2. Brief Outline of the Geology of the Great Dyke
The Great Dyke of Zimbabwe is one of the world’s largest and well studied intrusions. It is a layered mafic-ultramafic intrusion that cuts across the Zimbabwe Archean craton (
Figure 1a) and it acts as host to the second largest resource of PGEs in the world. The Great Dyke is composed of two major chambers, the North and South Chambers (
Figure 1b) which Wilson and Prendergast (1989) [
25] further subdivided into subchambers on the basis of continuity of layering, style, and thickness of cyclic units (
Figure 1b). A third and relatively small chamber, the Mvuradonha Chamber, occurs at the extreme north of the dike [
25].
The structure of the Great Dyke is synclinal with inwardly dipping layers—the dips increasing from the central axis towards the margins but decreasing again near the walls [
25,
27,
28]. At depth, the Great Dyke has a dike-like feeder which in places has been interpreted to be connected to deep-seated magma chambers [
29]. The transverse section of the Great Dyke (
Figure 2a) has been interpreted to be trumpet-shaped with individual layers thinning away from the axis and eventually becoming incorporated in the Border Group which rests against to the Great Dyke walls [
25]. The Mafic Sequence and the upper layers of the Ultramafic Sequence may have extended beyond the present margins of the Dyke by several kilometers, and these lateral extensions are thought to have been entirely eroded [
29].
The longitudinal section of the layers of the Great Dyke plunges gently towards the center of each subchamber, to form an overall boat-like structure [
27,
28]. The Mafic Sequence, which reaches its maximum thickness in the Darwendale Subchamber, is preserved as remnants in the center of each of the subchambers (
Figure 1b) and has been extensively eroded. Here, 2000 m of the Ultramafic Sequence is overlain by 1150 m of the Mafic Sequence [
30]. The stratigraphic section in the vicinity of Unki Mine in the Shurugwi Subchamber from just above the Mafic Sequence/Ultramafic Sequence contact to the footwall of the mineralized zone is shown in
Figure 2b. Plagioclase pyroxenite is overlain by a 6m-thick layer of plagioclase websterite which is capped by a very thin chromitite layer (or chromitite stringer) (
Figure 2b). Within and above this chromitite stringer occur gabbronorites of the Mafic Sequence (
Figure 2b). The BMSZ occurs ~3 m below the plagioclase websterite/plagioclase pyroxenite contact (
Figure 2b).
The Ultramafic Sequence in all subchambers is divided into a lower Dunite Succession (1000 m) of dominantly dunite and harzburgite and an upper Bronzitite Succession (1000 m) of mainly harzburgites and orthopyroxenites (
Figure 3) [
31]. Cyclic layering is well developed in the Ultramafic Sequence, and in the Shurugwi Subchamber, a total of twelve cyclic units have been recognized [
31]. The cyclic units in the Dunite Succession are defined by narrow chromitite layers overlain by dunite layers (Fig. 3) [
31].
In the Bronzitite Succession, however, cyclic units are defined by a complete progression from a basal chromitite layer, through dunite, harzburgite and olivine orthopyroxenite layers to a well-developed orthopyroxenite layer at the top, and as such have been interpreted by Wilson (1982) [
32] to represent the ideal cyclic unit of the Great Dyke. Cyclic units such as those occurring in the Dunite Succession have, therefore, been interpreted to be incomplete [
25]. Six major cyclic units are recognized in the Bronzitite Succession of the Darwendale Subchamber. Chromitite layers occurring in the Ultramafic Sequence which were originally numbered consecutively downwards [
27,
28], for example, are now numbered according to the cyclic unit in which they occur [
13,
32]. The cyclic units are interpreted to have been formed by repeated injections of magma followed by fractionation and differentiation [
32]. In all the three chambers of the Dyke, the topmost cyclic units 1 and 2 of the Bronzitite Succession are similar [
13,
33].
Both horizontal and vertical tectonics have been invoked to explain the colinearity of the Great Dyke, its satellites as well as the associated fracture pattern [
26,
34,
35,
36]. According to Wilson (1990) [
26], the Great Dyke fracture pattern (
Figure 1) can be explained in terms of a progressive craton-wide main deformation model involving SW movement of the Zimbabwe Province relative to the Limpopo Province. Continued deformation of an early crossfold pattern led to intra-block movements on different scales involving major strike-slip deformation zones and thrusting [
26]. Continuation of the stress field, which is inferred to have effected the post-Chirumhanzu Suite overthrusting of the Limpopo Province onto the Zimbabwe Province with variation in the direction of maximum compressive stress, is thought to have led to the development of the Great Dyke fracture pattern [
26]. Fracturing, overthrusting, and late Achaean main deformation were interpreted by Wilson (1990) [
26] to be cratonization responses to a protracted collision between the Zimbabwean and Kaapvaal cratons. Based on U-Pb dating of zircon and rutile from orthopyroxenites of the P1 Pyroxenite layer, the age of the Great Dyke is 2575.4 ± 0.7 Ma [
37]. Further SHRIMP U-Pb studies on the Great Dyke and its satellites by Wingate (2000) [
38] yielded a comparable Neoarchean emplacement age of 2574 ± 2 Ma for baddeleyite.
3. Occurrence of the Sulfide Mineralization in the Great Dyke and in the Shurugwi Subchamber
The 90 km-long Shurugwi Subchamber has a shape that has been controlled by its proximity to the Shurugwi greenstone belt (
Figure 1b and
Figure 2), resulting in its deflection and constriction in places [
28,
39]. At a width of 3.5 km in its central part, the Shurugwi Subchamber is the narrowest of all the subchambers of the Great Dyke. The Shurugwi Subchamber is important in terms of its petrology because it is the only subchamber that is extensively bounded (25 km) by greenstones as opposed to the dominantly Archean granites for the other subchambers [
28]. The Shurugwi Subchamber abuts against Archean granites on its eastern margin whereas on the western margin, it forms a steep contact mostly with rocks of the Shurugwi greenstone belt (
Figure 4a,b). Thus, irregularities in the wall-rocks may have affected the marginal zones of the Shurugwi Subchamber magma chamber and its style of layering [
39]. Fine-grained feldspathic pyroxenite, including olivine websterite, forms the Border Group on the west side. The east marginal rocks are largely in contact with granitic rocks and the dip of the contact is 40° to the west. The eastern margin of the Shurugwi Subchamber is characterized by abundant auto-intrusive granite dykes and pegmatoids resulting from melting of wall-rock and hydrous crystallization. Hybrid lithologies, such as pyroxene-bearing amphibolites, amphibole-bearing pegmatoids, amphibole gabbros and magnetite diorites, are common in the marginal zone as irregularly developed, steeply dipping, fine-grained layers [
39].
Cyclic Unit 1 of the Darwendale Subchamber contains several zones of sulfide mineralization [
28,
40]. Cyclic Unit 1 (
Figure 3) is 420 m thick at the axis of the Dyke in the Darwendale Subchamber and comprises a number of subunits defined on the basis of narrow chromitite layers, repeated stratigraphy, and reversals in mineral compositions [
32]. Bichan (1970) [
40] identified sulfide mineralization within the lower subunits, with the main metal zones occurring in Sulfide Zones l (i), l (ii), l (iii), 2, and 3 numbered with increasing distance downwards from the Mafic Sequence/Ultramafic Sequence contact [
41]. The base of Sulfide Zone 1, the MSZ [
42], occurs at the top of the main orthopyroxenite layer and the MSZ extends into the overlying websterite (
Figure 2b and
Figure 3).
PGE-enriched layers of the Great Dyke also include most of the chromitite layers [
37], as well as several silicate horizons in the upper portions of the Ultramafic Sequence [
13,
14,
37,
43,
44,
45]. According to Oberthür (2002) [
46] and Oberthür (2011) [
14], up to 2 ppm PGE (Pt/Pd = 0.1) occur in the C1d chromitite and its host rocks. PGEs in the Shurugwi Subchamber, as in the rest of the Great Dyke, occur mainly in the tabular and stratabound MSZ. The MSZ is located in orthopyroxenites close to, and overlapping with, the websterite layer and contains up to 6% sulfide. The more disseminated Lower Sulfide Zone (LSZ), occurring lower in the succession, has lower PGE concentrations than the MSZ and contains up to 1.5% sulfide [
41]. Having been discovered over a century ago [
47], the MSZ is similar in form in all the chambers of the Great Dyke [
25,
27,
28,
45]. In the field, a prominent feature of the MSZ is that preferential weathering of the plagioclase oikocryst margins results in the formation of the characteristic nodular texture, or ‘potato reef’, which is commonly associated with MSZ exposures [
47]. The MSZ was originally traced by copper staining, but trenching revealed that the staining was only a surface feature which does not extend at depth [
28]. Thus, this copper staining was most likely formed as a result of supergene enrichment.
The width of the MSZ varies from 2 to 8 m and it contains up to 5 ppm PGE and up to 8% sulfides over 2–3 m [
14,
43,
46,
48,
49]. Within the MSZ are the so-called “offset” metal distribution patterns defined by peak Pd levels which occur near the base of the reef, whereas peak Pt and Cu levels are reported to occur at progressively higher stratigraphic levels within the MSZ [
14,
43]. Both the mineralization and the offset pattern are continuous along the strike of the Great Dyke although considerable down-dip variation in the thickness of the MSZ and the amount of sulfides occurs [
44]. Mining activity in the MSZ is currently being undertaken at several localities such as at Mimosa Mine in the Wedza Subchamber, at Unki Mine in the Shurugwi Subchamber, and at Ngezi Mine in the Darwendale Subchamber. Sulfides occur as interstitial phases to the cumulus orthopyroxene crystals, and the sulfides are closely associated with late-stage minerals and, where abundant, leading to the formation of a net texture locally [
49]. In addition, Wilson (2001) [
49] observed that the sulfides have a tendency to concentrate around plagioclase oikocrysts margins, accentuating further the weathering of the nodules.
The thicker LSZ occurs approximately 35 m stratigraphically below the MSZ, but also within the uppermost P1 Pyroxenite of the Ultramafic Sequence. No petrological break occurs between the MSZ and LSZ. Both the MSZ and LSZ occur in all the subchambers of the Great Dyke in all remnants of the P1 Pyroxenite layer (e.g.,
Figure 3 and
Figure 4). The P1 Pyroxenite consists of tabular to prismatic medium-grained cumulus orthopyroxenes which occur as inclusions in postcumulus plagioclase and clinopyroxene [
25,
44,
49]. The LSZ has been investigated from boreholes that intersected it in the Shurugwi Subchamber [
49].
Interstitial plagioclase and late-stage minerals such as phlogopite, K-feldspar, quartz, apatite, magnetite, and rarely rutile and zircon also occur in the P1 Pyroxenite [
49]. Further, low-grade PGE mineralization has been described in the Middle Mafic Unit of the Mafic Sequence and also in the Mineralized Marginal Zone of the extreme marginal facies of the Border Group [
45]. Characteristics of mineralization, rock types, and the different magma chamber structures that occur in the Great Dyke have been described elsewhere [
39,
41,
43,
45,
50].
The MSZ has been classified as a sulfide hosted, magmatic PGE deposit [
39]. The regular distribution pattern of the PGEs in all areas where the MSZ has been investigated have been interpreted to indicate a primary, sulfide controlled fractionation pattern whereby the PGEs have been scavenged from the magma due to their strong partitioning into primary sulfide [
41,
43]. Analogues of this process were reported in the Munni Munni intrusion of Western Australia [
51,
52]. Although the MSZ has been identified in all subchambers of the Great Dyke [
25,
33], lithological details, position, and petrological characteristics are variable even within the same subchamber and they also differ between subchambers. Chromium-spinel occurring in the MSZ from the Darwendale Subchamber is enriched in TiO
2 by up to 2.1 wt % [
53].
A common occurrence in the central and western zones of the Shurugwi Subchamber is a suite of xenoliths which range from a meter to several tens of meters in size (
Figure 4a,b), with some exceeding 100 m in length [
28,
39]. The distribution of the xenoliths in the subchamber is also shown in
Figure 4b. The xenoliths are commonly composed of metasedimentary quartzite and banded iron formation possibly derived from the Wanderer Formation of the Shurugwi greenstone belt [
54]. Xenoliths of ultramafic rocks are inferred to be derived from the Shurugwi greenstone belt whereas blocks of chromitite and serpentinite were possibly derived from the Shurugwi greenstone belt [
31]. Mafic fragments which occur within the gabbroic rocks and the upper Ultramafic Sequence represent rock types encountered in the Great Dyke with a wide range of mineral compositions that were considered to have been derived from an early formed Border Group of the Great Dyke magma chamber [
39].
6. Discussion
From petrographic descriptions provided, it is proposed that the paragenetic sequence for the MSZ at Unki Mine is orthopyroxene plagioclase → clinopyroxene → chromite → sulfides + PGEs → amphibole/chlorite/quartz, although the documented hydrothermal activity may complicate this sequence. The observations made of sulfides occurring in association within cumulus minerals such as plagioclase, clinopyroxene, and orthopyroxene as well as sulfides occurring in association with minerals like chlorite and amphibole which are interpreted to be hydrothermal alteration phases (e.g.,
Figure 8 and
Figure 9) point to two generations of sulfide mineralization at Unki Mine. The sulfides that were encountered during this investigation are pyrrhotite, chalcopyrite, pentlandite, and pyrite. Pyrite was only observed in the two topmost samples (MusCa02 and MusCr01), and it is inferred to be an alteration product of pyrrhotite as it has a tendency of occurring in association with the hydrothermal alteration phases of amphibole and chlorite.
All primary sulfide minerals of pyrrhotite, chalcopyrite, and pentlandite tend occur as inclusions in cumulate minerals and generally tend to be coarse-grained in comparison to sulfides which occur in association with alteration minerals. Some sulfides inferred to have been formed due to hydrothermal alteration activities, however, may also be coarse-grained, but these tend to occur only in association with chlorite and amphibole. Some primary sulfides also occur in association with hydrothermal minerals, suggesting that primary mineralization in the MSZ may have been subsequently enriched due to magmatic hydrothermal fluids. Hydrothermal alteration processes may also have led to the formation of pyrite. Sulfides occurring as inclusions in chromite and silicate crystals are interpreted to be primary and are not associated with any fractures [
71], and were thus likely formed during the crystallization history of the intrusion. The relatively finer-grained sulfides, which have a tendency to occur in association with hydrothermal alteration phases of chlorite and amphibole and which are also associated with fractures and veinlets, were probably formed later during hydrothermal alteration. Further, the occurrence of fine-grained sulfides laths in association with hydrothermal alteration minerals like chlorite (e.g.,
Figure 8f) lends support to these sulfides having been formed during hydrothermal processes. The fine laths appear to have the same orientation as that of the basal cleavage in the chlorite host, probably an indication that fluid flow may have been responsible for the precipitation of the sulfide laths during chlorite formation. Thus, remobilization of earlier formed magmatic mineralization by hydrothermal processes appears to have occurred within the MSZ at Unki Mine. The difference
The predominance of chlorite in the sample just above the footwall fault (PxaFwt01) may have been promoted by hydrothermal fluids which likely permeated through the fault helping form the chlorite crystals above this fault (
Figure 9e) in particular, and possibly throughout the entire MSZ. In contradistinction, in the sample just below the footwall fault (MusPabFwt01), no evidence of hydrothermal processes is observed here unlike in the sample just above the footwall fault, since the footwall fault is located immediately above this sample (
Figure 9f). The timing of this fault is not constrained at present; it may have been synchronous with magmatism and it was later re-activated.
Pyroxene compositions from the present study are broadly consistent with magmatic pyroxenes and are comparable to those from other layered intrusions elsewhere (
Figure 17a). A few clinopyroxenes analyses, however, are diopsides enriched in calcium relative to most magmatic clinopyroxenes (
Figure 17a). Magmatic clinopyroxenes have a tendency to be uniform in rock successions where no fresh batches of magma were injected into the magma chamber, but secondary clinopyroxenes such as diopsides form during hydrothermal alteration due to calcium enrichment of magmatic clinopyroxenes caused by the transfer of components through the hydrothermal fluid [
72,
73].
The majority of pyroxene analyses have totals that are less 100% (
Table 3 and
Table 4,
Supplementary Materials), suggesting that these pyroxenes may contain small amounts of H
2O or other volatiles as submicroscopic inclusions [
77]. Further, the absence of non-quadrilateral components in diopside compared to igneous clinopyroxene, for example, has been interpreted to indicate formation of the diopside in a relatively lower-temperature hydrothermal environment of >300 °C [
72,
78,
79]. The diopsides in this investigation are also characterized by low concentrations of non-quadrilateral components (
Table 3 and
Table 4,
Supplementary Materials). Great Dyke MSZ clinopyroxene samples are thus more magnesian than those from most other layered intrusions [
32,
80].
Feldspars similarly have totals that are less 100% (
Table 5,
Supplementary Materials), again suggesting that they may contain small amounts of H
2O or other volatiles such as volatile submicroscopic inclusions [
77]. Plagioclase analyses from the Merensky Reef of the Bushveld Complex mostly plot in the labradorite and andesine fields, with few analyses plotting in the bytownite and oligoclase fields (
Figure 17b). Thus, just like pyroxene Mg#s (e.g.,
Figure 17a; [
81]), Merensky Reef plagioclase analyses are less calcic than those from the MSZ. Although some analyses for quartz crystals were obtained in this investigation, these were all much lower than 100, possibly due to interference from other minerals due to the very fine-grained nature of the quartz crystals.
The Pozanti-Karsanti ophiolite has amphiboles ranging in composition from magnesiohornblende to actinolite [
82] (
Figure 18a). The aluminous amphiboles are thought have formed during high pressure, high temperature hydrothermal alteration [
82]. In this ophiolite, calcic amphiboles such as actinolite commonly occur rimming clinopyroxenes or high-Al and high-Fe hornblende [
82]. Further, Stakes and Taylor (1992) [
83] suggested that high-temperature metamorphic hornblendes in the Semail ophiolite of Oman (pargasites and edenites; [
84]) were replaced by low-temperature actinolitic hornblendes due to the extensive moderate to low-temperature hydrothermal alteration.
Unlike amphiboles from ophiolites, those from the present investigation do not show any zoning, suggesting that they were formed from replacement of orthopyroxenes during hydrothermal alteration [
67]. Further, the observation that amphibole Mg#s from the MSZ mimic those of both pyroxenes (
Figure 11), lend support to the interpretation that Unki Mine MSZ amphiboles were derived from the alteration of pyroxenes in-situ and were not formed from cations and fluids derived from outside the intrusion. Magnesiohornblende and tschermakite are consistent with amphibolite facies hydrothermal alteration, whereas tremolite and actinolite are consistent with greenschist facies hydrothermal alteration. Hydrothermal alteration of orthopyroxenes formed Fe-Mg-Mn amphiboles whereas calcic amphiboles were formed from the hydrothermal alteration of clinopyroxenes [
67]. An application of the Zane and Weiss (1998) [
70] method of classifying chlorites based on microprobe analyses reveals that MSZ chlorites and chlorites from other intrusions such as the Stillwater Complex and the Mann Complex are Mg-rich Type I chlorites (
Figure 18b), likely pointing to their derivation from hydrothermal alteration of magnesium-rich minerals such as pyroxenes. Chlorite occurring in igneous rocks such as those under investigation, and its association with BMS, lends support to a hydrothermal origin for some of the mineralization as chlorite in such rocks tends to be derived from hydrothermal alteration of primary ferromagnesian minerals [
67] such as pyroxenes which form the dominant minerals in the Cyclic Unit 1 of the Great Dyke.
Compositions of chlorite from the MSZ at Unki Mine have total Fe {=Fe
2+ + Fe
3+ (apfu)} values which range from a low of 1.09 to a high of 3.43 with an average value of 2.34 (apfu) (
Table 7,
Supplementary Materials). Such low total Fe (apfu) values from the MSZ at Unki Mine are comparable to those from the Stillwater Complex and Mann Complex of 0.45–4.79 (average of 2.58) apfu and of 1.72–2.54 (average of 2.04) apfu, respectively. The low Fe
2+/(Fe
2+ + Mg) ratios of MSZ chlorites, which vary from 0.26 to 0.5, as well as the very low Fe
3+ values characterize these MSZ chlorites as Fe-rich and unoxidized [
24].
Equations for estimating the temperature of chlorite formation make use of the relationship between Fe in the octahedral site and temperature [
61,
62,
63]. The three equations by these investigators were utilized in this investigation and comparisons of the compositions of chlorites from these equations are shown in
Figure 19a,b. Fe
2+/(Fe
2+ + Mg) is plotted against Al
iv (apfu) (
Figure 19a) and Si/Al (in apfu) (
Figure 19b) reveal that samples under investigation compare better with the chlorite composition fields of Cathelineau (1988) [
62]. Consequently, high temperature calculations for chlorites utilized in this study only made use of the equations of Cathelineau (1988) [
62].
The estimated temperatures of the hydrothermal fluid(s) that affected the MSZ at Unki Mine are shown in
Figure 19c. These estimated temperatures of the hydrothermal event(s) range from 241 to 390 °C and from 491 to 640 °C (
Figure 19c). This bimodal temperature distribution could indicate that two different hydrothermal alteration events affected the MSZ at Unki Mine: one in the relatively higher temperature range of 491–640 °C, and another one in the relatively lower temperature (241–390 °C) range. These inferred temperature ranges of the hydrothermal event with is thought to have occurred during the cooling history of the MSZ at Unki Mine are broadly comparable to temperatures obtained from chlorite thermometry from the Duluth Complex (Gál et al., 2011). The lack of any stable isotope compositions consistent with meteoric fluids in the Great Dyke makes the involvement of meteoric fluids in mineralization in this intrusion unlikely. The observed bimodal temperature distribution is also consistent with amphibolite facies hydrothermal alteration for formation of magnesiohornblende and tschermakite, and the formation of tremolite-actinolite under greenschist facies conditions.
These calculated temperatures of hydrothermal fluids which affected the MSZ at Unki Mine fall within, and at the slightly higher end, of the range of hydrothermal fluids [
89]. The variation of temperatures obtained by chlorite thermometry with stratigraphic height across the investigated are is shown in
Figure 20a and there appears to be a general increase from the base to the top of the MSZ, with a reversal to lower temperatures. Such a trend of increasing temperatures with stratigraphic height probably indicates the involvement of magmatic temperatures as temperature due to circulating heated fluids would likely result in a decrease in temperature with stratigraphic height. The drop to lower temperatures would be due to cooling of the intrusion before temperature increase again resulting in heating probably due to new magma injections.
It must be pointed out that several chlorite analyses from the MSZ have compositions that fall outside of the compositional calibrations of various calculations of chlorite thermometry (
Figure 19a,b) and these results must be interpreted with caution. However, use of other chlorite thermometry equations suitable for low temperatures [
64] indicates that the MSZ chlorites may have been formed at very low to medium temperatures except in a few samples (
Figure 20a). This may have happened long after cooling of the intrusion. Chlorite thermometry temperatures lower than 200 °C, for example, would indicate more enriched Great Dyke silicate δ
18O values [
90] than measured values which are consistent with high temperatures [
91,
92]. Further, low chlorite thermometry temperatures would also be inconsistent with the observations made in this work that chlorites occur in association with primary sulfides which would require higher temperatures of formation.
Both of the documented hydrothermal fluid events that affected the MSZ at Unki Mine must have been of magmatic origin, since oxygen isotope [
91,
92], H and S isotope [
92] compositions and thermometry are consistent with high temperature, magmatic values for the Great Dyke. Indeed, Li et al. (2008) [
92] also suggested hydrothermal alteration was responsible for affecting the mineralization of the MSZ at the Hartley Platinum Mine located in the Darwendale Subchamber of the Great Dyke. Prendergast (1990) [
93] also reports on hydrosilicate “alteration” in the Wedza-Mimosa Platinum Deposit located in the Wedza Subchamber, the southernmost subchamber of the Great Dyke (
Figure 1).
In comparison to other layered intrusions such as the Bushveld Complex, isotope data from the interval straddling the contact between the Ultramafic and Mafic Sequences of the Great Dyke indicate a less enriched composition of initial
87Sr/
86Sr ratios (0.7024–0.7028) and ε
Nd (−1 to +1) [
94]. Sulfur isotope studies (δ
34S) of Great Dyke samples carried out by Li et al. (2008) [
92] on pyrite, pyrrhotite, and pentlandite range from 0.1‰ to 1‰ and Maier et al. (2015) [
94] obtained δ
34S on bulk-rock samples which range from −0.3‰ to 0.3‰ all fall within the 0 ± 5‰ range of mantle values [
95]. Thus, relatively moderate amounts of contamination of the Great Dyke parent magma must have occurred, an indication that the Great Dyke crystallized from a single magma type. Unlike in the Bushveld Complex where mixing of compositionally distinct magmas is proposed to have caused sulfide melt saturation, formation of both PGE mineralization in the MSZ and the uppermost cyclic unit (Cyclic Unit 1) chromitite layers in the Great Dyke may have been triggered by silicate fractionation and magma mixing between resident magma and unevolved replenishing magma [
14,
25,
43,
45,
48,
49,
74,
94]. Li and Ripley (2005) [
96] have suggested that sulfide melt saturation in the hybrid magma can be triggered only if both mixing magmas themselves are nearly saturated in sulfide melt, which seems unlikely to be the case in the Great Dyke in relation to the new replenishing magma.
However, unlike at Hartley Platinum Mine in the Darwendale Subchamber of the Great Dyke where alteration minerals such as epidote, talc, calcite and ankerite are reported to occur, these minerals were not observed in the MSZ at Unki Mine. Studies of radiogenic isotopes, also from Hartley Platinum Mine, of Sm-Nd and Rb-Sr (ε
Nd values of mostly −1 to +1 and initial
87Sr/
86Sr isotope ratios of 0.7024–0.7028; [
94]) and Re-Os isotopes (narrow initial
187Os/
188Os ratios range of between 0.1106 and 0.1126; [
97]) are consistent with magmatic signatures for the Great Dyke as only moderate amounts of crustal contamination must have occurred in the Great Dyke. Results from these radiogenic isotope studies by Maier et al. (2015) [
94] and Schoenberg et al. (2003) [
97], thus, help rule out the possible involvement of significant amounts of external fluids in the mineralization of the Great Dyke, further lending support to the findings from this investigation that the hydrothermal fluids which interacted with the MSZ at Unki Mine are of magmatic origin.
The occurrence of hydrothermal alteration events of different temperatures observed in this investigation seems not to be a feature only associated with the MSZ of the Great Dyke, as the J-M Reef of the Stillwater Complex (
Figure 19d) and the Mann Complex (
Figure 19e) are also characterized by chlorites with record distinct temperature ranges. This has implications for the involvement of multiple hydrothermal alteration events in the concentration of magmatic PGEs in layered intrusions.
Mogessie et al. (1991) [
98], Ripley et al. (1993) [
99], Severson (1994) [
100], and Gál et al. (2011, 2013) [
23,
101] have all documented hydrothermal processes and their roles in the concentration of PGE remobilization in the troctolitic intrusions of the Duluth Complex in Minnesota, USA. In this complex, Mogessie et al. (1991) [
98] observed that Cu and PGEs were remobilized from the primary magmatic mineralization by C, O, H, S, and Cl-enriched fluids along fracture zones whereas Gál et al. (2011) [
23] described vein-type, hydrothermal Cu-mineralization associated with actinolite–chlorite–prehnite–pumpelleyite–calcite alteration assemblage in the hanging-wall of the SKI at the Filson Creek deposit in the same intrusion. In the Babbitt Cu–Ni deposit of the Bathtub intrusion, Ripley et al. (1993) [
99] interpreted that hydrothermal remobilization of primary ores occurred on the basis of hydrogen and oxygen isotope studies Ripley et al (1993) [
99] concluded that fluids both from magmatic and metasedimentary sources were responsible.
Two-pyroxene thermometry yielded temperatures that range from 850 to 981 °C (
Figure 20a,b;
Table 8), using the equations of both Putirka (2008) [
66] and of Brey and Kohler (1990) [
65], with the exception of a relatively lower temperature of 670 °C which was obtained using the equation of Brey and Kohler (1990) [
65]. These pyroxenes are inferred to be been initially characterized by magmatic temperatures that were subsequently affected by the hydrothermal alteration as oxygen isotope thermometry yielded magmatic temperatures in the Mafic Sequence which range from 1016 to 1155 °C, with an average temperature of 1091 °C [
91]. The Mafic Sequence, with the exception of a 2.3 m-wide zone of alteration located between 13 and 16 m above the Mafic Sequence/Ultramafic Sequence contact [
102], is characterized by fresh plagioclase, clinopyroxene, and orthopyroxene cumulates that do not show any evidence of hydrothermal alteration [
100,
102], implying that the hydrothermal event(s) which affected the Great Dyke did not affect much of the gabbroic rocks overlying the Ultramafic Sequence. Studies of the Ultramafic Sequence beneath the mineralized P1 layer can address the extent of this hydrothermal activity within the Ultramafic Sequence.
Use of the calculated pressures (1.8–4.8 kbar) (
Figure 20c,d;
Table 8) obtained from using the temperatures calculated from the equations of Putirka (2008) [
66] and of Brey and Kohler (1990) [
65] allows one to estimate the thickness of the overburden above the MSZ. If one assumes lithostatic conditions (3000 kg/m
3) the overburden during the hydrothermal alteration of the MSZ ranges from 6.1 to 12.6 km and 8.8–16.3 km using Equations (38) and (39) of Putirka (2008) [
66], respectively. These thicknesses are much higher than the current 200 m depth of the MSZ at Unki Mine, and even higher than the estimated total thickness of the Mafic Sequence of 1150 m in the Darwendale and Sebakwe Subchambers [
30,
103]. Thus, at the level of the MSZ at Unki Mine, the thickness of the overburden above the 1150 m maximum thickness of the Great Dyke (Mafic Sequence) ranged from 5 to 11 km using Equation (38) of Putirka (2008) [
66] and from 7.7 to 15 km using Equation (39) of Putirka (2008) [
66].
Metamorphic conditions experienced in some greenstone belts in the Zimbabwe craton, for example, the Shamva greenstone belt, is low pressure but high temperature type, and range from upper amphibolite-facies (600–650 °C, 3–4 kbar) conditions near the margins of the greenstone belt to greenschist facies (450–550 °C, 1.5–2.5 kbar) within the core of the greenstone belt such as at Shamva Mine [
104]. These are pressures and temperatures comparable to those obtained in this work. Thus, mineralization in the MSZ of the Great Dyke may have occurred at depths comparable to those under which metamorphism of some of the greenstone belts in the Zimbabwe craton occurred.
Other investigators have also documented the involvement of hydrothermal alteration is other large layered intrusions elsewhere. Evidence for hydrothermal alteration that locally altered magmatic silicates, recrystallized BMS aggregates and remobilized sulfides and PGEs in the J-M reef of the Stillwater Complex has been documented by Polovina et al. (2004) [
10]. The alteration assemblage in this Complex is dominated by chlorite, clinozoisite, serpentine, calcite, talc, white mica, magnetite, tremolite, Cl-rich ferropargasite, and quartz [
10]. Hydrothermal alteration of magmatic Ni-Cu-PGE deposits has also been reported to occur in the Tootoo and Mequillon magmatic sulfide deposits located in the Cape Smith Belt, Canada [
22]. Olivine melagabbronorites, the host rocks in the Tootoo and Mequillon magmatic sulfide deposits, were pervasively altered to hydrous mineral assemblages dominated by chlorite, tremolite-actinolite, and relict late-magmatic hornblende that are consistent with greenschist facies metamorphism [
22]. Large-scale and very saline hydrothermal fluid activity is also interpreted to have affected rocks of the Sudbury Complex and in the process caused remobilization of both base metals and PGEs [
105,
106].
Hydrothermal alteration by meteoric water is thought to be responsible for decoupling of S and Cu occurring in the Sonju Lake intrusion of the of the 1.1 Ga Midcontinent rift-related Beaver Bay Complex in north-eastern Minnesota, USA [
92,
107]. Thermal migration zone refining is a process whereby igneous activity at convergent margin builds a thick volcanic pile which becomes a barrier to further magma ascent, leading to magma underplating by injection of sills at the base of the volcanic pile [
108]. When magma arrives at the location of underplating, it reacts and releases heat and water to the overlying materials (i.e., previously intruded sills) and this results in a downward moving zone having a near-steady-state temperature gradient [
108]. This leads to compositional differentiation by wet thermal migration then occurs in the middle of the underplated region but not on the more rapidly cooled edges of the sills, and this is thought by Lundstrom (2009) [
108] to take place over time scales of millions of year. If the Sonju Lake intrusion was formed by a top-down process of sill injection and reaction (thermal migration zone refining) [
109], then the PGE reef (PGE-Cu-S interval) in this intrusion was interpreted to have been formed as a moving sulfide band passed downward through a mineral-melt mush at the particular temperature of sulfide saturation [
110]. Further, in the Sonju Lake intrusion, disseminated sulfides have δ
34S values that range from −2.2‰ to +3‰ (V-CDT), within the range of mantle-derived minerals and rocks, suggesting that contamination by country rock sulfur was not an important process in the formation of the metal-rich interval [
104]. δ
18O values of plagioclase from the Sonju Lake intrusion range widely from 5.6‰ to 12.0‰ (V-SMOW), an indication that a relatively low-
18O fluid (δ
18O ~3‰–5‰) interacted with the rocks of the intrusion at temperatures less than ~275 °C based on oxygen isotope thermometry and the prehnite-pumpellyite to greenschist facies transition [
92,
104]. Some of alteration assemblages such as chlorite and actinolite, as well as magmatic stable isotope compositions reported by Park et al. (2004) [
107] for the Sonju Lake intrusion are broadly comparable to those reported in the present investigation and by previous investigators on the Great Dyke.
PGE-bearing horizons, or reefs, in layered intrusions are generally thought to have been formed during upward accumulation of mineral deposits. Within the context of the cumulate model (where crystals accumulate from the bottom upwards through crystal accumulation from a large volume of magma), the origin of PGE reefs is currently being debated between a number of models [
111]. In the orthomagmatic model, the PGE deposits are thought to have been formed when immiscible sulfide melts that scavenged PGEs from the silicate melt as they settle due to gravity [
112]. In the hydromagmatic model, PGE deposits are thought to have been formed when a magmatic hydrothermal fluid generated in the already formed crystal pile moves upward, scavenges and concentrates the PGEs and deposits them in the reef [
12,
113]. In the micro-nuggets model, metallic micro-nuggets are thought to segregate from magma and concentrate PGEs [
114]. The assumption in these three models describing PGE mineralization in layered intrusions is that layered intrusions were formed during the upward accumulation of crystals settling in a magma chamber (i.e., the cumulate model). PGE reefs have also recently been proposed to form if smaller layered intrusion formed incrementally by top-down sill accumulation [
109,
115] where no magma chamber may have existed. Findings from this work lend support to both an orthomagmatic model as well as a hydromagmatic imprint for the origin of mineralization in the MSZ at Unki Mine. Since this study was on silicate mineral compositions and did not focus on PGEs, no evaluation of the micro-nugget model can thus be undertaken.
In the Great Dyke, all three models (orthomagmatic, hydromagmatic, and micro-nugget models) have been proposed for the origin of the MSZ [
41,
43,
49,
50,
113]. Stratigraphic offsets in peak concentrations of PGEs and BMS which occur in the MSZ [
14,
43] have, in part, been attributed by Li et al. (2008) [
92] to the interaction between magmatic PGE-bearing BMS assemblages and hydrothermal fluids. Li et al. (2008) [
92] presented mineralogical and textural evidence from the Hartley Platinum Mine located in the Darwendale Subchamber of the Great Dyke which they interpreted to indicate that alteration of BMS and mobilization of metals and S occurred during hydrothermal alteration. Sulfur isotope data of pyrite, pyrrhotite, and chalcopyrite ranging from 0.1% to 0.8‰, as well as O isotope data for orthopyroxene ranging from 5.1‰ to 6.5‰ from the MSZ suggest that the fluids involved in the alteration were of magmatic origin [
92]. Further, actinolite has both O and H isotope data ranging from 5.0‰ to 5.6‰ and 64‰ to 73‰, respectively, which Li et al. (2008) [
92] interpreted to be consistent with magmatic fluids.
According to Oberthür et al. (2003) [
17], variations in Pt/Pd ratios in the MSZ may be related to hydrothermal alteration due to the observation that most of the Pd and Rh are hosted in pentlandite, whereas Pt dominantly occurs in the form of discrete minerals. According to Li et al. (2008) [
92], minor (<5 vol %) to significant (20 vol %) alteration is present in the MSZ. Actinolite, epidote, carbonate, talc, magnetite, and pyrite are the most common secondary minerals. Actinolite alteration occurs throughout the MSZ, but it is most intense in the sample which occurs immediately below the peak in Pd content and also occurring immediately below the peak in Pt content in the PGE subzone [
17,
92]. Boudreau and Meurer (1999) [
116] argued that a chromatographic process that involved magmatic fluids was responsible for the Pt/Pd offsets in the MSZ.