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JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 1
PAGES 103–128
2002
Platinum-group Elements and
Microstructures of Normal Merensky Reef
from Impala Platinum Mines, Bushveld
Complex
SARAH-JANE BARNES1∗ AND WOLFGANG D. MAIER2
1
SCIENCES DE LA TERRE, UNIVERSITÉ DU QUÉBEC, CHICOUTIMI, QUÉ., G7H 2B1, CANADA
2
CENTRE FOR RESEARCH ON MAGMATIC ORE DEPOSITS, DEPARTMENT OF EARTH SCIENCES,
UNIVERSITY OF PRETORIA, PRETORIA 0002, SOUTH AFRICA
RECEIVED NOVEMBER 13, 2000; REVISED TYPESCRIPT ACCEPTED JULY 9, 2001
The Merensky Reef of the Bushveld Complex contains one of the
world’s largest concentrations of platinum-group elements (PGE).
We have investigated ‘normal’ reef, its footwall and its hanging
wall at Impala Platinum Mines. The Reef is 46 cm thick and
consists from bottom to top of leuconorite, anorthosite, chromitite
and a very coarse-grained melanorite. The footwall is leuconorite
and the hanging wall is melanorite. The only hydrous mineral
present is biotite, which amounts to 1%, or less, of the rock. All
of the rocks contain 0·1–5% interstitial sulphides (pyrrhotite,
pentlandite and chalcopyrite), with the Reef rocks containing the
most sulphides (1–5%). Lithophile inter-element ratios suggest that
the magma from which the rocks formed was a mixture of the two
parental magmas of the Bushveld Complex (a high-Mg basaltic
andesite and a tholeiitic basalt). The Reef rocks have low incompatible element contents indicating that they contain 10% or
less melt fraction. Nickel, Cu, Se, Ag, Au and the PGE show good
correlations with S in the silicate rocks, suggesting control of the
abundance of these metals by sulphides. The concentration of the
chalcophile elements and PGE in the silicate rocks may be modelled
by assuming that the rocks contain sulphide liquid formed in
equilibrium with the evolving silicate magma. It is, however, difficult
to model the Os, Ir, Ru, Rh and Pt concentrations in the chromitites
by sulphide liquid collection alone, as the rocks contain 3–4 times
more Os, Ir, Ru, Rh and Pt than the sulphide-collection model
would predict. Two possible solutions to this are: (1) platinumgroup minerals (PGM) crystallize from the sulphide liquid in the
chromitites; (2) PGM crystallize directly from the silicate magma.
To model the concentrations of Os, Ir, Ru, Rh and Pt in the
chromitites it is necessary to postulate that in addition to the 1%
sulphides in the chromitites there is a small quantity (0·005%) of
cumulus PGM (laurite, cooperite and malanite) present. Sulphide
liquids do crystallize PGM at low fS2. Possibly the sulphide liquid
that was trapped between the chromite grains lost some Fe and S
by reaction with the chromite and this provoked the crystallization
of PGM from the sulphide liquid. Alternatively, the PGM could
have crystallized directly from the silicate magma when it became
saturated in chromite. A weakness of this model is that at present
the exact mechanism of how and why the magma becomes saturated
in PGM and chromite synchronously is not understood. A third
model for the concentration of PGE in the Reef is that the PGE
are collected from the underlying cumulus pile by Cl-rich hydrous
fluids and concentrated in the Reef at a reaction front. Although
there is ample evidence of compaction and intercumulus melt migration
in the Impala rocks, we do not think that the PGE were introduced
into the Reef from below, because the rocks underlying the Reef are
not depleted in PGE, whereas those overlying the Reef are depleted.
This distribution pattern is inconsistent with a model that requires
introduction of PGE by intercumulus fluid percolation from
below.
∗Corresponding author. E-mail: [email protected]
 Oxford University Press 2002
Merensky Reef; platinum-group elements; chalcophile elements; microstructures
KEY WORDS:
JOURNAL OF PETROLOGY
VOLUME 43
INTRODUCTION
The Merensky Reef of the Bushveld Complex, South
Africa, together with the UG-2 chromitite, constitutes
the largest resource of platinum-group elements (PGE)
on Earth (Vermaak, 1995). In spite of its economic
importance, the process by which the Reef became
enriched in PGE and formed is still poorly understood.
Three models have been proposed for the origin of reeftype PGE deposits: (1) the PGE were collected by sulphide
droplets and settled onto the crystal pile to form a PGErich layer (e.g. Campbell et al., 1983); (2) the PGE are
pre-concentrated in the silicate liquid by clusters of PGE
ions, which are then incorporated in sulphide droplets
or crystallize directly from the magma as platinum-group
minerals (PGM) (Tredoux et al., 1995; Cawthorn, 1999;
Ballhaus & Sylvester, 2000); (3) the PGE are collected
by rising intercumulus liquid (Boudreau & Meurer, 1999b;
Willmore et al., 2000) and precipitated at a reaction front
during compaction and cementation of the crystal pile.
To place some constraints on these three models we
collected 24 oriented samples from borehole core from the
farm Reinkoyalskraal 278 JQ (Fig. 1), covering 186 cm of
uninterrupted stratigraphy from the footwall into the
hanging-wall rocks of ‘normal’ Reef. The term ‘Reef’ is
a mining, rather than a lithological term, and refers to
the layer within the Merensky Unit that is enriched in
PGE (Leeb-du Toit, 1986). This layer may contain more
than one rock type. On the scale of the Bushveld Complex, the Reef has a number of distinct facies: ‘normal
reef’, ‘rolling reef’, ‘pothole reef’ and ‘wide reef’ (Viljoen,
1999). The ‘normal’ reef shows the least degree of transgression of the silicate stratigraphy, and is considered to
be the facies least altered by subsolidus processes. In this
study we first present petrographic descriptions of the
samples with the aim of understanding what primary
textures are preserved and how much subsolidus modification and hydrothermal alteration the rocks have experienced. PGE and chalcophile element concentrations
are then considered to evaluate the degree of sulphide
liquid control over the PGE. Chrome concentrations are
used to investigate whether chromite controlled the PGE
distribution, and the lithophile elements are used to
evaluate the role of the liquid component.
STRATIGRAPHY
The ultramafic and mafic rocks of the Bushveld Complex
are referred to as the Rustenburg Layered Suite (South
African Committee for Stratigraphy, 1980). The Suite is
generally subdivided into five zones (Hall, 1932): the
basal Marginal Zone (gabbro–norites), overlain by the
Lower Zone (peridotites and pyroxenites), the Critical
Zone (pyroxenites and chromitites), Main Zone (gabbro–
NUMBER 1
JANUARY 2002
norites) and Upper Zone (anorthosites, diorites and magnetitites). The platinum-bearing reefs (UG-2 and
Merensky) occur towards the top of the Critical Zone
(Fig. 1). The Bushveld Complex crops out as three lobes,
the western, northern and eastern lobes. Impala Platinum
Mines are located in the western lobe. The general
stratigraphy of the Bushveld Complex in the area of the
farm Reinkoyalskraal has been described by Cousins &
Feringa (1964) and by Vermaak (1976). At most localities
the Upper Critical Zone consists of eight cyclic units
(Scoon & Teigler, 1994). In most cases, each unit consists
of chromitite, overlain by harzburgite or pyroxenite,
norite and anorthosite. The economically important platiniferous horizons (UG-2 chromitite and Merensky Reef )
are located at the base of cyclic units 5 and 7. At the
beginning of each cyclic unit there is a prominent reversal
in the trend of Fe enrichment, both in the whole-rock
and mineral compositions; this feature is interpreted, by
many researchers, to indicate magma replenishments
(e.g. Kruger & Marsh, 1982; Eales & Cawthorn, 1996).
PETROGRAPHY
A detailed description of the Merensky Reef and its
immediate footwall and hanging wall at the Impala Mines
has been given by Leeb-du Toit (1986), and its platinumgroup mineralogy has been described by Mostert et al.
(1982). Leeb-du Toit (1986) showed that the Reef displays
considerable regional and local lithological variation.
The section of core studied here corresponds to normal
Merensky ‘A’ type Reef in the Leeb-du Toit (1986)
nomenclature, i.e. the Merensky unit does not transgress
into its footwall. Using the Impala cut-off grade of 1·5
ppm Pt + Pd to define the Reef, it begins 23 cm below
the lower chromitite in the leuconorite and ends 23 cm
above the chromitite in the melanorite (Table 1 and Fig.
2). Reef lithologies observed are leuconorite, anorthosite,
chromitite and a coarse-grained melanorite (Fig. 3). In
the Bushveld literature and local mining terminology,
the melanorite is referred to as a pegmatoidal pyroxenite,
based on the interstitial nature of the plagioclase and
following Irvine’s (1982) nomenclature. In naming our
rocks we have used international standard nomenclature
(IUGS) so as to use the same type of nomenclature for all
rock types. The melanorite does not have a pegmatoidal
texture and the grain size (5–20 mm) is slightly less
than that required for the use of the term pegmatite;
furthermore, many of the large grains are composite, so
we find the term pegmatoid misleading and will use
coarse-grained instead.
In this study the core was sampled continuously. The
length of the core used for each sample depended on
whether the sample contained visible sulphides and
whether there was any change in lithology. In the vicinity
104
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Fig. 1. Location of Impala Platinum Mine and of the borehole (bh). RSA, Republic of South Africa.
of the chromitites, which are normally considered to
mark the Reef, the samples taken were 5 cm in length,
except where the lithological unit was narrower than
5 cm, in which case the sample was the width of that
particular unit (see Table 1). Above and below the Reef
the samples analysed were 10 cm in length. The borehole
was vertical and the thin sections were oriented so that
long axis of the section represents the palaeovertical.
Oriented polished thin sections were made of each
sample.
The top 30 cm of the core is norite (Fig. 2). The
contact with the underlying melanorite is gradational.
Medium-grained (3–6 mm) subhedral orthopyroxene
makes up >50% of the rock. Subhedral medium-grained
(2–4 mm) plagioclase laths are the second most common
phase (>40%). The grains show prominent rectilinear
twinning and some are compositionally zoned. Large
clinopyroxene oikocrysts (1–2 cm) enclose both the plagioclase and orthopyroxene. Small interstitial grains of biotite make up >1% of the rock. The sulphide phases
present in all the rock types of the core consist of
intergrowths of the three phases pyrrhotite, pentlandite
and chalcopyrite in the proportions 5:3:2 (Fig. 4a) and
occur interstitially to the cumulus phases. Hereafter they
will be referred to collectively as sulphides. In the norite
sulphides 0·5–1 mm in size make up >0·5% of the rock.
Downwards, the next 80 cm of core is a coarse-grained
melanorite (Fig. 3). Subhedral orthopyroxene (1–7 mm)
is the principal phase (>65%). The orthopyroxene crystals are densely packed, many are interpenetrating (indented) on faces oriented at a high angle to the vertical
(Fig. 4b), show kinking (low-angle grain boundaries) and
undulose extinction, and locally develop subgrains. In
some samples (7, 9, 11) orthopyroxene crystals are broken
by subvertical fractures that are filled with undeformed
plagioclase. Plagioclase (>20%) generally occurs as large
oikocrysts (up to 10 mm). Like the orthopyroxene, the
plagioclase oikocrysts are deformed and exhibit undulose
extinction, and characteristic tapered or spindle-shaped
deformation twins. Where the plagioclase fills the cracked
orthopyroxene crystals, however, and where it occurs in
the pressure shadows of large clinopyroxene oikocrysts,
it does not show evidence of plastic deformation (Fig. 5).
Clinopyroxene (>10%) forms large (10–20 mm) anhedral
grains that poikilitically include the orthopyroxene. Biotite (1–2%) and quartz (p1%) are interstitial to pyroxene and do not appear to be deformed (no undulose
extinction or kinks in the biotite). Chromite (<1%) is
present as small (0·1 mm) cubic grains. Sulphides form
interconnecting networks with their longest axes parallel
to the vertical (Fig. 4b). The highest concentration of
sulphides is found at the base of melanorite (>5%). The
105
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 1
JANUARY 2002
Table 1: Composition of rocks from the Impala borehole and from the margins of the intrusion
Sample:
1
3
5
7
9
11
13
15
17
18
Height (cm):a
108·00
10·00
98·00
10·00
88·00
10·00
78·00
10·00
68·00
10·00
58·00
10·00
48·00
10·00
38·00
10·00
28·00
10·00
23·00
5·00
norite
norite
norite
Length (cm):b
Rock type:
mela-
mela-
mela-
mela-
mela-
mela-
mela-
norite
norite
norite
norite
norite
norite
norite
wt %
SiO2
TiO2
Al2O3
51·99
0·21
52·79
0·20
53·11
0·18
53·18
0·23
52·73
0·41
53·13
0·21
53·88
0·26
54·25
0·26
53·83
0·24
53·62
0·22
12·50
9·60
FeO∗
7·65
8·41
8·03
6·00
6·15
6·62
6·21
5·93
6·22
6·26
8·93
10·07
9·49
9·33
9·93
9·97
9·65
Fe(S)
MnO
MgO
CaO
Na2O
0·20
0·16
17·11
7·20
9·58
0·23
0·17
19·82
6·36
0·30
0·19
21·86
4·99
0·48
0·20
23·63
3·94
0·87
0·19
22·42
4·86
0·55
0·20
23·05
4·03
0·19
22·72
3·92
0·19
23·04
4·23
0·19
22·78
5·04
0·17
0·20
23·55
4·15
K2O
P2O5
S
Cr2O3
LOI
Total
ppm
Ba
Cs
Hf
Rb
Sc
Ta
Th
U
V
La
Ce
Nd
Sm
Eu
Tb
Yb
Lu
ppm
Ag
As
Co
Cu
Ni
Sb
Se
ppb
Os
Ir
Ru
Rh
Pt
Pd
Au
0·93
0·74
0·58
0·46
0·52
0·49
0·52
0·51
0·52
0·49
0·21
0·02
0·16
0·29
0·13
98·75
0·21
0·02
0·19
0·34
0·08
99·16
0·17
0·01
0·26
0·35
0·34
99·29
0·15
0·02
0·45
0·36
0·34
99·50
0·24
0·02
0·73
0·38
0·3
99·31
0·22
0·02
0·46
0·36
0·39
99·04
0·33
0·02
0·09
0·36
0·38
98·79
0·34
0·09
0·37
0·34
99·60
0·23
0·14
0·07
0·38
0·3
99·59
0·17
0·09
0·14
0·37
0·14
99·15
49
<0·4
0·35
5
23·8
0·02
0·31
<0·3
106
2·28
5·28
1·80
0·55
0·29
0·10
0·61
0·101
36
0·23
0·34
5
26·4
0·04
0·45
0·11
107
2·64
5·50
2·00
0·66
0·25
0·13
0·74
0·104
28
0·26
0·32
1
25·4
0·05
0·66
0·17
102
2·70
3·08
1·70
0·58
0·24
0·10
0·63
0·093
39
<0·1
0·60
3
26·8
0·06
0·87
0·23
108
3·08
5·72
2·30
0·61
0·21
0·11
0·71
0·110
59
0·21
0·72
9
30·2
0·18
1·06
0·20
130
3·70
7·96
3·60
0·88
0·24
0·15
0·83
0·118
43
0·23
0·71
7
26·6
0·04
0·74
0·20
106
2·96
6·05
2·30
0·62
0·20
0·09
0·68
0·113
74
0·32
1·33
17
28·5
0·12
2·02
0·30
112
3·10
6·72
2·80
0·85
0·19
0·12
0·81
0·120
69
0·41
1·03
16
28·8
0·13
1·88
0·35
119
5·14
13·09
5·70
1·29
0·22
0·19
0·82
0·120
43
0·39
0·82
13
30·3
0·10
1·49
0·46
119
7·23
15·40
6·20
1·60
0·29
0·23
0·92
0·133
40
0·24
0·49
3
27·1
0·05
0·97
0·35
110
5·34
11·00
4·20
1·04
0·22
0·13
0·74
0·122
<0·5
<0·5
76
303
905
<0·02
0·6
<0·5
<0·5
86
359
1105
<0·02
0·7
<0·5
<0·5
95
437
1354
0·02
1
<1
1·08
<5
2·7
35
21
51
<1
1·07
8
1·8
50
20
75
1·6
1·50
<5
2
53
17
93
<0·5
<0·5
120
793
2300
<0·02
1·6
0·6
<0·5
126
941
2736
0·05
2
<0·5
<0·5
109
649
1858
0·02
1·6
<0·5
<0·5
93
164
789
0·05
0·6
<0·5
<0·5
93
208
875
0·03
<0·4
<0·5
<0·5
88
169
765
0·03
<0·4
<0·5
<0·5
93
315
834
<0·02
<0·4
2·5
3·9
27
4·6
169
42
235
4·0
5·5
18
7
213
44
266
4·6
5·3
35
11
144
30
165
2·9
2·1
9
3·8
39
8
23
<1
2·0
11
2·9
41
27
35
1·4
2·8
17
6·1
68
53
36
4·7
4·7
27
15
136
147
44
106
BARNES AND MAIER
Sample:
Height (cm):a
Length (cm):b
Rock type:
PGE IN THE MERENSKY REEF
19
18·00
20
13·00
21
8·00
22
6·00
23
3·00
24
0·00
25
−3·00
26
−8·00
27
−13·00
28
−23·00
5·00
5·00
5·00
2·00
3·00
3·00
3·00
5·00
5·00
10·00
mela-
mela-
mela-
chrom-
mela-
chrom-
anortho-
leuco-
leuco-
leuco-
norite
norite
norite
itite
norite
itite
site
norite
norite
norite
wt %
SiO2
TiO2
Al2O3
52·73
51·52
50·41
34·10
48·84
17·87
48·48
47·34
49·66
48·70
0·30
5·71
0·21
5·98
0·31
5·01
0·88
13·09
0·24
6·64
1·00
20·12
0·08
28·82
0·09
21·93
0·08
22·87
0·09
22·85
FeO∗
9·56
10·51
9·57
17·31
11·22
21·24
1·49
3·58
3·74
3·19
Fe(S)
MnO
MgO
CaO
Na2O
0·46
0·20
22·64
5·91
1·27
0·20
23·11
3·92
2·28
0·20
21·89
6·28
0·31
0·18
14·14
4·72
0·63
0·19
23·72
3·40
0·44
0·15
8·98
4·39
0·45
0·03
2·92
14·08
2·11
0·08
8·29
10·72
0·50
0·07
8·52
11·36
1·12
0·07
7·90
11·50
0·44
0·14
0·53
0·11
0·44
0·09
1·78
0·09
0·37
0·06
0·60
0·16
1·85
0·27
1·33
0·16
1·42
0·14
1·41
0·13
0·06
0·47
0·42
0·27
99·30
0·04
1·21
0·38
0·45
99·44
0·02
2·05
0·92
0·78
100·25
0·01
0·34
12·35
<0·1
99·29
0·02
0·60
3·03
0·17
99·13
0·01
0·41
22·31
<0·1
97·67
0·02
0·41
0·05
0·71
99·64
0·01
1·83
0·14
1·16
98·77
0·01
0·45
0·14
0·55
99·51
0·01
0·97
0·13
0·89
98·97
38
0·37
0·64
5
32·8
0·09
0·70
0·33
128
4·99
9·35
4·00
1·13
0·27
0·21
0·83
0·129
40
<0·1
0·09
3
27·0
0·09
0·64
0·30
95
3·59
6·49
2·70
0·71
0·19
0·10
0·68
0·100
48
<0·1
0·38
2
35·3
<0·05
0·35
<0·2
163
3·07
6·16
3·30
1·12
0·22
0·16
0·86
0·140
<40
<0·1
<0·1
<3
15·8
<0·1
<0·2
<0·2
755
1·25
<2
<1
0·19
0·11
0·04
0·23
0·050
51
<0·1
0·16
4
25·6
<0·07
0·16
<0·1
213
1·22
3·03
<1
0·32
0·08
0·06
0·35
0·035
<40
<0·1
0·31
<5
10·9
0·06
<0·3
<0·5
995
1·30
<2
<1
0·27
0·07
0·06
<0·15
0·040
93
0·09
0·07
1
3·7
<0·02
0·12
<0·1
27
2·81
5·17
1·70
0·35
0·45
0·04
0·12
0·016
107
<0·1
<0·1
1
9·5
<0·03
0·06
<0·1
37
1·98
3·50
1·00
0·24
0·35
0·04
0·30
0·030
80
<0·1
0·05
1
9·6
<0·07
0·06
<0·08
51
1·90
3·18
1·00
0·25
0·36
0·04
0·20
0·030
84
<0·1
0·11
<1
9·3
<0·09
0·11
<0·08
37
2·01
3·89
1·40
0·28
0·37
0·05
0·18
0·031
K2O
P2O5
S
Cr2O3
LOI
Total
ppm
Ba
Cs
Hf
Rb
Sc
Ta
Th
U
V
La
Ce
Nd
Sm
Eu
Tb
Yb
Lu
ppm
Ag
As
Co
Cu
Ni
Sb
Se
ppb
Os
Ir
Ru
Rh
Pt
Pd
Au
<0·5
<0·5
119
1412
2468
0·03
2·3
1·4
<0·5
156
2231
5424
0·02
4
23
27
144
68
1062
1381
366
82
31
116
89
6418
3064
1250
<0·5
<0·5
195
2452
8143
<0·04
6·7
267
322
1628
682
18987
6330
1268
0·6
<0·5
167
1215
2130
<0·04
1·5
413
552
2708
1141
22827
2056
315
0·5
<0·5
142
936
3060
0·02
2·1
254
308
1692
830
13369
2807
1076
107
<1·6
<0·5
216
887
3400
<0·04
1·3
715
1264
6140
2665
37320
4173
339
0·3
<0·5
36
691
1411
0·03
1·3
97
105
600
286
5126
3025
389
1·4
<0·5
132
2442
6448
<0·02
6
262
307
1752
770
18276
9871
2000
0·9
<0·5
61
836
2117
<0·02
1·8
0·5
<0·5
82
1318
3680
<0·02
3
105
94
704
263
5660
3699
754
166
79
1051
394
4302
3032
1153
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 1
JANUARY 2002
Table 1: continued
Sample:
Height (cm):a
Length (cm):b
Rock type:
wt %
SiO2
TiO2
Al2O3
FeO∗
Fe(S)
MnO
MgO
CaO
Na2O
K2O
P2O5
S
Cr2O3
LOI
Total
ppm
Ba
Cs
Hf
Rb
Sc
Ta
Th
U
V
La
Ce
Nd
Sm
Eu
Tb
Yb
Lu
ppm
Ag
As
Co
Cu
Ni
Sb
Se
ppb
Os
Ir
Ru
Rh
Pt
Pd
Au
30
−33·00
10·00
leuconorite
32
−48·00
15·00
leuconorite
34
−58·00
10·00
leuconorite
36
−68·00
10·00
leuconorite
BC 5/.5
Wildebeest
BC 6
BC 25
Onverwacht Dieberg
Ax90
Basaltic
andesite
Tholeiitic
basalt
internal
standard
49·87
0·09
22·78
3·50
50·06
0·09
23·25
3·54
50·52
0·11
22·69
3·69
50·18
0·09
23·11
3·54
54·05
0·37
12·71
9·54
50·53
0·79
15·82
11·10
51·05
0·80
15·61
11·12
36·50
0·23
4·44
13·72
0·5
5·9
1·5
2·8
0·07
8·34
11·84
1·43
0·20
0·02
0·04
0·14
1·21
99·54
0·07
8·43
11·62
1·45
0·13
0·01
0·02
0·14
0·79
99·59
0·08
8·69
11·18
1·43
0·29
0·01
0·02
0·14
0·65
99·50
0·07
8·56
11·79
1·45
0·13
0·01
0·02
0·13
0·44
99·52
0·26
12·66
7·44
1·07
0·66
0·08
0·07
0·10
<0·1
99·01
0·18
7·18
10·32
2·14
0·20
0·15
0·02
0·05
<0·1
98·48
0·19
7·21
10·58
1·89
0·22
0·16
0·01
0·04
<0·1
98·88
0·17
27·69
5·22
0·05
0·05
0·01
3·44
0·65
9·26
4
1·6
3
3
10
6·3
4·4
5
2·5
59
<0·1
0·10
1
10·2
0·02
0·11
<0·08
42
2·14
4·11
1·00
0·28
0·33
0·05
0·19
0·032
68
0·11
0·20
2
10·0
0·01
0·13
<0·06
38
2·14
3·74
1·30
0·30
0·32
0·05
0·19
0·025
49
<0·1
0·13
2
10·4
0·02
0·13
<0·1
49
2·07
4·07
1·10
0·28
0·33
0·04
0·19
0·030
51
<0·1
<0·1
<1
9·5
0·02
0·13
<0·03
37
2·07
3·50
1·10
0·25
0·33
0·04
0·19
0·027
212
1·96
1·13
24
38·4
0·11
2·43
0·70
171
11·2
21·0
10·8
1·89
0·54
0·31
0·85
0·14
165
<0·1
1·63
<3
38·2
0·3
0·66
0·17
199
16·4
33·7
19·5
3·66
1·29
0·56
2·10
0·294
167
<0·1
1·54
<3
34·5
0·3
0·62
<0·1
214
14·7
30·4
16·7
3·51
1·22
0·56
1·90
0·273
30
0·12
0·29
5
17·1
<0·04
<0·03
<0·1
97
0·43
1·4
1·1
0·45
0·17
0·12
0·53
0·09
5
17
11
20
2
30
8·3
30
5·2
5
9
15
3·5
10
21
2·5
6
Tholeiitic
basalt
0·4
<0·5
34
105
298
<0·02
<0·4
<0·5
<0·5
31
69
223
<0·02
<0·4
<0·5
<0·5
32
46
205
<0·02
<0·4
<0·5
<0·5
30
46
203
<0·02
<0·4
<1
1·0
183
55
720
0·07
<0·4
<1
<1
59
71
116
<0·05
<0·4
<1
<1
55
76
162
<0·05
<0·4
<1
<1
222
864
7207
<0·02
2
<1
2·4
<5
6·4
91
52
21
0·9
0·42
<5
1·9
8
<4
2
<0·6
0·37
<5
1·2
<4
<2
2
<1
0·57
7
1·1
6
<3
1
0·47
0·51
<2
1·89
18·6
11·2
3·41
0·19
0·15
1·8
0·76
9·6
10
2·89
<0·2
0·14
<2
0·49
9·9
6
1·9
2·7
3·2
15·8
11·9
140
358
5·4
RSD
(n = 10)
26
38
3
7
5
25
14
15·7
7·6
20·6
7·3
10·9
8·3
14
∗All Fe was originally reported as Fe2O3. This has been recalculated in samples with >0·1% S to allow for Fe in sulphides
[Fe(S)]. Fe(S) = 1·527S − 0·6592Cu/10 000 − 0·5285(Ni/10 000 − 0·07). FeO = [Fe2O3 × 0·699 − Fe(S)] × 1·29, which
assumes that the sulphides consist of pyrrhotite, pentlandite and chalcopyrite and that the silicates contain 700 ppm Ni.
a
Height of the sample relative to the base of the Merensky Reef chromitite.
b
Length of the core used.
108
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Fig. 2. Variations in chalcophile elements with height. It should be noted that the vertical scale is not linear, the samples are spaced out evenly.
The reef as defined by rocks containing >1·5 ppm Pt + Pd starts 23 cm below the lower chromitite and ends 23 cm above it. The Ni, Cu, Se,
Au and Pd concentrations closely follow that of S, indicating that sulphides are the main phase controlling these elements. In contrast, although
Os, Ir, Ru, Rh and Pt follow the S content in the silicate rocks, the highest values for these elements are found in the chromite-bearing rocks,
indicating that some phase is important in controlling their distribution in chromite-bearing rocks. An, anorthosite; Chr, chromite; Mn,
melanorite.
Fig. 3. Polished slab through the Merensky Reef. The expression ‘reef’ is a mining term and simply refers to a narrow layer of the rocks
containing mineable grades of a metal. It does not necessarily refer to a particular rock type. The group of rocks containing >1·5 ppm Pt +
Pd are referred to as Reef. In this case, the Reef (normally referred to as pegmatoidal pyroxenite) consists of an anorthosite, overlain by chromite,
overlain by coarse-grained melanorite, overlain by a narrow chromite, overlain by melanorite.
quantity of sulphides decreases up-section to 0·3% at
50 cm from the base and then increases in the final 30 cm
to 1–2% (Fig. 2).
The melanorite is underlain by a thin chromitite layer
of irregular thickness (Fig. 3). The upper boundary is
planar, but the lower boundary undulates. There is a
109
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 1
JANUARY 2002
Fig. 4. Photomicrographs of the rocks. All photos are oriented such that up is towards the top of the page. (a) Typical sulphides; field of view
8 mm. (b) Melanorite overlying the upper chromitite; field of view 20 mm. (Note sulphides forming a vertical network.) (c) Contact between
lower chromitite and coarse-grained melanorite; field of view 20 mm. (d) Contact between anorthosite and lower chromitite; field of view 20 mm.
(e) Coarse-grained melanorite showing suturing of orthopyroxenes and spindle twins; field of view 15 mm. (f ) Leuconorite; field of view 20 mm.
(Note sulphides forming a vertical network.) Chr, chromite; Cp, chalcopyrite; OPX, orthopyroxene; Pl, plagioclase; Pn, pentlandite; Po, pyrrhotite;
Sul, sulphides.
second chromitite seam 3 cm below this and a very
coarse-grained melanorite lies between them (Fig. 3).
The boundary between the lower chromitite and the
overlying melanorite undulates on a centimetre scale
(Fig. 4c). The orthopyroxene in the overlying melanorite
appears to have been pushed into the chromitite, and
the chromitite crystal mush seems to be escaping upwards
between the orthopyroxene crystals into the overlying
layer (Fig. 5). The chromitite consists of >40 modal %
chromite, which ranges in shape from cubic to amoeboidal (Fig. 4c and d) with individual grains of >2 mm
in diameter. Plagioclase (>40%) in the chromite layer
takes the form of large oikocrysts up to 10 mm in size.
These oikocrysts have spindle-shaped twins, low-angle
grain boundaries, deformation bands and exhibit undulose extinction (Fig. 4c); all features that indicate that
plagioclase has been plastically deformed. Orthopyroxene
makes up >15% of the rock. Many of the grains are
indented on crystal faces oriented subhorizontally and
many grains are kinked. Sulphides (>1%) form a linked
interstitial network oriented parallel to the vertical. Biotite
is rare (p1%).
The coarse-grained melanorite between the two chromite seams (Fig. 3) contains >70% orthopyroxene,
>23% plagioclase, >6% chromite, >1% sulphides and
trace amounts p1% of biotite. The orthopyroxene forms
110
BARNES AND MAIER
PGE IN THE MERENSKY REEF
and e). The orthopyroxenes have undulose extinction,
are indented when two orthopyroxenes are in contact,
and show shearing along the mutual contacts (Fig. 4e).
The plagioclase forms large oikocrysts that partially enclose the pyroxene. The plagioclase shows evidence of
deformation in the form of spindle-shaped twins trending
parallel to the interpenetrating contact between orthopyroxenes (Fig. 4e). Chromite and sulphides occur at the
margins of the orthopyroxene grains.
The lower chromitite is underlain by a thin (3 cm)
layer of medium-grained anorthosite (Fig. 3) with large
(up to 8 mm) oikocrysts of orthopyroxene (>9% modal).
The plagioclase (>90%) has a cumulus texture in the
form of 2–4 mm subhedral laths. The long axis of the
plagioclase defines a shape-preferred orientation that is
parallel to the contact with chromite layer (Fig. 4d). A
few of the plagioclase grains show compositional zoning
(Fig. 4d) and the twins are rectilinear (i.e. not deformed)
suggesting that the preferred orientation may be an
igneous lamination. The sulphides (>1%) occur between
the plagioclase laths in the form of vertical networks
perpendicular to the layering. Biotite is rare (p1%). The
presence of the concentric compositional zoning, the
igneous lamination and rectilinear twins suggests that the
anorthosite has preserved some of its original igneous
texture.
The anorthosite is underlain by 60 cm of mediumgrained (1–5 mm) sub-ophitic textured leuconorite. The
plagioclase (60–70%) is a cumulus phase and occurs
as medium-grained (2–4 mm) subhedral laths with a
moderate alignment of the long axes of the laths parallel
to the horizontal. Many grains show rectilinear twins,
but with some signs of deformation in the form of
undulose extinction, spindle-shaped twins and indented
grains. The grain boundaries between the plagioclase
grains are sutured, suggesting some grain boundary migration has occurred. A few grains are enclosed by
orthopyroxene. These grains tend to be smaller
(0·5–1 mm) and do not show signs of deformation. This
indicates that the orthopyroxene grew shortly after the
initial crystallization of plagioclase and before the plagioclase had finished crystallizing. The orthopyroxene (15–
20%) forms large subhedral to anhedral grains (4–8 mm).
Clinopyroxene (5–10%) occurs as small exsolution blebs
in the orthopyroxene and as interstitial grains between
the plagioclase grains or rimming the orthopyroxene.
Just below the anorthosite layer sulphides make up 4%
of the rock. The sulphides form an interconnected network in the vertical direction (Fig. 4f ). The sulphide
content decreases downwards and by 26 cm below the
lower chromitite sulphides are present only as tiny patches
(<0·1 mm) in the interstitial material. Nevertheless, all of
the samples studied contain some sulphides.
Fig. 5. Sketches showing microstructural features of the core samples.
Field of view 40 mm in diameter. (a) Microstructures in the orthopyroxene framework, indicating deformation of the crystal mush during
compaction: (1) indented contacts; (2) fractured orthopyroxene crystals
surrounded by optically unstrained plagioclase; (3) shortened crystals
surrounded by unstrained plagioclase; (4) chromite grain forced between
orthopyroxene; (5) sulphide domain along pyroxene crystal faces oriented subvertically parallel to the core axis. (b) Microstructures indicative of deformation after the crystal mush is largely solid: (1)
clinopyroxene oikocryst with undulose extinction; (2) small strainfree subgrains at clinopyroxene–orthopyroxene contacts; (3) undulose
extinction and kinks (low-angle grain boundaries) in orthopyroxene; (4)
relict igneous cumulus plagioclase with concentric zoning and rectilinear
twins preserved in oikocryst and in a pressure shadow shadows; (5)
deformed recrystallized plagioclase with wedge-shaped twins, undulose
extinction and low-angle grain boundaries; (6) plagioclase with wedgeshaped deformation twins and shear plane; (7) sulphide along subvertical
contact in matrix of strained grains. Orientation of shear planes with
orthopyroxene and plagioclases, indented contacts and chromite bulges
between orthopyroxene grains indicate subvertical shortening of the
crystal mush, i.e. compaction. Sulphide has collected on grain contacts
corresponding to extension fractures parallel to the principal compressive stress 1.
large (1 cm) complex grains that appear to be made of
a number of smaller grains that have been forced together,
and developed into one large grain during static recrystallization. Small chromite grains mark the sutures
between original, smaller orthopyroxene grains (Fig. 4c
111
JOURNAL OF PETROLOGY
VOLUME 43
Evidence for deformation during
crystallization
NUMBER 1
JANUARY 2002
and some have triple junctions suggesting some grain
boundary migration.
Most of the samples studied show evidence of deformation
during crystallization. Many samples contain a similar
suite of mineral deformation features. In the melanorite
and the chromitite the cumulus orthopyroxene crystals
have a weak undulose extinction, deformation bands or
contain kink bands indicating dislocation creep. Many
orthopyroxene grains have indented contacts (Fig. 4b
and e) where they impinge on other orthopyroxene
grains. Generally cleavage planes and exsolution lamellae
and blebs remain straight around indents, suggesting that
dissolution or perhaps diffusion creep may have occurred
rather than bending of the lattice, but undulose extinction
and the development of small polygonal subgrains around
indentations in some rocks indicate that part of the
deformation was by dislocation creep. The interstitial
(porosity-filling) plagioclase oikocrysts in these rocks are
characterized by the widespread development of spindleshaped deformation twins, bent rectilinear twins, undulose extinction, deformation bands (Fig. 4c) and locally
intracrystalline shear planes offsetting the rectilinear
twins. However, texturally late biotite and quartz, which
occur interstitially as anhedral crystals in the melanorite,
generally do not exhibit evidence of lattice deformation
such as undulose extinction and, in biotite, kinking. Thus,
the textures in these rocks are interpreted as indicating
dislocation creep after framework orthopyroxene and
infilling plagioclase formed but before biotite and quartz
grew.
The microstructural relationships in a few melanorite
samples (7, 9, 11) preserve a still earlier stage of deformation. In these rocks many orthopyroxene crystals
are deformed, but the infilling plagioclase around them
is not deformed and retains rectilinear twins and lacks
undulose extinction or characteristic spindle-shaped deformation twins. Some orthopyroxene crystals are fractured and the fragments rotated a few degrees; the
fractures are filled with unstrained plagioclase. Other
orthopyroxene crystals contain shear planes along which
the crystals have shortened. Thus, in these rocks the
orthopyroxene framework records a much larger strain
than the infilling plagioclase, consistent with deformation
of an interlocking matrix of orthopyroxene containing
melt, i.e. deformation pre-dates the crystallization of the
infilling plagioclase.
In the leuconorite and anorthosite the roles of the
orthopyroxene and plagioclase are reversed. Small undeformed plagioclase crystals are included in the orthopyroxene oikocrysts, whereas larger deformed and
oriented plagioclase grains occur around the oikocrysts.
This suggests that orthopyroxene grew shortly after
plagioclase began crystallizing. Some of the grain boundaries of the plagioclase outside the oikocrysts are sutured
Compaction
The orientation of deformation structures in each sample
can be used to infer the orientation of the principal
compressive stress, as the oriented thin sections come
from a vertical borehole. Indented orthopyroxene–
orthopyroxene contacts are predominantly at ±30° to
the horizontal; conversely, crystal faces are best developed
on orthopyroxene faces that are subvertical (Fig. 5).
Intracrystalline shear planes, developed in either plagioclase or orthopyroxene, form two conjugate trends with
the vertical core axis in the acute angle between them
(Fig. 5). Therefore, the orientation of the deformation
structures suggests that the principal compressive stress
was subvertical, and as the microstructural sequence
indicates that deformation began in a crystal mush consisting of an orthopyroxene framework, the earliest deformation recorded appears to be related to gravitational
compaction of the mush. Further, evidence for compaction structures occurring when melt was still present
can be seen in the chromitite layers forced up between
orthopyroxene crystals (Fig. 4c). Sulphides generally
occur in net-like arrays along the grain boundaries (Fig.
4b and f ) and the longest sulphide-filled boundaries are
typically oriented parallel to the core axis, that is, they
lie parallel to the principal compressive stress, and can
therefore be interpreted as extension veins formed synchronously with compaction (Fig. 5). Clinopyroxene oikocrysts have subhorizontal pressure shadows around them
in which randomly oriented rectangular plagioclase crystals with rectilinear twins and concentric zoning are
preserved. In contrast, plagioclase outside the pressure
shadows has undulose extinction, and strongly developed
spindle-shaped deformation twins and a subhorizontal
shape preferred orientation. Clinopyroxene oikocryst
faces perpendicular to the vertical have locally developed
zones of fine-grained polygonal recrystallized grains at
contacts with orthopyroxene (Fig. 5). The contrast between microstructures preserved in the oikocryst pressure
shadows and outside the pressure shadows indicates that
recrystallization processes, although extensive in our normal samples, did not destroy all of the earlier textural
history.
GEOCHEMISTRY
Analytical methods
The samples were crushed in an Al-ceramic mill at the
University of Quebec at Chicoutimi (UQAC), yielding
sample masses of between 50 and 200 g. Major oxides,
112
BARNES AND MAIER
PGE IN THE MERENSKY REEF
nickel and V were determined by X-ray fluorescence
(XRF) at McGill University, Montreal. All other elements
were determined at UQAC: Cu by atomic absorption
spectrometry (AA), S by combustion iodometric procedure using a Laboratory Equipment Company (LECO)
titrator, the other trace elements by instrumental neutron
activation analysis (INAA) using the method of Bedard
& Barnes (1990), with the modification that corrections
on Ta, Ba and Lu were made for Pt and Ru interferences.
For most samples the PGE and Au were determined by
INAA after using the Ni-sulphide fire assay method on
50 g of sample. It was noted by Borthwick & Naldrett
(1984) that chromitites do not fully dissolve in a standard
flux mixture. Following their technique, 30 g of sample
were used in analysing the chromitites, the sodium borate
was replaced with lithium tetraborate, and 6 g of SiO2
was added. The blank contained 0·3 ppb Au, but all
other elements fell below the detection limits (Os <0·4
ppb, Ir <0·02 ppb, Ru <2 ppb, Rh <0·3 ppb, Pt <2
ppb, Pd <2 ppb). The detection limit is defined as 3
times the background and the relative standard deviation
on the background for the last 10 blanks is >50%.
Because of the unusual nature of the samples the fire
assay results were crosschecked against the results from
the whole-rock powder for Os, Ir, Ru, Pt and Au. The
results agreed within analytical error and we therefore
feel that there was good recovery on the fire assay for
all elements including Au.
The results for the 24 samples from the borehole, plus
three samples from the margins of the intrusion, are
presented in Table 1. The three samples from the margins
were obtained from sites 23.16, 23.18 and 23.25 of
Sharpe (1986). In addition, the relative standard deviation
(standard deviation × 100/average) for the last 10 runs
on our internal standard (Ax90) are listed in the last
column of Table 1.
How representative is a single borehole through the
Reef? We can investigate this question by considering
the similarity between the international standard SARM7 (a 7500 kg Merensky Reef composite taken from five
mines; Steele et al., 1975) and the weighted average for
our samples. As can been seen in Table 2 the weighted
average is very similar for most elements, thus it would
appear that our borehole is indeed representative of the
Merensky Reef in general. Two interesting exceptions to
this observation are Sb and As, which are higher in
SARM-7 than in the Impala sample. The reason for this
is not understood.
Table 2: Whole-rock compositions of Bushveld
rocks
wt %
SiO2
TiO2
Al2O3
FeO
Fe
MnO
MgO
CaO
Na2O
K2O
P2O5
S
Cr2O3
LOI
Total
ppm
Ba
Cs
Hf
Rb
Sc
Sr
Ta
Th
U
V
Zr
La
Ce
Nd
Sm
Eu
Tb
Yb
Lu
ppm
Ag
As
Co
Cu
Ni
Sb
Se
ppb
Os
Ir
Ru
Rh
Pt
Pd
Au
Merensky Reef
Initial magmas
SARM-7∗
Basaltic‡
andesite
Impala
wt av.†
51·80
0·21
8·30
10·15
0·49
0·20
20·10
5·40
0·80
0·11
0·10
0·40
1·02
0·9
99·99
51·09
0·22
10·60
8·62
0·53
0·16
18·89
6·20
0·75
0·20
0·03
0·51
1·08
0·39
99·27
50
0·36
0·53
9
31
61
0·09
0·57
6
24
0·04
1·00
0·20
60
0·04
0·75
0·11
100
4·2
7·9
4·5
1·06
0·24
0·10
0·75
0·10
Mixed
magma
Tholeiitic‡ 60:40
basalt
55·87
0·37
12·55
9·15
50·48
0·71
15·79
11·61
53·71
0·51
13·85
10·13
0·21
12·65
7·29
1·53
0·77
0·10
0·08
0·14
0·18
7·26
10·86
2·20
0·16
0·16
0·04
0·05
0·20
10·49
8·72
1·80
0·52
0·12
0·06
0·10
100·70
99·50
100·22
3·29
6·81
2·68
0·73
0·24
0·11
0·62
0·10
310
2·1
2·0
27
41
170
0·16
3·80
0·75
179
80
13·2
26·6
15·9
2·57
0·71
0·32
1·06
0·14
206
0·13
1·50
3
35
340
0·22
0·70
0·19
182
60
15·5
32·0
18·1
3·58
1·25
0·56
2·00
0·26
268
1·31
1·80
18
39
238
0·18
2·56
0·53
180
72
14·1
28·8
16·8
2·97
0·93
0·42
1·44
0·19
0·4
1·9
143
700
3100
2·52
1·41
0·16
<0·5
106
811
2300
0·02
1·69
<0·5
1·1
73
58
329
0·07
0·27
<0·5
<0·5
53
62
128
0·03
0·15
<0·5
0·67
65
60
249
0·06
0·22
64
75
434
242
3771
1543
313
71
79
476
201
3688
1402
422
0·4
0·35
2·0
1·4
19·0
11·2
3
0·2
0·17
1·5
0·6
13·1
7·8
3
0·29
0·28
1·80
1·08
16·6
9·8
3
∗Potts et al. (1992) and this work.
† Weighted average of samples 1–28.
‡ Estimates based on Davies & Tredoux (1985), Harmer &
Sharpe (1985), von Gruenewaldt et al. (1989) and the three
chills in Table 1.
Chalcophile elements and PGE
Nickel, Cu, Se, Au and Pd correlate well with S (Table
3, Fig. 2). In the basal 45 cm of the core these elements
occur in low concentrations, but in the next 23 cm, to
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Table 3: Correlation coefficients for the whole-rock samples
Os
Ir
Ru
Rh
Pt
Pd
Au
Ni
Cu
Cr
Ir
0·975
Ru
0·987
0·994
Rh
0·986
0·995
0·999
Pt
0·985
0·955
0·960
0·961
Pd
0·626
0·506
0·552
0·552
0·697
Au
0·424
0·285
0·322
0·333
0·518
0·912
Ni
0·509
0·384
0·417
0·417
0·589
0·838
0·885
Cu
0·460
0·320
0·354
0·353
0·539
0·825
0·893
0·959
Cr
0·878
0·946
0·923
0·922
0·829
0·224
0·005
0·137
0·094
Se
0·391
0·250
0·289
0·287
0·480
0·846
0·906
0·975
0·962
−0·016
S
0·379
0·239
0·280
0·278
0·467
0·841
0·901
0·976
0·944
−0·030
below the lower chromitite layer, they are relatively
enriched. In the chromitite layers and the coarse-grained
melanorite, the S and metal values drop slightly, but
are again elevated in the basal 20 cm of the overlying
melanorite. Above that they occur in relatively low concentrations for 30 cm, before their values rise at the top
of the melanorite and then fall back in the norite to form
a third peak. The overall strong correlation (Table 3)
between the metals and S suggests that sulphide is the
main carrier for these elements.
As in the case of Ni, Cu, Se, Au and Pd, the Os, Ir,
Ru, Rh and Pt values start to rise 23 cm below the lower
chromitite layer. The high values continue for another
15 cm above the upper chromitite layer and show a slight
peak in the same area as the Ni, Cu, Se, Au and Pd
values. However, in contrast to those elements, the main
peak values in Os, Ir, Ru, Rh and Pt are observed in
the chromitite layers (Fig. 2), which may suggest that
chromite could be a carrier of these elements. However,
the PGE show a stronger correlation with each other
than with Cr (Table 3) suggesting that the metals occur
as discrete phases rather than as solid solution in chromite.
Thus, Os, Ir, Ru, Rh and Pt appear to have been
influenced by at least two phases; the sulphides and
chromite or PGM associated with chromite. Lee (1983),
in his study of the Merensky Unit at Rustenburg, also
found that the metals could be divided into two groups,
one consisting of Ni, Cu, Au and Pd, and the other of
Ir, Rh, Pt and Cr, based on their covariance; he concluded
that Pd and Au were controlled by sulphides and Pt, Ir
and Rh were controlled by PGM or spinel.
In most rocks As and Sb are present at very low
concentrations and are close to their detection limits of
0·5 and 0·02 ppm, respectively. In most hydrothermal
Se
0·962
deposits concentrations of these elements are high; their
absence here suggests a lack of hydrothermal activity. In
the Reef rocks Ag is present at 0·5–1·5 ppm and mimics
Cu. In most other rocks it is present at less than detection
level.
The mantle-normalized metal patterns for the rocks
may be divided into four groups based on their different
shapes. The footwall leuconorites (Fig. 6a) show Ni and
Ir at approximately the same level, 0·1 times mantle.
The metal concentrations rise to around one times mantle
at Rh and then flatten out between Rh and Cu, with no
obvious enrichment of PGE over Cu. The silicate rocks
of the Reef, the leuconorite and anorthosite under the
chromitites and the melanorites above the chromitites
show a strong enrichment in PGE relative to Ni and Cu
(Fig. 6b), giving the arch-shaped patterns characteristic
of reef-type deposits from around the world (Barnes et
al., 1988). From Os to Pt there is a steady increase in
concentration but from Pt to Au the patterns are almost
flat. This is followed by a distinct fall in concentrations
from Au to Cu. The chromite-bearing rocks of the
Reef (the lower chromitite, the coarse-grained melanorite
between the chromitites, and the upper chromitite) have
metal patterns very similar to those of the other reef
rocks (Fig. 6d) with a strong enrichment in PGE over
Cu and Ni. However, the overall concentrations of Os,
Ir, Ru, Rh and Pt are 3–4 times higher (Table 1) and
the Pt/Pd ratios are also much higher (10 vs 1·5–3 for
the silicate rocks). The mantle-normalized metal patterns
of the melanorite and norites of the hanging wall do not
show an enrichment of PGE relative to Ni and Cu (Fig.
6c). In fact, Pt and Pd are depleted relative to Cu with
Cu/PdMN of three. The patterns are flat from Ni to Ir,
they rise from Ir to Pt and they show strong positive Au
114
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Fig. 6. Comparison of mantle-normalized metal patterns (continuous black lines) obtained for the various rock types with model patterns
[shaded areas in (a)–(c); dashed lines in (d)]. The footwall leuconorites (a) may be modelled by considering them to contain 0·05–0·1% sulphides
formed at moderate N factors. The silicate rocks of the reef (b) may be modelled as containing 1–5% sulphides formed at very high N factors.
This gives the reef patterns an arch shape. The hanging-wall melanorite (c) may be modelled as containing 0·5–2% sulphides formed from a
magma depleted in PGE. The chromite-bearing rocks (d) cannot be modelled by sulphide accumulation (dashed line) or mss accumulation
(dotted line). They can be modelled by accumulation of 1% sulphide liquid and 0·005% PGM (bold dashed line). Normalization values from
Barnes & Maier (1999).
anomalies. This is not an analytical problem, because
the Au concentrations from the whole-rock powder are
within 15% of Au concentrations obtained by the fire
assay method.
and mixes vigorously with the resident magma. The
mixed magma becomes saturated in an immiscible sulphide liquid and the sulphide liquid droplets collect the
PGE. The sulphide droplets then settle through the
magma onto the crystal pile with other dense phases,
such as chromite, to form a chromite layer rich in PGE
(the Reef ).
According to this model these rocks are cumulates with
a melt component. To model the trace elements the
equation
Lithophile elements
It is not the purpose of this paper to address the details
of the geochemistry of the lithophile elements. This has
been done recently (Wilson et al., 1999), and our results
are broadly similar. None the less, certain aspects of the
lithophile elements concentrations are relevant to the
understanding of the chalcophile element and PGE distributions and need to be considered in any model for
the origin of the Reef. The model most commonly used
to explain the origin of the Reef suggests that before the
formation of the Merensky Unit there is an unconsolidated crystal pile in the Upper Critical Zone,
consisting of a leuconorite and anorthosite mush overlain
by mafic magma (Campbell et al., 1983). Subsequently,
a new batch of magma is introduced into the chamber
CC = Cl(1 − F) [D (1 − T ) + T ]/(1 − F)
(where CC is the concentration of the element in the
cumulate, Cl is the concentration of the element in the
liquid, F is the mass fraction of residual magma, D is the
bulk partition coefficient and T is the mass fraction melt
in the cumulate) should be used (Allègre & Minster,
1978). To apply this equation we need to consider what
the possible compositions of the resident and incoming
magma were, so as to estimate Cl. Further, we need to
115
JOURNAL OF PETROLOGY
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estimate how much melt component ( T ) is present in
the rocks.
Magma compositions
At this level in the Bushveld Complex there are two
possible silicate liquids from which the rocks could have
formed (Davis & Tredoux, 1985; Harmer & Sharpe,
1985). These are a Mg-rich basaltic andesite resembling
a boninite and a tholeiitic basalt resembling a low-Ti
continental flood basalt. We present three new analyses
of these in Table 2 and Fig. 7. Our results are similar
to the previously published results, but cover a wider
range of elements. The basaltic andesite has slightly
radiogenic initial Sr isotopes (0·703–0·705); in contrast,
the tholeiitic basalt has strongly radiogenic initial Sr
isotopes (0·706–0·708) (Harmer & Sharpe, 1985). Stratigraphic sections through the Bushveld Complex show
that there is a change in Sr, Nd and Os isotopic compostion and incompatible element ratios between the
Lower Critical Zone and the Main Zone (Kruger &
Marsh, 1982; Maier & Barnes, 1998; McCandless et al.,
1999; Schoenberg et al., 1999; Maier et al., 2000). The
Lower Zone and Lower Critical Zone magmas have
isotopic and inter-element ratios similar to the basaltic
andesite. The Main Zone has isotopic and inter-element
ratios similar to the tholeiitic basalt. The Merensky Unit
occurs in the Upper Critical Zone, where inter-element
ratios and isotope ratios oscillate between the values for
the basaltic andesite and the tholeiitic basalt and it is
thought to represent a zone of mixing of the two magma
types. To test how much of each magma is required to
form the Impala samples we need two elements that (1)
are incompatible with the cumulus phases, (2) have
different ratios in the basaltic andesite and the tholeiitic
basalt, and (3) are both immobile. The basaltic andesite
is strongly enriched in the large ion lithophile elements
(LILE) but has similar concentrations of the middle rare
earth elements (MREE), Hf and Zr and is depleted in
heavy REE (HREE) relative to the tholeiitic basalt (Fig.
7, Table 2). The REE and Ba cannot be used because
these elements partition slightly into orthopyroxene or
plagioclase. Tantalum is close to the detection limit in
many samples. The Th/Hf ratio appears to fulfil the
criteria best. The basaltic andesite magma has a Th/Hf
value of >1·9 and the tholeiitic basalt magma has a
Th/Hf ratio of 0·4. The Th/Hf values of the Impala
rocks are between those of the two magmas (Fig. 8) with
a weighted average of 1·4. A mixture of approximately
60% basaltic andesite and 40% tholeiitic basalt would
be required to achieve this ratio. Rubidium, U and Cs
to Ta or Hf ratios give similar results. The ratios above
and below the chromitite are similar; thus the melt from
which the leuconorite and anorthosite formed appears
Fig. 7. Mantle-normalized incompatible element plots for the magmas
from which the Bushveld Complex formed. Sources are as given in
Table 2. It should be noted that the basaltic andesite has much higher
LILE to HFSE (high field strength element) ratios than the tholeiitic
basalt. Normalization values from McDonough & Sun (1995).
to have the same Th/Hf ratio as the melt from which
the melanorite formed.
Percentage melt component present
To investigate how much melt component is present we
can consider the concentrations of incompatible elements
in the cumulate divided by those in the mixed magma.
The ratios of elements such as Rb, Hf, U, Th and Ta
in the weighted average of the Impala samples to the
mixed magma (Table 2) are 0·21–0·33 implying 21–33%
melt component. A framework of cumulus crystals will
start to behave as a rigid matrix at >55% solids. Therefore initially the cumulate would contain >45% melt.
The rocks now contain 10–30% oikocrysts which would
have displaced some of the melt leaving 15–35% melt,
consistent with whole-rock geochemistry. During
solidification, compaction and cementation occur, and
the melt component may migrate. If we consider the
individual samples we can see that throughout the
leuconorite, the anorthosite, the chromitites and the
melanorite between the chromitites, the incompatible
element concentrations are very low. The weighted average of the Reef (samples 19–28) contains 0·18 ppm
Hf, which divided by the Hf concentration in the mixed
magma suggests the Reef contains on average 10% melt
component. Wilson et al. (1999) obtained similarly low
figures for Reef of the same thickness as the Impala
section. Inspection of their data shows that the thickness
of the Reef is directly proportional to the amount of melt
present. Thin reef sections, such as ours, have a low melt
component; thicker reefs have a larger melt component.
116
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Fig. 8. Variations in element ratios vs stratigraphic height. (a) The Th/Hf ratio of the cumulates is intermediate between those of the two
parental magmas, implying that the reef formed from a mixture of the two magmas. (b) Hf/(mixed magma) × 100 has been used to estimate
the percentage of melt component in the rocks. This shows that the reef contains <10% melt component. The hanging wall has a large melt
component. (c) The Cu/Hf ratio of all of the rocks is greater than that of the parental magmas, indicating that cumulate sulphides are present
in all of the rocks. The ratios are highest in the vicinity of the reef, indicating that these rocks contain the most sulphides.
This suggests that thin reef facies have experienced more
compaction that thick reefs. The hanging-wall melanorite
(samples 7–18) contains higher incompatible element
contents (e.g. 0·5–1·4 ppm Hf ) implying either a larger
melt component (30–70%) or that more evolved melt
is present. The melt fractions we have calculated are
maximum values as the intercumulus melt could have
been more evolved than the initial magma, especially in
the leuconorite and anorthosite.
The light REE (LREE) and MREE contents of most
of the Impala rocks are a quarter to a third of those of
the melt (Table 2), indicating that these elements are
behaving in an incompatible fashion. The samples containing cumulus plagioclase (the leuconorite, the anorthosite and the norite) show positive Eu anomalies (Fig.
9a and b). The leuconorites and anorthosite show a
stronger degree of LREE enrichment (La/SmN of 5–6)
than the melt (La/SmN 3·5) (Tables 1 and 2). This may
be explained by the fact that although the partition
coefficients for LREE into plagioclase are low, they are
still higher than the partition coefficients for the MREE
(Table 4). The leuconorite and norite show a stronger
degree of HREE enrichment (Tb/YbN of 0·7–1) than
the melt (Tb/YbN of >1·4) (Tables 1 and 2). The
difference may be due to the presence of the orthopyroxene, because partition coefficients for HREE are
higher than those for the MREE (Table 4).
The REE patterns of the melanorite and of the chromite-bearing rocks of the Reef show a similar degree of
LREE enrichment to that of the melt (La/SmN in the
melanorite is in the range of 2–4 compared with 3·5 for
the melt; Tables 1 and 2). The HREE are slightly
enriched relative to the melt with Tb/YbN of 0·6–1·2
compared with >1·4 in the melt. As in the case of the
leuconorites and norites, the HREE enrichment may be
explained by the presence of cumulus orthopyroxene.
Some samples have small negative Eu anomalies (Fig.
9c and d).
117
JOURNAL OF PETROLOGY
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Table 4: Modelling of the whole-rock compositions—lithophile elements
Rock type:
Reef
HW
Lower
FW
melanorite
melanorite
chromitite
leuconorite
Anorthosite
Initial
Ortho-
magma
pyroxene
F
0·999
0·999
0·999
0·8
0·999
See
wt frac melt
0·1
0·25
0·07
0·05
0·04
Table 2
wt frac opx
0·72
0·68
0·07
0·25
0·09
wt frac plag
0·09
0·05
0·25
0·69
0·85
wt frac chr
0·07
wt frac sul
0·0155
0·001
0·01
Compare
IM 23
Plagioclase
Chromite
0·6
0·011
IM 7
0·011
IM 24
IM 30
IM 25
with
wt %
SiO2
50·21
54·10
19·65
50·20
48·73
53·71
TiO2
0·25
0·23
0·86
0·08
0·04
0·51
0·15
Al2O3
6·61
6·15
19·17
23·63
28·48
13·85
1·49
FeO
10·65
9·54
20·46
3·55
1·52
10·13
FeS
0·65
0·46
0·46
0·04
0·42
MnO
0·22
0·21
0·30
0·08
0·03
0·20
MgO
23·24
23·02
7·93
7·37
2·85
10·49
CaO
3·47
3·88
6·46
11·65
13·69
8·72
Na2O
0·38
0·56
0·68
2·01
2·43
1·80
S
0·60
0·43
0·42
0·04
0·38
0·06
Cr2O3
3·22
0·37
24·46
0·11
0·04
0·10
56·3
47·8
1·36
33·5
10·3
31·72
0·24
0·45
30
1·29
16·2
8·5
16·4
2·77
2·2
0·5
D
40·7
D
D
K2 O
0·06
0·13
0·06
0·11
0·11
0·52
0
0·2
0
P2O5
0·01
0·03
0·01
0·01
0·01
0·12
0
0
0
ppm
Ba
0
0·23
0
Cs
0·13
0·33
0·09
0·08
0·06
1·31
0
0
0
Hf
0·18
0·45
0·13
0·11
0·08
1·80
0
0
0
Rb
1·90
4·45
1·63
2·35
2·24
17·52
0
0·1
0
5
39
Sc
32
268
1
0
0
0·01
0·01
0·01
0·18
0
0
0
Th
0·26
0·64
0·18
0·15
0·11
2·56
0
0
0
U
0·05
0·13
0·04
0·03
0·02
0·53
0
0
0
0·5
0
10
1022
12
64
0·05
91
5
62
0·02
190
29
33
Ta
V
29
69
34
15
180
La
1·76
3·74
1·60
2·79
2·80
14·13
0·015
0·18
0
Ce
3·44
7·54
2·85
4·40
4·21
28·76
0·015
0·12
0
Nd
1·95
4·38
1·50
2·06
1·88
16·78
0·015
0·08
0
Sm
0·34
0·77
0·25
0·32
0·28
2·97
0·015
0·06
0
Eu
0·13
0·25
0·14
0·29
0·31
0·93
0·02
0·34
0
Tb
0·06
0·12
0·04
0·05
0·04
0·42
0·05
0·06
0
Yb
0·24
0·44
0·13
0·19
0·15
1·44
0·1
0·06
0
Lu
0·04
0·07
0·02
0·03
0·02
0·19
0·2
0·06
0
HW, hanging wall; FW, footwall; opx, orthopyroxene; plag, plagioclase; sul, sulphide; chr, chromite.
Partition coefficients from Rollinson (1993), mineral compositions from Kruger & Marsh (1985).
118
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Modelling the whole-rock composition of
the Impala rocks
The major elements have been modelled by mass balance
using the mineral compositions determined for the Reef
by Kruger & Marsh (1985). The trace elements have
been modelled using the equation of Allègre & Minster
(1978). The mixed magma was used as the initial melt.
The Reef melanorite can be modelled as a mixture
of approximately 72% orthopyroxene, 10% melt, 7%
chromite, 9% plagioclase and 1·5% sulphides (Table 4,
column 1). The hanging-wall melanorite may be modelled
in a similar way, but requires a larger melt component
(Table 4, column 2). The chromitites consist mainly
of chromite (60%) and plagioclase (25%) with minor
orthopyroxene and melt components (Table 4, column
3).
The initial melt would have had orthopyroxene and
possibly olivine on the liquidus. Depending on the pressure and oxygen fugacity, 5–10% crystal fractionation is
required for the magma to crystallize plagioclase. The
leuconorite and anorthosite can be modelled as cumulates
from this fractionated melt containing a small amount
of melt (5%) and 95% plagioclase plus orthopyroxene
(Table 4, columns 4 and 5).
Modelling the role of cumulus sulphides
An essential element of Campbell et al.’s (1983) model is
the collection of the PGE by a sulphide liquid, which
settles onto the crystal pile. To test whether cumulus
sulphides are present in the core we may examine Cu
to incompatible element ratios. If no cumulus sulphide
is present the Cu to incompatible element ratio should
reflect that of the silicate liquid, but if cumulus sulphide
is present the ratio will be greater than that of the silicate
liquid. The Cu/Hf ratios for the basaltic andesite and
tholeiitic basalt magmas are 30 and 50, respectively
(Table 2). The Cu/Hf ratios in the core samples vary
from 100 to 30 000 (Fig. 8) and closely follow the S
concentrations, indicating that cumulus sulphides are
present in all of the rocks.
To investigate whether these sulphides are the phases
controlling the PGE the composition of the sulphides in
each of the rock types has been calculated by first
calculating how much Fe is located in the sulphides [Fe(S)
in Table 1] and assuming that the sulphides consist of
pyrrhotite, pentlandite and chalcopyrite. The formula
for calculating Fe(S) is given in the footnotes to Table 1.
The concentrations of all the chalcophile elements and
the PGE were then multiplied by the recalculation factor
100/[S + Cu + (Ni – 0·07) + Fe(S)]. This assumes
that all (except 700 ppm Ni) of these elements are located
in the sulphides. Some Ni is present in orthopyroxene
and chromite and 700 ppm is to allow for this. The value
Fig. 9. Chondrite-normalized REE plots. Normalization values from
Taylor & McLennan (1985).
119
JOURNAL OF PETROLOGY
VOLUME 43
700 is based on the Ni concentrations in the melanorites
containing <0·1% S (samples 13, 15 and 17 in Table 1)
and on the observation that the orthopyroxenes of the
Reef contain >600 ppm Ni (Kruger & Marsh, 1985).
The compositions of the sulphides calculated from the
weighted average for Impala and for SARM-7 [Table 5,
(a)] are very similar (except for Sb) indicating that the
Reef has a fairly uniform composition on a large scale.
The average Reef sulphide contains approximately 6%
Cu, 12% Ni, 30 ppm Ag, 1·2 ppm Sb, 130 ppm Se and
460 ppm PGE. The compositions of sulphides from
individual rock types contain similar amounts of Ni, Cu,
Ag, Sb and Se, but the PGE contents vary from unit to
unit [Table 5, (a)]. The sulphides from the reef
leuconorite, anorthosite and melanorite contain similar
PGE content to the weighted average. In contrast, sulphides from the hanging-wall norite are depleted in PGE
and contain only 15 ppm PGE. The sulphides from the
chromitites are much richer in Os, Ir, Ru, Rh and Pt
than the weighted average.
A feature of the sulphide collection model is that
the PGE concentrations in the sulphides are critically
dependent on the volume of magma with which the
sulphide liquid interacted. The ratio of silicate to sulphide
liquid is referred to as the R factor in closed systems
(Campbell & Naldrett, 1979). More recently Brügmann
et al. (1993) suggested that in the case where a sulphide
droplet is sinking through a magma column a variation
of the zone refining equation would be more appropriate
to model the sulphide composition;
Cs=Cl[D − (D − l)e–(1/DN ) ]
(where Cs is the concentration of an element in the
sulphide liquid, Cl is the concentration of the element in
the silicate liquid, D is the partition coefficient for the
element between silicate and sulphide liquid and N is the
number of volumes of magma with which the sulphide
liquid interacted) is used in this work, but the results
using the R-factor equation are similar.
The compositions of the sulphides calculated using this
equation are listed in Table 5, (b), for N = 40 000,
20 000, 15 000 and 1000 assuming that the magma was
a mixture of 60% basaltic andesite and 40% tholeiitic
basalt. The composition of the sulphides for the weighted
average is similar to that of the model sulphide at N =
20 000 [Table 5, (b)]. The composition of the leuconorite
and melanorite sulphides are similar to the model sulphides at N = 40 000 and N = 15 000, respectively (Table
5). The composition of the whole rocks can be modelled
by assuming they contain 1–5% sulphides of these compositions (Fig. 7b).
It is not possible to calculate the composition of the
sulphides in the footwall leuconorite because of the very
low abundance of sulphides in these rocks. However, the
shape of the metal patterns indicates that PGE are not
NUMBER 1
JANUARY 2002
enriched or depleted relative to Ni and Cu in these rocks.
Using a sulphide liquid formed in equilibrium with the
melt at N = 1000 (Table 5b) and allowing 0·05–0·1%
cumulus sulphides it is possible to model the whole-rock
compositions (Fig. 7a).
The sulphides in the hanging-wall melanorite and
norite are depleted in PGE relative to Ni and Cu. The
silicate liquid that remained after the segregation of the
reef sulphides would have been depleted in PGE. Thus,
the hanging-wall sulphides could have formed by segregation from this PGE-depleted liquid. The composition of the sulphide in the hanging wall may be
modelled as having formed from the melt after 0·006%
sulphides have been removed (Table 5b). The composition of the whole rocks may be modelled as assuming
that they contain 0·5–2% sulphides of this composition
(Fig. 7c).
It is not possible to model the composition of the
sulphides from the chromite-bearing rocks as a sulphide
liquid. The sulphide liquid at N = 40 000 has approximately the correct Cu, Ni, Se, Pd and Au values
[Fig. 7d, Table 5, (b)], but not enough Os, Ir, Ru, Rh
and Pt. Chromitites throughout the Critical Zone are
enriched in Os, Ir, Ru and Rh (Scoon & Teigler, 1994;
von Gruenewaldt et al., 1986). Monosulphide solid solution (mss) accepts Os, Ir, Ru and Rh in preference to
Pt, Pd and Au, and Maier & Barnes (1999) used this
observation to suggest that mss crystallized from sulphide
liquid in the chromite layers. During compaction the
fractionated sulphide liquid was squeezed out leaving a
chromite cumulate with a small mss component rich in
Os, Ir, Ru and Rh. We can apply this model to the
Merensky chromitites. An mss cumulate formed from
the sulphide liquid (N = 40 000) would have sufficient
Os, Ir, Ru and Rh [Table 5, (b), Fig. 7d] to match the
chromite sulphides, but it would contain too little Cu,
Pt, Pd and Au. Therefore, some other phase must be
involved.
Experimental studies have found that at low fS2 metal
alloys and platinum-group minerals enriched in Os, Ir,
Ru and Pt may crystallize directly from sulphide liquids
(Fleet & Stone, 1991; Li et al., 1996; Jana & Walker,
1997). It is possible to model the composition of the
chromite-bearing rocks by considering them to be mixture
of 1% sulphides liquid plus 0·005% PGM [Table 5, (b),
Fig. 7d].
DISCUSSION
There are, broadly speaking, three models currently
favoured for the formation of the Merensky Reef: (1)
collection of the PGE from the silicate magma by an
immiscible sulphide liquid; (2) crystallization of the PGE
directly from the silicate magma as PGM; (3) collection
120
121
34·67
14·9
31·8
0·089
40·28
38·46
38·38
38·09
38·43
37·29
38·29
38·05
13·75
6·01
9·70
3·50
1·02
0·006
0·025
0·006
0·024
1000
550
0·2
1
13·75
13·75
13·75
11·52
13·30
13·75
12·36
12·33
12·30
14·36
10·99
15·68
15·16
13·06
Ni
(%)
6·00
6·00
6·00
3·79
5·55
1·20
6·46
5·40
5·99
7·66
5·94
13·32
6·01
8·27
Cu
(%)
3·93
9
40·90
43·81
43·32
39·89
44·64
33·71
40·54
40·62
Fe
(%)
0·034
0·032
1000
0·2
34
34
34
34
21
31
6·8
14
31
38
44
32
66
32
<150
Ag
(ppm)
0·050
0·050
25
0·1
1
1·25
1·25
1·25
1·25
1
0·13
194
2·82
1·68
2·44
<5·5
1·28
<5·6
Sb
(ppm)
91607
6407
4233
3423
286
180
19220
5453
4895
7536
3150
263
45291
16317
66679
Os
(ppb)
7·1
%
0·13
0·29
0·122
0·048
1000
30000
0·8
3
130
130
130
130
82
120
104
125
109
130
146
132
164
135
121
Se
(ppm)
551366
39766
26276
21248
1772
1114
119299
36668
33414
48707
15279
1178
296968
108697
572597
Ru
(ppb)
184660
23860
15766
12749
1063
669
47719
15730
18650
20267
7042
502
125126
53321
248529
Rh
(ppb)
6·8
%
42·63
%
1·23
0·3
13·7
%
0·28
1·8
1·08
0·046
0·298
0·179
30000
30000
30000
3
3
2
87786
6186
4087
3305
276
173
18558
6371
5730
6864
3080
323
60534
19787
117877
Ir
(ppb)
249505
216505
143060
115686
9648
6068
43301
108755
118891
231767
108819
2747
225467
180327
389160
Pd
(ppb)
50151
50151
36664
30604
2831
3167
10030
32045
24089
47137
30869
15531
34544
69124
31614
Au
(ppb)
79·33
36·1
%
1·10
%
17
9·8
2·9
2·74
1·62
0·87
30000
30000
20000
0·2
0·2
0·2
3107733
366733
242327
195958
16342
10278
73347
296337
290622
379628
214818
11440
2503280
858850
3480342
Pt
(ppb)
∗Composition of sulphides calculated by assuming that they consist of pyrrhotite, pentlandite and chalcopyrite and all the chalcophile elements are located in
sulphides.
†Peach et al. (1990) and Fleet et al. (1999), except for Ag and Sb. These were adjusted so that the modelled Ag and Sb matched the observed Ag and Sb.
‡Barnes & Maier (1999) and references therein plus Theriault & Barnes (1998).
sul, sulphide.
Composition of PGM used in
model
laurite (Mostert et al., 1982)
cooperite (Mostert et al., 1982)
malanite (Hey, 1999)
(c) Parameters of the models
60:40 mix of initial magmas
depleted magma, −0·006% sul
D silicate liquid/sulphide liquid†
D mss/sulphide liquid‡
(b) Model compositions
N = 40000 undepleted magma
N = 20000 undepleted magma
N = 15000 undepleted magma
N = 1000 undepleted magma
N = 4000 depleted magma
mss cumulate from N = 40000
sul
N = 40000 sulphide plus 0·52
PGM%
In the proportions 0·12% laurite,
0·3% cooperite, 0·1% malanite
(a) Observed compositions∗
Weighted average Impala,
samples 1–28
SARM-7
Leuconorite n = 4
Reef melanorite n = 3
Hanging-wall melanorite n = 4
Upper chromitite n = 1
Chr-bearing melanorite n = 1
Lower chromite n = 1
S
(%)
Table 5: Comparison of the observed composition of the sulphides in the Merensky Unit with model compositions
BARNES AND MAIER
PGE IN THE MERENSKY REEF
JOURNAL OF PETROLOGY
VOLUME 43
of the PGE from below the Reef by a Cl-rich fluid. The
implications of our data for each of these models will
now be considered.
Sulphide collection model
As outlined in the section on geochemistry, in the sulphide
collection model the mixing of the resident magma and
a new injection of magma resulted in the mixed magma
becoming saturated in an immiscible sulphide liquid.
This liquid collected the PGE and settles through the
magma onto the top of the crystal pile to form a thin
layer enriched in sulphides and PGE (i.e the Reef )
(Campbell et al., 1983; Naldrett et al., 1986).
Maier & Barnes (1999) used metal to incompatible
element ratios to argue that many of the Lower and
Critical Zone rocks contain too much Cu and PGE to
be accommodated in the melt fraction of these rocks,
and that a small quantity of cumulus sulphides rich in
PGE must be present throughout the Lower and lower
Critical Zones. If this is true, then the Bushveld magma
would appear to have been sulphide saturated throughout
the formation of much of the Lower and Critical Zones
and the timing of sulphide saturation is not the critical
factor in forming a reef. The reason for the Merensky
and UG-2 reefs being economically exploitable is that
they contain more of these PGE-rich sulphides than the
other Lower and Critical Zone rocks; however, collection
of metals by sulphides formed at high R factors occurred
throughout the formation of the Lower and Critical
Zone.
Cawthorn (1999) has criticized the sulphide collection
model on the grounds that many of the Critical Zone
rocks, and in particular those overlying the Reef, contain
insufficient S for the melt to be saturated in sulphides,
assuming that sulphide saturation is achieved at >1000
ppm S. Maier & Barnes (1999) showed that the Lower
Zone and upper Critical Zone samples contain sufficient
S for the melt fraction to have been sulphide saturated,
but many of the lower Critical Zone samples do not.
Therefore, if the presence of cumulus sulphide is the
source of enrichment of PGE and Cu in the lower Critical
Zone it is argued that many of the samples have lost S.
This could have happened when the cumulus pile was
reheated by a fresh injection of magma, or possibly when
the percolating intercumulus fluids partially dissolved the
sulphides (e.g. Willmore et al., 2000).
In the case of the Impala samples, they contain sufficient S to be considered sulphide saturated (Table 1).
Furthermore, sulphides are present in all samples as
interstitial material. In the Reef, sulphides appear to have
filled the open pore spaces and extensional fractures that
developed in the partially consolidated cumulus pile.
Thus, the sulphide collection model is supported by the
petrographic observations.
NUMBER 1
JANUARY 2002
The sulphide collection model may be applied to the
Impala data as follows. Before the Merensky Unit was
formed there was a crystal pile of plagioclase and pyroxene (the proto-leuconorite and anorthosite). These
contained a small quantity of sulphides formed at a
moderate N (1000) overlain by fractionated liquid that
would have been sulphide saturated (Fig. 10a). This part
of the magma’s history is preserved in the footwall
leuconorite. Then a fresh injection of magma mixed
vigorously with the resident magma. The sulphide droplets interacted with a large volume of magma and thus
they have a high N factor (Fig. 10b). The mixed magma
was also saturated in chromite. The sulphide liquid, and
chromite, settled onto the crystal mush to form the lower
chromite seam (Fig. 10c). Later the amount of chromite
accumulating decreased and a layer consisting of a chromite-bearing melanorite formed (Fig. 10d). As there are
two chromite seams it is necessary to argue that the
process occurred twice (Viljoen, 1999). The melanorite
higher in the section contains sulphides that formed from
a PGE-depleted magma and therefore crystallized from
the evolved magma (Fig. 10e). The crystal pile would
behave as a solid framework once 55% crystals were
present. The sulphide liquid could have percolated downwards into the pore space. Furthermore, during compaction some extensional fractures may have opened up
and the PGE-rich sulphide from the chromite-bearing
melanorite layer could have percolated into the underlying anorthosite and leuconorite, resulting in the presence of PGE-rich sulphides in this layer (Fig. 10d). This
model is satisfactory for the silicate rocks; however, the
chromite-bearing rocks contain too much Os, Ir, Ru, Rh
and Pt to be modelled in this fashion.
To explain the concentrations of the PGE in the
chromite layers some post-crystallization modification of
the sulphides is necessary. In modelling the whole-rock
concentrations of the chromitite layers in addition to the
1% cumulus sulphides it was necessary to invoke the
crystallization of 0·005% PGM. The crystallization of
PGM from a sulphide liquid requires that the S concentration of the liquid be very low, which raises the
question of why the S concentration of the sulphide liquid
in the chromitites should be lower than in the silicate
rocks. It has been noticed that sulphides from many of
the chromitite layers of the Bushveld are very rich in
Cu and Ni. This led Naldrett & Lehmann (1988) and
Mathez (1999) to suggest that sulphide liquid in the
chromite layers interacts with the chromite by the
reaction
4/3 Fe2O3 (chr) + 1/3 FeS = Fe3O4 (chr) + 1/6S2.
The S would then leave the system and remaining
sulphide liquid could become S poor. Possibly during
compaction of the chromite-bearing rocks the sulphide
liquid reacted with the chromite, which resulted in loss
122
BARNES AND MAIER
PGE IN THE MERENSKY REEF
Fig. 10. The collection by sulphide liquid followed by re-equilibration of the sulphides model. (a) Formation of the leuconorite as a
plagioclase–orthopyroxene cumulate. (b) Injection of new magma results in vigorous mixing of the sulphide droplets to form PGE-rich sulphide.
(c) Chromite, orthopyroxene and PGE-rich sulphide liquid settles onto the crystal mush and forms the basal chromitite. (d) The proportion of
chromite accumulating decreases and a melanorite containing chromite forms; the PGE-rich sulphide percolates through the crystal mush. (e)
The overlying magma is PGE depleted and a PGE-poor sulphide segregates. (f ) The sulphide droplets in the chromitite layers lose Fe to the
chromite and release S; this lowers the fS2 of the sulphide liquid and PGM crystallize. (g) During compaction some of the fractionated sulphide
liquid escapes from the chromitites.
of Fe and S from the sulphide liquid, and the S content
of the liquid dropped sufficiently for PGM to crystallize
from the sulphide melt (Fig. 10f ). Finally, some of the
fractionated sulphide melt was squeezed out of the
chromitites during compaction (Fig. 10g).
Crystallization of PGE from the magma as
PGM
As mentioned above, Cawthorn (1999) has argued that
many of the Lower Zone and Critical Zone rocks contain
insufficient S to be considered sulphide liquid saturated.
Further, the Lower and Critical Zone samples have
different PGE patterns from proposed initial magmas
(the basaltic andesite and tholeiitic basalt). In general,
the Lower and Critical Zone rocks are enriched in Os,
Ir, Ru and Rh relative to Pd and Pt, whereas the melts
are enriched in Pt and Pd relative to Os, Ir, Ru and Rh.
This is contrary to the sulphide collection model, as
the partition coefficients for PGE between silicate and
sulphide liquids are similar and thus the sulphide collection model predicts that all the PGE should be concentrated equally. It has long been known that Os, Ir
and Ru tend to concentrate in olivine and chromite-rich
rocks (e.g. Crocket, 1981). Some workers have suggested
that these elements crystallize directly from magmas
as laurite and Os–Ir alloys (Keays & Campbell, 1981;
Tredoux et al., 1985). Hiemstra (1979) suggested a model
for the Bushveld in which the Pt alloys and laurite
crystallized from the magma and then were settled out
by inclusion in chromite. Cawthorn (1999) suggested that
throughout the formation of the Critical Zone the PGE
crystallize continuously, directly from the magma as
PGM. Normally, the accumulating PGM would be diluted by the crystallization of silicates and oxides and no
economically significant concentrations would form. In
the case of the Merensky Reef, Cawthorn suggested that
123
JOURNAL OF PETROLOGY
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NUMBER 1
JANUARY 2002
Fig. 11. The cluster model. In a silicate melt PGE ions and chalcophile elements could cluster together. (a) Once the magma becomes saturated
in a sulphide liquid the cluster would dissolve in the liquid to form a PGE-rich sulphide liquid. (b) If the magma becomes saturated in chromite,
Ru could partition into the chromite and this may destabilize the cluster, resulting in the co-precipitation of chromite and PGM.
the mixed magma was up to 100°C above the silicate
and oxide liquidus. As a result, oxide and silicate crystallization temporarily ceased, but PGM crystallization
continued. In this model, the Reef did not form by
enrichment of PGE, but rather because of the absence
of other phases that normally dilute the PGE.
A weakness of the direct PGM crystallization model is
that it requires that the magma becomes saturated in
PGM when the PGE are present at only ppb levels (e.g.
Mathez, 1999). To circumvent this problem, Tredoux et
al. (1995) have suggested that PGE and other nonlithophile elements could aggregate together in a silicate
magma as clusters (Fig. 11). These clusters would then
pre-concentrate the PGE. Theoretically, the heavy PGE
(Os, Ir and Pt) should have a tendency to form clusters
more readily than the light PGE (Ru, Rh, Pd) (Tredoux
et al., 1995). Thus the PGM formed from these clusters
would be Os, Ir, Pt rich and this would explain the
enrichment of the Lower and Critical Zone rocks in these
elements.
The main weakness of the cluster model is that the
existence of the PGE clusters has not been demonstrated
experimentally and therefore, their existence is strongly
questioned (e.g. Mathez, 1999). On the other hand, Fleet
et al. (1999) found that there is a wide range in partition
coefficients for PGE into sulphide liquid, from 1000 to
100 000, and that the partition coefficients are positively
correlated with PGE concentrations. The low partition
coefficients were obtained from runs that contained PGE
levels in the ppb range, i.e. similar to mafic magmas; in
contrast, the high partition coefficients were obtained for
runs that contained PGE at the ppm level, which is higher
than most magmas contain. Tredoux et al. (1995) suggested
that clusters of PGE ions could have formed in the PGErich experiments and that the clusters could have been
precipitated out by the sulphide liquid giving an apparently
higher partition coefficient than would occur in nature.
Ballhaus & Sylvester (2000) have applied this idea to the
Merensky Reef and suggested that the magma from which
the Reef formed contained clusters and that when sulphide
saturation occurred these clusters were included in the
sulphide liquid and precipitated along with the sulphides
to form the Reef. As the clusters should be Os, Ir and Pt
rich (Tredoux et al., 1995) this would explain the high
Os, Ir and Pt concentrations of the Merensky Reef. This
may be true, but the similarity between the empirical
partition coefficients calculated from sulphide droplets in
mid-ocean ridge basalt (MORB) and ocean island basalt
(OIB) (Peach et al., 1990; Helz & Rietz, 1988) and the
experiments yielding the high partition coefficients is hard
to ignore and suggests that sulphide liquid–silicate liquid
partition coefficients in the 20 000–50 000 range do indeed
occur in nature and are not an experimental artefact.
There is, however, an alternative explanation for the
variations observed in the Fleet et al. (1999) experiments.
The experiments that give low partition coefficients were
performed at low-S activities, so that alloys would have
been present in the experimental runs; this is inappropriate
for most lithospheric conditions. In contrast, the high
124
BARNES AND MAIER
PGE IN THE MERENSKY REEF
partition coefficients were determined in S-rich experimental runs, conditions that would more closely
simulate the natural case.
As evidence for the existence of clusters, Ballhaus &
Sylvester (2000) cited the irregular distribution of PGE
in the sulphides of the Merensky Reef; they suggested
that the PGE-rich patches represent frozen clusters. However, it is also possible that the PGE-rich patches in the
sulphides represent proto-exsolution phases, because the
sulphide liquid would have crystallized as monosulphide
solid solution and then undergone exsolution to pentlandite, chalcopyrite and pyrrhotite at 600°C. Makovicky
et al. (1986) showed that as temperature drops, the solubility of PGE in mss decreases dramatically. Hence it
seems unlikely to us that clusters would survive the
solidification and exsolution process.
The cluster model could be applied to our dataset as
follows. The first stage is the same as in the sulphide
collection model; a leuconorite containing a sulphide
formed at a moderate N is overlain by crystal mush and
fractionated liquid. A new batch of magma is emplaced
and mixes with the fractionated magma. Platinum-group
element clusters form in the mixed liquid (Fig. 11). The
magma becomes sulphide saturated and the sulphides
incorporate any clusters they encounter and thus became
PGE rich. These sulphides settle through the magma
onto the crystal mush and percolate into the underlying
leuconorite.
To explain the enrichment of Os, Ir, Ru, Rh and Pt
in the chromitites an additional process is necessary.
Capobianco et al. (1994) showed that Ru and Rh partition
into oxides, but Pd does not. They suggested that oxide
precipitation provokes the crystallization of the Os, Ir
and Pt clusters whereas Ru and Rh are incorporated
into the oxide structure. Adopting this, one could suggest
that when chromite segregated from the hybrid magma
Ru, and to a lesser extent Rh, partitioned into the
chromite and that destabilized the cluster, resulting in
the co-precipitation of chromite and the clusters to form
chromite-rich layers enriched in all the PGE except for
Pd (Fig. 11).
The crystallization of chromite destabilizes the cluster
and allows the precipitation of the PGE clusters as PGM.
The presence of the PGM enriched Os, Ir, Ru, Rh and
Pt in the chromitite layer. All of the rocks contain some
cumulus sulphide; in the chromite layers it could be
argued that in addition to the sulphides, Ru–Rh-bearing
chromite and Os–Ir clusters precipitated.
Collection of PGE from below by hydrous
intercumulus fluid model
In the collection-from-below model, PGE, base metals
and S are collected from the cumulus pile by a Cl-rich
fluid that exsolved from the intercumulus fluid (Boudreau
& Meurer, 1999a,b; Willmore et al., 2000). This fluid rises
into the overlying magma, mixes with it and causes an
immiscible sulphide liquid to form, which collects up the
PGE to form the Reef. Many workers have remarked
on the coarse-grained nature of the Merensky Reef,
referring to it as a pegmatoid, and used this as evidence
of a high fluid content for the magma from which it
crystallized. Nicholson & Mathez (1991) and Mathez et
al. (1997) have presented a fairly complex model for
the formation of the Merensky Unit rocks involving
replacement of the original semi-consolidated cumulate
by a ‘hydration–melting front’. Our sampled section
contains very little mica and quartz and no K-feldspar,
and has low contents of incompatible trace elements.
Thus, a relatively high proportion of late-stage components, including fluid, is not apparent. The ‘pegmatoid’
is coarse grained rather than a rock with pegmatite
texture, and need not necessarily indicate the presence
of fluids. All that is required for a coarse grain size to
develop is that the rock be held close to the solidus
temperature for longer than usual. As Cawthorn & Barry
(1992) pointed out, this could have occurred when the
chamber was replenished. In the case of the Merensky
Reef it could be argued that two batches of hot primitive
magma were introduced in rapid succession. These could
have held the crystal pile at close to the solidus for longer
than normal, thus allowing larger than normal crystals
to grow. Further, the large crystals in our reef melanorite
are not single crystals, as would be expected to form
from a hydrous fluid, but they appear to be composite
grains formed by static recrystallization when the orthopyroxenes were forced into each other during compaction.
The concentrations of LILE and other incompatible
elements in our samples are low and can be modelled by
assuming they are hosted in the silicate liquid (Table 4).
Similarly, Cawthorn (1996) showed that the REE-rich
clinopyroxenes cited by Mathez (1995) as evidence of
interaction with metasomatic fluid could have crystallized
from a fractionated liquid. None the less, the presence of
10–40% oikocrysts (plagioclase in the chromitite and
melanorite and orthopyroxene in the anorthosite and
leuconorite) could be used as evidence that the crystal
pile originally had a porosity at least this high and thus
melt migration during compaction appears inevitable.
The oikocrysts could have crystallized from the liquid
and, during compaction, the fractionated intercumulus
liquid could have risen through the crystal pile, eventually
entering the overlying magma. This model may be a little
too simple, as it assumes that the migrating liquid does
not react with the crystal pile through which it migrates.
It is possible that the crystal pile was layered: a bottom
layer consisting of a framework of plagioclase crystals
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JOURNAL OF PETROLOGY
VOLUME 43
overlain by a layer consisting of a framework of orthopyroxene crystals. During compaction the framework
deformed and the intercumulus fluid rose through the
pile and crystallized either plagioclase or orthopyroxene
oikocrysts, which cemented the pile. The chromite could
crystallize from a hydrous intercumulus melt as suggested
by Mathez et al. (1997). The critical question, originally
raised by Barnes & Campbell (1988) in their review of
the problem for PGE reefs in general, is: ‘Does the possible
migration of intercumulus liquid affect the distribution of
the PGE and other chalcophile elements?’
Boudreau & Meurer (1999a,b) suggested that migrating
hydrous fluid derived from the underlying intercumulus
melt dissolves the sulphides, and the accompanying PGE,
and transports them into the overlying cumulates which
contain vapour-undersaturated interstitial fluids. The two
fluids mix and sulphides are deposited from the fluids to
form the Reef. We argue that in the silicate rocks the PGE
and chalcophile element concentrations are controlled by
the presence of sulphides of a composition that represents
a sulphide liquid in equilibrium with the silicate magma.
The sulphides below the Reef are not depleted in PGE
in this section. The sulphides overlying the Reef are
depleted in PGE relative to Ni and Cu. The observation
that sulphides underneath the Reef are not depleted in
PGE and the sulphides above the Reef are depleted in
PGE is also true on a much larger scale. Maier & Barnes
(1999) found that throughout the Lower and Critical
Zones the rocks are largely enriched in PGE and rocks
overlying the Reef are depleted in PGE. These observations are a strong argument, in our opinion, for a
model whereby the PGE are extracted from the overlying
magma rather than being obtained from the underlying
intercumulus liquid. Thus, although there is evidence for
migration of intercumulus silicate liquid, this has not
changed the PGE distributions.
NUMBER 1
JANUARY 2002
distribution of all the metals can be accounted for by
assuming that they were controlled by a sulphide liquid
that formed in equilibrium with a large volume of magma.
The composition of the sulphide liquid evolved from
PGE enriched in the lower portions of the section to
PGE depleted in the upper portions, probably reflecting
the evolution of the silicate liquid. The chromitites and
chromite-bearing melanorite between them appear to
have 3–4 times too much Os, Ir, Ru, Rh and Pt to be
accounted for by accumulation of sulphide liquid. This
could be because the sulphide liquid trapped between
the chromite grains reacted with chromite resulting in
the loss of S. As a result, the fS2 of the sulphide liquid
dropped sufficiently for PGM (laurite, cooperite and
malanite) to crystallize from the sulphide liquid. Some
fractionated sulphide liquid was squeezed out of the
chromite layers into the overlying magma, leaving the
chromitites enriched in Os, Ir, Ru, Rh and Pt.
ACKNOWLEDGEMENTS
The authors wish to thank Impala Platinum Ltd. for
providing the sample material and granting permission
to publish the results. Professor E. W. Sawyer (UQAC)
is thanked for reading the early drafts of the manuscript
and for advice on interpreting the microstructures. Miss
Valerie Becu is thanked for her help with the PGE
analyses. This research was supported by an individual
operating grant from the Natural Science and Engineering Research Council of Canada (to S.J.B.) and by
a research development grant of the University of Pretoria
(to W.D.M.).
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CONCLUSIONS
The studied section of the Merensky Reef consists of
leuconorite overlain by anorthosite, chromitite, coarsegrained melanorite, chromitite and melanorite. Therefore
lithology does not appear to be important in controlling
the PGE distribution. All rock types contain 1–4% interstitial sulphides. The rock type beneath the Reef is
leuconorite and above the Reef it is norite. All the rocks
consist of a framework of a cumulus phase (plagioclase,
chromite or orthopyroxene depending on the rock type)
with 10–40% oikocrysts. The microstructures indicate
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by compaction while the rock was crystallizing. The
sulphides fill the interstitial space and form a network
parallel to the vertical, and appear to have percolated
down into the cumulus pile. In the silicate rocks the
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