Introduction

The Pyroxenite Marker is a thin, orthopyroxene-dominated marker horizon occurring towards the top of the Main Zone of the Bushveld Complex. It crops out in the Eastern Limb of the Bushveld Complex, where it was first described by Von Gruenewaldt (1973) and Molyneux (1974). The Pyroxenite Marker is sandwiched between a succession of rocks exhibiting a sustained reversal in mineral compositions (i.e. increasing plagioclase An% and mafic silicate Mg# upwards), which starts > 100 m below the layer, and an overlying sequence of rocks exhibiting a normal differentiation trend (i.e. decreasing plagioclase An% and mafic silicate Mg# upwards). The layer also marks the level within the Main Zone where primary orthopyroxene replaces inverted pigeonite as low-Ca pyroxene. The Pyroxenite Marker appears to be present in all three of the major exposed limbs of the Bushveld Complex, having been described in the Bierkraal (BK2) drill core in the Western Limb of the complex (Cawthorn et al. 1991) and in the Bellevue (BV1) drill core in the Northern Limb of the complex (Cawthorn 2020).

The Pyroxenite Marker interval is commonly interpreted as having formed in response to the addition of a large volume of a less-differentiated melt and subsequent mixing with the resident melt, a view which is strongly supported by Sr-isotopic investigations (Kruger et al. 1987; Cawthorn et al. 1991; Kruger 1994). Sharpe (1985) suggested an alternative model, proposing that a dense magma, added at the level of the Merensky Reef, underplated the resident magma without undergoing significant mixing. The underplated magma then partially crystallised to form the Main Zone below the Pyroxenite Marker. Mixing occurring at the top of the underplated magma and the upwards displaced magma resulted in the development of the Pyroxenite Marker interval. This model, however, received criticism from several authors on the basis of thermal, density and compositional considerations (Kruger et al. 1987; Cawthorn et al. 1991).

Gabbronoritic rocks below the Pyroxenite Marker typically exhibit initial 87Sr/86Sr ratios of > 0.708, with the sequence of rocks above the Pyroxenite Marker displaying a near-constant initial 87Sr/86Sr ratio of 0.7073 all the way to the top of the layered sequence (Kruger 1994). The melt that was added to the chamber at the level of the Pyroxenite Marker is, therefore, regarded as the last major influx of melt into the Bushveld Complex, a view that not only finds support when considering the isotopic evidence (Kruger et al. 1987; Kruger 1994), but also mineral compositions that vary regularly towards lower temperature compositions above the level of the Pyroxenite Marker (e.g. Mangwegape et al. 2016).

The composition of the melt added at the level of the Pyroxenite Marker has been rather well-constrained, particularly its trace element and Sr-isotopic compositions. Cawthorn et al. (1991) suggested that the added melt had an initial 87Sr/86Sr ratio of 0.7063 and showed, on the assumption, that the resident and incoming melts had similar Sr contents, that the volume of melt added should have been approximately equal to the volume of the resident melt. Vantongeren and Mathez (2013) calculated an initial 87Sr/86Sr ratio of 0.7066–0.7068 for the incoming melt, using equilibrium Sr contents calculated for the melts from the trace element contents of the minerals present across the Pyroxenite Marker interval. Their calculations suggest that the added melt volume was ~ 33–44% of the volume of the resident melt, depending on whether melt was lost from the chamber or not.

In this contribution, we attempt to constrain the Nd-isotopic composition and nature of the magma that was added at the level of the Pyroxenite Marker through analysis of the whole-rock isotopic compositions of rocks across the Pyroxenite Marker interval as sampled from borehole BH7771 on the Eastern Limb of the Bushveld Complex.

Driekop drill core: geology and sampling

Borehole BH7771 is a 2696.38 m deep hole drilled at Driekop (Fig. 1), on the Central Sector of the Eastern Limb, by Impala Platinum in 2008. It is collared in the Main Zone and extends down to the footwall of the Merensky Reef in the Upper Critical Zone. A total of 45 samples were collected as part of this study, covering the depth interval between 474.09 m and 665.10 m. The Pyroxenite Marker is represented by a 1.91 m-thick orthopyroxenite layer with a sharp lower contact and a gradational upper contact, occurring at a depth of 488.26 m. The base of the Pyroxenite Marker was used as the reference datum for the profiles presented in this study.

Fig. 1
figure 1

Geological map showing the extent of the Rustenburg Layered Suite (RLS) of the Bushveld Complex and the location of borehole BH7771. The inset shows the location of the complex in South Africa. (Map modified after Scoon and Mitchell 2009, Cawthorn et al. 2002 and Roelofse and Ashwal 2012). “f” indicating major faults

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Methods

Petrography

Polished thin sections were prepared and petrographically examined using transmitted light microscopy. The Olympus BX53M/DP74 microscope at the University of the Witwatersrand was used to take high-resolution scans of each thin section. These scans were then used for point counting by placing a digitised point grid on the scan and assessing 300 points. Minerals were assigned concentrations of less than 0.5% by volume if they were observed under the microscope, but not encountered during point counting.

Whole-rock trace element geochemistry

Trace element analyses were performed on 45 samples in the Earth Lab of the University of the Witwatersrand on a Thermo Scientific iCAP RQ Inductively Coupled Plasma Mass Spectrometer (ICP-MS). A portion of each of the 45 samples collected was crushed and milled using a jaw-crusher and swing mill, respectively. 50 mg of sample was weighed and digested in either a microwave digester (MARS from CEM) or using the open beaker/hotplate method depending on the sample type and digestibility. For the microwave digester method, the sample powder was added to a Teflon vessel along with 6 ml of ultra-high purity 2:1 HF:HNO3. Digestion vessels were then microwaved for 40 min at 180 °C and 400 PSI. The mixtures were transferred to Savillex beakers, capped and placed on a hotplate for 24 h at 70 °C. The acid was thereafter allowed to evaporate. 2 ml of HNO3 was then added to the Savillex beakers, which were capped and placed on a hotplate for 24 h at 70 °C, before being allowed to dry down. This was repeated once more, before the samples were removed from the hotplate and 300 µl HNO3 was added. Samples were stored in this state until they were ready for analysis. For the open beaker/hotplate method, 50 mg sample was placed into a Savillex beaker with 3 ml of 2:1 HF:HNO3 and placed on a hotplate at 70 °C and evaporated off. Once dried, 3 ml of HF:HNO3 was added and capped and placed on the hotplate at 70 °C for 72 h. The sample was then dried down at 70 °C and 3 ml HNO3 added and evaporated off. This was repeated three times.The prepared samples were diluted to 50 ml (dilution factor 1:1000) with 5% HNO3 with 100 ppb Re and Rh, as well as 50 ppb In and Bi as internal standards.

The Thermo Scientific iCAP RQ was optimised for maximum counts on In and both oxide and doubly charged ion levels set to < 2%. All measurements were done in triplicate and averaged. Certified Reference Materials (BCR-1 and BHVO-2) were digested and analysed along with all unknowns for quality control. The Certified Reference Materials had to return less than 10% deviation from known concentrations for all elements to pass the quality controls. Deviations were usually within 5% (see Electronic Appendix A). A Total Procedural Blank (TPB) was analysed along with all unknowns.

Whole-rock major element geochemistry

A portion of each of the 45 samples collected was crushed and milled using a jaw-crusher and swing mill, respectively. An amount of 10 g of sample powder was dried overnight at 110 °C and then subjected to 1000 °C for 4 h, to gravimetrically determine the loss on ignition (LOI). An amount of 0.28 g of powder was mixed with 1.5 g Alfa Aesar spectroflux (La2O3 (47%), Li2B4O7 (37%), Li2CO3 (16%)) and 0.02 g NaNO3. This mixture was then fluxed at 1000 °C (± 5 min) in a Pt crucible. The melt was moulded and pressed to produce fusion disks. The prepared fusion disks were analysed by wavelength dispersive X-ray Fluorescence (XRF) spectrometry using a Rigaku Primus IV spectrometer equipped with a 4 kW Rh tube. Certified Reference Materials were used to monitor the accuracy and precision of the analyses (see Electronic Appendix A). NIM-N (SARM 4) and NIM-P (SARM 5) were prepared following the same sample preparation method as mentioned above and analysed during subsequent analytical batches. Repeatability for the major elements obtained on replicates and expressed as percentage relative standard deviations (%RSDs) were all less than ~ 3% for SARM 5 (except for K2O, 10.2%) and less than 5% for SARM 4 (except for K2O, 6.9% and P2O5, 17.4%). Relative uncertainties for all oxides compared to the accepted values for the SARM 5 standard were all below 3%, except for Na2O (+ 9.5%), K2O (− 11.1%), and P2O5 (− 7.5%). For SARM 4, relative errors were also all below 3%, except for TiO2 (− 5.5%), Na2O (− 10.6%) and P2O5 (− 17.3%).

Mineral chemistry

Twenty-one samples were used for quantitative mineral chemical analyses of feldspar and pyroxene using four wavelength dispersive spectrometers on a JEOL JXA-8230 electron probe micro-analyser at Rhodes University. Analyses were performed at an accelerating potential of 15 kV and a beam current of 20 nA. A 10 µm beam was used on feldspar to avoid Na migration, with a 1 µm beam used for the analysis of pyroxenes. All elements were measured on K-alpha peaks except for Sr and Ba, which were measured on L-alpha peaks. Counting times on peak were 10 s, and 10 s total on background for all elements. Commercial “SPI” standards were used for intensity calibration. The standards for feldspar were orthoclase (Al, Si, K), albite (Na), rhodonite (Mn), MgF2 (F), plagioclase (Ca), barite (Ba), SrTiO3 (Sr) and kaersutite (Mg, Ti). Further standards for pyroxene included pyrope (Mg, Si), Ni-metal (Ni), Cr2O3 (Cr), and kaersutite (Al). Unknown acquisitions were peaked on the calibration standards. The data were collected with JEOL software (PC EPMA 1.9.2.0), and its ZAF matrix correction algorithm (Heinrich/Duncumb-Reed with FFAST-2005 MACs) was applied to correct for differential matrix effects. Oxygen was calculated by stoichiometry. Potential analytical drift was monitored for by interspersing secondary standards during the analyses.

Isotopic determinations

A total of 21 samples were chosen for whole-rock Rb–Sr and Sm–Nd isotopic determinations. Approximately 100–350 mg of sample was weighed and different volumes of 85Rb–84Sr and 149Sm–150Nd spikes were added and then dissolved in a mixture of HF–HNO3. Sr and Rb separation took place in 2 ml BioRad AG50 × 8, 200–400 mesh resin columns. Sm and Nd were separated in 1 ml Eichrom Ln Spec 100–150 µm resin columns. The isotopic compositions of samples were measured in low resolution mode on a Nu Instruments Plasma II MC-ICP-MS, housed at the Spectrum Analytical Facility (University of Johannesburg). Backgrounds were measured and subtracted from the measurements. Rb analyses were collected for 20 cycles and Sr for 60 cycles, giving an internal precision of ˂ 0.001% (1 standard error) on fractionation-corrected 87Sr/86Sr. Sm measurements were collected for 20 cycles and Nd for 60 cycles, giving an internal precision of ˂ ± 0.002% (1 standard error). Total procedural blanks and two Certified Reference Materials (BCR-2 and BHVO-2) were included in the analyses to ensure quality and reproducibility (see Electronic Appendix B). Initial 87Sr/86Sr and 143Nd/144Nd ratios were calculated using decay constants of 1.42 × 10–11 y−1 (Steiger and Jäger 1977) and 6.54 × 10–12 y−1 (Begemann et al. 2001), respectively, for an age of 2.055 Ga (Zeh et al. 2015). ƐNdi values were calculated using values for CHUR of 143Nd/144Nd = 0.512638 (Goldstein et al. 1984) and 147Sm/144Nd = 0.1967 (Jacobsen and Wasserburg 1980).

Results

Petrography

A summary of the mineral modes, grain sizes and additional petrographic observations is provided in Electronic Appendix C. The variation in the modal mineralogical compositions of samples with depth is graphically depicted in Fig. 2. The petrography of samples from the Main Zone within the Eastern Limb has been extensively documented (e.g. von Gruenewaldt 1970, 1971 and 1973; von Gruenewaldt and Weber-Diefenbach 1977; Molyneux 1974; Marais 1977; Klemm et al. 1985). Microscopic observations reveal that the studied samples consist of 26 gabbronorites, 5 pyroxenites, 5 leucogabbronorites, 3 melagabbronorites, 3 anorthosites and 1 leucogabbro (Fig. 2). Two samples of iron-rich ultramafic pegmatoid (IRUP) were also encountered at depths of − 40.74 m and − 43.69 m, respectively. Lithologies in the Driekop drill core (BH7771) are dominantly composed of plagioclase, high-Ca and low-Ca pyroxene (orthopyroxene and /or inverted pigeonite). Accessory phases include mica (biotite and phlogopite), quartz and amphibole throughout the studied interval. Phlogopite and amphibole are especially abundant within the pyroxenites. Other minerals encountered include talc (as an alteration product of pyroxene) and chlorite (as an alteration product of plagioclase and pyroxene).

Fig. 2
figure 2

Simplified lithological log on the left, indicating depths in metres for the various lithologies and the positions of samples. Variation in modal proportions and mineral compositions (plagioclase An% and orthopyroxene Mg#) with depth are shown to the right. Error bars indicate 1σ standard deviations. Arrows show reversed differentiation trend beneath the Pyroxenite Marker. Variations in the Mg# for clinopyroxene are not shown, but parallels the orthopyroxene trend. Data for the Western Limb from Cawthorn et al. (1991) and for the Eastern Limb from VanTongeren and Mathez (2013). *Numbers refer to depth in metres in borehole relative to the Pyroxenite Marker

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Gabbroic and anorthositic rocks

The majority of the rocks in the drill core are gabbroic to anorthositic as per mineral mode. These rocks generally have a spotted appearance resulting from the presence of subhedral to anhedral crystals of pyroxene in an intergranular texture (Fig. 3a) or sub-ophitically enclosing plagioclase. Plagioclase is the main phase in these rocks with a modal abundance of 16–91%. It occurs as abundantly twinned laths (Fig. 3b) frequently displaying preferential orientation. Wedge-shaped and bent plagioclase twin-lamellae (Fig. 3b), indicative of post-crystallisation deformation (von Gruenewaldt 1971) are common. Plagioclase grain sizes vary from 1.2 mm to 5.9 mm. Ophitic to sub-ophitic textures are present in some samples, where plagioclase occurs enclosed by either orthopyroxene or clinopyroxene. In some samples, plagioclase chadacrysts are finer grained than unenclosed plagioclase crystals. Myrmekite is very common in the gabbronoritic lithologies and increases to up to 5 modal % in some parts. It generally occurs along plagioclase grain boundaries and in some cases as patches within plagioclase crystals (Fig. 3c).

Fig. 3
figure 3

a Subhedral to anhedral crystals of pyroxene occurring intergranularly between plagioclase laths. Photomicrograph under cross-polarised light of a gabbronorite from depth -74.49 m. b Photomicrograph of anorthosite (depth of − 96.83 m) under cross-polarised light displaying wedge-shaped and bent twin-lamellae. c Myrmekite occurring at the grain boundaries and as patches within plagioclase. Photomicrograph under cross-polarised light of a leucogabbro (depth − 65.86 m). d Clinopyroxene containing orthopyroxene exsolution lamellae, showing a distinct schiller structure. Photomicrograph under cross-polarised light of a gabbronorite (depth − 115.91 m)

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Orthopyroxene modally constitutes 1–38% of samples below the Pyroxenite Marker and 14–37% of samples above the Pyroxenite Marker. It occurs as subhedral to anhedral crystals (Fig. 3d) throughout most of the studied interval. Orthopyroxene grain sizes vary considerably within the gabbronorites, with an average grain size of ~ 2 mm below the Pyroxenite Marker and ~ 4 mm above the Pyroxenite Marker. Orthopyroxene normally contains an abundance of thin and closely spaced clinopyroxene exsolution lamellae. Below the Pyroxenite Marker, orthopyroxene containing thick, bleb-like exsolution lamellae of clinopyroxene is abundant (Fig. 3d). This inverted pigeonite (Poldervaart and Hess 1951) is abundant in the rocks below the Pyroxenite Marker and is replaced by primary orthopyroxene as the low-Ca pyroxene at approximately 2 m above the base of the Pyroxenite Marker. Its modal abundance within the Pyroxenite Marker itself is ~ 1%.

Clinopyroxene modally constitutes between 0 and 38% of the samples studied and occurs as anhedral to subhedral crystals between plagioclase laths. In some samples, ophitic textures were observed, with clinopyroxene oikocrysts enclosing plagioclase. Clinopyroxene displays the presence of orthopyroxene exsolution lamellae and is frequently twinned (Fig. 3d). Clinopyroxene exhibits distinct schiller structures throughout the investigated interval (Fig. 3d).

Pleochroic (brown to colourless) amphibole modally constitutes 0 to 21% of the samples studied and occurs mostly as patches around or near pyroxene crystals. It is particularly abundant within two altered samples of gabbronorite at depths of − 16.69 m and − 143.25 m. The amphibole seems to replace the pyroxene in these two samples. Amphibole also occurs in the vicinity of minerals such as quartz and biotite.

Quartz constitutes up to 13 modal % and occurs mostly interstitially. Biotite and phlogopite constitute up to 9 modal % and also tends to occur interstitially. A detailed study of the opaque minerals was not done, due to their limited occurrence. Opaque minerals constitute up to 1 modal % of the gabbronoritic samples. Other minerals encountered in the gabbroic samples include chlorite (as alteration product of pyroxenes and plagioclase) and talc (as alteration product of pyroxene).

Pyroxenitic rocks

Three prominent pyroxene-enriched horizons were observed in the studied interval. The lower occurrence (at depths 574.96–576.75 m) is an ~ 1.8 m-thick zone surrounded by gabbronorite. The second (514.70–518.53 m) is an ~ 3.8 m-thick zone also surrounded by gabbronorite. The Pyroxenite Marker will be discussed separately below. The sampled pyroxenite horizon (the lower occurrence at depth − 86.85 m) is dominated by poikilitic textures. Orthopyroxene modally constitutes 8% of this sample. Some of the orthopyroxene contains an abundance of thick clinopyroxene exsolution lamellae consistent with it being inverted pigeonite, modally constituting about 45% (Fig. 4a). The unit consist of orthopyroxene crystals that are up to ~ 6 mm in size. Orthopyroxene locally encloses small chadacrysts of clinopyroxene (Fig. 4b). Clinopyroxene modally constitutes about 20% of the sample and occurs as subhedral to anhedral crystals (Fig. 4b). Clinopyroxene displays the presence of orthopyroxene exsolution lamellae and is frequently twinned (Fig. 4a). Greenish pleochroic amphibole modally constitutes about 6% of the pyroxenite horizon and occurs mostly as patches in or as alteration rims on pyroxene crystals. Intercumulus plagioclase (Fig. 4b) modally constitutes about 11% of the pyroxenite sample. Brown to colourless pleochroic phlogopite (Fig. 4b) modally constitutes ~ 1% of the pyroxenite horizon and usually occurs as patches along the edges of orthopyroxene crystal boundaries.

Fig. 4
figure 4

a Inverted pigeonite and twinned clinopyroxene. Photomicrograph under cross-polarised light of pyroxenite (depth of − 86.85 m). b Photomicrograph of pyroxenite (depth of − 86.85 m) under cross-polarised light showing orthopyroxene enclosing small clinopyroxene crystals. c Intercumulus plagioclase in the Pyroxenite Marker and clinopyroxene enclosing plagioclase. Photomicrograph under cross-polarised light (depth of + 0.22 m). d Photomicrograph of the Pyroxenite Marker (depth of + 0.22 m) under cross-polarised light showing orthopyroxene enclosing blebs of clinopyroxene

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Pyroxenite Marker

The Pyroxenite Marker as exposed in BH7771 is broadly similar to that described by von Gruenewaldt (1973), Mitchell (1990), Cawthorn et al. (1991), Maier et al. (2001) and Nex et al. (2002). Orthopyroxene modally constitutes 40–41% of the Pyroxenite Marker and occurs as subhedral to anhedral crystals (8.6–8.9 mm) (Fig. 4c). The orthopyroxene often encloses plagioclase or blebs of clinopyroxene (Fig. 4d). Clinopyroxene (38–44%), with grain sizes of ~ 6–8 mm, appears cumulus in nature. Clinopyroxene displays distinct schiller structure throughout the Pyroxenite Marker (Fig. 4c). Inverted pigeonite was noted in two of the samples of the Pyroxenite Marker at a modal abundance of ~ 1%. Brown to colourless pleochroic phlogopite modally constitutes 1–2% of the samples studied and usually occurs as patches at the grain boundaries of clinopyroxene crystals. Greenish pleochroic amphibole (~ 2%) occurs as patches or as alteration rims on pyroxene crystals. Intercumulus plagioclase (Fig. 4c) modally constitutes between 6 and 11% in individual samples of the Pyroxenite Marker.

Mineral chemistry

Compositional data for plagioclase, clinopyroxene and orthopyroxene are reported in Electronic Appendix D. Between three and four analyses of plagioclase, orthopyroxene and clinopyroxene were performed on each sample, with one analysis spot per crystal centred on the apparent core of each crystal. Due to the absence of visible zonation during petrographic examination of the samples, no attempt was made to study the presence of zonation through the analysis of multiple spots on individual crystals.

Plagioclase

The composition of plagioclase varies across the studied interval (Fig. 2), with an average An% [An% = 100 × molar Ca/ (Ca + Na + K)] of 64 below the Pyroxenite Marker, 61 in the Pyroxenite Marker and 70 above the Pyroxenite Marker. The average anorthite contents for individual samples, with within-sample variation expressed as 1 σ-standard deviations, are between 55.1 ± 0.7 and 73.9 ± 0.2 below the Pyroxenite Marker, 59.7 ± 0.9–61.9 ± 1.3 in the Pyroxenite Marker and 64.2 ± 2.6%–72.8 ± 2.68% above the Pyroxenite Marker. Below the Pyroxenite Marker, an overall increase in the An% can be seen from 57.6 at − 176.84 m to 73.9 at − 0.78 m below the Pyroxenite Marker. Localised decreases in An%, below the Pyroxenite Marker, are coincident with the pyroxenite (− 86.85 m) and the more mafic (melagabbronorite) horizons (− 26.44 m), which is probably a reflection of the intercumulus nature of plagioclase within these samples. The anorthite content of plagioclase decreases at the base of the Pyroxenite Marker to a value of 59.8 and then increases again towards the top of the Pyroxenite Marker to a value of 61.9. In the uppermost portion of the studied section (above the Pyroxenite Marker), an overall increase can be seen from + 1.58 m to + 14.17 m with An% values increasing from 65.6 to 72.8. Very little variation is seen in the last 5 m of the uppermost portion of the studied interval, with An% values averaging ~ 72.

Pyroxene

The average Mg# [molar Mg/(Mg + Fe)] of orthopyroxene below the Pyroxenite Marker is 0.65 ± 0.04, 0.70 ± 0.01 in the Pyroxenite Marker and 0.69 ± 0.01 above the Pyroxenite Marker (Fig. 2). The average Mg# of orthopyroxene for individual samples ranges between 0.58 and 0.71 for the whole succession. Within sample variation of Mg# expressed as 1σ-standard deviations does not exceed 0.05. Below the Pyroxenite Marker, an overall increase in the Mg# can be seen from a value of 0.62 at − 176.84 m to a value of 0.71 at − 1.82 m below the Pyroxenite Marker. The Mg# of orthopyroxene in the Pyroxenite Marker increases marginally from the base (0.69) to the top (0.71), with similar values encountered above the Pyroxenite Marker (0.69 at + 1.81 m to 0.71 at + 14.17 m).

The Mg# of clinopyroxene parallels the orthopyroxene trend and is on average ~ 0.08 higher than that of the co-existing orthopyroxene. The average Mg# for clinopyroxene below, within and above the Pyroxenite Marker is 0.74 ± 0.03, 0.78 ± 0.01 and 0.77 ± 0.01, respectively.

Whole-rock major and trace element geochemistry

Whole-rock major and trace element data are reported in Electronic Appendix A. A significant correlation exists between the modal mineralogy (reported in Electronic Appendix C) and the CIPW norm. Normative An% below, within and above the Pyroxenite Marker varies from 50.2 to 77.7, 60.3 to 69.4 and 59.3 to 74.9, respectively. Normative An% is relatively constant (64.5) from − 177 m to − 96.83 m below the Pyroxenite Marker. A sudden decrease to a value of 50.2 is associated with a pyroxenite layer at -86.85 m. Normative An% increases again from − 74.4 m to − 26.44 m with an average value of 66.7. Normative An% then increases significantly towards the base of the Pyroxenite Marker, from 71.1 to 77.7. Within the Pyroxenite Marker, the normative An% decreases to an average value of 64.0. Above the Pyroxenite Marker (+ 1.58 m to + 14.17 m), normative An% increases from 64.9 to 71.0.

Whole-rock Mg# over the depth interval of − 176.84 m to − 106.01 m fluctuates between 0.68 and 0.73. At a depth of − 96.83 m below the Pyroxenite Marker, there is a significant decrease in the Mg# to 0.43, coincident with the presence of an anorthosite layer. From − 86.85 m to − 50.39 m, the Mg# decreases from 0.74 to 0.62 and at − 43.69 m, the Mg# decreases to 0.58 coincident with the presence of the IRUP. Between − 30.52 m and − 0.78 m, Mg# fluctuates between ~ 0.71 and ~ 0.77. The anorthosite occurring at the base of the Pyroxenite Marker (at depth − 0.38 m) has a whole-rock Mg# of 0.56. Within and above the Pyroxenite Marker, whole-rock Mg# remains relatively constant, with an average of 0.75.

Chondrite-normalised REE patterns are presented in Fig. 5. All samples, apart from one from the Pyroxenite Marker, show enrichment of the LREE relative to the HREE. (Ce/Sm)N values of samples below, within and above the Pyroxenite Marker vary between 1.3 and 3.7, 1.1 and 1.7 and 1.4 and 1.7, respectively. The HREE show limited fractionation with (Tb/Yb)N being on average 1.2 for samples below the Pyroxenite Marker, 0.8 for the Pyroxenite Marker itself and 1.0 for samples above the Pyroxenite Marker. The majority of samples exhibit positive Eu anomalies [EuN/((SmN + GdN)/2)] varying between 1.0 and 10.1. Samples showing negative Eu anomalies are the pyroxenite and IRUP samples, with values of 0.7–0.9.

Fig. 5
figure 5

Chondrite-normalised REE patterns of rocks across the Pyroxenite Marker interval as exposed in BH7771. Normalisation factors from Lodders (2003). The inset shows the range in REE concentrations for the Main Zone of the western Bushveld Complex at Union Section (Maier and Barnes 1998) and the range in REE concentrations of B3 magma (Harmer and Sharpe 1985).

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Selected incompatible and compatible trace element concentrations and ratios are presented in Fig. 6 as a function of stratigraphic height. Zr is fairly constant, at an average value of 9.0 ppm, in the gabbronoritic and anorthositic rocks covering the depth interval − 176.84 to + 14.17 m. A localised maximum in the Zr concentration, with a value of 23.9 ppm, is observed in the pyroxenite layer covering the depth interval − 86.72 to − 86.85 m. Fluctuations in the Zr concentrations are associated with the pyroxenite layers as well as the IRUP samples. In the Pyroxenite Marker, the Zr is fairly constant with an average value of 11.8 ppm.

Fig. 6
figure 6

Simplified lithological log and diagrams showing variation in modal mineralogy, whole-rock Zr, (Sr/Al2O3) × 104 and Cr/Sc against depth for BH7771. *Numbers refer to depth in metres in borehole relative to the base of the Pyroxenite Marker

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The Sr/Al2O3 ratio as a proxy for differentiation of plagioclase shows a step-wise decrease towards the Pyroxenite Marker, from a value of ~ 0.0015 at the base of the studied interval, to an average of ~ 0.0013 within and above the Pyroxenite Marker. The Cr/Sc ratio as a proxy for clinopyroxene differentiation shows very little variation from the base of the studied interval, with values << 5 and up to ~ 40 m below the base of the Pyroxenite Marker. It then increases rather abruptly to values > 20 towards the Pyroxenite Marker, a value which is also retained in samples above the Pyroxenite Marker. The abrupt increase in Cr/Sc ratios occurs across the level where inverted pigeonite is replaced by primary orthopyroxene as the dominant low-Ca pyroxene.

Isotope geochemistry

87Sr/86Sri ratios and ƐNdi values are presented in Table 1, and their variation with depth is displayed graphically in Fig. 7. Values for 87Sr/86Sri show little variation over the depth interval − 106.01 m to − 65.86 m, averaging 0.7081. An abrupt decrease to a value of 0.7071 is observed in the sample from a depth of − 50.39 m, with the overlying sample at depth − 26.44 m again showing a value of 0.7081. A trend of decreasing 87Sr/86Sri values is observed upwards between the aforementioned depth and the lowermost sample within the Pyroxenite Marker, which returned a value of 0.7069. An abrupt increase in 87Sr/86Sri to a value of 0.7076 is encountered within the next sample (depth 0.22 m), with values showing a scattered increase to a local maximum of 0.7079 in the sample from a depth of + 9.17 m.

Table 1 Rb–Sr and Sm–Nd isotopic data for whole-rock samples from BH7771
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Fig. 7
figure 7

Variations in 87Sr/86Sri and ƐNdi (at 2.055 Ga) with depth across the Pyroxenite Marker interval in BH7771. *Numbers refer to depth in metres in borehole relative to the base of the Pyroxenite Marker. Data for the Western Limb from Cawthorn et al. (1991) and for the Eastern Limb from Sharpe (1985). Insets show close up of the area around the Pyroxenite Marker

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Values of ƐNdi show relatively little variation from the base of the studied succession to a depth of − 50.39 m, varying between − 6.2 and − 6.7. The next sample, at depth − 26.44 m, returned an ƐNdi value of − 4.9. This is followed by a sustained decrease in the ƐNdi values towards depth − 2.77 m, where ƐNdi is − 6.8. From − 1.82 m to the lowermost sample of the Pyroxenite Marker, ƐNdi increases from − 6.7 to − 6.0. From the base of the Pyroxenite Marker to the top of the studied interval, significant inter-sample variations appear to be the norm, with values fluctuating between − 5.8 and − 6.7.

Discussion

Three main models have been proposed to explain the observed compositional and isotopic variations across the Pyroxenite Marker interval. Sharpe (1985) attributed the Pyroxenite Marker interval to mixing occurring between underplated Main Zone magma and upwards displaced resident magma of Upper Critical Zone lineage. Cawthorn et al. (1991) suggested that the Pyroxenite Marker interval formed in response to mixing between approximately equal volumes of resident melt (of Main Zone lineage) and less-differentiated melt that was added to the chamber at the approximate level of the Pyroxenite Marker. Vantongeren and Mathez (2013) proposed that the melt that was added at the level of the Pyroxenite Marker was added as numerous, “small” injections that became homogenised, cooled and partially crystallised between successive injections.

The implication of all of the cited models is that the rocks occurring across the Pyroxenite Marker interval, which is thought to represent the zone of mixing between resident and incoming melts, should exhibit isotopic compositions that may be modelled through simple two-component melt-melt mixing curves. As will become clear from the subsequent section, in which our data across the Pyroxenite Marker interval are compared to several melt-melt mixing models based on different assumptions from the published literature, none of the mixing models provide a particularly good fit to our data, necessitating a re-assessment of the mechanism by which mixing took place and the nature of the end-members involved.

Constraints on the Nd-isotopic composition and amount of melt/magma added at the level of the Pyroxenite Marker

To constrain the Nd-isotopic composition of the magma added at the level of the Pyroxenite Marker, the following additional parameters need to be assessed: (i) the Nd-isotopic compositions of the resident and mixed magmas, (ii) the Nd concentration of the resident and incoming or mixed magmas and (iii) the ratio in which the two magmas mixed, defined as:

$$f=\frac{\mathrm{Mass\, of\, resident\, magma}}{\mathrm{Mass\, of\, resident\, magma}+\mathrm{Mass\, of\, incoming\, magma}}$$
(1)

For the purposes of our calculations, the Nd-isotopic composition of the resident magma is taken as ƐNdi = − 7.01 (the average value for the lower Main Zone as reported by Maier et al. 2000) and that of the mixed magma as ƐNdi = − 6.43 (Tanaka and Masuda 1982). It should be noted that the latter value is not well constrained and is based on the sole available whole-rock Nd-isotopic composition currently available for the upper Main Zone (Vantongeren et al. 2010). As shown by our own data across the Pyroxenite Marker interval (Fig. 5), the REE contents of the resident and incoming magmas were likely rather similar. Similar Nd contents for the resident and incoming magmas allow for a simplification of the two-component mixing equation from:

$${\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Mix}={\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Res}.\left(\frac{{Nd}_{Res}.f}{{Nd}_{Mix}}\right)+{\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Inc}.\left(\frac{{Nd}_{Inc}.\left(1-f\right)}{{Nd}_{Mix}}\right)$$
(2)

to

$${\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Mix}={\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Res}.\left(f\right)+{\left(\frac{{}^{143}Nd}{{}^{144}Nd}\right)}_{Inc}.\left(1-f\right)$$
(3)

The mixing ratio may be constrained with reference to the available Sr-isotopic data for the Pyroxenite Marker interval. Cawthorn et al. (1991) argued that the magma added at the level of the Pyroxenite Marker had an 87Sr/86Sri ratio of 0.7063, similar to that of a magma intruded into the lower reaches of the western Bushveld Complex (Davies and Cawthorn 1984). On this basis, and on the assumption that the resident magma and the incoming magma had similar Sr contents, these authors could show that the volume of magma added at the level of the Pyroxenite Marker was approximately equal to that of the resident magma. If it is assumed that the Sr contents of the resident and incoming magmas were indeed similar, and setting 87Sr/86SrMix = 0.7073 (Kruger et al. 1987), 87Sr/86SrRes = 0.7084 (Maier et al. 2000) and 87Sr/86SrInc = 0.7063 (Cawthorn et al. 1991), equation 3 can be applied to the Rb–Sr system to show that the mass of the incoming magma was approximately equal to that of the resident magma (\(f=0.48)\). The Nd-isotopic composition of the incoming magma is then constrained to ƐNdi = − 5.9 (Fig. 8a).

Fig. 8
figure 8

Whole-rock isotopic mixing models for the Pyroxenite Marker interval. a This mixing scenario is based on the assumption that the resident and incoming magmas had similar Sr contents. A is the resident magma: 87Sr/86Sri = 0.7084 and ƐNdi = − 7.01 (Maier et al. 2000); and B is the incoming magma: 87Sr/86Sri = 0.7063 (Cawthorn et al. 1991) and ƐNdi = − 5.9. b This mixing scenario is based on the assumption that the Sr content of the incoming magma (319 ppm) was higher than that of the resident magma (181 ppm) as per Vantongeren and Mathez (2013). A is the resident magma: 87Sr/86Sri = 0.7084 and ƐNdi = -7.01 (Maier et al. 2000); and B is the incoming magma: 87Sr/86Sri = 0.7067 (Vantongeren & Mathez 2013) and ƐNdi = − 5.9. c This mixing scenario is based on the assumption that the Sr content of the incoming magma (319 ppm) was higher than that of the resident magma (181 ppm) and that the Nd content for the resident magma (16 ppm) is lower than the incoming magma (31 ppm), as per Vantongeren and Mathez (2013). A is the resident magma: 87Sr/86Sri = 0.7084 and ƐNdi = − 7.01 (Maier et al. 2000); and B is the incoming magma: 87Sr/86Sri = 0.7067 (Vantongeren & Mathez 2013) and ƐNdi = − 6.1

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Vantongeren and Mathez (2013), using equilibrium Sr contents calculated from trace element contents of the minerals present across the Pyroxenite Marker interval, estimated an 87Sr/86Sri ratio of 0.7066–0.7068 for the magma that was introduced at the level of the Pyroxenite Marker. Their calculations, performed on the assumption that no magma was lost from the chamber, show that the Sr content of the incoming magma (319 ppm) was likely significantly higher than that of the resident magma (181 ppm), such that the incoming magma had a much greater effect on the 87Sr/86Sri ratio of the mixed magma than the resident magma. Using 87Sr/86Sri ratios of 0.7067 (Vantongeren and Mathez 2013) and 0.7084 (Maier et al. 2000) for the incoming and resident magmas, respectively, and the aforementioned Sr contents a mixing ratio (\(f\)) of 0.49 is required to produce an 87Sr/86Sri ratio of 0.7073 for the mixed magma, a result that is not significantly different from that produced when relying on the values of Cawthorn et al. (1991) as presented in the preceding paragraph. In the present scenario, the Nd-isotopic composition of the incoming magma is still constrained to ƐNdi = − 5.9 when assuming similar Nd contents for the resident and incoming magmas (Fig. 8b). When using the calculated Nd contents for the resident (16 ppm) and incoming (31 ppm) magmas as calculated by Vantongeren and Mathez (2013), an ƐNdi value of − 6.1 is obtained for the incoming magma (Fig. 8c). This value does not change significantly when using the equilibrium liquid compositions calculated by Vantongeren and Mathez (2013) for the resident and incoming magmas in the involving where magma loss.

Mechanism of mixing: mixing of minerals or melts?

As is clear from Fig. 8, none of the mixing models provide a particularly good fit to our data across the Pyroxenite Marker interval, with the majority of data points scattered about the mixing curves. We propose that this is the result of minerals having crystallised from the resident and incoming magmas, respectively, and having been incorporated into individual rock layers across the Pyroxenite Marker interval. The notion that rocks occurring within the Bushveld Complex might represent mixtures of minerals accumulated from different parts of a stratified magma column is not entirely novel. This was shown to be the case in the Merensky Reef (Prevec et al. 2005; Seabrook et al. 2006), in noritic blocks enveloped by anorthosite in the upper Main Zone (Bourdeau et al. 2022) and also across the Pyroxenite Marker interval, where Cawthorn et al. (1991) reported the existence of plagioclase primocrysts from both magmas occurring within single samples and envisaged that “composite packets of liquid plus crystals” from the incoming magma (intruded at an intermediate level) plunged through the column of resident magma to accumulate at the base of the chamber. It seems feasible that pyroxenes that crystallised from the resident and incoming magmas might have co-accumulated in a similar manner, although compositional evidence for this is likely to have been destroyed due to the rapidity of solid-state diffusion of Fe and Mg in pyroxene relative to that of plagioclase, in which coupled substitution of CaAl and NaSi has to take place (Cawthorn et al. 1991).

Figure 9 shows the mixing model as presented in Fig. 8a with two additional mixing curves that constrain all of our data apart from two data points that we consider to be aberrant. The assumptions on which the upper mixing curve (curve C) is based are as follows: (i) rocks represent mixtures of plagioclase crystallised from the resident magma and clinopyroxene crystallised from the incoming magma, (ii) plagioclase and clinopyroxene have Sr concentrations of 457 ppm and 11.6 ppm, respectively and (iii) plagioclase and clinopyroxene have Nd concentrations of 1.4 ppm and 9.2 ppm, respectively. The Sr and Nd concentrations are averages from the work of Vantongeren and Mathez (2013) for rocks across the Pyroxenite Marker interval. The lower mixing curve is based on the assumption that rocks represent mixtures of plagioclase crystallised from the incoming magma and clinopyroxene crystallised from the resident magma. Orthopyroxene was omitted from our calculations for the sake of simplicity, but would have little effect on both the Sr and Nd isotopic compositions of the mixtures due to the low abundance (< 0.2 ppm) of both elements in it. Our data across the Pyroxenite Marker interval are not fully constrained if the incoming magma had an ƐNdi value of -6.1 (Fig. 8c), and for that reason we propose that the incoming magma had an ƐNdi value of -5.9 (Fig. 8a, b). This would imply that the Nd contents of the resident and incoming magmas must have been rather similar, in contrast to what has been proposed by Vantongeren and Mathez (2013).

Fig. 9
figure 9

Mixing models for the Pyroxenite Marker interval in BH7771, showing that data may be constrained through the incorporation of minerals derived from the resident and incoming magmas, respectively, into rocks occurring across this interval, as opposed to only through the mixing of aphyric magmas. The central mixing curve is as per Fig. 8a. Mixing curve C represents mixtures of plagioclase crystallised from the resident magma and clinopyroxene crystallised from the incoming magma. Mixing curve D represents mixtures of plagioclase crystallised from the incoming magma and clinopyroxene crystallised from the resident magma. See text for details

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Where and how did mixing take place?

Most models for the development of the Pyroxenite Marker interval assume that mixing of the resident and incoming magmas happened within the confines of the presently exposed Bushveld Complex magma chamber and that the incoming magma was added gradually to the chamber, either forming a layer that was stabilised at an intermediate level within a stratified magma column before being thoroughly mixed with the resident magma (Cawthorn et al. 1991) or as numerous, “small” injections that became homogenised, cooled and partially crystallised between successive injections (Vantongeren and Mathez 2013). A possibility that has not been explored is that the rocks occurring across the Pyroxenite Marker interval may represent mixtures of minerals and melts that underwent mixing during ascent and subsequent emplacement from a sub-compartmentalised staging chamber as envisaged by Roelofse and Ashwal (2012). In this scenario, the Main Zone beneath the Pyroxenite Marker interval could represent the product of the repeated intrusion of variably contaminated crystal mushes with an average 87Sr/86Sri and ƐNdi on the order of 0.7084 and − 7.01, respectively. Input and mixing during ascent between this magma and another crystal-laden magma with fundamentally different isotopic composition (e.g. 87Sr/86Sri = 0.7063; ƐNdi = − 5.9), resident in a different sub-compartment, could equally account for the present dataset, a finding which is consistent with that of Yao et al. (2021), who argued that the Rustenburg Layered Suite below the level of the Pyroxenite Marker was emplaced as numerous crystal slurries that obtained a variety of isotopic compositions as a result of interaction with chemically and isotopically distinct assimilants as they turbulently rose through transcrustal conduits before entering the magma chamber and being deposited on the chamber floor as masses of cumulates when magma velocity slowed.

The distribution of rocks dominated by plagioclase from the incoming magma and pyroxene from the resident magma versus plagioclase from the resident magma and pyroxene from the incoming magma appears difficult to reconcile with straightforward gravity-driven processes. Rocks occurring more than 10 m below the Pyroxenite Marker appear to be dominated by plagioclase from the resident magma and pyroxene from the incoming magma, whereas rocks within 10 m below the Pyroxenite Marker appear to be dominated by pyroxene from the resident magma and plagioclase from the incoming magma. Above the Pyroxenite Marker, rocks again appear to be dominated by plagioclase from the resident magma and pyroxene from the incoming magma. This apparently random distribution of minerals constituting the different layers across the Pyroxenite Marker interval is entirely consistent with the model of Yao et al. (2021) as explained above, which would also explain the lateral and vertical differences in isotopic composition across the Pyroxenite Marker interval as is apparent from Fig. 7, where our own data across the interval are compared with that of Sharpe (1985) and Cawthorn et al. (1991).

Irrespective of the veracity of these arguments, which would need additional analytical work (e.g. in situ isotopic determinations or isotopic determinations on mineral separates) to further evaluate, the magma that was responsible for the observed isotopic shifts across the Pyroxenite Marker interval had a unique Sr and Nd isotopic composition (Fig. 10) not exhibited by any of the rocks occurring below this level of the intrusion. The similarity in ƐNdi values for the Upper Critical and Main zones, as is apparent from Fig. 10, along with the observed Nd-isotopic variations across the Pyroxenite Marker interval as presented in this study, argue against the model of Sharpe (1985), who attributed the Pyroxenite Marker interval to mixing between underplated Main Zone magma and displaced resident magma of the Upper Critical Zone.

Fig. 10
figure 10

The results of this study and the proposed Sr–Nd isotopic composition of the magma added at the level of the Pyroxenite Marker compared to the Sr–Nd isotopic composition of rocks occurring below the Pyroxenite Marker. (Data for the Lower, Critical and Main zones come from Maier et al. 2000)

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Conclusions

The results of our modelling suggest that:

  • The magma added at the level of the Pyroxenite Marker had a unique Sr and Nd isotopic composition not seen in any of the layered rocks occurring below the level of the Pyroxenite Marker, with an 87Sr/86Sri ratio of 0.7063–0.7067 and an ƐNdi value on the order of − 5.9. The unique isotopic composition of the added magma argues against the model of Sharpe (1985), who attributed the Pyroxenite Marker interval as having formed in response to mixing between displaced Upper Critical Zone and underplated Main Zone magma.

  • The magma added at the level of the Pyroxenite Marker records evidence for a lower degree of crustal contamination compared to the resident magma.

  • The Pyroxenite Marker interval records evidence for the accumulation of minerals derived from both the resident and incoming magmas, which is inconsistent with the injection, homogenisation, cooling and crystallisation of numerous small pulses of magma as envisaged by Vantongeren and Mathez (2013).

  • The inferred presence of minerals with isotopic compositions pointing to derivation from either the resident or incoming magma co-existing in rocks of the Pyroxenite Marker interval could be explained either by the mixing of minerals settling through a stratified magma column, as envisaged by Cawthorn et al. (1991), or potentially through the intrusion and mixing of crystal-laden magmas with unique isotopic compositions from a sub-compartmentalised staging chamber during ascent, as envisaged by Roelofse and Ashwal (2012). Detailed in situ isotopic investigations, which fall beyond the scope of the present investigation, may assist in differentiating between these fundamentally different alternatives.