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Composition of the ultramafic-mafic contact interval of the Great Dykeof Zimbabwe at Ngezi mine: Comparisons to the Bushveld Complex andimplications for the origin of the PGE reefs

W.D. Maier, S. Maatta, S. Yang, T. Oberthur, Y. Lahaye, H. Huhma, S.-J. Barnes

PII: S0024-4937(15)00327-8DOI: doi: 10.1016/j.lithos.2015.09.007Reference: LITHOS 3689

To appear in: LITHOS

Received date: 10 June 2015Accepted date: 12 September 2015

Please cite this article as: Maier, W.D., Maatta, S., Yang, S., Oberthur, T., La-haye, Y., Huhma, H., Barnes, S.-J., Composition of the ultramafic-mafic contact in-terval of the Great Dyke of Zimbabwe at Ngezi mine: Comparisons to the BushveldComplex and implications for the origin of the PGE reefs, LITHOS (2015), doi:10.1016/j.lithos.2015.09.007

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Composition of the ultramafic-mafic contact interval of the Great Dyke of Zimbabwe at Ngezi

mine: Comparisons to the Bushveld Complex and implications for the origin of the PGE reefs

WD Maier1, S Määttä

2, S Yang

2, T Oberthür

3, Y Lahaye

4, H Huhma

4, S-J Barnes

5

1 School of Earth and Ocean Sciences, Cardiff University, UK, Tel: +44 7554645684, Fax:+44 29

20874326, Email: maierw@cardiff.ac.uk

2 Oulu Mining School, Oulu University, Oulu, Finland

3 Bundesanstalt für Geowissenschaften und Rohstoffe (BGR), Hanover, Germany

4 Geological Survey of Finland (GTK), Espoo, Finland

5 Université du Québec à Chicoutimi, Canada

Abstract

The Great Dyke contains the world’s second largest platinum resource after the Bushveld Complex.

Isotopic and trace element data from the interval straddling the contact between the Ultramafic and

Mafic Sequences of the Great Dyke indicate a less enriched composition than in the Bushveld

Complex (Great Dyke: Sri 0.7024-0.7028, εNd mostly -1 to +1, Ce/Sm 2-6; Bushveld: Sri 705-

0.709, εNd -5 to -7, Ce/Sm 5-15). These data are interpreted to reflect relatively moderate amounts

of contamination of the Great Dyke parent magma. All analysed isotopes show little variation

across the Main Sulphide Zone and the ultramafic-mafic contact. This corroborates earlier work by

other researchers that the Great Dyke crystallized from a single magma type. Mixing of

compositionally distinct magmas, proposed to have caused sulfide melt saturation in the Bushveld

Complex, seemingly played little or no role in the formation of the PGE mineralization in the Main

Sulphide Zone, and neither did enhanced crustal contamination of specific magma batches. Instead,

sulfide melt saturation of the magma was likely triggered by silicate fractionation. The mechanism

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of concentration of the sulfide melt remains uncertain, but theoretical considerations suggest that

phase sorting in response to slumping of crystal mushes, possibly caused by chamber subsidence,

played an important role. This model is consistent with the highly irregular, undulating nature of the

contact between the mafic and ultramafic zones of the intrusion, in the hanging wall of the Main

Sulfide Zone.

1. Introduction

The Great Dyke of Zimbabwe is one of the world’s largest layered mafic-ultramafic intrusions and

hosts the second largest PGE resource globally, i.e., the Main Sulfide Zone (MSZ, Prendergast and

Wilson, 1989; Prendergast and Keays, 1989; Wilson and Prendergast, 1989; Wilson, 1996, 2001).

The mineralization occurs in the transition interval between the Ultramafic and Mafic sequences of

the Great Dyke, analogous to most other PGE reefs elsewhere. In the richest layered intrusions

globally, namely the Bushveld Complex of South Africa and the Stillwater Complex of Montana,

there is a sharp change in isotopic signature and incompatible trace element ratios across this

transition zone. To many workers this suggests that the PGE reefs formed through mixing of two

types of compositionally distinct magmas (magnesian basalt and Al-tholeiite; Campbell et al., 1983;

Naldrett et al., 1986; Scoon and Teigler, 1993; Lambert et al. 1994; Naldrett et al., 2012), or via

enhanced contamination of the magma (Arndt et al., 2005). In the present paper, the transition

interval of the Great Dyke is examined to test whether these models could also apply to the MSZ.

The results have implications for the general relevance of the magma mixing and contamination

models in PGE reef formation.

2. Geological overview

2.1 Geology and structure of the Great Dyke

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The Great Dyke forms a ~550 km long and 4 to 11 km wide, north-northwest trending, elongated

igneous body that intruded Archaean granitoids and greenstone belts of the Zimbabwe craton at

about 2575 Ma (Mukasa et al., 1989; Wingate, 2000; Oberthür et al., 2002). The intrusion consists

of two major magma chambers, the North and the South Chamber which can be further subdivided

into five sub chambers, the Musengezi, Darwendale and Sebakwe sub-chambers in the north, and

the Selukwe and Wedza sub-chambers in the south (Fig. 1).

Borehole data and gravity studies have shown that the Great Dyke has a boat-like structure

with a deep keel (Fig. 2), interpreted to represent a near-continuous feeder dyke along most of the

length of the intrusion (Wilson, 1982; Prendergast and Wilson, 1989, Wilson and Prendergast,

1989, Wilson, 1996). Based on gravity profiles (Podmore and Wilson, 1987) it is estimated that the

North Chamber has a considerably larger volume than the South Chamber. The layers thin and

steepen towards the margin, eventually merging into the fine grained Border Group that dips at

about 20-35° (Wilson and Prendergast, 1989).

The lower Ultramafic Sequence has a discrete stratigraphy in the different sub-chambers,

suggesting that the sub-chambers were not linked at the early stage of magma emplacement. In

contrast, the Mafic Sequence is of similar composition and stratigraphy in all sub-chambers,

suggesting that the sub-chambers had become linked at this stage in the filling of the chamber

(Prendergast and Wilson, 1989, Wilson and Prendergast, 1989, Wilson, 1996).

In comparison to most other layered intrusions, the central portions of the Great Dyke are

relatively well exposed. This raises the interesting question whether certain structural features of the

Great Dyke, such as the presence of central feeder dykes and cumulate layers that thicken and

flatten towards the centre of the intrusions also occur in other intrusions, e.g., the Bushveld

Complex.

2.2 Stratigraphy

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The exposed stratigraphy of the Great Dyke amounts to about 3500 m (Prendergast and Wilson,

1989; Wilson and Prendergast, 1989; Wilson, 1996). The sequence is sub-divided into a lower

Ultramafic Sequence and an upper Mafic Sequence (Fig. 3). The Ultramafic Sequence is further

sub-divided into the basal Dunite Succession which is overlain by the Bronzitite Succession (Fig.

3). Both host a number of chromitite layers generally occurring at the base of cyclic units

comprising, from the base to the top, dunite, harzburgite and, in some cases, orthopyroxenite. The

Mafic Sequence is sub-divided into the Lower, Middle and Upper Mafic Successions, all consisting

of gabbronorite and subordinate gabbro and norite. Due to erosion and the doubly-plunging

synclinal nature of the subchambers, only remnants of the Mafic Sequence are preserved at four

locations (Fig. 1).

The uppermost portion of the Ultramafic Sequence is of particular interest as it hosts the bulk

of the economically important PGE mineralization. It consists mainly of orthopyroxenite cumulates

grading into websterite with height. These lithologies constitute the approximately 200 m thick P1

pyroxenite layer situated at the top of Cyclic Unit 1 (Fig. 3), which hosts the Main Sulphide Zone

(Prendergast and Keays, 1989; Wilson, 1992).

As has been briefly mentioned earlier, the Great Dyke shows distinct changes in thickness and

composition of layers with position between the margin and the axis of the synclinal structure (Fig.

2). Most layers, including the P1 pyroxenite, thicken in the down-dip direction, the amount of

interstitial material decreases, and mineral compositions become less evolved, analogous to the

Jimberlana intrusion (Campbell, 1986). For example, the P1 pyroxenite layer measures up to 220 m

in the axial domains of the Wedza sub-chamber, gradually thinning towards the margins of the

intrusion (Prendergast, 1991). The websterite layer is also thickest in the axis (Prendergast and

Keays, 1989; Prendergast and Wilson, 1989; Prendergast, 1991, Wilson, 1996, 2005), e.g., in the

Darwendale sub-chamber where it has a thickness of 33 m in the axis and thins to 7 m at the

margins. The bulk amount of sulfides within, and the thickness of, the MSZ is greatest in the centre

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(Prendergast and Wilson, 1989; Wilson and Tredoux, 1990). Some of these features can potentially

be explained by the enhanced cooling rate of the magma near the walls and floor of the intrusion,

preventing adcumulate growth and trapped melt expulsion at the margin of the intrusion

(Prendergast and Wilson, 1989; Wilson, 1992).

The Mafic Sequence rests with a sharp, often non-planar contact on top of the Ultramafic

Sequence (Fig. 4). In the Darwendale sub-chamber the Lower Mafic Succession (LMS) comprises

approximately 700 m of gabbroic rocks (Wilson and Chaumba, 1997). The mineral compositions of

the LMS show subtle trends of progressive differentiation with height suggesting closed-system

fractionation within the chamber, in stark contrast to the multiple reversals in the cyclic units of the

Ultramafic Sequence, interpreted to represent multiple magma replenishment (Wilson and

Prendergast, 2001).

The Middle Mafic Succession (MMS) consists of relatively less evolved rock-types than the

LMS. In the Darwendale sub-chamber it has a thickness of approximately 100 m, forming narrow

layers of fine- to medium-gained gabbro and feldspathic (ol)orthopyroxenites (Wilson, 1996). The

Upper Mafic Succession (UMS) is approximately 300 m thick in the Darwendale Sub-chamber and

consists mainly of norite. Inverted pigeonite is common and primary magnetite is found near the

top.

2.3 Nature of the Ultramafic-Mafic contact

At Ngezi, the contact between the Ultramafic and Mafic Sequences is generally sharp (Fig. 4). The

websterite tends to become coarse grained or pegmatoidal in its uppermost 1-2m (see also

Prendergast and Keays, 1989), although medium grained contact examples also exist (notably in

drill core MLF45, Fig. 4f). Coarse poikilitic olivine may occur in the top portion of the websterite

(cf. Wilson and Chaumba, 1997) and within the basal few 10s of meters of the gabronorite (Fig. 4c,

d, f). At some localities in the Wedza and Selukwe sub-chambers a thin layer of chromitite is found

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at the websterite-gabbro contact (Prendergast, 1991). In the rare exposed outcrops at Ngezi mine,

the knife sharp contact is undulous and highly irregular, showing apophyses of websterite within the

overlying gabbronorite (Fig. 5). In addition, small irregular fragments of websterite may occur

within the gabbronorite. Whether the fragments represent autoliths or are connected in 3-D to the

main websterite remains uncertain. At Minosa mine in the Wedza sub-chamber Prendergast and

Keays (1989) and Prendergast (1991) observed interdigitating wedges of websterite and gabbro

towards the margin of the intrusion and channels of gabbro within websterite. These were

interpreted to represent “erosional channels orientated perpendicular to the margin of the intrusion,

caused by cascades of plagioclase saturated magma from higher up the walls of the magma

chamber”. The depth of the channels is variable, but in one case the entire MSZ has been removed,

suggesting considerable thickness of the unconsolidated mush zone at the top of the crystal pile.

The base of the websterite may show interdigitating relationships with the underlying

orthopyroxenites. At Hartley mine in the Darwendale subchamber, websterite and gabbronorite are

interlayered in the axis of the intrusion, whereas at the margin websterite is sharply overlain by

homogenous gabbronorite (Wilson and Brown, 2005). In summary, the available evidence indicates

mixing of websterite and gabbronorite crystal mushes.

2.4 PGE mineralization

Analogous to many other layered intrusions, the Great Dyke contains several PGE enriched layers,

including most of the chromitites (Germann and Schmidt, 1999; Oberthür, 2002), and several

silicate horizons in the uppermost portion of the ultramafic sequence (Prendergast and Keays, 1989;

Prendergast and Wilson, 1989; Wilson and Tredoux, 1990; Wilson et al., 1990; Wilson and

Prendergast, 2001; Oberthür, 2002, 2011). Sporadic PGE mineralization with peak grades of about

1 ppm and unknown lateral continuity also occurs in the Middle Mafic Sequence (Wilson and

Wilson, 1981; Podmore and Wilson, 1987).

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The most significant accumulations of PGE occur in the P1 layer, namely within the C1d

chromitite, and the Lower and Main Sulphide Zones (Prendergast and Keays, 1989; Prendergast and

Wilson, 1989; Wilson, 2001; Oberthür, 2002, 2011). The C1d chromitite and its hostrocks contain

up to 2 ppm PGE (Pt/Pd 0.1)(Oberthür, 2002, 2011) and is currently uneconomic for PGE. The

same applies to the Lower Sulphide Zone which forms a broad 30-85 m wide zone containing up to

~1 ppm PGE (average <0.5 ppm) and mostly < 0.5% sulfide (average 0.3%, Wilson, 2001). The

Main Sulphide Zone (MSZ) is a narrow, 2-8 m wide layer in the uppermost portion of the P1 layer

that contains up to 5 ppm PGE and up to 8% sulfides over 2-3 m (Prendergast and Keays, 1989;

Prendergast, 1991; Wilson, 2001; Oberthür, 2011). A characteristic feature of the MSZ are the so-

called “offset” metal distribution patterns, whereby the peak Pd levels occur near the base of the

reef, whereas peak Pt and Cu levels occur at progressively higher stratigraphic positions within the

reef (Prendergast and Keays, 1989; Oberthür 2011). The mineralization, including the offset pattern,

appears to be relatively continuous along strike throughout the Great Dyke, but there is considerable

down-dip variation in the thickness of the MSZ and the amount of sulfides (Wilson and Tredoux,

1990). It is currently mined at several localities, namely Mimosa, Unki, and Ngezi (Fig. 1).

3. Methods

The bulk of our samples are from drill core MLF45 at Ngezi mine (Fig. 1, Table 1). One additional

sample from the Lower Mafic Sequence is from drill core DP13, and two samples underlying the

MSZ are from drill core HN009, both at Ngezi mine. One sample of magnetite gabbronorite located

approximately 800 m above the MSZ was collected from outcrop in the east Selous area.

The MLF45 and DP13 samples were milled in an aluminium vessel at the University of

Quebec at Chicoutimi (UQAC), whereas the remainder were milled in agate at Cardiff University.

The PGE contents in the MLF45 and DP13 samples were determined at UQAC, using ICP-MS after

Ni sulfide fire assay (see Savard et al., 2011 for analytical details). The PGE for the other 3 samples

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were determined at Cardiff University, using the same method (McDonald and Viljoen, 2006). The

lithophile major and trace elements were determined by ICP-OES and ICP-MS at Cardiff

University, and Se by TCF-INAA at UQAC (see Savard et al., 2006, for a description of the

method). Strontium and S isotopes were determined by laser ablation ICP-MS at the Geological

Survey of Finland (GTK) in Espoo, and Nd isotopes by mass spectrometry, also at GTK. Analytical

details are given in Maier et al. (2014).

4. Sample description

The analysed orthopyroxenites contain > 90% orthopyroxene, 2-5% clinopyroxene, 5-7%

plagioclase and 1-2% subhedral or anhedral chromite (Fig. 6a). Most of the rocks are thus

orthopyroxenite mesocumulates. Orthopyroxene forms mostly subhedral cumulus grains with an

average grain-size of 1-2 mm, slightly coarsening with height. Clinopyroxene exsolution lamellae

and blebs are common. Clinopyroxene grains are subhedral with an average grain size of 1 mm, or

form large intercumulus oikocrysts several mm across. Plagioclase is a minor constituent interstitial

to the pyroxenes and has an average grain-size of 0.5 mm. Phlogopite is a trace mineral. Alteration

is relatively minor, but increases somewhat in abundance with height. Sulfides (pyrrhotite,

pentlandite, chalcopyrite, rare pyrite) are mainly confined to samples MLF45- 263m and 261 which

contain around 1% sulfide. Most of the sulfides occur in the intergranular spaces between the

silicates (Fig. 7a). In places, sulfides have been partially replaced, resulting in the formation of

coronas of actinolite, epidote, carbonate, talc, magnetite and pyrite (cf Li et al., 2007) (Fig. 7b-d)

that contain small grains, veinlets and fracture fillings of sulfides, predominantly chalcopyrite.

PGM tend to be concentrated at the margins of the sulfides and within the alteration halo and

consist mainly of cooperite–braggite and moncheite, with laurite, merenskyite, sperrylite, and Fe–Pt

alloys being subordinate (Li et al., 2007).

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The websterite at the top of the P1 layer contains 10-70% clinopyroxene, 80-20%

orthopyroxene, 5-10% plagioclase, up to 4% opaque minerals and minor phlogopite (up to 3 vol. %)

and hornblende. The rocks are mostly mesocumulates, but sample MLF45-242m is an adcumulate,

and the uppermost portion below the contact to the Mafic Sequence is a pegmatoidal orthocumulate

(Supplementary data). The proportion of orthopyroxene decreases with height, whereas that of

clinopyroxene (Fig. 6) and sulfide increases. Plagioclase is strongly enriched in the pegmatoid at the

upper contact of the Ultramafic Sequence, such that the rock is a gabbronorite. Orthopyroxene

forms subhedral grains with an average grain size of 2-4 mm. Cumulus clinopyroxene forms sub- or

anhedral grains averaging 2-3.5 mm in width. Plagioclase is an interstitial phase with an average

grain-size of 1 mm, except in the pegmatoid where grains are > 1 cm wide. The basal sample of

websterite is noticeably enriched in primary magmatic phlogopite (3 vol. %) that may form large

laths. Disseminated sulfides occur either in the intergranular spaces or in veins and cleavage planes

(Fig. 7e-f). The former have lobate shapes, whereas the latter form small veinlets and specks. In

both types, the sulfides consist mainly of pyrrhotite, pentlandite, and chalcopyrite and make up 1-3

vol.% of the rock. Notably, sulfides are closely associated with pyroxene, whereas intercumulus

plagioclase is largely free of sulfides. Apart from the occurrence of pegmatoid at the upper contact,

the grain size of the websterite shows no significant change with height. However, alteration

increases with height.

Gabbronorite contains mostly plagioclase (57%), with pyroxenes forming much of the

remainder of the rock (Supplementary data). Clinopyroxene is up to seven times more abundant

than orthopyroxene. Ortho and clinopyroxene tend to form sub- or andehral grains with an average

grain-size of 1-1.5 mm. Plagioclase forms subhedral tabular grains, with an average grain size of

1.5 mm. Only a few small (<0.1mm) opaque grains are present. Sulfides (po, pent, cp) are mostly

trace phases (<<1%), except in sample MLF45-47m which contains approximately 0.1% very fine

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grained sulfide. The rocks can be classified as adcumulates, consistent with their low Zr contents (<

5 ppm) (Fig. 6).

The gabbro (DP13, 852m) consists mostly of plagioclase (60 %) and pyroxenes (opx:cpx =

1:10) as well as trace amounts of opaque minerals (0.4 %). The abundance of sulfides is markedly

higher (0.5%) than in the gabbronorites, concentrated in a 1cm thick band of disseminated

intercumulus grains up to 1 mm in size. The sulfide mineralogy is similar to the other samples,

comprising pyrrhotite, pentlandite and chalcopyrite. The pyroxenes and plagioclase are mostly

anhedral or subhedral, forming a granular texture with 120° grain boundaries. The average grain-

size is markedly lower than in the gabbronorites, at about 0.1 mm for the pyroxenes, and slightly

larger for plagioclase. The analysed sample is a microgabbro adcumulate.

Magnetite gabbronorite comprises just one of our samples, collected in the Selous area. The

rock consists of cumulus plagioclase (~65%), orthopyroxene (~25%), clinopyroxene (~5%),

magnetite (~5%), and traces of biotite.

5. Analytical results

5.1 Mineral chemistry

The composition of the main minerals shows complex variation across the contact between the

Ultramafic and Mafic sequences, as reported previously by Prendergast and Keays (1989) and

Wilson and Chaumba (1997). The anorthite content of plagioclase increases, reflecting, at least in

part, the change from intercumulus crystals in the websterites to cumulus crystals in the

gabbronorites. Clinopyroxene shows a sharp drop in Mg# and Cr content, whereas Mg# and Cr

content of orthopyroxene are virtually unchanged across the contact (Fig. 6b). However, in a

detailed profile from the Sebakwe subchamber (i.e. the same subchamber that hosts Ngezi mine)

analysed by Wilson and Chaumba ( 1997), there is a clear reversal towards higher Mg#, Cr and Ni

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contents of orthopyroxene across the contact. This is not evident in our data, possibly because our

lowermost gabbroic sample is situated 3 meters above the contact (226.38m).

5.2 Lithophile whole rock geochemistry

The major element and compatible trace element contents of our samples are controlled by the

modal proportions of the main rock forming minerals, i.e., orthopyroxene, clinopyroxene and

plagioclase (Fig. 8a). Figure 8b shows that the orthopyroxenites and websterites have different

Cr/MgO ratios, likely due to relatively higher Cr contents in clinopyroxene than orthopyroxene. The

orthopyroxenites cluster around 3000 ppm Cr, whereas the websterites have between 2500 – 4000

ppm Cr. The gabbros of the Mafic Sequence have much lower Cr contents, due to the larger

proportion of plagioclase, but also the Cr-poor nature of relatively evolved ortho- and

clinopyroxene, consistent with the data of Wilson and Chaumba (1997). The incompatible trace

elements show broadly positive inter element correlations, but the websterites are characterized by

pronounced heterogeneity in most incompatible elements, with the uppermost sample (229.6a)

having strongly enriched incompatible trace element contents. Whether this trend is continued into

the contact pegmatoid remains unknown, as the latter has not been analysed. Notably, of the

analysed lithologies the gabbronorites are most depleted in incompatible trace elements reflecting

their adcumulate nature, whereas most of the ultramafic rocks are mesocumulates.

Lithophile multi-element (“spider”) diagrams show broadly similar patterns for all rock types,

but the gabbroic rocks have lower abundances for most elements (Fig. 9). An exception is the

stratigraphically highest sample (DP13 852.1m) which has much less fractionated incompatible

trace element patterns than the other gabbroic samples, possibly reflecting a different magma

composition. All rock types have pronounced positive Pb anomalies, indicating the presence of a

crustal component. There are only weak negative Nb-Ta anomalies, notably in the ultramafic rocks.

The gabbronorites have positive Sr and Eu anomalies, due to the abundance of cumulus plagioclase.

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Compared to equivalent Bushveld Complex lithologies, most Great Dyke rocks have markedly

lower trace element contents, possibly reflecting lower trapped melt contents or a less enriched

magma. The Great Dyke rocks also tend to have less fractionated patterns (expressed by, e.g., lower

Ce/Sm) with less pronounced negative Nb-Ta and, in the case of the ultramafic rocks, negative Ti

anomalies. Furthermore, the mafic Great Dyke rocks lack the pronounced positive Eu anomalies

that are typical of Bushveld mafic rocks.

Whole rock compositional variation with height (Fig. 10) is consistent with previous profiles

across the transition zone of the Ultramafic-Mafic sequences (Wilson and Chaumba, 1997), in that

the P1 pyroxenite shows upward-decreasing values in MgO and Cr, whereas the Mafic Sequence

shows low and relatively constant MgO and Cr contents. The P1 layer shows little variation in Mg#,

but a reversal occurs at the base of the Mafic Sequence, as previously noted by Wilson and

Chaumba (1997). This is followed by a progressive decrease in Mg# with height through the Mafic

Sequence. Notably, the reversal is not evident in Cr/V, which shows a broadly continuous decrease

from the orthopyroxenites of the MSZ into the Mafic Sequence.

There is a pronounced drop in Ce/Sm ratio across the ultramafic-mafic transition (Fig. 10),

somewhat analogous to that seen at the contact between the Critical and Main Zones of the

Bushveld Complex (Maier and Barnes, 1998). However, in stark contrast to the Bushveld Complex,

Sr isotopes show no systematic variation across the ultramafic-mafic contact interval (Fig. 10); All

samples have Sri 0.7024-0.7028, markedly less enriched than in the Bushveld Complex (Sri 0.704-

0.709). The same applies to Nd isotopes (Great Dyke: εNd around 0 to +10,; Bushveld: εNb -5 to -

7,). The Sr and Nd isotope ratios measured by us are similar to those reported by Mukasa et al.

(1998) and Oberthür et al. (2002), but higher than those of Nebel et al. (2008; Sri 0.7021).

5.3 Chalcophile whole rock geochemistry

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The concentrations of PGE in the rocks show a marked bimodality, with most gabbronorites and

websterites having < 10 ppb Pt+Pd, whereas the orthopyroxenites have 20->1000 ppb (Fig. 11a).

However, as Cu/Zr is > 10 throughout the analysed sequence, all rocks likely contain cumulus

sulfide implying that they crystallised from S saturated magma. The low PGE contents of the

websterites are particularly notable in view of the fact that the rocks contain up to 1% sulfide, with

similar relative abundances of pentlandite to pyrrhotite and chalcopyrite as in the MSZ. As a result,

Cu/Pd and Cu/Pt are significantly above the level of primitive mantle in the websterites and

gabbronorites. Notably, Cu/Pt in the basal gabbronorite samples is lower than in the websterite (Fig.

11b), i.e. the gabbronorites are less PGE depleted. There is no significant correlation between PGE

and Cu, but the PGE correlate very strongly with each other. Being that the IPGE tend to be

immobile in fluids, this suggests that none of the PGE was mobile during late magmatic or

hydrothermal processes.

The PGE spider patterns of the rocks (Fig. 12) are bell-shaped for the orthopyroxenites, but

trough shaped for the websterites and most gabbroic rocks. The patterns of the websterites show a

peculiar positive Rh anomaly, analogous to, e.g., magnetite gabbros of the Koitelainen intrusion

(unpublished data of authors).

6. Discussion

6.1 Constraints on mantle sources and crustal contamination

Schönberg et al. (2003) proposed that Os, Pb and Nd isotope data of the Great Dyke and the gneiss

suites intruded by the Great Dyke indicate around 10-25% crustal contamination of the magma by

gneiss. The authors further argued that the superchondritic Os isotope ratios (initial 187

Os/188

Os

0.1106 - 0.1126) rule out an SCLM source for the Great Dyke magma and instead imply an

asthenospheric mantle source. Because the east dyke and Umvimeela dyke, believed to represent

the parent magma to the Great Dyke (Wilson, 1982; Stubbs et al., 1999), have depleted Nd isotopes

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(εNd +4), the Great Dyke magma was proposed to have been contaminated during ascent into the

upper crust, rather than at the base of, or below the crust. Schönberg et al. (2003) further argued that

the similarity of Os isotope signatures along strike in the Great Dyke suggests that contamination

mainly occurred in a staging chamber, rather than the Great Dyke itself. This model is consistent

with the homogeneity of the isotopic and trace element ratios seen in the rocks of the transition

interval between the Ultramafic and Mafic Sequences of the Great Dyke.

In trace element ratio diagrams the Great Dyke data plot along a trend between plume derived

picrite and partially molten country rock gneiss (Fig. 13), consistent with the model of Schönberg

et al. (2003). These results suggest an important role for a mantle plume in magma generation, but

they do not exclude the presence of an SCLM component in the Great Dyke magmas.

6.2 Magmatic evolution of the Main Sulfide Zone

The petrogenesis of the PGE bearing zones of Cyclic Unit 1 has been discussed in considerable

detail by Prendergast and Keays (1989), Prendergast and Wilson (1989), Wilson and Prendergast

(1989, 2001), Prendergast (1991), Wilson (1992, 2001), Wilson et al. (1989) and Oberthür (2011).

These authors proposed that the PGE mineralization of the C1d and C1c chromitites formed by

mixing between resident magma that was nearly saturated in sulfide melt following the

crystallization of the thick pile of underlying ultramafic cumulates, and relatively unevolved,

replenishing magma. The mixing triggered both chromite and sulfide melt saturation. This model is

consistent with the presence of a major reversal and complex layering at the contact of Subunits 1e

and 1d. However, more recent work on the controls on sulfide melt saturation in basaltic magmas

by Li and Ripley (2005) has shown that magma mixing can only trigger sulfide melt saturation in

the hybrid if both mixing partners themselves are nearly saturated in sulfide melt, which seems

unlikely in the case of the unevolved replenishing magma.

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Wilson and Prendergast (2001) further proposed that after the formation of the C1c chromitite

the magma in the chamber reverted to S undersaturation due to a major influx of unevolved and S

undersaturated magma, consistent with data from the Selukwe sub-chamber showing that the

footwall cumulates of the LSZ have as little as 150-200 ppm S (Wilson et al., 2000). The new

magma would also have replenished the PGE budget of the magma. Sulfide saturation was again

reached at the base of the LSZ, but the absence of chromite and olivine and the decreasing Mg# of

orthopyroxene suggest that this occurred mainly due to fractionation of the magma, rather than

major magma replenishment or mixing (Wilson and Prendergast, 2001). The persistence of elevated

PGE contents through > 100 m of LSZ stratigraphy may indicate additional magma replenishment

during the formation of the LSZ (Wilson, 2001), or a very large magma reservoir that was buffered

against PGE depletion through sulfide saturation. Wilson and Prendergast (2001) argue that a

further replenishment with S undersaturated magma at the top of the LSZ was responsible for

termination of sulfide melt saturation in the chamber. This may also have replenished the PGE

budget of the resident magma after the formation of the LSZ. However, Prendergast and Keays

(1989) show that, whereas the rocks overlying the LSZ tend to be relatively sulfide poor (<100 ppm

S), most contain several 10s of ppb PGE, suggesting the presence of a small proportion of cumulus

sulfide. Wilson and Prendergast (2001) go on to propose that the MSZ formed by renewed sulfide

melt saturation in response to in situ silicate fractionation. In summary, fractionation and magma

replenishment are considered key to the formation of the LSZ and the MSZ.

Our trace element and isotope data are in broad agreement with certain aspects of the above

model, namely the importance of fractionation in triggering sulfide melt saturation. There is no

strong indication for a compositionally distinct magma replenishing the chamber in the ultramafic-

mafic transition interval, including the level of the MSZ. The sharp drop in Ce/Sm from the

ultramafic to the mafic rocks is likely to be a modal effect; The ultramafic rocks have much higher

trapped liquid contents than the gabbroic rocks, as indicated by significantly higher Zr contents

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(Fig. 6). There is thus no evidence for mixing of magmas of different lineage in the MSZ. This

finding is significant because magma mixing has long been considered to be critical in the

formation of other important PGE reefs, namely the Merensky Reef of the Bushveld Complex

(Campbell et al., 1983; Naldrett et al., 2012). The present study implies that this mechanism is not

essential for the formation of PGE reefs in layered intrusions and we question its relevance in

general. Our data also provide no evidence that the PGE reefs formed from particularly

contaminated magma batches, as proposed by Arndt et al. (2005) for the Merensky Reef; Not only

are the Sr isotopes homogenous, but all our δ34

S values, and the δ18

O values determined by Li et al.

2007 (δ18

O is 5.1-6.4) are essentially in the range of the mantle.

If fractionation is the key factor in triggering sulfide melt saturation, the question arises what

caused sulfide concentration in the 2-m wide MSZ? Based on the slope of the calculated S solubility

curve for SHMB magmas (Ripley and Li, 2013), the cotectic ratio between sulfide melt and silicate

crystals is around 0.5-1%. This proportion broadly overlaps with the amount of sulfides in the LSZ,

but the MSZ contains 1-8% sulfides (Wilson and Tredoux, 1990). One possibility is that relatively

dense sulfide melt segregated preferentially from the magma. However, mass balance indicates that

the PGE in the MSZ must be derived from a magma column of several 100m, assuming 10-20 ppb

Pt and Pd in the magma, i.e., levels that are comparable to Bushveld magmas. A large volume of

silicate magma may also be required to achieve the high R factors (on the order of at least 104)

proposed by many authors to be necessary for generating the high tenors of the sulfides in the PGE

reefs (Campbell et al., 1983). Effective concentration of early forming high tenor sulfides from such

a large magma volume without dilution by low tenor sulfides and silicates is difficult to model

(Naldrett et al., 2009). In the case of the Merensky Reef, Naldrett et al.’s solution to the problem

was to propose a PGE rich magma, having on the order of 200 ppb PGE. This would allow a much

reduced magma column from which to extract the sulfides (meters to 10s of meters as opposed to

100s of meters). However, the Bushveld magmas exposed in the Marginal Suite have a maximum

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of ~15 ppb Pd and 20 ppb Pt, and we know of no magmas globally that have more than

approximately 20-25 ppb Pt and Pd each.

Maier et al. (2013) have proposed an alternative model for the formation of the PGE reefs,

whereby the sulfides (and other dense phases such as olivine, pyroxene and chromite) were

concentrated during magmatic sorting of crystal slurries sliding towards the subsiding centre of

intrusions in response to seismic events related to magma replenishment. The variable sulfide

contents of the reefs and their host rocks are attributed largely to variation in the efficiency of the

sorting process, which is related to rate of subsidence of the chamber. Apart from providing a

mechanism to form supercotectic sulfide accumulations, a key advantage of this model is that it

provides a better explanation for the location of sulfide reefs at the ultramafic base of cyclic units

than the model of magma replenishment and mixing, in view of the likely S undersaturated nature

of the replenishing magma.

The model also offers an explanation for the “offset” distribution of PGE in the MSZ and

other silicate reefs (Munni Munni, Barnes et al., 1993; Stella, Maier et al., 2003), with Pd peaking

stratigraphically below Pt and Cu: These patterns represent the preserved metal zoning of the proto

cumulates (i.e., prior to sorting), resulting from relatively higher sulfide melt-silicate liquid partition

coefficients of Pd than Pt and Cu (Barnes and Lightfoot, 2005). Sorting merely led to concentration

of the sulfides and vertical compression of the original metal distribution patterns.

Is the slumping and sorting model consistent with the large scale consistency of PGE reefs

such as the Merensky Reef and the MSZ? Firstly, we note that the data accumulating from the

increasing number of open pit operations of the reefs as well as careful observations of

underground developments indicates that the reefs are far less consistent than generally

perceived (Maier et al., 2013; Latypov et al., 2015; Mitchell et al., 2015). The Merensky Reef has

abundant potholes, sill-like apophyses, and shows significant lateral and down-dip variation in

thickness and stratigraphic disposition. The frequency of these features has been underestimated

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from drill core studies. Secondly, the envisaged sorting process is not chaotic, but gentle, triggered

by seismicity related to magma replenishments. This kinetic sieving process (the Brazil nut effect)

leads to concentration of the small and relatively dense particles towards the base of the

reservoir, but broad preservation of the original stratification.

6.3 Formation of the ultramafic-mafic contact interval

Wilson and Chaumba (1997) also considered whether the compositional reversal from the

websterites to the gabbronorites could be due to relatively more efficient equilibration of evolved

trapped liquid with the cumulus minerals in the websterites, thereby lowering their Mg#, Cr and Ni

contents. This model would be consistent with the low Zr content and thus trapped liquid

component of the gabbronorites relative to the underlying websterites (Fig. 6), and it could

potentially explain the decoupling of Mg# and Cr contents of orthopyroxene across the ultramafic-

mafic contact, as Cr equilibrates less readily with evolved trapped melt. The model was ultimately

rejected by Wilson and Chaumba (1997) who argued that trapped liquid shift can only account for a

2-3% shift in Mg#, whereas the reversal is by up to 8% Mg#. It could be argued that the websterites

could have initially contained higher trapped melt contents than what is now observed, followed by

some compaction or melt expulsion. However, the drop in Cu/Pt across the ultramafic-mafic contact

is consistent with influx of a compositionally distinct, less PGE depleted magma at the base of the

Mafic Sequence (Fig. 11). On the basis of relatively elevated An content of plagioclase and Mg# of

pyroxene (An80, Mg#88), the new magma was less evolved than the resident magma The relatively

low Cr contents of pyroxene (0.2%) may be due to fractionation of pyroxene (Wilson and

Chaumba, 1997) or chromite at depth.

The field relationships documented in the present paper suggest that the websterites intruded

the base of the Mafic Sequence (Fig. 5). At the same time, observations by Prendergast (1991) on

drill core and underground exposure at Mimosa mine indicate that the basal gabbronorite of the

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Mafic Sequence locally erodes and intrudes into websterite. The undulating and locally flame-like

nature of the contact, the absence of evolved and metal depleted rocks at the top of the websterite,

and the pegmatoidal nature of the uppermost portion of the websterite suggest that both the

websterite and gabbronorite crystallised from crystal mushes that mingled rather than mixed. Based

on mineral and whole rock composition, the websterite could be the residual melt of the underlying

orthopyroxenite cumulates, but the gabbronorite appears to represent a major magma replenishment

emplaced on top of the semi consolidated cumulate pile of websterite.

6.4 Synthesis and petrogenetic model

As in most other PGE mineralized layered intrusions, the main PGE reefs of the Great Dyke are

located near the contact from the ultramafic to the mafic sequence. Several models have been

proposed to explain this pattern: (i) replenishment and mixing of S saturated or nearly S saturated

magma of a different lineage (as has been proposed for the Bushveld by, e.g., Campbell et al.,

1983), (ii) sulfide saturation in response to enhanced assimilation-fractional crystallization (AFC) of

specific magma batches (Bushveld: Arndt et al., 2005), (iii) emplacement of a PGE rich magma

with or without entrained sulfides (Bushveld: Naldrett et al., 2009), or (iv) sulfide melt saturation of

the magma essentially in response to in situ fractionation (Great Dyke: Prendergast and Wilson,

1989; Rincon del Tigre, Prendergast et al., 1998).

The present data, namely the homogeneity of incompatible trace elements and isotopes across

the reef interval indicate that the Great Dyke PGE reefs are unlikely to have formed through models

(i) –(ii). Model (iii) is difficult to test without access to rocks that may represent the parent magmas

to the intrusion. In addition, the model would imply that the general spatial association of PGE reefs

with the mafic-ultramafic transition interval of layered intrusions is a coincidence. Model (iv)

presents a relatively simple and thus elegant solution as long as fractionation was accompanied by

intermittent magma replenishment, to explain the occurrence of several reefs in most intrusions.

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Furthermore, the replenishing magmas need to have been relatively fractionated so that the hybrid

magmas remain close to S saturation.

The location of most PGE reefs and chromitites in strongly layered intervals near the

ultramafic-mafic transition of large and super-large intrusions, and often below thick, relatively

homogenous mafic segments of the intrusions (Bushveld, Great Dyke, Penikat, Stillwater) is also

unlikely to be a coincidence and indicates an important role for chamber dynamics. We propose that

the evolution of magma chambers is controlled by a positive feedback mechanism. Filling of the

chambers caused crustal loading and subsidence, which triggered seismic pumping of magmas from

the staging chambers, which further inflated the upper chamber causing further subsidence, and so

on. Progressive subsidence led to sorting of crystal slurries at the top of the cumulate pile and thus

layering (Fig. 14). This positive feedback mechanism may ultimately lead to roof collapse of the

staging chamber, and ascent of a major pulse of crystal charged magma forming the thick, relatively

poorly layered, mafic portions of the Bushveld, Stillwater, and Great Dyke complexes (Fig. 15).

7. Conclusions

In the present paper we have provided new compositional and field data that contribute to a better

understanding of the petrogenesis of the PGE mineralization in the Great Dyke. (i) The magmas are

at least partly asthenosphere-derived, and underwent significant (up to 15-20%, Schönberg et al.,

1999) contamination during ascent through the crust. There is no evidence for interaction of the

magma with the SCLM. This implies that the formation of PGE reefs does not require a magma

component derived from the SCLM. (ii) The interval comprising relatively high-tenor PGE

mineralization in the LSZ and MSZ is up to > 100 m thick. This suggests that the chamber must

have been replenished with fertile magma after the formation of the LSZ. (iii) The resident and

replenishing magmas were of broadly similar composition, as suggested by the lack of prominent

compositional reversals and breaks. Evidence for enhanced crustal contamination of specific

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magma batches, e.g., those hosting the reefs, is lacking. This implies that the formation of PGE

reefs is not dependent on mixing of compositionally different magmas or on crustal contamination

of specific magma batches. (iv) Based on theoretical considerations, it is considered likely that the

replenishing magma was undersaturated in sulfide melt. This may prohibit sulfide melt saturation in

the chamber, even if the resident and new magmas mixed, and thus we favour a model whereby

sulfide melt saturation was triggered by in situ fractionation. (v) The super-cotectic abundance of

the sulfides in the MSZ is proposed to be the result of sorting of sulfide-bearing cumulate mushes

sliding towards the interior of the subsiding chamber in response to crustal loading . (vi) The

websterite represents the residual magma of the P1 orthopyroxenites. (vii) The Mafic Sequence

formed from a major influx of crystal-charged magma from a staging chamber, possibly triggered

by seismic pumping induced by subsidence of the Great Dyke itself. (viii) Gabbronorite and

websterite crystal mushes underwent magma mingling and local erosion. The pegmatoid at the top

of the websterite formed in response to damming up of magmatic fluids below the gabbronorite

mush.

Acknowledgements

Sincere thanks go to Andrew Du Toit and Howard Gumindoga of ZIMPLATS who provided rock

and core samples as well as additional information for this study. The thin sections were produced

by Sari Forss (Oulu University) and Tony Oldroyd (Cardiff University). Funding for this project

was provided by the Finnish Renlund Foundation. Martin Prendergast and Chusi Li provided

constructive and thorough reviews that greatly improved the paper.

References

Arndt, N., Jenner, G., Ohnenstetter, M., Deloule, E., Wilson, A.H., 2005. Trace elements in the Merensky

Reef and adjacent norites Bushveld Complex South Africa. Mineralium Deposita 40, 550–575.

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Barnes, S.-J., Maier, W.D., 1999. The fractionation of Ni, Cu and the noble metals in silicate and sulfide

liquids: Geological Association of Canada Short Course Volume 13, p. 69−106.

Barnes, S-J, Maier, W.D., Curl, E., 2010. Composition of the marginal rocks and sills of the Rustenburg

Layered Suite, Bushveld Complex, South Africa: Implications for the formation of the PGE deposits.

Economic Geology 105, 1491-1511.

Barnes, S.J., 1993. Partitioning of the platinum group elements and gold between silicate and sulphide

magmas in the Munni Munni Complex, Western Australia, Geochimica et Cosmochimica Acta 57, 1277-

1290.

Campbell, I.H., 1986. Distribution of orthocumulate textures in the Jimberlana intrusion. Journal of Geology

95, 35-54.

Campbell, I.H., Naldrett, A.J., Barnes, S.J., 1983. A model for the origin of the platinum-rich sulfide

horizons in the Bushveld and Stillwater Complexes. Journal of Petrology 24,133–165.

Capobianco, C.J., Drake, M.J. 1990. Partitioning of ruthenium, rhodium, and palladium between spinel and

silicate melt and implications for platinum-group element fractionation trends. Geochimica et Cosmochimica

Acta. 54, 869-874.

Germann, K., Schmidt, F., 1999. Platinum-group element concentrations in chromitites of the Great Dyke,

Zimbabwe: Hannover, Bundesanstalt für Geowissenschaften und Rohstoffe (BGR), Berichte zur

Lagerstätten- und Rohstoffforschung 34, 68 p.

Lambert, D.D., Walker, R.J., Morgan, J.W., Shirey, S.B., Carlson, R.W., , Zientek, M.L., Lipin, B.R., Koski,

M.S., Cooper, R.L., 1994. Re-Os and Sm-Nd Isotope Geochemistry of the Stillwater Complex, Montana:

Implications for the petrogenesis of the J-M Reef. Journal of Petrology 35. 1717-1753.

Latypov, R., Chistyakova, S., Page, A., Hornsey, R., 2015. Field Evidence for the In Situ Crystallization of

the Merensky Reef. Journal of Petrology, doi: 10.1093/petrology/egv023

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Li, C., Ripley, E.M., 2005. Empirical equations to predict the sulfide content of mafic magmas at sulfide

saturation and applications to magmatic sulfide deposits. Mineralium Deposita 40, 218-230.

Li, C., Ripley, E.M., Oberthür, T., Miller, J.D., Joslin, G.D., 2007. Textural, mineralogical and stable isotope

studiesof hydrothermal alteration in the main sulfide zone of the Great Dyke, Zimbabwe and the precious

metals zone of the Sonju Lake Intrusion, Minnesota, USA. Mineralium Deposita 43, 97–110.

Luais, B., Hawkesworth, C.J., 1994. The Generation of Continental Crust: An Integrated Study of Crust-

Forming Processes in the Archaean of Zimbabwe. Journal of Petrology 35, 43-93.

Maier, W.D., Eales, H.V., 1997. Correlation within the UG2 - Merensky Reef interval of the Western

Bushveld Complex, based on geochemical, mineralogical and petrological data. Geological Survey of South

Africa Bulletin 120, 56 pp.

Maier, W.D., Barnes, S-J., 1998. Concentrations of rare earth elements in silicate rocks of the Lower, Critical

and Main Zones of the Bushveld Complex. Chemical Geology 150, 85-103.

Maier, W.D., Barnes, S.-J., Gartz, V., Andrews, G., 2003. Pt-Pd reefs in magnetitites of the Stella intrusion,

South Africa: a world of new PGE exploration opportunities. Geology 31, 885-888.

Maier, W.D., Peltonen, P., McDonald, I., Barnes, S.J., Barnes, S-J., Hatton, C., Viljoen, F., 2012. The

concentration of platinum-group elements and gold in southern African and Karelian kimberlite-hosted

mantle xenoliths: Implications for the noble metal content of the Earth’s mantle. Chemical Geology 302–

303, 119–135.

Maier, W.D., Barnes, S-J., Groves, D.I., 2013. The Bushveld Complex, South Africa: Formation of

platinum-palladium, chrome and vanadium- rich layers via hydrodynamic sorting of a mobilized cumulate

slurry in a large, relatively slowly cooling, subsiding magma chamber. Mineralium Deposita 48, 1-56.

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Maier, W.D., Rasmussen, B., Li, C., Barnes, S-J., Huhma, H., 2013. The Kunene anorthosite complex,

Namibia, and its satellite bodies: geochemistry, geochronology and economic potential. Economic Geology

108, 953–986.

Mavrogenes, J.A., O’Neill, H.StC., 1999. The relative effects of pressure, temperature and oxygen fugacity

on the solubility of sulfide melts in mafic magmas. Geochimica et Cosmochimica Acta 63, 1173-1180.

McDonald, I., Viljoen, K.S., 2006. Platinum-group geochemistry of mantle eclogites: a reconnaissance study

of xenoliths from the Orapa kimberlite, Botswana. Transactions of the Institution of Mining and Metallurgy

115, 81-93.

Mitchell, A.A., Henckel, J., Mason, A., 2015. Factors affecting PGE mineralization processes in the Upper

Critical Zone of the Swartklip Sector, north western Bushveld Complex, on the farm Wilgespruit 2JQ. South

African Journal of Geology (submitted).

Mukasa, S. B., Wilson, A. H., Carlson, R. W., 1998. A multielement geochronologic study of the Great

Dyke, Zimbabwe: significance of the robust and reset ages. Earth and Planetary Science Letters 164, 353 –

369.

Naldrett, A.J., Wilson, A., Kinnaird, J., Chunnett, G., 2009. PGE tenor and metal ratios within and below the

Merensky Reef, Bushveld Complex: Implications for its genesis. Journal of Petrology 50, 625-659.

Nebel, O., Mezger, K., 2008. Timing of thermal stabilization of the Zimbabwe Craton deduced from high-

precision Rb–Sr chronology, Great Dyke. Precambrian Research 164, 227–232.

Oberthür, T., 2011. Platinum-group element mineralization of the Main Sulfide Zone, Great Dyke,

Zimbabwe. Reviews in Economic Geology 17, 329–349.

Oberthür, T., Davis, D. W., Blenkinsop, T. G., Höhndorf, A., 2002. Precise U-Pb mineral ages, Rb-Sr and

Sm-Nd systematics for the Great Dyke, Zimbabwe: constraints on late Archean events in the Zimbabwe

Craton and Limpopo Belt. Precambrian Research 113, 293–305.

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Podmore, F., Wilson, A.H., 1987. A reappraisal of the structure, geology and emplacement of the Great

Dyke. In HC Halls, WF Fahrig (eds) Mafic Dyke Swarms. Geological Association of Canada, Special Paper

34, 317-330.

Prendergast, M.D., 1988a. An investigation of the stratigraphy and petrology of the Pyroxenite No. 1 layer of

the Great Dyke, Zimbabwe, with special reference to the characteristic features and origin of the platinum-

group element-bearing Main Sulfide Zone. PhD thesis (unpublished), Department of Geology, University of

Zimbabwe, 412 pp.

Prendergast, M., 1991. The Wedza-Mimosa platinum deposit, Great Dyke, Zimbabwe: layering and

stratiform PGE mineralization in a narrow mafic magma chamber. Geological Magazine 128, 235-249.

Prendergast, M.D., 1998. The Great Dyke of Zimbabwe: International Platinum Symposium, 8th, Excursion

Guidebook to the Great Dyke, Zimbabwe, Johannesburg, South Africa, 47pp.

Prendergast, M.D., Wilson, A.H., 1989. The Great Dyke of Zimbabwe - II: Mineralisation and mineral

deposits. In Magmatic sulphides - the Zimbabwe volume. Edited by M.D. Prendergast and M.J. Jones. The

Institution of Mining and Metallurgy, London, p. 21-42.

Prendergast, M., Bennett, M., Henicke, G., 1998. Platinum exploration in the Rincon del Tigre Complex,

eastern Bolivia. Transactions of the Institute of Mining and Metallurgy 107, B39–47.

Ripley, E.M., Li, C., 2013. Sulfide Saturation in Mafic Magmas: Is External Sulfur Required for Magmatic

Ni-Cu-(PGE) Ore Genesis? Economic Geology 108, 45–58.

Savard, D., Bedard, L.P., Barnes, S-J., 2006. TCF selenium preconcentration in geological materials for

determination at sub-μg g-1 with INAA (Se/TCF-INAA). Talanta 70, 566-571.

Savard, D., Barnes, S-J., Meisel, T., 2011. Comparison between Nickel-Sulfur Fire Assay Te Co-

precipitation and Isotope Dilution with High-Pressure Asher Acid Digestion for the Determination of

Platinum-Group Elements, Rhenium and Gold. Geostandards and Geoanalytical Research 34, 281-291.

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Schönberg, R., Kruger, F.J., Nagler, T.F., Meisel, T., Kramers, J.D., 1999. PGE enrichment in chromitite

layers and the Merensky Reef of the western Bushveld Complex; a Re-Os and Rb-Sr isotope study. Earth

and Planetary Science Letters 172, 49-64.

Schönberg, R., Nägler, Th. F., Gnos, E., Kramers, J. D., Kamber, B. S., 2003. The Source of the Great

Dyke, Zimbabwe, and Its Tectonic Significance: Evidence from Re-Os Isotopes. Journal of Geology 111,

565-578.

Stubbs, H.M., Hall, R.P., Hughes, D.J., Nesbitt, R.W., 1999. Evidence for a high Mg andesitic parental

magma to East and West satellite dykes of the Great Dyke, Zimbabwe: a comparison with the continental

tholeiitic Mashonaland sills. Journal of African Earth Sciences 28, 325-336.

Sun, S.-s., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for

mantle composition and processes. In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins,

vol. 42. Geological Society Special Publication, pp. 313–345.

Wilson, A.H., 1982. The geology of the Great Dyke, Zimbabwe: The ultramafic rocks. Journal of Petrology

23, 240−292.

Wilson, A. H., 1992. The Geology of the Great Dyke, Zimbabwe: Crystallization, Layering, and Cumulate

Formation in the P1 Pyroxenite of Cyclic Unit 1 of the Darwendale Subchamber. Journal of Petrology 33,

611-633.

Wilson, A. H., 1996. The Great Dyke of Zimbabwe. In: Cawthorn, R. G. (ed.), Layered Intrusions. Elsevier

Science B.V., Amsterdam, 365-402.

Wilson, A.H., 2001. Compositional and lithological controls on the PGE-bearing sulfide zones in the

Selukwe subchamber, Great Dyke: A combined equilibrium- Rayleigh fractionation model. Journal of

Petrology 42, 1845−1867.

ACC

EPTE

D M

ANU

SCR

IPT

ACCEPTED MANUSCRIPT

Wilson, A. H., Chaumba, J. B., 1997. Closed system fractionation in a large magma chamber: mineral

compositions of the websterite layer and lower mafic succession of the Great Dyke, Zimbabwe.

Mineralogical Magazine 61, 153-173.

Wilson, A.H., Prendergast, M.D., 1989. The Great Dyke of Zimbabwe - I: Tectonic setting, stratigraphy,

petrology, structure, emplacement and crystallisation. In Magmatic Sulphides - The Zimbabwe Volume.

Edited by M.D. Prendergast and M.J. Jones. The Institution of Mining and Metallurgy, London, p. 1-20.

Wilson, A.H., Prendergast, M.D., 2001. Platinum-group element mineralization in the Great Dyke,

Zimbabwe, and its relationship to magma evolution and magma chamber structure. South African Journal of

Geology 104, 319-342.

Wilson, A.H., Tredoux, M., 1990. Lateral and vertical distribution of platinum-group eleme nts and

petrogenetic controls on the sulfide mineralization in the P1 Pyroxenite Layer of the Darwendale

Subchamber of the Great Dyke, Zimbabwe. Economic Geology 85, 556-584.

Wilson, A.H., Wilson, J.F., 1981. The Great Dyke, in DR Hunter (ed), Precambrian of the Southern

Hemisphere, Amsterdam, Elsevier, 572-8.

Wilson, A.H., Brown, R.T., 2005. Exploration and mining of the Main Sulfide Zone of the Great Dyke,

Zimbabwe: case study of the Hartley platinum mine. in JE Mungal (Ed), Exploration for Platinum-group

elements, Mineralogical Association of Canada, Short Course Series 35, p. 409-430.

Wilson, A.H., Murahwi, C.Z., Coghill, B.M., 2000. Stratigraphy, geochemistry and platinum-group

mineralisation of the central zone of the Selukwe Subchamber of the Great Dyke, Zimbabwe.

Journal of African Earth Sciences 30, 833-853.

Wingate, M. T. D., 2000. Ion microprobe U-Pb zircon and baddeleyite ages for the Great Dyke and its

satellite dykes, Zimbabwe. South African Journal of Geology 103, 74-80.

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Yang, S., Maier, W.D., Lahaye, Y., O'Brien, H., 2013. Strontium isotope disequilibrium of plagioclase in the

Upper Critical Zone of the Bushveld Complex: evidence for mixing of crystal slurries. Contributions to

Mineralogy and Petrology DOI 10.1007/s00410-013-0903-4

Zhang, M., O’Reilly, S.Y., Wang, K-L., Hronsky, J., Griffin, W.L., 2008. Flood basalts and metallogeny:

The lithospheric mantle connection. Earth Science Reviews 86, 145–174.

Figure Captions

Fig. 1. General map showing the location of the Great Dyke in Zimbabwe, and the study locality

(from Oberthür, 2011).

Fig. 2. Transverse section of the Darwendale Subchamber of the Great Dyke. Note that the dip of

the layers decreases, and the layers get gradually thinner towards the margins. (modified after

Wilson and Prendergast, 1989).

Fig. 3: Stratigraphic sub-division of the Great Dyke (modified after Wilson and Prendergast, 1989)

Fig. 4: Lithologies of the ultramafic-mafic contact interval. a) Contact interval between strongly

altered pegmatoidal websterite at top of Ultramafic Sequence and gabbronorite of the Mafic

Sequence. b+c) Occurrence of olivine oikocrysts within basal portion of Mafic Sequence. d+e)

Detailed view of sharp contact between Ultramafic and Mafic Sequences. gb=gabbronorite,

web=websterite. e) View of a drill core intersection in which websterite is absent and instead Mafic

Sequence directly overlies olivine-bearing bronzitite. f) Olivine enrichment in bronzitite (bro)

where websterite is absent. g) Bronzitite just below contact to websterite, showing patches of

quartz.

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Fig. 5: Contact between websterite and gabbronorite at Portal 2, east open pit, Ngezi mine, showing

apophyses of websterite in gabbronorite.

Fig. 6. (a) Mineral modes in the analysed sequence. Whole rock concentrations of Zr are a proxy for

the trapped liquid component. (b) Mineral compositional variation. See text for discussion.

Fig. 7: Sulfide petrography. (a) Interstitial magmatic sulfides of the MSZ (GD263.05), (b-d)

Partially replaced sulfides of the MSZ (GD261.1), (e-f) Sulfides within websterite at top of P1 layer

(GD 254). Note irregulat shape of magmatic sulfide in e, and fine sulfides in cleavage planes within

pyroxene, but absence in intercumulus plagioclase in (f).

Fig. 8: Binary variation diagrams vs MgO of (A) CaO, (B) Cr, (C) Zr, (D) Nb (see text for

discussion).

Fig. 9: Primitive mantle normalized lithophile variation patterns. Normalisation factors from Sun

and McDonough (1989)

Fig. 10 Compositional variation of lithophile elements and isotopes across the analysed sequence.

See text for discussion.

Fig. 11 (a): Binary variation diagrams of chalcophile elements. (b) Concentration of Pt+Pd, Cu,

Cu/Pt and Cu/Zr with height. Note sharp increase in Cu/Pt above the MSZ, persisting throughout

the gabbroic rocks, indicating a relatively PGE depleted magma. PM=primitive mantle (data from

Barnes and Maier, 1999).

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Fig. 12. Primitive mantle-normalised chalcophile metal patterns of analysed samples (normalization

factors from Barnes and Maier, 1999)

Fig. 13: Plot of (A) La/Sm vs La/Nb and (B) La/Sm vs Th/Nb for Great Dyke rocks. Solid line

represents mixing line between picrite (with trace element contents assumed to be 4x primitive

mantle, i.e. equivalent to ~25% partial mantle melting) and a contaminated magma produced by

AFC (r=0.8, f=0.8) of picrite with Shabani gneiss (data from Luais and Hawkesworth, 1994). Great

Dyke data can be modeled by up to 20% AFC with granite. There is little evidence for involvement

of SCLM (data for Kaapvaal peridotite and MARID xenoliths are from Maier et al., 2012).

Fig. 14: Schematic model illustrating envisaged origin of transition interval between the Ultramafic

and Mafic Sequences. (a-c) Emplacement of magmas leads to progressive crustal loading and

subsidence of central portion of intrusion. Incompletely consolidated cumulates at the top of the

crystal pile undergo some slumping towards the centre, resulting in crystal sorting and formation of

layers that progressively thicken towards the centre of the intrusion. Stippled line indicates level at

which the residual magma in chamber reached sulfide melt saturation through fractionation. Sorting

of crystal mushes that crystallized from such magma may lead to formation of PGE reefs. (d)

Intrusion of large, possibly crystal laden magma influx of the Mafic Sequence, leading to

accumulation of gabbronorite. (e) Gabbronorite and websterite mushes mingle, with websterite

locally injecting into roof and floor rocks, and gabbronorite locally eroding websterite (and MSZ).

Pegmatoidal texture at top of websterite is due to building up of volatiles below gabbronorite mush.

Fig. 15: Schematic model illustrating sequence of magma emplacement, triggered by progressive

crustal loading and deformation of a staging chamber. (A-B): Filling of magma chamber by influxes

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of unevolved Mg-basaltic magma. (C) Progressive subsidence leads to the formation of the

layering, including the main sulfide reef. Crustal subsidence leads to roof collapse of staging

chamber and (D) ascent of a large pulse of relatively evolved magma crystallizing the Mafic

Sequence.

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Figure 3

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Figure 4

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Figure 5

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Figure 6

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Figure 7

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Figure 8

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Figure 9

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Figure 10

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Figure 11

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Figure 12

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Figure 13

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Figure 14

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Figure 15

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Table 1. Whole rock compositions of analysed Great Dyke samples

Sample

50214

0 852.1 46.95 98.05 146.0

2 172.7

5 223.0

5 229.6

A 229.6

B 242 254 261-1 263.0

5 272.9

5 GD19 GD27

drill

core

DP13 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 MLF

45 NH00

9 HN00

9

Height*

770 202.2

8 179.3

3 128.2

3 80.26 53.53 3.23 -3.23 -3.33 -

15.72 -

27.72 -

34.82 -

36.77 -

46.67 -

71.43 -81

Rock

type

gn gb gn gb gb gn gn web web web web bron bron bron bron bron

SiO2 wt

% 49.97 49.71 48.94 49.98 49.31 49.02 49.72 50.78 51.99 51.66 54.13 52.69 52.95 54.53 55.12 53.86

TiO2

1.06 0.16 0.10 0.09 0.09 0.10 0.09 0.22 0.18 0.14 0.20 0.10 0.14 0.11 0.12 0.13

Al2O3

15.46 15.45 19.82 17.84 17.40 19.39 17.63 3.41 4.29 2.24 3.36 2.15 2.54 2.09 2.55 2.24

Fe2O3

14.68 6.68 4.98 4.99 4.29 3.98 4.37 8.62 7.07 8.63 11.44 12.22 11.90 11.41 11.97 11.69

MnO

0.16 0.14 0.11 0.11 0.11 0.10 0.11 0.18 0.18 0.22 0.25 0.25 0.25 0.25 0.21 0.21

MgO

5.43 9.87 9.07 10.17 10.78 10.42 12.25 17.65 18.28 22.48 25.99 27.82 27.12 27.91 26.43 27.02

CaO

8.67 15.27 13.55 13.76 15.17 15.68 15.32 17.16 17.59 13.40 4.87 2.84 2.83 2.56 2.95 2.85

Na2O

3.07 1.76 1.95 1.63 1.39 1.34 1.08 0.51 0.51 0.24 0.42 0.14 0.24 0.14 0.27 0.19

K2O

0.32 0.08 0.12 0.10 0.07 0.06 0.07 0.24 0.10 0.05 0.15 0.07 0.08 0.04 0.07 0.05

P2O5

0.03 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.02 0.00 0.01 0.00 0.00 0.00 0.02 0.01

LOI

0.51 0.45 0.43 0.27 0.47 0.34 0.60 0.48 0.31 0.14 0.27 1.32 1.11 0.37 -0.22 0.71

Total

99.36 99.56 99.07 98.93 99.08 100.4

2 101.2

5 99.25 100.5

3 99.21 101.0

9 99.60 99.17 99.41 99.50 98.97

Sc pp

m 23 43 22 24 25 21 24 57 55 49 31 30 30 28 24 23

V

290 162 84 102 97 82 94 215 203 203 131 112 125 116 143 137

Cr

77 605 133 266 362 424 887 3084 2341 3787 3839 3027 3114 3270 2681 2703

Co

74 52 39 43 38 34 44 105 57 86 101 119 105 103 99 100

Ni

15 309 149 152 191 186 283 2800 536 853 779 1232 592 570 1099 955

Cu

27 145 50 30 123 26 77 928 199 257 182 534 178 64 409 230

Zn

96 61 40 18 14 42 32 27 36 56 111 81 81 76 54 92

Ga

20.5 10.3 11.1 11.1 9.6 10.3 9.3 6.3 5.4 4.9 5.9 5.0 5.2 5.3 2.6 2.5

Rb

5.76 0.31 1.12 1.36 1.03 0.90 1.59 8.88 2.27 1.24 4.68 1.85 2.82 1.46 3.49 3.02

Sr

305 190 230 205 190 196 175 28 37 15 32 13 24 14 17 14

Y

4.05 3.78 1.85 1.68 1.27 1.08 1.73 5.46 4.11 3.57 3.56 2.05 2.59 2.03 2.31 2.37

Zr

12.33 7.13 11.00 10.83 6.13 7.38 7.28 17.94 52.16 11.81 19.76 15.74 16.29 7.17 10.09 7.53

Nb

0.67 0.37 0.39 0.32 0.29 1.12 0.25 1.37 7.46 0.84 0.91 0.52 0.90 0.40 1.37 1.18

Mo

0.60 0.20 0.06 0.11 0.11 0.08 0.04 0.28 0.25 0.32 0.21 0.15 0.07 0.09 na na

Cs

0.37 0.07 0.08 0.11 0.10 0.19 0.26 0.58 0.22 0.12 0.30 0.21 0.26 0.15 0.48 0.46

Ba

108 25 28 26 20 18 20 46 28 12 34 10 20 11 25 25

La

2.19 0.44 0.74 0.60 0.56 0.52 0.52 2.49 1.39 0.73 1.45 0.35 1.00 0.52 1.05 1.17

Ce

4.21 1.18 1.52 1.27 1.19 1.22 1.24 5.11 3.01 1.84 3.24 0.78 1.86 1.11 2.33 2.42

Pr

0.53 0.19 0.22 0.17 0.18 0.18 0.19 0.69 0.43 0.26 0.45 0.09 0.24 0.12 0.29 0.32

Nd

2.15 1.15 0.93 0.94 1.01 0.97 0.95 2.91 2.02 1.24 2.02 0.43 1.02 0.55 1.17 1.30

Sm

0.48 0.40 0.29 0.29 0.27 0.25 0.29 0.65 0.50 0.34 0.45 0.14 0.25 0.16 0.30 0.30

Eu

0.53 0.20 0.17 0.14 0.11 0.11 0.15 0.20 0.17 0.11 0.15 0.05 0.07 0.06 0.10 0.09

Gd

0.50 0.59 0.27 0.32 0.31 0.25 0.34 0.78 0.62 0.48 0.50 0.18 0.26 0.20 0.34 0.30

Tb

0.10 0.11 0.05 0.06 0.05 0.04 0.05 0.16 0.12 0.08 0.08 0.03 0.05 0.04 0.06 0.06

Dy

0.59 0.75 0.36 0.36 0.39 0.31 0.39 0.96 0.74 0.62 0.52 0.24 0.32 0.27 0.40 0.38

Ho

0.10 0.15 0.07 0.07 0.07 0.06 0.07 0.18 0.14 0.13 0.13 0.06 0.07 0.06 0.08 0.08

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Er

0.34 0.45 0.21 0.22 0.18 0.16 0.23 0.56 0.45 0.39 0.36 0.17 0.25 0.17 0.25 0.24

Tm

0.05 0.06 0.03 0.03 0.03 0.03 0.03 0.08 0.06 0.05 0.05 0.03 0.04 0.03 0.04 0.04

Yb

0.35 0.44 0.24 0.23 0.22 0.15 0.22 0.49 0.43 0.38 0.35 0.22 0.28 0.20 0.27 0.28

Lu

0.04 0.06 0.03 0.03 0.03 0.03 0.02 0.07 0.07 0.06 0.05 0.03 0.04 0.04 0.04 0.04

Hf

0.30 0.16 0.19 0.26 0.13 0.16 0.13 0.42 1.23 0.26 0.43 0.32 0.40 0.19 0.20 0.17

Ta

0.06 0.01 0.05 0.02 0.02 0.05 0.03 0.09 0.21 0.06 0.07 0.01 0.04 0.06 0.06 0.06

Pb

2.24 2.89 1.90 1.52 2.61 1.62 2.09 4.99 2.26 3.63 4.22 2.00 4.74 2.22 1.99 2.92

Th

0.47 0.12 0.24 0.21 0.15 0.16 0.15 1.27 0.59 0.25 0.55 0.21 0.40 0.27 0.31 0.28

U

0.10 0.02 0.05 0.04 0.03 0.03 0.02 0.37 0.12 0.06 0.19 0.06 0.09 0.06 0.07 0.08

S

285 2329 535 411 607 351 573 9830 1760 2068 1638 4055 853 477 1599 1294

C

na 1882 655 547 421 428 332 870 656 432 790 1813 2344 815 na na

Se pp

b na 670 91 197 117 107 148 2592 431 572 584 1544 257 175 na na

Os

<0,16 0.11 <0,07 <0,07 <0,07 <0,07 <0,07 0.23 <0,07 <0,07 <0,07 20 9.50 3.15 0.25 2.72

Ir

0.01 0.23 0.04 0.05 <0,03 <0,03 <0,03 0.23 <0,03 0.06 0.03 84 43 7.11 0.91 8.5

Ru

<0,03 0.79 0.12 <0,12 0.13 <0,12 0.14 1.19 0.25 0.37 0.44 171 111 33 1.88 22

Rh

0.01 1.06 0.31 0.45 0.23 0.31 0.29 2.08 1.34 0.71 1.92 195 108 19 0.77 13

Pt

<0,83

6 18.4 0.65 1.96 0.64 0.88 5.57 4.74 0.75 2.26 0.99 2571 639 42 23.3 212

Pd

0.38 20.3 <0,47 0.72 <0,47 <0,47 <0,47 4.68 0.49 1.46 0.83 1546 998 41 2.28 17

Au

<1,61

4 2.00 0.72 <0,48 <0,48 1.66 0.56 9.97 1.91 2.75 9.34 172 5.49 0.61 35 85

Isotopic

data

Sri

0.702

92 0.702

56 0.702

49 0.702

42 0.702

53 0.702

56 0.702

59 0.702

61 0.702

61 0.702

41 0.702

52 0.702

51 0.702

56 0.702

73 0.702

77 0.702

77

εNd

na na -0.3 na na -1.5 4 -0.2 na na na na na 1.2 na na

δ34S

na 0.2 0 -0.3 na -0.1 na -0.2 na 0.2 0.2 0 0.3 0.2 na na

Modal

proporti

ons

Opx

20 5 17 4 5 12 15 20 na 44 79 86 87 92 92 95

Cpx

5 38 23 36 35 31 28 67 na 52 12 5 5 2 5 2

Pla

68 54 59 59 60 56 57 9 na 2 4 4 7 5 3 3

Opaques

5 3 1 1 0.4 0.5 0.2 4 na 2 2 5 2 1 1 2

Mica

2 trace trace trace trace trace trace trace na trace 3 trace trace trace <1 <1

Notes: Height* = height relative to base of Mafic Sequence Notes: gn=gabbronorite; gb=gabbro; web=websterite; bron=bronzitite

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Highlights for review:

(i) The Great Dyke magmas are asthenosphere-derived, with up to 15-20% contamination

during ascent through the crust. This implies that the formation of PGE reefs does not

require the involvement of the SCLM.

(ii) The resident and replenishing magmas were of broadly similar composition, implying

that the formation of PGE reefs is not dependent on mixing of compositionally different

magmas or AFC.

(iii) Sulfide melt saturation was triggered by in situ silicate melt fractionation.

(iv) The super-cotectic abundance of the sulfides in the reefs is due to sorting of semi

consolidated sulfide-bearing cumulates sliding towards the interior of the chamber.