto

S ilic •/•'"" '"" '•':: •/'"::: :•!i ..' .'•• i-' ..' • :' .::,: .•½lz•x/'• •

X ', .............. 7;' Hornblende- tourmaline schist

ALPINE •YPE PERIDOTITE now predominantly massive to schistose talc and serpentinite

Trevorite (Ni-Fe oxide) Bunsenite (Ni oxide) Liebenbergite (Ni- Oiivine) Bonaccordite (Ni borate) Ni, Co, Cr SpineIs

S2 foliation

Ni- serpentines (Nepouite) Ni-taics (Wiilemseite) Ni-chlorites (Nimite) Bonaccordite Minor trevorite

Fig. 3. (a) Schematic section across the BA occurrence and its host rocks. (b) Expanded section of the BA ore body, showing simplified distribution of the nickel minerals with respect to the overprinted D2 schistosity.

crust formation and hydrothermal metamorphism, as expected at ocean ridge spreading centres; de Wit et al. refer to this Archean ophiolite as the Jamestown ophiolite complex.

The ocean floor type metamorphism predates a dynamothermal overprint (in amphibolite facies) along the margins of the greenstone belt, which is syntectonic with thrusting [Fripp et al., 1980; de Wit et al., 1987b]. For example, the Stentor tonalite, which forms part of a granitoid terrain in contact with the northern margin of the greenstone belt (see Figure 2) and to which the local amphibolite "aureole" can be related, is reported to be a strongly tectonized gneissic border unit of the Nelspruit granitoid (Figure 2) [Fripp et al., 1980]. It has yielded U-Pb zircon ages of 3347 + 67/-60 Ma and 3250 _+ 30 Ma [Tegtmeyer and Kroner, 1987] and is thus appreciably younger than the adiacent ultramafic rocks.

Geology of the Ore Deposit and Surrounding Rocks The BA deposit is entirely mined out. Its original geo-

logical structure can only be deduced from historic descriptions (given by de Weel [1978]) and from blocks found on the mine dump along the slopes below the deposit. These relics indicate that the deposit consisted of an undeformed core of Ni-rich spinels, oxides, and hydrated silicates, sur- rounded by a schistose envelope containing preferentially orientated Ni-rich minerals (Figure 3). The mined-out area measured approximately 18 m 2, with a depth of 0.35 m, and is estimted to have had a mass of between 20 and 25 tons [de Weel, 1978].

B A formed a structural boudin within altered ultramafic

rocks, close to the boundary between the Barberton green-

stone belt and the granitic gneisses of the Stentor pluton (Figures 1-3). The contact between these two terrains is tectonic. It has been shown that the original transition between the greenstone belt and the Stentor gneisses is contained within a thick thrust complex [Fripp et al., 1980]. Thrusts dip between 20 ø and 40 ø south (see section in Figure lb), and their surfaces are mylonitic. Detailed strain analyses close to the granite-greenstone interface [Gay, 1969] (locations shown on Figure la) indicate that the maximum extensional direc- tion is subparallel to the northward tectonic transportation, as deduced from the thrust and associated shear zone geo- metries [Fripp et at., 1980].

Field observations in the vicinity of the BA deposit indicate that the rock sequence of this area is part of the same southerly dipping thrust complex. Both the upper and lower contacts of the ultramafic sheet which hosts BA are tectonic.

These contacts are bound by cherty to siliceous mylonitic layers up to 10 m thick. Syntectonic mineral deformations define a well-developed northwest-southeast stretching lin- eation in the mylonites, perpendicular to the fold axis of isoclinal intrafolial folds (Figures 2 and 4). The ultramafic sheet is therefore allochthonous. Similar ultramafic allochthons

have now been recorded from at least four other areas close

the margins of this greenstone belt [de Wit et al., 1987a; Bar- ton, 1986; Lamb, 1987].

Figure 3 shows a section through the BA boudin and its ultramafic host rocks. The ultramafics are entirely serpentinized, and serpentinite and talc-carbonate schists occur along the contacts between the mylonites and the BA body. Relic textures, preserved in massive serpentinite, indicate that the original ultramafic rock type was a coarse-grained dunite with minor (ortho?) pyroxenite. The textures are similar to those described in the REE-depleted Stolzburg peridotites [de Wit et al., 1987a] (for location see Figure la) and are comparable to the high-temperature tectono metamorphic rocks described from the lowermost ultramafic ex- posures of Phanerozoic ophiolites. Such ultramafics are generally believed to represent residual upper mantle material [Gasset al., 1984; Prinzhoffer and Allegre, 1985; de Wit et al., 1987a]. Accordingly, the BA host rocks are interpreted as "hot tectonites" formed at the Archean crust- mantle transition zone, which represent depleted upper mantle residue from which the overlying igneous rocks of the ophiolite were extracted during partial melting beneath a mid-ocean-type ridge [de Wit and Tredoux, 1988].

On Figure 2, a small outcrop of a deformed hematite-magnetite (HM) rock (marked H) is indicated close to BA. An elongate zone of disseminated Ni-sulphide mineralization, which lies within a similar sequence of serpetinized ultramafic rocks as the BA deposit [Keenan, 1986], is also shown in Figure 2. It should be noted that BA, the HM body, and the sulphide mineralization all lie within the metamorphic aureole of the Stentor pluton. Note also the tourmaline-rich zone along the upper margin of BA (Figure 2).

MINERALOGY AND PETROGRAPHY

The unusual mineralogy of B A has been reviewed in detail elsewhere [de Weel 1978, 1979]; Table 1 summarizes this work. De Waal subdivided BA into five zones. In view of

the uncertainty associated with the exact positioning within the original structure, samples in this study were subiected to a much simpler classification: the only criterion used was whether the sample was visibly deformed (i.e., from the


Fig. 4. D2 isoclinal folds in the siliceous mylonite which outcrops above the BA deposit. The position of one lithologi- cal unit is enhanced by a dotted trace.

schistose outer rim) or not. Some new petrographic data are presented in this section.

In the undeformed core of the deposit, subangular domains of up to 0.8 cm in diameter pseudomorph a coarse peridotitic texture (see Figure 5a). Two mineralogically distinct domains are present: (1) The first type (type A), which is predominant, consists mainly of trevorite euhedra, inter- grown with subordinate silicates (nepouite, willemseite and nimite), bonaccordite, and a little gaspeite. The trevorite crystals often show complex zoning (Figure 5b), with cores which are often composed of bunsenite. In addition, relatively intact crystals of bunsenite are observed, frequently enveloped, and crosscut by the trevorite. These textures are interpreted as indicating that trevorite replaced the bunsenite. The inner cores of some of the trevorite crystals are Ni- and Co-rich spineIs. An intermediate chrome-rich zone with silicate inclusions is also present (Figure 5b). Compositions of these assemblages are given in Table 2. The compositions of chromite grains in the B A samples are similar to those found in the Stolzburg ultramafic tectonites but dissimilar from chromites reported from nonchondritic meteorites [Bunch and Keil, 1971] (Figure 6). The bonaccordite occurs as fine needles which occasionally crosscut the trevorite (Fig- ure 5c); (2) The less abundant domains (type B) consist mainly of the hydrated silicates, i.e., nepouite, willemseite, and nimite. The nepouite occasionally contains remnants of its precursor, the Ni-olivine liebenbergite. Only minor amounts of trevorite are seen. Finely dispersed gaspeite and bonaccordite also occur. The grain boundaries between the pseudomorphed domains are filled with a finer-grained groundmass of trevorite and/or hydrated silicates. Minute spherical and hexagonal particles of millerite (NiS) are seen in trevorite and chrome spineIs, but no evidence for a major replacement of sulphides has been observed.

In the envelope of the deposit the relic peridotitic texture that can be identified in the core is not observed; a strong

schistosity prevails. Willemseite and nimite dominate the assemblage and impart the schistosity. They occasionally show kinking and alteration to goethite. Syntectonic bonaccordite also occurs along the schistosity. Trevorite is less abundant and occurs mainly as dispersed euhedra, sometimes over- growing the schistosity. Thus, unlike the core of the body, which is dominated by the trevorite-rich assemblage (type A), the rim is dominated by the silicate-rich assemblage (type B).

In summary, the earliest identifiable mineral assemblage appears to have contained liebenbergite, bunsenite, and Co- Ni-Cr spineIs. The rare millerite is intimately associated with this assemblage. The nepouite and trevorite, which now form the bulk of the mineralogy, are secondary. Bonaccor- dite, nimite, willemseite, and the schistosity all postdate the trevorite-nepouite paragenesis. Small grains of stibnite (Sb2S3), antimony metal, and breithauptite (NiSb) are occasionally encountered, but their placement within the sequence of events is unclear. No PGE minerals were observed. This is not surprising as PGE minerals are usually submicroscopic.

The mineralogy of the HM body is similar to that of the deformed BA material, except that the Ni analogue is Fe. Thus, instead of trevorite one observes magnetite and hematite. The hydrated minerals are talc and chlorite. Chromite forms a much more substantial part of the mineralogy than is the case for BA, but the exact composition of the chromite has not yet been determined.

GEOCHEMISTRY

General Geochemical Character

A suite of 17 elements was determined in nine whole rock

samples from BA, two samples of the HM body, and a sample which is representative of the ultramafic host rock. Data are listed in Table 3, with graphic representation-in Figure 7.


Fig. 5a

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Fig. 5. (a) Photomicrograph showing the pseudomorphed domains in the undeformed core of BA. (b) Photomicro- graph of complex zoning in a spinel from an undeformed BA sample: trevorite rims (pale gray) with cochromite cores (dark); for analysis see Table 2. Scale bar and units in micrometers. (c) Late elongate crystals of bonaccordite, over- growing a trevorite crystal. Scale as in Figure 5b.

TREDOUX ET AL.: PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE 801

TABLE 2. Composition of Minerals From the Bon Accord Deposit and Chrome-Spinels From the Stolzburg Complex

Olivine Nichromite Cochromite Chromite Trevorite Magnetite Chromite**

SiO2 30.33 0.40 0.27 0.23 0.33 0.00 0.39 TiO2 0.04 1.31 2.04 0.43 1.26 0.07 1.21 A1203 0.06 8.17 13.58 10.55 0.65 0.00 10.25 Fe203 NA 9.94'v 0.88 3.24 + 63.18 + 65.51 NA FeO 4.36 5.84 9.60 26.66 1.14 31.10 36.15 MnO NA NA 0.74 3.09 + 0.08 + 0.10 1.90

MgO 7.98 0.47 0.69 1.91 0.21 0.12 4.80 CaO NA NA 0.04 0.01 + 0.04* 0.02 0.05 Na20 NA NA NA 0.00 + NA 0.00 0.49 K:O NA NA NA 0.00 + NA 0.00 0.02 NiO 54.57 13.95 9.24 0.09 x 31.41 0.00 0.14 CoO NA 14.61 17.02 NA 0.8 + NA 0.26 Y20 NA NA 1.60 NA 0.09* NA 0.03 Cr203 0.03 43.55 44.77 50.16 1.72 3.97 44.80 ZnO NA NA 1.08 NA 0.28 x 0.04 NA Total 97.37 98.23 101.53 100.44 99.87 100.93 100.44

Number of grains 1 1 3 7 7 1 11 analyzed (+ :4 only) ( + :4 only)

(x :6 only) (': 3 only) (*: 2 only)

Analyses were done by electron microprobe, at Imperial College, London. NA, not analyzed. W is calculated from stochiometry. **Chromites from the Stolzburg Complex.

Short descriptions of the analytical techniques are given in the footnotes of Table 3. Additional major and trace element data are given by de Waal [1978]. To date it has not been possible to obtain a full set of major element analysis by XRF, as the BA material does not fuse under standard laboratory conditions. Of the 17 elements plotted in Figure 7, only those listed in Table 3a will be discussed here. The PGE geochemistry is addressed in a separate section.

When compared to average undepleted mantle [Jagoutz et al., 1979], the ultramafic host rock (STD 2) shows very similar concentrations of Fe, Ni, Co, and As but is depleted in Cr and the LREE. The REE depletion is consistent with the interpretation of this rock being part of the depleted mantle section of an ophiolite [de Wit et al., 1987a]. STD 2 is enriched by an order of magnitude in Sb, relative to average mantle. It is not yet known whether this Sb enrichment is a

AI

CHROME-SPINEL ß STOLZBURG

? ø2.

/ Fe 3+ / Cr

Fig. 6. Chromian-spinel compositions from the BA deposit, compared with those from the Stolzburg complex. Also plotted are chromites from some pallasites, mesosiderites, and achondrites [Bunch and Keil, 1971].

primary or secondary feature of the rock; Sb enrichment is common in all greenstone belts of the Kaapvaal craton. (T. N. Pearton, personal communication, 1986).

Relative to the host rock (and to undepleted mantle), the BA body is enriched in Co, very enriched in Fe, and extremely enriched in Ni, As, and Sb. These trends are illustrated by the high Ni/Fe, Ni/Co, and Sb/(Ni + Fe) ratios of BA (see Table 3c). The strong relative depletion of Cr in BA is shown by the high Co/Cr ratios. BA is enriched in LREE in comparison with its host rock and could have approximately mantlelike LREE concentrations. However, this observation must remain speculative until the BA data base is improved.

In overall chemical character, BA is clearly different from C-1 chondrites and iron meteorites. The BA ore is notably enriched in Ni, relative to iron meteorites, but slightly depleted in Co and very depleted in Fe (compare, for example, the Ni/Fe and Ni/Co ratios in Table 3c). Cr is slightly depleted and Co/Cr ratios are high when compared to both C-1 and iron meteorites (Table 3c). Au is extremely depleted relative to iron meteorites; the values in BA are more akin to those of chondrites. As and Sb are very enriched in BA relative to both iron meteorites and C1 chondrite. The

LREE of BA is very enriched relative to iron meteorites; group B is even enriched in LREE relative to chondrites.

A consistent feature of the BA samples is that they form two distinct chemical groups. The differences between averages of these groups are shown in Figure 8: group A is enriched in Sb and (to a lesser extent) Ni and Fe, and group B is enriched in As and the LREE. Cr, Co, and Au are not appreciably differentiated. Semiquantitative XRD analyses of the samples have shown that group A samples are dominated by trevorite (i.e., by the type A petrographic domains described in the previous section) and group B by the Ni silicates (i.e. type B domains). It would therefore seem that the chemical differences between the two groups reflect mineralogical control of Ni, Fe, and Sb by trevorite-bunsenite and of the REE by the Ni silicates. Also noteworthy is

802 TREDOUX ET AL ' PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE

TABLE 3a. Geochemistry of BA and HM Samples

Fe% Ni% Co As Sb Cr La Ce Sm Eu Au Group

R-4.5

BA-83.B BA-83.1 DBS-3 BA-84.3 L-4.5 BA-84.1* BA-84.2' D-4.5 BAE-1 BAE-2

$TD-2

32.3 30.5 3900 60 1600 201 0.28 1.19 0.48 0.08 0.17 A 32.3 29.6 3600 70 5100 53 0.36 1.47 0.61 0.11 0.25 A 33.3 17.7 3700 40 4200 55 0.44 1.45 0.49 0.05 nda A 34.5 31.3 3800 60 2300 171 0.26 1.08 0.49 0.05 0.19 A 33.9 27.9 4100 30 2400 24 0.38 1.49 0.58 0.08 0.34 A 33.8 27.7 4100 40 4300 111 0.64 2.26 0.68 - 0.12 A 30.1 18.4 3300 90 15 76 0.98 1.11 0.46 0.19 0.18 B 18.9 16.8 4000 400 24 112 2.30 2.30 2.44 0.19 0.21 B 30.7 23.8 3300 50 180 nda 2.58 2.46 0.62 - 0.20 B

41.1 0.59 673 1723 704 17300 .... 0.25 HM body 45.0 0.60 386 1217 2615 18200 .... 0.16 HM body

6.7 0.27 130 0.3 2.2 1326 0.03 - 0.12 0.14 0.002 Host ---8 5000 9.3 0.28 150 0.0004 nda 0.00005 0.0003 1.20

1.65 760 2.87 0.23 3975 0.367 0.96 0.231 0.087 0.22

0.21 105 0.1 nda 3140 0.24 0.82 0.33 0.13 0.005

3.2 568 nda nda 1138 nda nda nda nda 0.11

Fe meteorite -•90

C-1 chondrite 27.8

"Average" mantle 6.08

Kambalda ore + 10.8

Compared to (1) a metadunite from the Stolzburg complex (STD 2); (2) the average of 13 iron meteorites [Crocket, 1978] of various classes (data selected only from those meteorites for which for all six PGE are available); (3) C1 chondrite (average) [Taylor and McLen- nan, 1985]; (4) "average" mantle [Jagoutz et al., 1979; Morgan, 1986]; and (5) average ore from Kambalda [Cowden and Woolrich, 1987; Cowden et al., 1986]. The following elements are listed: Contains (1) major (Fe,Ni) and trace elements, done by (INAA) instrumental neutron activation analysis, using the method described by Erasmus et al. [1977]; (2) Cr, done by standard XRF at the Dept of Geology, Uni- versity of the Witwatersrand; (3) REE done by radiochemical NAA (RNAA), using the method of Loo [1975]. Sm was also determined by isotope dilution (see Table 4); comparison between the two sets of REE data show that the RNAA data are too high by +40%. This is a constant error, so that the difference between the groups remain unaffected; dashes indicate not detected; nda, no data currently available. Values given in ppm, except for Fe and Ni.

*Deformed (schistose) sample. +Calculated back from 100% sulfides, using an approximate factor of 4.4 [Keays et al., 1982].

that the group A samples all represent matedhal from the undeformed core of the deposit. Two of the group B samples consist of schistose rim matedhal, while the third (D4.5) has characteristics of both groups. It would therefore appear that the chemical variation may be related to the position of

the samples within the body, but this is difficult to verify since the ore is no longer in situ.

The major chemical difference between BA and the HM body is (as preempted by the mineralogical investigation) the relative depletion of Ni in the latter (see the Ni/Fe ratios

TABLE 3b. PGE Data, ppb

Os Ir Ru Rh Pt Pd Group

R-4.5 (n=5) 488 _+ 112 651 _+ 79 1556 _+ 210 478 2166 _+ 565 1825 _+ 82 BA-83.B 1114 932 1830 515 1764 1537 BA-83.1 550 730 1314 nda 1794 1793 DBS-3 584 501 1771 493 1478 1845 BA-84.3 441 380 1722 426 1694 1533 L-4.5 278 384 1123 nda nda 1561

BA-84.1 (n=4)* 108 _+ 9 150 _+ 29 684 _+ 66 221 1321 _+ 255 896 _+ 142 BA-84.2' 73 72 617 225 1078 833 D-4.5 66 97 849 242 1583 461

BAE-1 (average) 119 204 130 nda 15530 <lee BAE-2 (average) 129 247 201 nda <lee <100 STD-2 10.5 6.9 8.7 4.4 21.1 < 10 Fe meteorite 3500 3000 6000 2000 8000 4000 C-1 chondrite 761 710 1071 201 1430 836

"Average" mantle 3.3 3.4 nda nda -•20 nda Kambalda ore + 109 52 223 73 375 466

Josephenite (average) <0.1 0.5 73 70 1107 6690

A A

A

A

A

A B

B

B

HM body HM body

Host

High background values in the gamma ray spectra of the BAE samples made it impossible to determine Pt and Pd at values of <100 ppb, although this exceeds the theoretical detection limit (usually <10 ppb) by an order of magnitude. Determined by NiS fire assay precon- centration, followed by NAA, as described by de Wit and Tredoux [1988]. Values listed for C-1 chondrite were taken from Taylor and Mc- Lennan [1985] for Ir, Ru, Rh, Pt, and Pd, while Os concentration was calculated from sources given by Anders and Ebihara [1982] and adjusted for volatile loss during accretion (i.e., multiplied by 1.5) as was done for the other PGE by Taylor and McLennan [1985].

*Deformed (schistose) samples. +Calculated back from 100% sulfides, using an approximate factor of 4.4 [Keays et al., 1982].

TREDOUX ET AL ' PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE 803

TABLE 3c. Interelemental Ratios

Ni/Fe Ni/Co Sb/(Ni+Fe) Co/Cr Ni/Ir Fe/Ir Ir/Au Pd/Ir Os/Ir Ru/Os Group

R-4.5 0.94 78.2 25.48 19.40 46.85 49.62 3.83 2.81 0.75 3.19 BA-83.B 0.92 82.2 82.39 67.92 31.76 34.66 3.70 1.65 1.19 1.64 BA-83.1 0.83 74.9 68.85 67.27 37.95 45.62 - 2.46 0.75 2.39 DBS-3 0.91 82.4 34.95 22.22 62.48 68.86 2.68 3.68 1.16 3.03 BA-84.3 0.82 68.0 38.83 170.83 73.42 89.21 1.11 4.04 1.16 3.91 L-4.5 0.82 67.6 69.92 36.94 72.14 88.00 3.15 4.07 0.72 4.04 BA-84.1* 0.61 55.8 0.31 43.42 12.27 20.07 0.84 5.98 0.72 6.34 BA-84.2' 0.89 42.0 0.67 35.71 23.33 26.25 0.34 11.61 1.01 8.43 D-4.5 0.78 72.1 3.30 - 36.06 46.51 0.49 4.74 0.67 12.96 BAE-1 0.01 8.77 16.89 0.04 2.89 201.47 0.82 - 0.58 1.09 BAE-2 0.01 15.54 57.35 0.02 2.43 182.19 1.56 - 0.52 1.56 STD-2 0.04 20.9 0.32 0.10 39.13 971.01 3.0 - 1.50 0.83 Fe meteorite 0.09 16 0.003 33.33 2.67 30.00 2.5 1.3 1.16 1.71 C-1 chondrite 0.06 21.7 0.008 0.19 2.32 39.15 3.26 1.18 1.07 1.41

"Average" mantle 0.03 20 - 0.03 61.76 1788.23 7.39 1.62 0.97 -

Kambalda ore 0.30 56.8 - 0.50 61.54 206.52 0.46 8.91 2.09 2.04

A

A

A

A

A

A

B

B

B

HM body HM body

Host

Dashes indicate lack of data for a relevant element.

*Deformed (schistose) sample.

in Figure 3c). Further, the HM body is enriched in Cr, As, and Fe relative to BA. Au shows no differentiation.

Isotope Geochemistry The Pb-Pb system. Seven whole rock Pb isotopic compo-

sitions have been determined on BA material, of which five were on group A samples (see Table 4). The data are

10'

10 •

10 2

10

1

10-"

10 -2

10

ß ' ß . a Iron meteodte

ß ' ' El Ultramafic host ß ß o ß

_ ' x - . ß ß Group A (ave.) i ß . : . ß o Group B (ave) -.:

ß .

ß ß + ß . . + BAE-1 ß ß ß : ß ß x BAE-2

ß ß ß

ß

: .X 0 ß ß .

: o : ß . ' ' ß .

ß . . ß .

- ß ' 0 El . . . ' A X ß

+X .•-/X - 0 ' 0 A ' 'A ' - A' A A 0 ' A ß ' ß .

' ß ' 0 ' ' ß - ß ' ß . +x' . ß . . ß ß .+ ß ß ß A ß ' ' 0 ' ß

:-- o-+ .... ß :ß- : 'o' :S x- ' 'x ' ß . :El ß. o. o ß .

' El ß i+x ' ' '

o o ß ß . o +• o + ' ß . .

ß . .

ß ß ' . O ß 0 ß O ' O- "- ' ß o' ' . o

ß . .

I ' ' . . , , , .'.', , ::. :., !.__i, . -

Cr fe Co Ni As Sb La Ce Sm Eu Os Ir Ru Rh Pt Pd Au

Fig. 7. Multielement diagram of BA samples, the HM (BAE) samples, and the metaperidotite (STD 2). For comparison, average iron meteorite concentrations [Crocket, 1978] (as described in Table 3) are also indicated. All values were normalized to C1 chondrite; the values of Taylor and McLennan [1985] were used for all elements except Os (see comment in footnotes of Table 3b).

plotted on conventional 2ø8pb/2ø4pb and 2ø7pb/2ø4pb versus 2ø•pb/2ø4pb diagrams (Figures 9a and 9b). Although the data scatter appreciably and do not conform to a simple linear trend which could be of geochronological significance, we can make the following statements:

1. The present-day Pb isotopic compositions are relatively radiogenic, with samples D4.5 and BA84.2 (both group B) showing uncommonly high values for a whole rock sample of this nature.

2. The data plot to the right of the geochron or meteoritic isochron shown in Figure 9b.

3. The data cannot be modeled according to the Stacey and Kramers [1985] two-stage terrestrial growth curve, as the present-day ratios indicate /z values of between 11 and 12, which are significantly greater than the 9.74 second-stage value adopted in the Stacey-Kramers model. The evolution

lOO

> o

•' •0

o

o

o

0.1

ß ß ß

ß ß ß

ß ß ß _

ß ß ß

ß

ß

:- -

, , , , I , , , , I , , , , I , ,

Cr Fe Co Ni As Sb La Ce Sm Eu Os Ir Ru Rh Pt Pd Au

Fig. 8. Comparison of the average concentrations of elements in BA group A samples with BA group B averages. The elements that plot above the line are relatively enriched in group A, those below the line in group B.

804 TREDOUX ET AL.: PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE

TABLE 4. Pb-Pb, Sm-Nd, and Os Isotopic Whole Rock analyses of Selection of BA Samples

Sample •Pb/eø4pb :ø7pb/eø4pb •Pb/:ø4pb •87Os/•Os Re, ppb Sm, ppm Nd, ppm •n7Sm/•44Nd •n3Nd/•½nNd

L-4.5 38.204 16.096 20.447 R-4.5 38.258 16.326 20.262 1.63_+0.10 BA-84.3 39.307 16.420 21.374

BA-83.B 39.090 16,451 21.549 DBS-3 38.202 16.272 19.900 BA-83.1 D-4.5 41.222 17.923 27.045 BA-84.1 1.83_+0.15 BA-84.2 53.943 19.464 39.168 STD 2 1.94_+0.10

Canyon Diablo 1.03_+ 0.10

0.23 0.87 0.162227 0,512084_+34 0.17 0.58 0.179967 0.512203_+39 0.22 0.51 0.186794 0.21 0.68 0.184033 0.512326__+29

0.21 0.75 0.168985 0.512189_+35

57 1.48 13.46 0.06656 0.509884 _+20 0.67 3.06 0.131587 0.511286_+20

Pb abundances were determined by isotope dilution, using the methods described by Hart et al [1981]. Methods used for the measure- ment of the Sm and Nd isotopes are described by Zindler et al., [1979]. Os isotopes were determined by TAMS, as described by Fehn et al. [1986]. One Re analysis, determined by S. R. Hart, is also listed.

of the U/Pb isotope systematics of the samples appear to in- volve a multistage process, with at least one stage of U enrichment. This U enrichment (also manifested by reduced Th/U ratios relative to the Stacey-Kramers model) must have been an ancient event in order to exert such an influ-

ence on the 2ø7pb.

The Sm-Nd system. The Sm and Nd contents and isotopic ratios have been measured on seven BA samples of which five were from group A. The data are listed in Table 4 and plotted on a Sm-Nd evolution diagram (Figure 9c), to- gether with the data obtained by Hamilton et al., [1979] for mafic-ultramafic rocks from the Onverwacht Group (Tjaka- stad Subgroup). On their own, the BA samples define an errorchron; which corresponds to an age of 3173 +_ 310 Ma (2s) and an initial 143Nd/144Nd ratio of 0.50853 + 16. This age is not considered geologically meaningful; first, because the data define an errorchron; and second, the slope of the regression line is largely controlled by the two group B samples (from the schistose rim) which are the most likely to have been affected by the syntectonic intrusion of the nearby Stentor granite. If we exclude the two group B samples, the remainder of the BA data scatter closely about the regression line for the Tjakastad Subgroup which define an age of 3510 + 60 Ma and an initial 143Nd/1•Nd of 0,50809 _.+ 4. The geological setting and the similarity of the BA data to that of the Tjakastad rocks suggest that the Sm/Nd isotopic system of the undeformed core has remained undisturbed for at

least 3500 m.y. It is noteworthy that there is a strong similarity between

the Sm-Nd and the Pb-Pb data (see Figure 9). In both cases, the data for group A (undeformed samples) scatter about a linear array defined by other Barberton rocks, while the schistose group B samples plot well off the regression lines.

Os isotopes. Os isotopic analyses were performed by the TAMS (Tandem Accelerator Mass Spectrometer) group of the University of Rochester on two BA samples as well as on the host rock; the results are given in Table 4. The measured 187Os/1•Os of the group A sample (R4.5) is lower than that of the group B sample (BA84.1). Re data for one sample (BA84.1; see Table 4) was obtained from the Massachusetts Institute of Technology. The initial Os isotopic ratio for this sample, using a decay constant for 187Re of 1.52 x 10 -11 year -1 [Luck and Allegre, 1982] and assuming 10% errors on the Re and Os measurements, is 0.783. This value is consistent with the proposed mantle evolution line of the Allegre

and Luck [1980] for an age of 3200 + 360 Ma. The age has no geological significance as it is based on one data point only and large errors are involved.

Platinum Group Element (PGE) Geochemistry

The entire suite of samples (BA, the HM body, and the host rock) was analyzed for their PGE contents. The data are listed in Table 3b. Two of the B A samples (R4.5 and BA84.1) were analyzed in replicate (n = 5 and 4, respecti- vely) for all the PGE except Rh. Average values and standard deviations are listed in Table 3b. Note that the values

all fall within the expected 20% error limit, except for Os and Pt in R4.5. This probably indicates that these elements, especially Pt, are inhomogeneously distributed in R4.5, as the check samples that were run with the replicates fall within the 20% limit. The samples of the HM body (BAE-1, BAE-2) were analyzed in duplicate and averages are given in Table 3b.

The chondrite-normalized PGE concentrations are plotted in Figure 10. Similar graphs, with the PGE values plotted in order of decreasing melting point, have proved useful in the discussion of the PGE geochemistry [Naldrett, 1981]. The data of this suite were not recalculated to 100% sulphide, as is often done for PGE ores [Naldrett, 1981], because of the low concentration of S in BA. Note that the PGE chemistry follows the distinct bimodality defined by the major and other trace element chemistry. The group A samples (i.e., those characterized by higher Ni and Sb concentrations) have relatively enriched PGE patterns with the highest Os and Ir values (0.5-1.5 x C1) and definite positive Rh and Negative Pt anomalies. The group B samples (enriched in LREE and radiogenic Pb) have more depleted PGE trends, especially for Os and Ir (0.06-0.11 x C1), with Rh values peaking near C1 and a weaker Pt trough, lying consistently below the group A values.

A comparison between BA and other rock types with elevated PGE concentrations leads to the following broad observations:

1. Although there are superficial resemblances between the group A patterns with those of "average" iron meteorite (see Figure 11), the total noble metal concentrations in BA are much lower than those in the iron meteorites: they are much more akin to those of chondrites. However, unlike the PGE patterns of group A, chondrites do not display a pronounced negative Pt anomaly [Crocket, 1978].


6O

3O

2O

14

(a)

ß

BA84/2

- (b)

• ß D4.5 o

•..½...--'""•• Stacey-Kramers growth curve ß • I '. I ; I : I : I : • : I '

.• . ,•o•

..,•o• •e•, e' ß

•x D4.5 2 1 o "•

.,••/ $ acey-Kramers growth curve I • I i I • I • I i I . I .

10 14 18 22 26 30 34 38

aoSpb/ao•pb

.511-

.510'

BAe84-2

.06 .1•2 t43Sm/t44Nd .18 Fig. 9. (a) Thorogenic Pb-Pb, (b) uranogenic Pb-Pb, (c) Sm-Nd isotope plots for the BA samples (circles). The Stacey and Kramers [1985] two-stage growth curves, calibrated in 109 years, are included for comparison in Figures 9a and 9b. Note that the BA data (excluding the group B samples, D4.5 and BA84.2) overlap with the data recently reported for mafic-ultramafic rocks of the Tjakastad Subgroup of the Onverwacht Group (triangles) [Brevart et al., 1986] in that they fall on the same linear array. The slope of this linear array (or secondary isochron) yields and age of 3460 _+ 70 Ma [Brevart et al., 1986]. The Sm-Nd data show a similar trend; that is, the BA data are in concordance with the data from the Tjakastad Subgroup (triangles in Figures 9c [Hamilton et al., 1979], with the group B samples showing a much greater enrichment in LREE.

2. No meteorite PGE trends [Crocket, 1978] match the group B patterns. Also, various interelement ratios (e.g., Ni/Ir, Pd/Ir, Ru/Os; see Table 3c) are much higher in both BA groups than those of any meteorites.

3. The relatively flatter PGE trends of group A are similar to those of Ni-Cu ore deposits associated with Archean komatiites (for example, Mount Edwards, Western Austra- lia, and Langmuir, Canada) [Naldrett, 1981; Keays et al.,


10 ' I ' I .-{- I '

/\ ..'

ß ........ • ......... '• ': O

"-' ß c- ..-" ' z.:' .." l• ..-' l.:'-. .'

0 10 -1 ....... + o

N

E - R-4.5 (undeformed) •_ 10-2 o ß ..... BA-85.B (uncleformed)

I, BA-85.1 (uncleformed) • e-- DBS-5 (undeformed) -o e-- BA-84.5 (undeformed) o 10-a ß L-4.5 (uncleformed) 0 o-- BA-84.1 (deformed)

0 ..... BA-84.2 (deformed) 0-- D-4.5 (hybrid) + ..... HM body (BAE-1)

10-+ , I , I , I

Os !r Ru Rh Pt Pd

MP=5050 2454 2554 1967 1768 1555"C

Fig. 10. Chondrite-norma[ized POE plots of BA and HM samples. The POE are plotted in order of decreasing melting temperatures [Westland, 1981] from left to fight [Naldrett, 1981]. Note the large range in Os and Ir values and the distinct negative Pt anomaly in most of the samples. C-1 chondrite values used for normalization come from Taylor and McLennan [1985] for It, Ru, Rh, Pt, and Pd. Os values were calculated as described in footnotes of Table 3b.

1982], although absolute concentrations of these ore deposits are an order of magnitude lower than those of BA (compare Figures 10 and 11). The Ni/Co and Ni/Fe ratios of B A and Ni-Cu ores from Kambalda are very similar, but the Co/Cr ratios are very different (see Table 3c).

10

r-

e c-

o 10 -• o

N

0

E ,._ 10-2 o

o 10-• o

10 -+

//

:...-: :'.'):::-): ::, ...•. ......... ..... ,.. •:\•?' . ..... ,'

t

/ .. ......

t ...-' '• .....

.... /...•..." .... / / ; __•1,7•... Field for BA

x- ..... Hont rock of BA o-- Average FeM

? ..... Komatiitic Ni$ ore -t-- Jonephinite

, I , I

Os Ir Ru Rh Pt Pd

Fig. 11. Chondrite-normalized PGE plots of "average" iron meteorite, the average of three analyses of josephinite and "average" komatiitic Ni-Cu sulphide ore. The BA field and the trend for the host rock are indicated for reference. Data are given in Table 3b.

10

10-2

o 10-: o

10-+

I Layered trend - upper x--Layered trend - lower o ..... Ophiolite trend - upper 0 ..... Ophiolite trend - lower

, I m I m I

Os Ir Ru Rh Pt Pd

Fig. 12. Chondrite-normalized compositions of PGE concentration fields in post-3.0 Ga mafic-ultramafic intrusions [Naldrett, 1981] and ophiolites [Page and Talkington, 1984]. Note that the Os and Ir in these post-3.0 Ga occurrences are all depleted relative to chondrite and correspond to the values in the BA group B samples (Figure •0).

4. The PGE patterns for group B are broadly similar to those exhibited by sulphides associated with layered mafic intrusions [Naldrett, 1981] (see Figure 12), except that the Pd values are much lower.

5. It is interesting to note that although PGE concentrations in ophiolites and layered complexes have very dissimilar patterns, in both cases the normalized Os-Ir values are of the same order as those of group B (Figure 12). In fact, the only other known terrestrial PGE occurrences which have Os values > 0.3 x CI chondrite are associated with the

depleted residual basal portions of the Phanerozoic Shetland ophiolite [Prichard et al., 1986].

The PGE patterns of the HM body are different from those of BA, although there is some similarity in Os and Ir concentrations between the HM body and the BA group B samples. Ru concentrations in the HM samples are much lower than in B A; the same is probably true for Pd. (Pd was not detected in either of the HM samples; see Table 3b for further comment.) The HM body appears to be even more heterogeneous with respect to Pt than the core of BA (R4.5): in the one sample (BAE-2), Pt is not detected, while the other (BAE-1) is very enriched in Pt relative to B A. The chondrite normalized plot of BAE-1 show a flat pattern with a positive Pt anomaly (see Figure 10), which complements the negative Pt anomaly of the BA samples.

ORIGIN OF BeN ACCORD

The key observations regarding BA which show its singu- larity in the present crustal setting are listed. Any model for the formation of the deposit, should attempt to explain all of the following: (1) the exceptionally high Ni contents of the ore and its unusual mineralogy; (2) the position of the body within an extensive sheet of coarse-grained, chemically depleted ophiolitic ultramafi½s; (3) similar compositions of the


chromite-spinels of BA and those of other local peridotites; (4) the unusual trace element geochemistry, including the very low S and Cu contents; (5) the fact that the Pb isotopes do not occur with meteoritic values and that they indicate an ancient U enrichment event; (6) the strong similarity between the Pb-Pb and Sm-Nd isotopic data of BA and those of mafic-ultramafic rocks at other localities in the greenstone belt; (7) the agreement of the available Os initial ratio with the Luck-Allegre mantle evolution line; (8) the consistent bimodal grouping of the chemistry and istotopic ratios, which appears to reflect the control of two distinct mineral assemblages; (9) the PGE geochemistry which is not only ex- traordinary in absolute concentrations of these elements but also in the two trends displayed between the groups; (10) the association of BA with hornblende-tourmaline schist along its upper margin; (11) the occurrence of a Pt-rich hematite- magnetite-chromite body, and low-grade Ni-sulphides in talc rocks, along strike from BA, as the well as complementary PGE patterns between B A and the HM body.

It is difficult to reconcile a meteorite origin with almost every one of the above observations; neither geological nor geochemical data support an extraterrestrial origin for the B A occurrence. Fresh chondrites and iron meteorites do not

contain major and trace element concentrations resembling those of the deposit. An altered chondrite would have had to gain an order of magnitude more Ni, while losing an order of magnitude of Cr, during secondary (terrestrial) processes. On the other hand, the most common types of iron meteorite (types IA, IIA, and IIIA) would have had to gain a substantial amount of Ni and lose large quantities of Fe and PGE. Some rare iron meteorites and the pallasites have sufficiently high Ni contents, but they usually have extremely low Ix (and probably other PGE) concentrations [Wasson, 1974] which would not match those of BA. Furthermore, the Pb-Pb and Sm-Nd isotopic data do not concur with the common meteorite age (4.55 Ga). The Os isotopic data also do not support an extraterrestrial origin for BA.

On the other hand, there is much evidence to support an endogenous relationship between BA and its ultramafic environment (e.g., points 3, 6, 7, 10 and 11). The siderophile element contents of BA therefore are presumed to be derived from the Earth's mantle. It now needs to be resolved

whether the extreme enrichment in Ni (and the PGE) of BA happened in the mantle (i.e., before emplacement in the crust) or whether it resulted from secondary redistribution of disseminated values in the crustal environment.

Possible Crustal Models for the Formation of BA There are several well-established models whereby sidero-

phile elements can be concentrated into massive bodies in the crust. Most of these require that sulphide minerals form the major part of the ore body. This is not the case in BA; there is no petrographic evidence of relic or pseudomorphed sulphide minerals. In spite of this serious objection, such models nevertheless are discussed in the following discussion to provide a balanced view.

1. Formation of B A due to secondary alteration of a massive Ni-sulphide ore body similar to those described in association with the Archean spinifex-textured komatiitic flows at Kambalda [Naldrett, 1981; Keays et al., 1982; Groves et al., 1986]. The Ni/Fe and Ni/Co ratios and the clustering of the Ni/Ir ratio of BA around the "Ni-sulphide

lOO

0.1 0.01

Volcanic'exhalative/ I I I) o •:•iii?•'"':•].'"•i-•'.. :. hydrothermal ores I J,;'

I I I I I I I I I I i LUl•l.•' '• • ..':' i [ i [ i! [ !: :?": :: :• i:•. :.:....:. [ !.j" .x.•'"':::?::: •••' .'"':J"

./.. ..... ½/,'.o ...:..•".;4,•:x

•'.:::::::::::.' •Hi!ii!!!:' .:::::::

./..:."/-'"i!•'""'":• *t'• ':!•••:•' ß GROUP A •.::::::::::..::::::.::::::.•, x...:•iiii'-"'"'•:::•'•:•••} :•' o GROUP B

0.1 I 10 100 I 000

Ir (ppb)

Fig. 13. Comparison of the Ni/Ir ratio of BA with the NiS-trend line of Keays et al. [1982] and with the distributions of that ratio in volcanic-exhalative/hydrothermal Ni ores from the Yilgarn [Keays et al., 1982].

trend line" of Keays et al., [1982] (see Figure 13) could support such a model. Also, PGE trends of BA group A resem- ble the flat patterns observed for the Ni-Cu sulphide ores associated with Archean komatiites [Naldrett, 1981; Keays et al., 1982] (see Figure 11).

The major problem with this model concerns the field re- lationships. At Kambalda the massive Ni-Cu ores are always associated with komatiitic flows and feeders to such flows

(the sulphide mineralization is attributed to contamination of the magmas by surrounding sediments [Groves et al., 1986]), while the underlying dunites carry only weakly disseminated sulphides [Groves and Keays, 1979]. The host rocks of BA are thought to represent deep-level ophiolitic peridotites [de Wit et al., 1987a; de Wit and Tredoux, 1988], stratigraphically placed below the above mentioned se- quences (see point 2 above), and would therefore be even less likely to host massive ore bodies. In general, residual ultramafics in ophiolites have very low concentrations of sulphides [Bird and Basset, 1980]. Further problems are (1) that some of the chemistry of B A, with Cr < 200 ppm and Cu < 200 ppm, does not compare well with that of the Archean sulphide deposits (Cr > 1000 ppm, Cu > 2% [Nal- drett, 1981]); and (2) major desulphurization would be necessary but is not indicated by the mineralogical evidence.

2. Formation of BA due to alteration of a chromitite

pod, such as are often associated with the lowermost peridotites of ophiolites [Page and Talkington, 1984]. The geological setting and morphology of BA would favor this model. However, although trevorite does replace chromite in BA (see Figure •b), it must be stressed that such replace- ments form a very small part of the mineralogy. Fur- thermore, the chromitite pods of ophiolites have very distinctive PGE patterns which do not match the patterns of BA (compare Figure 12 with Figure 10).

$. Formation of BA due to serpentinization during mid- ocean-ridge-type metamorphism. It has been shown that the rocks of the Jamestown ophiolite complex underwent extensive hydration and serpentinization closely following their formation [de Wit et al., 1987a]. These processes could be invoked to explain the association of bunsenite (originally Ni-Fe metal?) with serpentinized liebenbergite, following the suggestion that awaruite (Ni•Fe) forms during serpentinization of olivine [Ramdohr, 19•0; Dick, 1974]. Although


serpentinization is pervasive in and around BA, the petrography of BA indicates that the liebenbergite-bunsenite assemblage predates the formation of nepouite from liebenbergite. The presence of the bunsenite therefore can not be attributed to the serpentinization of the Ni-olivine. The ad- vantage of this model, i.e., that a desulphurization process is not needed, is offset by the problem that awaruite is not known to occur in a massive form but only as finely disseminated granules. Furthermore, one analysis of awaruite (M. Tredoux, unpublished data, 1986) shows that this mineral is barren of PGE and three analyses of josephinite (which contains a Ni-Fe metal phase and has been likened by some to awaruite [Dick, 1974]), as listed in Table 3b, show that the PGE trend for josephinite is very different to the patterns encountered in BA (see Figure 11). It should be noted that at mid-ocean ridges, large quantities of hydrothermal fluids do not penetrate extensively as far down as the cumulate and residual portions of the oceanic crust (as also seen in ophiolites) [Stakes et al., 1986; Dewey en Kidd, 1977; Stern et al., 1976]. Fluid/rock ratios in the BA host environment would therefore have been low, as supported by the fact that the oxygen isotope ratios of the serpentinized ultramafics of the nearby Stolzberg complex approach mantle values [Hoff- man et al., 1986; de Wit et al., 1987a]. Under such conditions, large-scale redistribution of the PGE is unlikely. Oshin and Crocket [1986] show that the mobility of PGE in this environment is negligible.

4. Formation of BA due to dynamothermal metamorphism during emplacement of the granites. There is ample evidence that BA was affected by the late metamorphic events in the greenschist-amphibolite fades and associated deformation recorded along margins of the greenstone belt. Metasomatism in the BA environment has undoubtedly occurred, as suggested by the dose association between BA and the tourmaline-hornblende schist and by the presence of a borate mineral (bonaccordite) in BA. The most likely source for the boron-beating fluids is the nearby syntectonic granitoid plutons, which preferentially exploited the thrusts and mylonites dose to BA [de Wit et al., 1987a]. The concordance of the Sm-Nd and Pb isotopic data of the BA group A samples with the 3510 Ma reference line would be difficult to explain if BA was formed at the time of the intrusion of the Stentor granite at circa 3300 Ma, but BA might have been altered by the later metasomatic fluids. Higher fluid/rock ratios in the more schistose margin of the body could account for the elevated radiogenic Pb and the disturbance of the Sm/Nd isotopic systematics of the group B samples. It can then be argued that preferential leaching of Os and Ir from this schistose rim might account also for the fractionation in PGE patterns. However, although some Pd might be mobilized, with Ni, during metamorphism, in general, the PGE appear to be more immobile than Ni in metamorphic environments [Keays et al., 1982; Groves and Keays, 1979]. PGE immobility has been challenged by Bowles [1986] and Barnes et al., [1985]; the latter suggested that the negative Pt anomalies in Archean Ni-Cu sulphide ores (Figure 11) were caused by metamorphic alteration. Metasomatic reworking of disseminated Ni-sulphides into massive ores has also been suggested to have occurred in the Yilgarn block [Groves and Keays, 1979]. The paucity of information regarding the mobility of the PGE in metasomatic fluids makes it virtually impossible to evaluate this

model. In the case of BA it is dear that the fractionation of

Ni between the two groups has been minimal (see Figure 8). Even if this can be entirely attributed to metasomatic redistribution, the amount of associated movement of Os and Ir would have been negligible. It is notable, too, that the work of Oshin and Crocket [1986] indicates that Ir, Pt, and Pd were not mobilized in the Thetford Mines ophiolite, "despite pronounced thermal metamorphism". The lack of evidence for desulphurization in BA should once again be noted, as sulphides are a major component of the rock in those cases where this mechanism is invoked.

5. Hydrothermal/volcanic-exhalative models for the formation of BA, such as have been proposed for some of the Ni-sulphide ores in the Yilgarn block, Western Australia [Keays et al., 1982], and elsewhere (Rathburn Lake deposit north-eastern Ontario [Rowell and Edgar, 1986]). Such ores tend to have distinctly different fields of Ni/Ir (and, although not plotted, also Ir/Pd) ratios than those of the BA deposit (Figure 13). Moreover, some of the general geochemistry of BA (Co > 5000 ppm, Cu < 100 ppm) does not compare well with those of the volcanic-exhalative/hydrothermal ores of the Yilgarn block which are generally depleted in siderophiles and enriched in chalcophiles (i.e., Co < 300 ppm, Cu > 1000 ppm) and have S --- 9% [Keays et al., 1982]. Although PGE data on such deposits are generally very incomplete, it appears that they are depleted in Os, Ir, and Ru relative to Pt and Pd [Rowell and Edgar, 1986; Keays et al., 1982]. The PGE patterns can therefore be expected to have much steeper positive slopes than the BA group B samples. The lack of evidence for desulphurization should again be noted, as well as the fact that BA is not associated with volcanic rocks.

To conclude, it is important to emphasize that the occurrence of coexisting Ni-Co-Cr spineIs, Ni-olivine, and Ni- oxide is, to our knowledge, unique to the BA occurrence. Although there is very little theoretical or experimental data on liebenbergite/Ni-Co-Cr spinel assemblages, it is not likely that these minerals would have equilibrated under the prevailing conditions (< 5 kbar, < 700øC, high P(H20)) of greenschist to lower amphibolite fades metamorphism which affected the ultramafic host rocks after their extraction from

the mantle. Also, none of the above models can account for the HM body along strike from BA.

Core Formation and Its Implications for BA An alternative to these "crustal" (redistribution) models is

that BA was already concentrated into a massive, Ni-PGE enriched form in the asthenospheric mantle prior to its emplacement into what was then lithospheric mantle (presently part of the cratonic crust). The extreme enrichment of Ni in BA, the absence of sulphides, and the lack of any evidence that BA ever had a sulphide mineralogy, all lead us to concur in part with de Waal [1978], i.e., that BA originally was a metallic mass. However, for those reasons outlined above, we do not accept the interpretion of B A as a paleometeorite, and therefore we are led to explore other mechanisms whereby a metallic mass could have formed.

Further observations that must be addressed are (1) the current mineralogy of BA shows that any postulated metallic part of the body has been completely oxidized; its metallic character can be indirectly deduced only from the bunsenite remnants; (2) the immobility of Os and Ir (especially in the


metallic state) in low-temperature geological environments is well established [Keays et al., 1982; Westland, 1981; Groves and Keays, 1979], and the observed fractionation of more than an order of magnitude for those elements between the two BA groups could thus point to initial conditions where the prevailing temperature was very high, perhaps • 2000øC; see Figure 10 for the melting points of the PGE.

In order to accommodate these constraints, and bearing in mind that a sulphide "precursor" is counter-indicated by the mineralogy and petrography, we suggest that BA might have originated from a metal-enriched mass, which was formed in the lower mantle during inefficient separation of Fe-Ni alloy (destined for the core) from the protomantle. Such a model offers more satisfactory explanations for both the unusual chemistry-mineralogy of BA and its endogenous relationship with the surrounding rocks than any of the possibilities discussed in the previous section.

The mechanism by which the Fe-alloy core of the Earth was formed is a matter of continuous debate in the literature

[e.g., Ringwoo& 1977; Stevenson, 1981; Brett, 1984; Cooper- man and Kaula, 1985; Jones and Drake, 1986]. The main contending theories, which all attempt to offer a feasible explanation for the chemical disequilibrium (of siderophile elements) between the core and mantle [e.g., Ringwood, 1977; Arculus, 1985; Jones and Drake, 1986], are (1) addition of a "late veneer" (+1% of the Earth's mass) of oxidized material which did not partake in the core-forming event [e.g., Chou, 1978; Newson and Palme, 1984; Morgan, 1986]; (2) equilibration between the upper mantle and an Fe-O-S liquid [Brett, 1984]; (3) incomplete separation of Fe-Ni metal from the mantle, i.e., that core-mantle fractionation was a disequilibrium event [e.g., Stevenson, 1981; Jones and Drake, 1986].

The merits and demerits of these models are discussed in

detail by Jones and Drake [1986], who conclude that although the current data base is not complete, disequilibrium during core formation seems to be most reasonable. An "endogenous" explanation was also favored by Arculus [1985] in a recent review. Following their arguments, the model for subsequent discussion will be therefore essentially the one described by Stevenson [1981], i.e., a multistage, disequilibrium process of continuous "downward refining", which has been active during a long period of geological time (see Figure 14).

The proposed model allows for addition of extraterrestrial material after the main core-forming event, though not in the way envisaged by the "late veneer" theory. Proto-BA is envisaged as alloy that separates out from surrounding silicates at depth (below the iron-wustite (IW)fO2 buffer) either (1) in the residual lower mantle after initial core formation (case 1, Figure 14d); or (2) from material which ac- creted and sank after the main core event formation (case 2, Figure 14d). Either of these will suffice for the argument of this paper, which only requires that a siderophile-enriched inhomogeneity [Stevenson, 1981] develops at depth. How- ever, it is a function of the model that the alloy-silicate separation never goes to completion, so that the Fe-Ni metal remains intimately associated with the surrounding silicate.

It must be noted that the nonchondritic siderophile element ratios of BA (e.g., Ni/Fe, Ni/Co, Niflr, Pd/Ir, Ru/Os; see Figure 3c) do not support an association of BA with the "late veneer" model: if BA represented an incompletely as-

similated remnant of a "late veneer", it should have retained chondritic ratios. On the other hand, the enrichment of Ni in BA (relative to Fe and Co) is compatible with the later stages of disequilibrium core formation. The oxidation state of the mantle is speculated to be inhomogeneous [Arculus, 1985] and to have changed (become more oxidizing) with time [Arculus, 1985; Jones and Drake, 1986]. Late stage "downward refining" metal, in a more oxidized mantle, would have been Ni-enriched relative to the earlier alloys [Jones and Drake, 1986; Pernicka and Wasson, 1987], and it is unlikely that chondritic ratios of elements with differing siderophile tendencies (e.g., Co and Ni) would have been preserved [Jones and Drake, 1986].

Although core material itself may be too dense to rise through the mantle and diffusivity of siderophiles from the core into the mantle is much too slow [Loper et al., 1988] to explain the formation of a proto-BA, such constraints are not relevant if long-lived siderophile-enriched heterogeneities are present within the lowermost layer of the mantle (the D"layer). Large seismic heterogeneities within this zone almost certainly do exist [Lay, 1987].

Silicate which contains a small amount of (interstitial?) metal could rise with an upwelling plume generated near the D"layer [cf. Stacey and Loper, 1983; Olson et al., 1987; Loper et al., 1988]. (If the alloy:silicate ratio is low (e.g., < 1% alloy), then the density of such siderophile-enriched material will be only fractionally heavier than that of the bulk of the D"material.) Loper et al. [1988] suggest, on experimental and theoretical grounds, that surges of hot material from the D"zone are likely to travel up preexisting plume traces at least as fast as 1-4 m/yr; such material would reach the Earth's surface in 1 m.y. or less. In this way, disseminated siderophiles in a significant volume of D"-derived mantle material may be moved toward the surface within such rising diapirs.

The siderophiles could be concentrated into pods during the ascent, particularly in the higher mantle regions, because of the large-scale partial melting of the plume material that is necessary for the formation of the simatic lithosphere. The chemical fractionation of such an evolving metal-silicate system would be affected by changes in temperature and JD2. Initially, at high temperature and low JD2 the system would have continued to equilibrate, with the highly siderophile PGE partitioning out of the silicates and into the Fe-Ni metal. With decreasing temperature the rate of equilibration of Os and Ir, which (in the metallic state) have very high melting points (see Figure 10), probably would have been the first to abate, so that differences in concentration between the original "metallic" (type A) and "silicate" (type B) parts would be greatest for Os and Ir and least for Pt and Pd (lowest melting points; Figure 10). It is difficult to speculate at what stage the Os-Ir retardation might have occurred. There is growing evidence that Os-Ir alloys are essentially unaffected by the temperature regime of the upper mantle [e.g., Davies and Tredoux, 1985; Bacuta et al., 1988], and experimental work of Blurn et al. [1988] indicates that at 1273 K, exsolution of Ir from Ni-Fe metal is much more sluggish than that of Ru. In BA, Ru (MP = 2334øC) equilibrated to a much greater extent than Os and It. We therefore tentatively suggest that the Os-It retardation began before BA passed through the 2300øC isotherm, i.e., at an approximate depth of 1000 km [Stacey, 1977; Ernaliani,


a

4.6 - 4.5 Ga

rockberg

c

•• Fe-rich alloy Cold primordial mantle

Convective mantle

•, Convecting currents • Oceanic lithosphere

• Continental lithosphere

case 2

•- impacting projectile

..•..,., ,.• :.•..• ,:. •.....:?• downward •:':!• •:r;'•, .'•' i refining blobs

Rayleigh-Taylor instability

d e

convecting mantle reservoir

/ \ V v

4.5 - 4.0 Ga 4.0 - 3.0 Ga

Fig. 14. Schematic model of core-mantle segregation, modified after Stevenson [1981] and subsequent mantle-crust segregation. This envisaged sequence of events is as follows: (a) separation of Fe alloy from the hot outer Earth; (b) development of gravity instability; (c) sinking of Fe alloy to the center to replace cold, undifferentiated material; this is as- sumed to happen after +80% of accretion; (d) the undifferentiated mantle starts to equilibrate with the core and overlying mantle and the remaining 20% of accretion continues. Fe ahoy (progressively more Ni-rich) continue to separate to the core, but equilibration does not occur; (e) start of mantle-crust segregation. Incomplete core formation results in anomalous enrichment of siderophiles in the lowermost mantle, which can be transported upward in mantle plumes.

1987]. The order of magnitude differences in Os and Ir concentrations between the BA groups are thus interpreted as a record of a high-temperature (deep mantle) phase in the evolution of BA.

We speculate that as the plume moved to higher mantle levels and the fO2 continued to increase, the metallic bodies became oxidized. The work of Blurn et al., [1988], on the oxidation of metallic Ni-Fe-Pt-Ru-Ir systems, shows that magnetite forms first, while Ni and the PGE move away from the oxidation front. As a corollary, that part of the metallic BA body which was farthest from the oxidation front (the centre) would have become progressively more enriched in Ni relative to Fe. Once the Ni-NiO buffer was

exceeded, the Ni-enriched central part would finally be oxidized, resulting in the formation of bunsenite (NiO), with rare Ni-rich silicate (liebenbergite) and spinels (nichromite, cochromite). At this mantle level (---180 km and less) the density of an oxidized BA-type pod (5.2 gcm -3, assuming a metal: silicate mix of 30: 70) would now be less than that of the mantle package that originally left the D" layer (5.59 gcm -3) [Ernaliani, 1987].

During the oxidation of the BA body the PGE also would

have migrated toward the center [Blum et al., 1988]. How- ever, it is possible that a portion of the Pt might have remained in the Fe-rich envelope. In geological systems, Pt has a greater affinity for Fe (rather than Ni) than the other PGE [Kinloch, 1982]; and Westland [1981] indicates that Fe- Pt alloys are thermodynamically very nonideal, thus making it possible for such alloys to coexist with FeO. In this way, the body could have developed a Ni-rich part, enriched in PGE (but with a negative Pt anomaly), and a magnetitic part, perhaps containing some trapped Fe-Pt alloy. This exemplifies the differences between BA and the HM body, especially regarding the complementary PGE trends.

To summarize the scenario described above, BA rep- resents a dynamically emplaced heterogeneity which evolved, at least in part, due to small degrees of continuous partial melting during ascent of a 3.5 Ga lower mantle dia- pit. Subsequently, BA was embedded in the residual lithospheric mantle during the formation of the Jamestown ophiolite complex; and the entire complex was later ob- ducted into the sialic environment [de Wit et al., 1987b]. BA therefore was, at all times of its evolution, intimately associated with its surrounding mantle. That BA was involved in

TREDOUX ET AL.'. PLATINUM GROUP ELEMENTS IN A 3.5 GA NICKEL-IRON OCCURRENCE 811

SEAFLOOR SPREADING/HYDROTHERMAL ALTERATION --•- GRANITOID - GREENSTONE - TERRAIN

•1• calc' alkaline [f • • • .......... SEA LEVEL ........ _vol_ca_nics - _x/• ..............

late chromitite

restite chromitite UPPER MANTLE

BA becomes oxidized

,., 25OOøC

PROTO- CRATON ENRICHED IN PGE

LOWER MANTLE BA probably fractionates

Os and Ir

SIDEROPHILE-RICH PODS RESULTING FROM INCOMPLETE CORE FORMATION (BON ACCORD PRECUSOR)

Spinel-oxide, enriched in siderophile elements Mantle peridetite Magma chamber MORB

Tonalite

Fig. 15. Schematic diagram of the proposed model wherein BA is derived from the deep mantle and transported into the lithosphere during mantle upwelling. The diagram also shows how the early thickening of cratonic "keels" might concentrate PGE-rich pods.

the simatic lithosphere-forming processes until at least 3.5 Ga, is borne out by the close affinity of the BA radiogenic isotope data (Pb-Pb, Sm-Nd) with other data arrays from that area. However, BA itself might have been too refractory to be chemically affected to any great extent by asthenosphere-lithosphere segregation processes. Mantle deformation or later crustal tectonic activity might have separated the Fe-rich part from the Ni-rich portion, thus forming at least two discrete bodies which we now refer to as the BA ore body and the HM body.

The proposed model of deep mantle origin for BA can explain much of the chemical fractionation within the BA body and between BA and the HM body, as well as the close association between BA and the other mantle-derived rocks

of the area. It further suggests that partial melting of the upwelling Archcan asthenosphere destined for spreading centers may have started at lower mantle depth (Figure 14e). If this is correct, spreading centers at > 3.5 Ga were perhaps fed from a one-mantle reservoir. This would be consistent

with a more vigorous early mantle convection related to much greater radiogenic heat production in the Archcan mantle than today.

CONCLUSIONS

It is suggested that the BA rocks were derived from a siderophile-silicate heterogeneity formed in the deep mantle during disequilibrium core growth. The geochemistry of BA and its subsequent alteration history may thus help to yield a better understanding of the geochemistry of the core and the D" layer. Should this model be correct, then at least the following two "predictions" should hold:

1. BA should not be unique, and similar deposits should occur in other, especially older, ophiolites. It is therefore in-

teresting to note that trevoritc-rich material with relatively flat PGE patterns has been reported in the Archcan Puddy Lake ultramafics in northwestern Ontario (C. T. Barrie, personal communication, 1987). Bird and Basset [1980] envisage a similar deep-mantle origin for Os-Ir-Ru alloys found with josephinite (a Pd-rich Ni-Fe alloy) in the Mesozoic Jose- phine ophiolite complex, western United States. The high PGE concentrations in the Shetland ophiolite are also interesting, as they may, in part, have been derived from previously depleted mantle, as indeed the suprasubduction model for this ophiolite implies [Prichard and Lord, 1987].

2. It has been suggested [de Wit, 1986; de Wit and Tre- doux, 1988] that during Mid-Archean tectonism, simatic lithosphere of southern Africa formed intraoceanic thrust stacks which continuously underplated granite-greenstone terrains to form mafic-ultramafic "keels" to Archcan conti-

nental lithosphere. Analyses of 3.3-3.2 Ga diamonds and the geochemistry of associated ultramafic xenoliths found in Mesozoic kimberlites yield independent evidence that in southern Africa, such Archcan "keels" reached depths of at least 150 km [Richardson et al., 1984] and retained this minimal thickness throughout further development of the lithosphere which underlies the southern African craton [Jordan, 1981; Nickel and Green, 1985; Haggetty, 1986; de Wit, 1986; MacGregor and Manton, 1987]. Inference from our work is that this thick lower (mantle) lithosphere could be a potential PGE store and hence a contributor to post-3.0 Ga igneous rocks, should their parent magmas pass through or originate within this region. Phanerozoic kimberlitic magmas have indeed penetrated this keel episodically between 600 and 60 Ma; some associated diamonds contain up to 160 ppm of iridium (no other PGE analysis have been quan- tified) [Sellschop, 1979]. This is consistent with an ocean-

812 TREDOUX ET AL.: PLATINUM GROUP ELEMENTS iN A 3.5 GA NICKEL-IRON OCCURRENCE

floor-derived, PGE-enriched, Archcan mantle lithosphere. Figure 15 summarizes this suggestion.

Acknowledgments. We have benefited by discussions with T. J. Ahrens, L. D. Ashwall, E. Barton, H. Bergh, J. M. Bird, K. Burke, M. J. Drake, R. A. Hart, S. R. Hart, E. Jagoutz, J. Keenan, K. Kenyon, T. Molyneux, T. Pearton, C. Stern, S. R. Taylor, S. A. de Waal, and C. J. Wasserburg. J. M. Bird generously gave us many samples of josephinite, and G. A. Challis kindly gave us a sample of awaruite. The Sm-Nd isotopic analyses were done by Rodger Hart while on sabbatical leave at the Department of Earth, Atmospheric, and Planetary Science, MIT, and S. R. Hart is thanked also for the Re data. The TAMS group of the University of Rochester (under the leadership of U. Fehn) is thanked for providing the Os isotopic data. Our thanks are due to D. Buchanan, P. Suddaby, N. Royale, T. McCarthy, P. Chipkin, S. Gates, and L. de Matos for analytical help and to D. Mthembu and M. Dickey for typing the manuscript. The paper was improved by constructive criticism from S. E. Hag- gerty, C. J. Hawksworth, J. Kruger, D. Malvin, J. Morgan, R. W. Nesbit, and L. O. Nicolaysen. The research was funded through the Foundation of Research Development and the University of the Witwatersrand. Part of the work was finished while Maarten de Wit

was a visiting scientist at the Lunar and Planetary Institute, which is operated by the Universities Space Research Association under con- tract NASW 4066 with the National Aeronautic and Space Adminis- tration. This is LPI contribution 658.

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R. A. Armstrong, Department of Geochemistry, University of Cape Town, Private Bag, Rondebosch, 7700, Republic of South Africa.

M. J. de Wit, Bernard Price Institute for Geophysical Research, University of the Witwatersrand, Johannesburg, P.O. Wits, 2050, Republic of South Africa.

R. Hart, N. N. Lindsay, J.P. E. Sellschop, and M. Tredoux, Wits-CSIR Schonland Research Centre for Nuclear Sciences, Uni- versity of the Witwatersrand, Johannesburg, P.O. Wits, 2050, Re- public of South Africa.

(Received December 3, 1987; revised July 1, 1988;

accepted July 11, 1988.)

Platinum group elements in a 3. 5 Ga nickel-iron occurrence: Possible evidence of a deep mantle...

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