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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
This is a PDF file of an unedited manuscript that has been accepted for publication.As a service to our customers we are providing this early version of the manuscript.The manuscript will undergo copyediting, typesetting, and review of the resulting proofbefore it is published in its final form. Please note that during the production processerrors may be discovered which could affect the content, and all legal disclaimers thatapply to the journal pertain.
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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: [email protected]
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.
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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 8
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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.