Greenstone belts at the northernmost edge of the Kaapvaal Craton: Timing of tectonic events and a possible crustal fluid source

Precambrian Research 253 (2014) 96–113

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Greenstone belts at the northernmost edge of the Kaapvaal Craton: Timing of tectonic events and a possible crustal fluid source Jan D. Kramers a,∗ , Michel Henzen b , Laurent Steidle b a b

Department of Geology, University of Johannesburg, P.O. Box 542, Auckland Park 2006, Johannesburg, South Africa Institute of Geological Sciences, University of Bern, Baltzerstrasse 3, 3012 Bern, Switzerland

a r t i c l e

i n f o

Article history: Received 9 February 2014 Received in revised form 3 June 2014 Accepted 9 June 2014 Available online 19 June 2014 Keywords: Geochronology Archaean tectonics Limpopo Complex Kaapvaal Craton Crustal heating Crustal fluids Crustal doming

a b s t r a c t The Hout River shear zone, a complex thrust system, marks the boundary between the low- to medium grade metamorphic granitoid-greenstone Northern Kaapvaal Craton province and the high grade Southern Marginal Zone of the Limpopo Complex to the north of it. The Giyani greenstone belt, located on the Northern Kaapvaal Craton immediately adjacent to the Hout River shear zone, is in a key position for studying the tectonics of this shear zone. We have mapped two central transects through the greenstone belt, dated metamorphic minerals by 207 Pb/206 Pb stepwise leaching and 40 Ar/36 Ar, and compared our results with data from the Pietersburg/Polokwane and Rhenosterkoppies greenstone belts, which are close to the Hout River shear zone further west. The SE domain of the NE–SW striking Giyani greenstone belt appears to constitute a NW-vergent nappe. Movement along the Hout River shear zone is consistently SW-vergent, varying in character between frontal and lateral ramps, and nappe. The NW domain of the Giyani greenstone belt is dominated by ultramafic schists with mafic and banded iron formation (BIF) intercalations; it is considered a separate tectonic unit from the SE domain in which mafic (intrusive and pillow-volcanic) rocks are prevalent, with intercalated clastic sediments. Correlations between this SE domain and the lower mafic unit of the Pietersburg greenstone belt and with the mafic-BIF association of the Rhenosterkoppies greenstone belt are possible. Ages for peak metamorphism in the immediate footwall of the Hout River shear zone from this and previous work yield a weighted mean age of 2717 ± 28 Ma (95% confidence), documenting a “hot-iron” metamorphism. A garnet age of 2833 ± 38 Ma in the SE domain of the Giyani greenstone belt overlaps with ages from 2854 to 2870 Ma for youngest zircons in syntectonic intrusions within the Giyani and the Pietersburg greenstone belts, as well as youngest detrital zircons in the (probably syntectonic) clastic sedimentary Uitkyk Formation of the latter. These ages constrain a tectonometamorphic event in both belts, termed here the PGB-Lwaji orogeny. Amphibole 40 Ar/39 Ar dates from the Hout River shear zone (from this and previous work) and shear zones within the Giyani greenstone belt yield disturbed age spectra with step ages mainly in the 2550–2670 Ma range. The amphiboles are chemically heterogeneous and the ages can be shown to reflect mixtures of older and younger amphibole generations, due to reactivation of the shear zones and fluid activity after the main event in the Hout River shear zone. Thermal modelling shows that, in spite of the very low content of radioactive heat producing elements in the Southern Marginal Zone, its lower crust, after moderate crustal thickening, could be heated to granulite facies conditions in 30–60 Ma, a much shorter time than the 120 Ma age difference between the PGB-Lwaji orogeny and the Hout River shear event. A causal link between the two tectonic events is thus precluded. On the other hand, previously low-grade supracrustal rocks from the Kaapvaal Craton buried at lower crustal levels during Hout River shear event could provide a source for fluids up to about 80 Ma after the event. © 2014 Elsevier B.V. All rights reserved.

∗ Corresponding author. Tel.: +27 11 559 4755; fax: +27 11 559 4702. E-mail address: [email protected] (J.D. Kramers). http://dx.doi.org/10.1016/j.precamres.2014.06.008 0301-9268/© 2014 Elsevier B.V. All rights reserved.

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

1. Introduction

2. The Giyani greenstone belt

Greenstone belts in the Kaapvaal Craton appear to mark the sutures of terrane accretion that occurred in the course of its assembly (de Wit et al., 1992a). This is shown by the fact that they separate granitoid-gneiss provinces having different zircon age ranges as well as Nd and Hf model ages (Jahn and Condie, 1995; Zeh et al., 2009; Schoene et al., 2009; Kramers and Zeh, 2011). Thus the branches of the Barberton Greenstone Belt (GB) separate the up to 3660 Ma Ancient Gneiss Complex of Swaziland, the up to 3450 Ma Stolzburg–Steynsdorp (or Barberton-South) Terrane, and the 3200–3100 Ma Barberton-North Terrane (Schoene et al., 2009). Likewise, the Murchison greenstone belt marks a suture, dated at ca. 2970 Ma, between the Barberton North Terrane and a back-arc system to the north of it (Zeh et al., 2009, 2013). While the Barberton greenstone belt has an extremely complex structure, the architecture of the two main greenstone belts of the Northern Kaapvaal Craton, the Murchison and Pietersburg GB’s can be characterized as moderately steeply north vergent thrust sequences. The metamorphic grade of supracrustals is up to amphibolite facies (Vearncombe, 1988 for the Murchison GB; de Wit et al., 1992a for the Pietersburg GB), and no high grade metamorphism is recorded in the gneissic basement adjacent to these greenstone belts on either side. The Hout River Shear Zone (HRSZ) system, which separates the Southern Marginal Zone (SMZ) of the Limpopo Complex from the Kaapvaal Craton proper (Fig. 1), has a completely different character. First, it marks a radical increase in metamorphic grade from the low-grade conditions in the Kaapvaal Craton to (albeit regionally retrogressed) granulite facies in the SMZ. Second, currently available geochemical characteristics, Pb isotope data and Nd model ages on either side of it are very similar, demonstrating that it is not a terrane boundary but the SMZ represents the northernmost Kaapvaal Craton at a deeper crustal level (Kreissig et al., 2000). Third, it is south-vergent, and in places has a strike strike-slip component (McCourt and van Reenen, 1992). Fourth, it is in most areas not associated with a supracrustal belt. Only the Giyani GB, the small Rhenosterkoppies GB and the very minor Loskop GB remnant are adjacent to it. In the Giyani GB, in the immediate footwall of the HRSZ, replacement of kyanite by sillimanite (McCourt and van Reenen, 1992) shows a prograde path at >4 kb, and a hairpin-shaped clockwise P-T loop was determined peaking at 580 ◦ C and 4.5 kb. These conditions are coincident with part of the decompression-cooling path of the SMZ granulitic gneisses (Perchuk et al., 1996; Smit et al., 2001). These mutually consistent observations are interpreted as reflecting “hot-iron” metamorphism within the Giyani GB, caused by the overthrust hot SMZ material (Roering et al., 1992a,b). Between the Giyani GB and a point approximately 20 km E of the Rhenosterkoppies GB, foliations and lineations within the SMZ are horizontal or subhorizontal, so that the HRSZ has the character of a nappe sole. The inferred hot-iron metamorphism is in accord with this character (Smit et al., 2014) The position of the Giyani and Rhenosterkoppies GB’s along the footwall of the HRSZ means that their study is highly relevant to the unravelling of tectonics along this shear zone. In this paper we present the results of mapping NW–SE transects through the Giyani GB along the Klein Letaba and Ntsami Rivers and geochronology of metamorphic minerals, and compare these with published data on the HRSZ (Kreissig et al., 2001), the Pietersburg GB (de Wit et al., 1992a,b, 1993) and the Rhenosterkoppies GB (Passeraub et al., 1999). We address the questions of (a) links between tectonic events, and (b) possible crustal fluid sources in the context of lower crustal heating following crustal thickening.

2.1. Geological setting and previous work

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The Giyani greenstone belt (Fig. 1) is contiguous in its eastern region, but W of the Klein Letaba River it consists of a northern (Khavagari) and a southern (Lwaji) arm, which are separated by a granitoid-gneiss zone (McCourt and van Reenen, 1992). Gravity and geoelectric profiles across the belt (de Beer and Stettler, 1992; Kleywegt et al., 1987) are interpreted to show a 4 km downdip extension into the crust for the Lwaji, and a mere 1 km depth for the Khavagari arm, whereas the central portion of the belt is very shallow. McCourt and van Reenen (1992) reported mafic and ultramafic metavolcanic schists, as well as banded iron formation and fine to medium grained clastic sediments in the Giyani GB. They found that, due to the structural complexity, no stratigraphic sequence can be defined. However, they noted that ultramafic schists dominate the NW zone, whereas the metavolcanic rocks in the central and SE domains are on the whole mafic. Various sedimentary units occur throughout the belt. The lithological strike as well as that of the schistosity is between NE–SW and ESE–WSW, with prominent shear zones developed particularly close to the NW margin of the belt. Here the boundary of the belt itself coincides with the HRSZ, and is intensely sheared, but subsidiary shear zones parallel to it occur within it, with the Khavagari shear zone as a prominent example (McCourt and van Reenen, 1992). These authors found lineations and shear sense indicators close to the NW edge indicating SW vergent thrusting, oblique to the margin of the Belt. This was confirmed by de Wit et al. (1992c), who further noted S to SW dipping foliations, pointing to NW vergent thrusting that predated the structures related to the HRSZ. Several gold mineralizations, notably the Klein Letaba, Frankie and Fumani mines and the Gemsbok prospect, occur within the Giyani GB in the immediate footwall of the HRSZ (Gan and van Reenen, 1996). The mineralization is associated with fluid-related retrogression of upper amphibolite facies lithologies. In contrast to the NW margin of the Giyani GB, no marked shear zones have been found to coincide with its SE boundary. A lit-par-lit intrusive contact with the granodioritic Klein Letaba Gneiss is reported here instead (Kröner et al., 2000 and references therein). A meta-andesite intercalated with ultramafic and mafic schists close to the NW margin of the belt has yielded a zircon 207 Pb/206 Pb evaporation age of 3203 ± 2 Ma (4 analyses, Kröner et al., 2000) which is regarded as the depositional age of this part of the succession. Although intrusion ages from zircon dating in the gneisses of the northern KC are mainly in the range 2780–2950 Ma (Kröner et al., 2000; Zeh et al., 2009; Laurent et al., 2013), there is evidence of reworked much older crust: The 2931 ± 8 Ma Meriri granite just south of the Khavagari arm of the Belt contains 3200–3230 Ma zircon xenocrysts that have Hf model mantle derivation (TDM ) ages between 3350 and 3500 Ma (Zeh et al., 2009; see also Kramers and Zeh, 2011). Also, Nd TDM model ages of ca. 3400 Ma were obtained from two metapelites in the Khavagari arm (Kreissig et al., 2000). The older zircon ages could thus reflect a crustal source for clastic sediments in the Giyani GB. A probable maximum age limit for the deformation is given by zircon ages of 2877 ± 33 and 2874 ± 2 Ma (by single grain ID-TIMS as well as 207 Pb/206 Pb evaporation) on a feldspar porphyry that intrudes talc-tremolite schists in the central part of the Belt, and is itself foliated (Kröner et al., 2000). To obtain a minimum age of tectonism and metamorphism along the Hout River shear zone (HRSZ), Barton and van Reenen (1992) dated muscovites in pegmatites crosscutting its foliation by Rb–Sr and obtained ages between 2630 and 2680 Ma. In accord with these results, stepwise leaching 207 Pb/206 Pb analyses on syntectonic, prograde garnet, staurolite and kyanite from schists in the

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Fig. 1. Overview map of northern Kaapvaal Craton and Southern Marginal Zone of the Limpopo Complex (SMZ), showing the mapped area and sample localities for published geochronological work. MGB: Murchison greenstone belt; PGB: Pietersburg greenstone belt; GGB: Giyani greenstone belt; RGB: Rhenosterkoppies greenstone belt; KL: Klein Letaba gneiss; GL: Groot Letaba Gneiss (referred to as Duiwelskloof Batholith by Laurent et al., 2013); HR: Hout River gneiss; Tu: Turfloop granite; Me: Meriri granite; M: Matok granite; Ma: Matlala granite; Mo: Moletsi granite; Ui: Uitloop granite (referred to as Mashashane granite by Laurent et al., 2013). Localities of dated samples – filled diamonds: U-Pb on zircon and monazite; white diamonds: Pb/Pb stepwise leaching on metamorphic minerals; while ellipses; Ar dating of amphiboles; white rectangles; Rb/Sr on pegmatitic muscovites. Sources: (1) this work; (2) Kreissig et al. (2001); (3) Passeraub et al. (1999); (4) Barton and van Reenen (1992); (5) Barton et al. (1992); (6) Zeh et al. (2009); (7) Henderson et al. (2000); (8) de Wit et al. (1993); (9) Kröner et al. (2000); (10) Zeh and Gerdes (2012); (11) Laurent et al. (2013). For a review of further published geochronology within the SMZ see Belyanin et al. (2014).

Khavagari Hills yielded 2728 ± 19 Ma, 2712 ± 37 and 2672 ± 53 Ma, respectively (Kreissig et al., 2001). These latter data constrain the age of oblique thrusting and prograde to peak metamorphism along the HRSZ, where it forms the NW boundary of the Giyani GB, close to 2700 Ma. Interestingly, 40 Ar/39 Ar dating of fabric-forming hornblende from amphibolite in the HRSZ of the same region yielded major step ages between 2580 and 2620 Ma, in disturbed patterns, a very similar result to those from 40 Ar/39 Ar analyses of two amphibolite samples taken 100 and 150 km further west along the HRSZ (Kreissig et al., 2001). 2.2. General geology and petrography of the mapped area Outcrop in the belt away from the ranges of hills marking the NW and SE margins is generally poor, but along the Klein Letaba and Ntsami rivers two complete NW–SE transects across strike could be mapped (Fig. 2). Along these transects, as also noted by McCourt and van Reenen (1992) and Kröner et al. (2000), the spectrum of rock types changes, although due to the structural complexity no sequence can be defined. Along the NW margin, chlorite–actinolite schists dominate, which are talc-bearing in the northernmost area. Further, an association of cherts and banded iron formation with banded amphibolites builds the marked Nangombe Hill NE of Giyani, as well as the Khavagari Hills. There are rare calc silicate and metapelitic lenses close to the NW edge of the belt, but no other fine-clastic metasediments occur in the NW domain. In the central and SE domains of the belt, a range of mafic rocks and diverse clastic metasediments are dominant although minor ultramafic schists are also present. Further, in the central and SE domains deformed dykes and minor stocks of porphyritic tonalite are widespread, and

these are absent in the NW domain. Within both domains what appears to be a lithostratigraphic sequence is locally repeated and inverted by open to isoclinal folding as well as (probably) thrusting, but the contact between the two domains occurs only once. Further, foliations near this contact are parallel to it. From these observations we conclude that the contact between the two domains is tectonic. The chlorite–actinolite–(talc) schists are strongly sheared in many areas, including near the contact between the domains, but this contact is nowhere exposed so that the hypothesis of a tectonic contact cannot be tested. In any case, no stratigraphic relationship can be established between the two domains. The NW contact of the Giyani GB supracrustal rocks with the gneisses of the SMZ is clearly tectonic. It coincides with the HRSZ and a number of subsidiary zones of intense shearing, as noted by McCourt and van Reenen (1992). In contrast, there is no increase in the intensity of shearing towards the SE contact against the gneisses. In the southernmost tip of the mapped area this contact cuts the local supracrustal sequence and foliation. Porphyritic tonalitic dykes are more common and larger close to the contact than elsewhere, and a number of them appear to emanate from the gneisses as lit-par-lit intrusions, noted also by Kröner et al. (2000), and references therein. Further, several large rafts of supracrustal units occur within the granitoid gneisses SE of the Giyani GB (Fig. 2), notably in the Ntsami River section. Since these dykes and the adjacent gneiss terrain are deformed and show a foliation similar to that of the supracrustals, the SE edge of the Giyani GB appears to be a syntectonic intrusive one. The chlorite–actinolite and talc/chlorite actinolite schists, dominant in the NW domain, are strongly foliated; plagioclase (occurring in the talc-actinolite schists, in the northernmost region) is the only

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Fig. 2. Geological map of the Klein Letaba and Ntsami River sections of the Giyani greenstone belt. Dated samples shown in red (garnet Pb/Pb) and green (amphibole 40 Ar/39 Ar).

recognizable relict magmatic mineral and is usually sericitized and saussuritized. From the texture, talc is seen to replace olivine and orthopyroxene. These schists are thus at least in part derived from intrusive feldspar peridotites or pyroxenites. The association with chert and BIF bands indicates that these are sills, rather than tectonically emplaced mantle peridotites. Chlorite defines the foliation or even predates it, while actinolite occurs in a syntectonic, fabric forming, as well as a posttectonic, garben-like mode. The talc-free chlorite–actinolite schists are typically rich in magnetite and have accessory rutile. There are no recognizable magmatic relict minerals and it cannot be established whether the protolith was extrusive or intrusive. Amphibolites in the NW domain are mostly strongly banded, particularly near shear zones, with hornblende defining the foliation and lineation. The banded amphibolites do not retain primary features pointing to the protolith, although the frequent association with BIF suggests that some of them were basalts. In the central – SE domain, spotted amphibolites predominate, which in some cases retain relict pyroxene mantled by hornblende (Fig. 3), as well as plagioclase. As the associated metasediments (see below) have not undergone high grade metamorphism, these relicts must be magmatic and these amphibolites are thus metagabbros. The mantling texture is strongly reminiscent of the features produced

by autometamorphism during hydration of gabbros in suboceanic magma chambers (Manning et al., 1996, 2000). In many instances, the pyroxenes and hornblendes are entirely replaced by actinolite. Very similar features were observed in the lower, mafic and ultramafic unit of the Pietersburg GB (de Wit et al., 1992a). Garnet was

Fig. 3. Thin section photograph of spotted amphibolite (Klein Letaba River at 23◦ 22 30 S, 30◦ 46 30 E) showing original igneous pyroxene mantled by amphibole.

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carry garnet and staurolite porphyroblasts that however predate the last deformation in which quartz stringers and muscovite define a foliation that wraps around them. This mineralogy and texture documents an amphibolite facies peak metamorphism, followed by syntectonic greenschist facies retrogression. 2.3. Overview of structure and metamorphism

Fig. 4. Clastic metasediments (Ntsami River, 23◦ 17 40 S, 30◦ 49 50 E). (A), outcrop photograph of undeformed laminated greywacke; white spots are sericitized andalusite (B) thin section photograph showing sericitized andalusite.

seen in amphibolites in three localities only: close to the HRSZ in Mangombe Hill, in the Klein Letaba River in the Central Domain, and in the southeast of the mapped area (the dated sample HL 108, see Fig. 2). The latter occurrence is in somewhat heterogeneous and relatively quartz-rich amphibolites that may contain assimilated sedimentary material. Meta-pillow basalts showing varying degrees of flattening of the pillow structures occur in the southernmost part of the mapped area along the Klein Letaba River (Fig. 2) where they are associated with the heterogeneous, relatively quartz rich variety of amphibolite. They are fine grained and the paragenesis hornblende, plagioclase ± biotite, quartz defines medium grade metamorphism, while relict magmatic pyroxene can be present in adjacent heterogeneous amphibolite. Banded iron formations occur in up to several metres thick units, particularly in the NW domain. Here they are associated with cherts that may include chemical sediments as well as silicified shales. The BIF consist of magnetite, quartz and actinolite, with individual magnetite-rich bands up to 10 mm in thickness. Actinolite is particularly developed in the magnetite free bands and constitutes 25% of the rock. No grunerite was seen, and its absence indicates that upper amphibolite facies conditions were not reached. However, no BIF from the immediate footwall of the HRSZ was studied. Clastic metasediments range from metapelites to metagraywackes, which can show clear lamination (Fig. 4a) and graded bedding. They are particularly abundant in the centralSE domain, where metapelites consist essentially of quartz, sericite, chlorite, and albite, and metagraywackes further contain minor K-feldspar, biotite, epidote and titanite. Metapelites mostly have a pronounced foliation, defined by the phyllosilicates and quartz stringers and commonly oblique to the bedding. Locally porphyroblasts of andalusite occur (Fig. 4a and b), which appear to predate the foliation and are generally fully sericitized. These are assumed to have resulted from contact metamorphism which could be associated with the gabbro sills. Rare metapelites in the NW domain close to the HRSZ (locality of sample HL 82, see Fig. 2)

Other than the primary bedding of the supracrustal rocks, also visible on the map scale as lithological contacts (Figs. 2 and 4a) there are two distinct generations of structures and textures present, reflecting deformation events here labelled D1 and D2 . Both are generally NE-SW trending, but that is where the similarity ends. In Fig. 5 the main structural features are shown. D1 structures are mainly visible in the central and SE domains (L1 , S1 ). In the NW domain they are mostly overprinted by the younger D2 event (L2 , S2 ). The S1 foliation is an amphibolite- and greenschist facies schistosity. It is defined by the alignment of amphiboles, biotite, quartz and feldspar ribbons, chlorite, white micas as well as flattened pillows. It dips moderately to steeply NW in the SE domain, but is generally vertical in the central domain along the Klein Letaba River. Large open to isoclinal folds are identified in the Central domain of the belt; their fold axes plunge mainly shallowly NE (some are horizontal or plunge SW) and S1 is folded by them. A mineral lineation L1 is broadly parallel to them. Similar to S1 , L1 is defined by elongated amphiboles, elongated quartz and feldspar aggregates, by elongated biotite crystals and by stretched pillows. Both S1 and L1 where thus formed during a metamorphic event (an exception is the local contact metamorphism mentioned above), which will be discussed further below. While S1 is present in the granitoids bordering the SE margin of the belt, L1 is absent. The structures defined as D2 are mainly found along the NW boundary of the Giyani GB, i.e., in the HRSZ. The schistosity S2 is broadly parallel to this boundary and dips subvertically. In the central and SE domains, this steep schistosity is only locally found, concentrated along specific shear zones. These, and the HRSZ itself, are seen to cut S1 as also noted by de Wit et al. (1992c) and most clearly apparent immediately E of Giyani. The foliation S2 and the prominent lineation L2 within it are defined by biotite, elongated quartz and feldspar rods, oriented amphibole, fuchsite and occasionally tourmaline. L2 plunges fairly uniformly ENE at 50◦ –70◦ . A large isoclinal fold in the NW domain (Fig. 2) appears to be associated with D2 , and small scale movement indicators show WSW-vergent thrusting, as also noted by McCourt and van Reenen (1992). Locally the coincidence of S1 and S2 /L2 , where S1 is not obliterated, has produced crenulation cleavage, particularly in talc-actinolite schists. The time-relationship of deformation and metamorphism in D2 is complex and not uniform. Mostly the retrograde paragenesis (actinolite, chlorite, or sericite) is syntectonic, but on the other hand garnet has been observed to overgrow an S2 shear zone foliation, and near the Klein Letaba Mine prograde sillimanite replacement of kyanite occurred posttectonically (McCourt and van Reenen, 1992). It thus appears that D2 shearing was diachronous and locally active in different zones throughout the period of metamorphism. While D2 can be firmly attributed to the HRSZ, the D1 structures present some problems. The steep foliations and folds (some isoclinal) with similar SW–NE strikes and fold axes suggest SE–NW compression. At first sight the steep structures are at odds with the shallowness of the central domain of the belt as indicated by geophysical data and by the spatial relationship of granitoid gneiss with supracrustals in the Klein Letaba River, which indicates a subhorizontal contact (Figs. 2 and 5). Particularly striking is a large anticline, defined by metasediments, mafic and ultramafic rocks, immediately above a shallow contact against gneissic basement in the Klein Letaba River. These features can be understood

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Fig. 5. Structural map of the Klein Letaba and Ntsami River sections of the Giyani greenstone belt with domains of regional metamorphism.

in the context of SE–NW compression of the supracrustals within a NW-vergent nappe, with the contact to the underlying gneiss representing the sole, or zone of décollement. In this scenario the supracrustals of the central-SE domain of the Giyani GB are

allochthonous, while the ultramafic + BIF series of the NW domain might be parautochthonous. The metamorphic grade of the supracrustals is variable (Fig. 5). If the formation of hornblende in the metagabbros is considered to

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be autometamorphic as suggested above, these rocks do not yield information on the regional metamorphism in the central domain. Along the Klein Letaba river section, the fabric forming hornblende in pillow basalts in the southernmost region, and garnet, biotite and plagioclase bearing metapelite in the northernmost outcrop of central domain supracrustals indicate amphibolite facies metamorphism. Along the Ntsami River, metasediments have greenschist facies parageneses. Towards the SW edge of the Belt, this increases to lower amphibolite facies as shown by fabric forming hornblende, biotite and a single occurrence of garnet. Towards the HRSZ in the NW, a “telescoped increase” (McCourt and van Reenen, 1992) in peak metamorphism up to upper amphibolite facies at pressures >4 kb and temperatures >500 ◦ C (sillimanite replacing kyanite, presence of staurolite and garnet) is seen, with a syn- to postkinematic retrogression to greenschist facies.

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2.4. Geochronology

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Age dating of garnet by the 207 Pb/206 Pb stepwise leaching procedure (Frei and Kamber, 1995; Frei et al., 1997) was carried out on two samples, HL 82 from the immediate footwall of the HRSZ in the Klein Letaba River, and HL 108 taken close to the SW edge of the BGB in the Ntsami River (See Fig. 2). Two further analyzed samples proved undatable by this method as they yielded unradiogenic lead only. 40 Ar/39 Ar dating of fabric forming amphiboles was done on four samples, ranging from the NW boundary of the Giyani GB with the HRSZ to amphibolite rafts in the gneiss bordering the Giyani GB to the SE (Fig. 2). The aim was to obtain time constraints on the actual growth of the metamorphic minerals, and thus of the fabric. All the geochronological work was carried out at the Institute of Geological Sciences of the University of Bern, Switzerland, using a VG Sector 5-collector thermal ionization mass spectrometer (TIMS) in static mode for the Pb isotope work, and an MAP 215–50 noble gas mass spectrometer with double-vacuum furnace for the 40 Ar/39 Ar dating. Methods used were as described in Kreissig et al. (2001), but the 40 K decay constant of Renne et al. (2010) was applied. Sample HL 82 is a metapelitic schist consisting of white mica, quartz, feldspar, staurolite, garnet and minor opaque phases. Garnet and staurolite are cogenetic and syntectonic. Both are coarse (up to 8 mm) and poikilitic, containing inclusions of quartz, feldspar and opaques. Microprobe analyses are given in Table 1 and show a high almandine component (>0.9) and variable MgO and MnO contents between cores and rims (in opposite senses in two grains) but otherwise unzoned. Sample HL 108 consists of hornblende, plagioclase, garnet and opaque phases. The garnet is much finer grained than in HL 82 (<1 mm), has an almandine component of 0.8 (Table 1), is unzoned and poikilitic with the same inclusions as in HL 82. Lead isotope data of the 6 leaching steps of the dated garnets are listed in Table 2 and plotted in Fig. 6A and B. Large spreads in 206 Pb/204 Pb ratios were obtained. For HL82, the 208 Pb/204 Pb vs 206 Pb/204 Pb does not reveal a significant contribution from zircon inclusions that might predate the garnet. Regression of the data of all steps yields an apparent age of 2726 ± 15 Ma (95% confidence limit) with an MSWD value of 9 and we consider this result to give the age of garnet growth. This is within error of the result obtained by Kreissig et al. (2001) on a staurolite from the same domain (2712 ± 37 Ma). For HL 108, two aliquots of the garnet separate were analyzed. Regression of all 10 results yields an age of 2833 ± 38 Ma (MSWD = 8). Low 208 Pb/204 Pb vs 206 Pb/204 Pb ratios of two HF leach data indicate low Th/U ratios, suggesting (submiscroscopic) inclusions of metamorphic zircon (Table 2). The most radiogenic fraction plots slightly above the main regression, and would yield a date of ca 2870 Ma if regressed with the least radiogenic Pb. The scatter of data indicated by the high uncertainty and MSWD value, indicates some disturbance of the system as ascribed

120

250

300

garnet HL 108 all data: 2833 ± 38 Ma MSWD = 8

B

100 80 60

steps 1A,2A,3A,4A,6A,1B,2B (omitting highly radiogenic fractions with large errors): 2829 ± 12 Ma MSWD = 1.4

207

Pb/204Pb

150 200 Pb/204Pb

40 20 0

0

100

200

300

400

500

600

Pb/204Pb

206

Fig. 6. Results of stepwise leaching Pb isotope analyses. Where error ellipses are not shown they are smaller than the symbols (see Table 3).

to interaction with fluids for similar cases in the polymetamorphic CZ of the Limpopo Belt (Kramers and Mouri, 2011). If the three most radiogenic fractions with large uncertainties are excluded, an age of 2829 ± 12 Ma (MSWD = 1.4) results. We prefer to use the allinclusive result of 2833 ± 38 Ma even if its uncertainty is greater. The difference in ages between HL 82 and HL 108 is clearly resolved. Metamorphism in the SE domain of the Giyani GB, associated with D1 structures, is thus clearly significantly older than that at the NW edge of the belt, which is related to the HRSZ and D2 . Microprobe analytical data for the dated amphiboles are given in Table 3. Sample HL80 from close to the HRSZ (Fig. 2) is a banded amphibolite. The dated amphibole is a magnesiohornblende with XMg = 0.65. The rock further contains actinolite, strongly sericitized plagioclase and epidote. HL 72 and 89 were taken from amphibolite rafts within the Klein Letaba Gneiss, whereas HL 105 is from a strongly sheared zone close to the SE Margin of the belt (Fig. 2). A strong lineation is defined by the hornblende, which in HL 72 and 105 is heterogeneous in composition, with XMg ranging from 0.56 to 0.66 (Table 3). In HL 89 anthophyllite is also present. Plagioclase with anorthite content between 0.28 and 0.38 is the second major phase. The Ar analytical results for the four samples are listed in Table 4, and apparent age and Ca/K ratio spectra are shown in Fig. 7. The latter, derived from 37 Ar/39 Ar ratios, are helpful in interpreting the results. Broadly the Ca/K ratios found are in agreement with the microprobe results listed in Table 3. The values are generally high for hornblendes, particularly in HL 89, and reflect the mafic chemistry of the rocks. All of the age spectra are complex. The low apparent ages for low temperature steps are quite common in amphibole and reflect alteration phases (often submicroscopic

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

103

Table 1 Microprobe analyses of garnets dated by Pb/Pb stepwise leaching (calculations based on 8 cations and 12 oxygens). Sample

HL 82-1 core

HL 82-1 rim

HL 82-2 core

HL 82-2 rim

HL 108-1 rim

HL 108-2 core

HL 108-2 rim

SiO2 TiO2 Cr2 O3 Al2 O3 Fe2 O3 FeO Mn2 O3 MgO NiO CaO Na2 O K2 O

36.48 0 0.08 21.20 0.04 40.01 0.14 1.58 0 0.35 0.04 0

36.58 0.04 0.06 21.16 0.31 39.33 2.00 0.99 0 0.49 0 0

36.69 0.01 0.15 20.96 0 39.67 1.70 0.98 0 0.42 0.03 0

36.96 0 0.19 21.22 0 40.49 0.20 1.56 0 0.35 0.03 0

37.34 0.02 0.02 21.32 0.33 35.80 0.27 2.16 0 3.62 0.02 0.00

37.29 0.09 0.12 21.28 0.15 35.01 0.74 2.17 0 3.85 0.02 0.01

37.24 0.09 0.16 21.10 0 35.29 0.33 2.00 0 3.99 0.01 0.01

Total

99.94

100.96

100.61

101.01

100.91

100.73

100.22

Si Ti Cr Al Fe3+ Fe2+ Mn3+ Mg Ni Ca Na K

2.979 0 0.005 2.041 0.003 2.733 0.010 0.193 0 0.031 0.006 0.001

2.973 0.002 0.004 2.028 0.019 2.673 0.138 0.120 0 0.043 0.001 0

2.992 0.001 0.01 2.015 0 2.706 0.118 0.119 0 0.036 0.004 0

2.989 0 0.012 2.023 0 2.739 0.014 0.189 0 0.030 0.005 0

2.985 0.001 0.001 2.009 0.020 2.394 0.018 0.258 0 0.310 0.003 0

2.984 0.006 0.008 2.007 0.009 2.344 0.050 0.258 0 0.330 0.003 0.001

2.998 0.006 0.010 2.002 0 2.376 0.022 0.240 0 0.344 0.001 0.001

Total

8.000

8.000

8.000

8.000

8.000

8.000

8.000

chlorite), in accord with the low Ca/K ratios in these steps. Cl/K ratios (obtained from 38 Ar/39 Ar ratios; Table 4) also show this alteration, being up to an order of magnitude higher in early steps than the main, very low value. Thus the low temperature step results have limited geochronological value. The HRSZ-adjacent sample HL 80 yielded no plateau, but a reasonably consistent age range was obtained for the higher temperature steps, which yields an integrated age of 2566 ± 20 Ma (an integrated age is obtained by adding up the gas from all included steps and then deriving an age from this). This sample was taken close to sample HL 82 which yielded a garnet Pb/Pb stepwise leaching date of 2726 ± 15 Ma, and thus, similar to the results of Kreissig et al. (2001) a hornblende 40 Ar/39 Ar age significantly younger than that of a peak metamorphic mineral from an adjacent outcrop was

found. It could be argued that this age discrepancy is a result of (very) slow cooling after the peak metamorphism. However, that scenario could not have produced the general conformity of the degassing spectra for Ca/K ratios and apparent ages. As discussed further below, this conformity more likely portrays a mixture of different amphibole phases that crystallized at different times, some much earlier than the “mix” age range, and some much later. The same close conformity of Ca/K ratios and apparent 40 Ar/39 Ar ages appears in the spectra for the other amphibole samples. It is particularly well expressed in sample HL 72, taken from a large sheared amphibolite xenolith in the gneiss terrain SE of the Giyani GB. In the medium and high temperature gas releases, this sample shows some quite young apparent ages, down to 2200 Ma, coupled with low Ca/K ratios. Sample HL 105, taken from the contiguous

Table 2 Lead isotope data of leaching steps from garnets of samples HL 82 and HL 108. Time

206

Garnet of sample HL 82 1.5 N HBr, 2 N HCl, 12:1 HL 82-1 4.4 N HBr HL 82-2 8.8 N HBr HL 82-3 HL 82-4 14.4 N HNO3 HF conc. HL 82-5 HF conc. HL 82-6

15 4h 12 h 24 h 48 h 65 h

18.02 18.10 34.82 68.07 168 254.2

± ± ± ± ± ±

0.05 0.14 0.11 0.51 26 3.1

15.33 15.61 18.73 24.77 44.6 59.89

± ± ± ± ± ±

0.04 0.12 0.06 0.19 7 0.72

35.83 37.91 72.69 104.1 141 372.8

± ± ± ± ± ±

0.06 0.30 0.24 0.8 22 4.5

0.986 0.994 0.992 0.999 0.999 1.000

Garnet of sample HL 108 1.5 N HBr, 2 N HCl, 12:1 HL 108 A1 4.4 N HBr HL 108 A2 8.8 N HBr HL 108 A3 14.4 N HNO3 HL 108 A4 HF conc. HL 108 A5 HF conc. HL 108 A6 1.5 N HBr, 2 N HCl, 12:1 HL 108 B1 HL 108 B2 4.4 N HBr 14.4 N HNO3 HL 108 B4 HF conc. HL 108 B6

15 4h 12 h 24 h 48 h 65 h 15 4h 24 h 48 h

34.98 175.9 37.44 87.79 213 31.60 148.9 113.7 180 541

± ± ± ± ± ± ± ± ± ±

0.35 1.5 0.33 1.13 83 0.11 3.8 1.6 39 57

18.85 47.34 19.50 30.02 53.4 18.43 42.35 34.66 47.96 125

± ± ± ± ± ± ± ± ± ±

0.2 0.42 0.17 0.39 20.8 0.06 1.11 0.48 10.4 13

54.51 228.4 55.31 75.29 62.7 40.01 193.9 141.8 95.3 110.9

± ± ± ± ± ± ± ± ± ±

0.56 2.0 0.49 0.97 24.4 0.14 4.9 2.0 10.7 11.7

0.977 0.999 0.995 0.996 0.999 0.986 0.958 0.998 1.000 1.000

Step #

a

Acid

Uncertainty limits are 1 standard deviation, absolute. Pb/204 Pb vs. 206 Pb/206 Pb error correlation.

b 207

Pb/204 Pba

207

Pb/204 ba

208

Pb/204 Pba

Rb

104

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

Table 3 Microprobe analyses of amphiboles dated by 40 Ar/39 Ar (calculations based on 13 cations and 23 oxygens). Sample Comment SiO2 TiO2 Cr2 O3 Al2 O3 Fe2 O3 FeO Mn2 O3 MgO NiO CaO Na2 O K2 O F Cl H2 O Total

HL 72-1 coarse-g

HL 72-2 coarse-g

HL 80-1 fine-g

HL 80-2 fine-g

HL 80-3 single-g

46.26 0.41 0.05 10.56 4.50 12.01 0.28 10.86 0 11.57 1.31 0.34 0.05 0.01 2.03

46.70 0.39 0.1 9.46 4.41 11.88 0.38 11.36 0 11.73 1.22 0.38 0.05 0.02 2.02

47.96 0.69 0.1 10.90 2.53 9.50 0.21 13.01 0 11.97 0.96 0.34 0.03 0.01 2.08

46.38 0.68 0.09 11.80 3.17 9.57 0.19 12.58 0 12.01 1.18 0.42 0 0.01 2.08

55.33 0.57 0 10.84 0 9.93 0.19 10.57 0 10.10 0.90 0.22 0 0.01 2.14

100.22

100.08

100.29

100.17

100.80

HL 80-4 single-g

HL 89-1 coarse-g

HL 89-2 coarse-g

HL 89-3 coarse-g

HL 89-4 coarse-g

HL105-1 single-g

HL105-2 single-g

54.17 0.48 0.27 7.38 0 9.74 0.26 13.96 0 10.76 0.65 0.19 0.02 0.02 2.12

48.88 0.44 0.12 9.02 5.91 7.19 0.15 13.86 0 11.33 1.15 0.14 0.03 0.03 2.08

48.98 0.43 0.08 9.17 6.38 6.80 0.20 13.85 0 11.27 1.11 0.13 0.04 0.01 2.09

55.17 0 0.03 0.51 24.34 0 0.59 18.63 0 0.76 0.03 0.02 0 0 2.36

56.38 0.03 0 4.74 20.40 0 0.46 15.36 0 1.30 1.69 0.02 0.12 0 2.26

47.22 0.26 0 9.35 2.99 12.67 0.39 11.28 0 11.99 0.81 0.53 0.26 0 1.92

51.84 0.17 0.08 4.60 2.56 11.5 0.27 14.03 0 12.24 0.49 0.21 0.31 0 1.93

99.99

100.32

100.55

102.45

102.79

99.66

100.32

Si Ti Cr Al Fe3+ Fe2+ Mn3+ Mg Ca Na K F Cl OH

6.764 0.046 0.006 1.821 0.459 1.468 0.034 2.366 1.812 0.372 0.063 0.021 0.002 1.978

6.843 0.043 0.012 1.634 0.486 1.455 0.047 2.481 1.841 0.354 0.017 0.021 0.006 1.973

6.864 0.075 0.012 1.839 0.273 1.137 0.026 2.775 1.836 0.266 0.062 0.011 0.004 1.985

6.687 0.074 0.011 2.005 0.344 1.154 0.023 2.703 1.856 0.331 0.078 0 0.002 1.998

7.754 0.060 0 1.790 0 1.164 0.023 2.209 1.516 0.246 0.039 0 0.002 1.998

7.607 0.051 0.030 1.215 0 1.144 0.031 2.921 1.620 0.178 0.034 0.007 0.004 1.989

6.970 0.047 0.014 1.515 0.634 0.858 0.018 2.945 1.731 0.317 0.026 0.015 0.007 1.978

6.961 0.046 0.009 1.536 0.682 0.809 0.024 2.934 1.716 0.304 0.023 0.019 0.003 1.978

7.006 0 0.003 0.077 2.325 0 0.064 3.526 0.103 0.008 0.003 0 0 2.000

7.283 0.003 0 0.722 1.983 0 0.051 2.958 0.181 0.424 0.003 0.050 0 1.950

6.924 0.029 0 1.619 0.330 1.558 0.049 2.473 1.889 0.230 0.099 0.121 0 1.879

7.471 0.018 0.009 0.781 0.277 1.398 0.033 3.014 1.890 0.137 0.039 0.124 0 1.858

Total

17.247

17.257

17.148

17.264

16.801

16.832

17.073

17.044

15.114

15.608

17.219

17.066

Giyani GB close to its SE margin in the Klein Letaba River (Fig. 2), is the only one of the four 40 Ar/39 Ar dated samples that has close to a plateau result, with an integrated age of 2654 ± 15 Ma, but high temperature steps yielding younger ages and lower Ca/K ratios. Finally, sample HL 89, from a large amphibolite raft in the Klein Letaba Gneiss SE of the belt, has some high gas release steps with old apparent 40 Ar/39 Ar ages between 2820 and 2970 Ma, but also here there are excursions with conforming variations in the Ca/K ratios. While it can be argued that unsupported (excess) 40 Ar is the cause of the old apparent age range, the conformism of Ca/K ratios can again be used as a counterargument. We conclude that this sample retained a ‘memory’ in the 2820–2970 Ma range, although it is not possible to narrow this down to a more precise age. The young age range indicated by samples HL 72 and HL 105 is surprising, as these are from a domain of amphibolite facies metamorphism, in which garnet (HL 108) yielded a robust if imprecise date of 2833 ± 38 Ma. Possible solutions for this and other apparent contradictions will be discussed below. 3. Attempt at a regional correlation of events Most of the existing geochronology along the HRSZ between longitudes 29◦ and 31◦ east, as well as in the adjoining section of the Kaapvaal Craton is summarized in Fig. 8. We first consider the supracrustal units of the Pietersburg and Rhenosterkoppies GB’s, then possible regional correlations in tectonism and metamorphism between these greenstone belts and along the HRSZ. 3.1. The Pietersburg and Rhenosterkoppies greenstone belts The Pietersburg greenstone belt is dominated by an association of ultramafic and mafic rocks, both intrusive and extrusive,

associated with banded iron formation (BIF), termed the “Simatic Basement” by de Wit et al. (1992a). In the western part of the Pietersburg GB this is unconformably overlain by the clastic sedimentary Uitkyk Formation (de Wit et al., 1992a). A deformational event (D1 ) in the simatic basement prior to the deposition of the Uitkyk formation is demonstrated by strong angular unconformities and the presence of folded BIF clasts (from the simatic basement) in the latter. Mafic and ultramafic rocks in the simatic basement show ocean-floor (auto-) metamorphism equivalent to greenschist- to lower amphibolite facies (de Wit et al., 1992a; van Schalkwyk et al., 1993), which appears pretectonic and similar in character to that described above for mafic rocks in the Giyani GB. N- to NNW vergent thrusting constituted a second deformational event (D2 ), affecting both the simatic basement and the overlying Uitkyk formation. D2 is however thought to be long-lasting, with the clastic sediments of the Uitkyk Formation being a direct result of the uplift and erosion it caused in its early stages (de Wit et al., 1992a). Therefore D1 could merely represent an early phase of D2 . The metamorphism M2 associated with D2 reached only greenschist facies in the western Pietersburg GB, increasing eastwards to amphibolite facies (de Wit et al., 1992a). The Rhenosterkoppies greenstone belt is made up of interlayered mafic (amphibolitic), minor ultramafic (actinolite-chlorite schists) and BIF units (Passeraub et al., 1999) and therefore reminiscent of the simatic basement of the Pietersburg GB, although it is assigned a separate formation name, Zandrivierspoort Formation, by SACS (1980). As in the Pietersburg and Giyani GB’s, clinopyroxene relics occur in some amphibolites, which are thus considered metagabbros, and their retrogression probably occurred in ocean floor metamorphism. The belt is an anticline, consisting of a southern, S-dipping limb, a mostly flat-lying central area and a NE-dipping northern limb, which are contiguous (Passeraub et al., 1999). The northern limb forms the footwall of the HRSZ and its

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

105

Table 4 Ar degassing results for dated amphiboles. Temp. ◦ C

39

Ar ml

% 39 Ar

Age Ma

Sample HL72 777 931 972 993 1007 1027 1046 1075 1088 1100 1113 1122 1154 1167 1174 1193 1272 1359 1396

8.04E−11 6.96E−11 1.99E−10 3.82E−11 1.53E−11 1.38E−11 1.39E−11 1.33E−11 1.35E−11 1.46E−11 6.92E−11 2.32E−11 2.12E−11 1.18E−11 1.02E−11 4.69E−11 4.74E−12 9.39E−12 3.92E−12

12.62 10.92 31.22 5.99 2.40 2.17 2.18 2.09 2.12 2.29 10.87 3.64 3.33 1.85 1.60 0.74 0.90 1.47 0.62

1674 2048 2593 2345 2249 2261 2345 2447 2519 2539 2540 2528 2489 2458 2361 2715 2646 2602 2765

4 5 2 7 17 23 17 22 25 24 3 10 11 21 28 48 43 25 51

Sample HL80 783 935 954 970 993 1010 1032 1048 1068 1082 1103 1114 1136 1152 1164 1197 1234 1274 1356 1403

4.80E−11 6.30E−11 4.21E−11 5.52E−11 5.57E−11 2.42E−11 2.48E−11 3.19E−11 3.63E−11 2.89E−11 2.82E−11 5.07E−11 6.62E−11 4.16E−11 4.22E−11 1.25E−11 2.10E−11 1.06E−11 3.12E−11 7.04E−12

6.66 8.73 5.84 7.65 7.72 3.36 3.44 4.42 5.03 4.01 3.91 7.03 9.16 5.77 5.85 1.73 2.92 1.48 4.33 0.98

1855 2091 2317 2593 2613 2393 2512 2542 2496 2549 2627 2546 2590 2583 2577 2517 2664 2461 2611 2267

Sample HL89 778 933 975 994 1012 1029 1051 1067 1084 1105 1124 1143 1156 1174 1195 1234 1276 1354 1436

8.53E−12 8.51E−12 5.98E−11 7.84E−11 7.20E−12 2.03E−12 3.95E−12 4.23E−12 5.59E−12 8.38E−12 2.15E−11 3.98E−12 4.13E−12 3.17E−12 4.68E−12 5.11E−12 6.26E−12 2.49E−12 2.67E−12

3.24 3.24 22.74 29.80 2.74 0.77 1.50 1.61 2.13 3.19 8.19 1.51 1.57 1.21 1.78 1.94 2.38 9.45 1.02

Sample HL105 777 931 972 993 1007 1027 1046 1075 1088 1104 1113 1124

6.26E−11 1.63E−10 6.79E−11 3.23E−11 2.57E−11 2.41E−11 2.43E−11 1.48E−11 1.55E−11 1.07E−11 9.83E−12 1.21E−11

9.05 23.58 9.83 4.68 3.72 3.49 3.51 21.36 2.25 1.53 1.42 1.75

1 sigma abs.

Ca/K

1 sigma abs.

Cl/K

1 sigma abs.

3.0 19.5 29.1 23.2 23.2 25.8 26.2 29.5 30.1 31.8 30.1 28.8 28.2 26.2 25.5 30.8 25.7 30.8 32.6

0.1 1.0 1.5 1.2 1.2 1.3 1.3 1.5 1.5 1.6 1.5 1.4 1.4 1.3 1.3 1.5 1.3 1.5 1.6

0.031 0.015 0.008 0.004 0.029 0.007 0.023 0.011 0.008 0.002 0.014 0.017 0.018 0 0 0.001 0 0 0

0.0016 0.0008 0.0004 0.0002 0.0015 0.0004 0.0012 0.0006 0.0004 0.0001 0.0007 0.0009 0.0009 0 0 0.0001 0 0 0

5 4 6 4 5 11 10 6 6 7 9 5 3 7 7 19 13 26 7 40

16.2 24.6 27.0 28.9 27.4 26.9 30.5 30.6 28.8 30.3 31.3 28.6 30.0 29.7 28.7 26.1 31.7 25.4 28.8 22.5

0.8 1.2 1.4 1.4 1.4 1.3 1.5 1.5 1.4 1.5 1.6 1.4 1.5 1.5 1.4 1.3 1.6 1.3 1.4 1.1

0.068 0.01 0.007 0.003 0.008 0.005 0.008 0.012 0.005 0.007 0.009 0.008 0.005 0.007 0.004 0.002 0 0.016 0.003 0.026

0.0034 0.0005 0.0004 0.0002 0.0004 0.0003 0.0004 0.0006 0.0003 0.0004 0.0005 0.0004 0.0003 0.0004 0.0002 0.0001 0.0000 0.0008 0.0002 0.0013

2832 2308 2964 2843 2739 3147 2407 2554 2640 2646 2695 2636 2714 3012 2303 2673 2657 2918 1443

42 32 4 4 28 110 60 51 35 23 9 64 54 74 47 47 30 8 80

13.5 58.5 79.2 76.0 70.1 102.1 61.5 62.1 63.6 65.4 68.3 64.9 65.9 75.0 49.7 63.6 62.5 73.5 22.1

0.7 2.9 4.0 3.8 3.5 5.1 3.1 3.1 3.2 3.3 3.4 3.2 3.3 3.8 2.5 3.2 3.1 3.7 1.1

0.249 0.009 0.019 0.017 0 0.008 0.03 0.059 0.044 0.018 0.015 0.032 0.027 0.035 0.021 0 0 0.009 0.144

0.0125 0.0005 0.0010 0.0009 0.0000 0.0004 0.0015 0.0030 0.0022 0.0009 0.0008 0.0016 0.0014 0.0018 0.0011 0.0000 0.0000 0.0005 0.0072

489 2631 2723 2601 2542 2669 2670 2674 2547 2722 2607 2588

6 2 4 9 8 10 17 2 21 24 22 26

2.7 27.8 24.5 24.3 24.1 26.1 26.3 25.9 22.4 26.3 23.8 22.3

0.1 1.4 1.2 1.2 1.2 1.3 1.3 1.3 1.1 1.3 1.2 1.1

0.24 0.003 0.003 0.011 0.009 0.023 0.005 0.004 0.009 0.002 0.027 0.015

0.0120 0.0002 0.0002 0.0006 0.0005 0.0012 0.0003 0.0002 0.0005 0.0001 0.0014 0.0008

106

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Table 4 (Continued) Temp. ◦ C

39

Ar ml

% 39 Ar

1156 1164 1173 1197 1234 1276 1351 1396

1.22E−11 1.09E−11 9.38E−12 8.56E−12 1.21E−11 8.41E−12 2.37E−11 1.03E−11

1.76 1.58 1.36 1.24 1.76 1.22 3.47 1.49

Age Ma 2491 2485 2200 2314 2549 2343 2676 2339

1 sigma abs. 16 23 23 21 24 40 13 25

western continuation is found intermittently over 40 km as a line of hills consisting of mainly steeply dipping BIF and mafic rocks along the HRSZ. Similarly to the N edge of the Rhenosterkoppies GB, a S-N section trough the HRSZ at the type locality (Hout River, 20 km W of the Rhenosterkoppies GB) shows a transition from mostly subhorizontal foliations and SW-SE trending lineations in the craton, to steep foliations and downdip lineations (with movement indicators showing a thrust sense) towards the SMZ (Smit et al., 1992). The dominant structural features in the Rhenosterkoppies GB are a foliation F2 , and a lineation L2 within it, assigned to a deformation D2 (Passeraub et al., 1999). F2 broadly follows the layering as described above; L2 trends SW-NE and is subhorizontal in the central part, while dipping NE at 30–60◦ in the northern arm. Preceding D2 is a deformation D1 , expressed as small-scale folding in the BIF. Finally, Passeraub et al. (1999) define D3 as a flexure which created the anticlinal form, folding both F2 and L2 . Peak metamorphism in the Rhenosterkoppies GB is defined at upper amphibolite facies (550–650 ◦ C and >4 kb) by sapphirine, cogenetic garnet + staurolite in metapelites, and garnet + kyanite in quartzose rocks, and persists throughout the central area and northern arm (Miyano et al., 1992; Passeraub et al., 1999; the southern arm was not investigated). These parageneses all precede the D2 fabric, which is defined by chlorite, sericite, zoisite and quartz ribbons, showing retrogression to greenschist facies during D2 . 3.2. Lithological and age comparisons between Giyani and Pietersburg greenstone belts

Ca/K 21.7 21.0 18.5 18.9 22.3 20.4 23.9 20.4

1 sigma abs.

Cl/K

1 sigma abs.

1.1 1.0 0.9 0.9 1.1 1.0 1.2 1.0

0.032 0.021 0.001 0.042 0.004 0.002 0.001 0.026

0.0016 0.0011 0.0001 0.0021 0.0002 0.0001 0.0001 0.0013

The central-SE domain of the Giyani GB has certain lithological similarities with the simatic basement of the Pietersburg GB: metagabbros with ocean floor metamorphism and banded iron formations (although these are less common in the Giyani GB). The clastic metasediments of the Giyani GB cannot be equivalent to the Uitkyk Formation, since they are intruded by metagabbro sills as shown by the contact metamorphism, while the Uitkyk formation rests unconformably on the unit containing metagabbros. There are no clastic metasediments in the simatic basement of the Pietersburg GB. Thus there is scant basis for proposing that the central-SE unit of the Giyani GB and the simatic basement of the Pietersburg GB were deposited at the same time and/or in the same basin. However, the inferred (NW) vergence of thrusting in the SE domain of the Giyani GB is similar to that found in the Pietersburg GB, and in view of the identical ages of pre- to syntectonic intrusives in both units and the marginally coeval metamorphism in the SE Giyani GB (Fig. 8, see above), we propose, as a working hypothesis, that D1 /M1 of the central-SE domain in the Giyani GB and D1 /M1 + D2 /M2 of the Pietersburg GB are expressions of the same orogenic event, which occurred between 2850 and 2870 Ma and led to the deposition of the Uitkyk Formation (which was subsequently affected by its later stages) in the Pietersburg GB. Following this proposal, a terrane collisional suture along the Pietersburg GB (de Wit et al., 1992a, 1993) continues ENE into the Lwaji arm and SE edge of the Giyani GB (Fig. 9). Below we refer to this as the PGB-Lwaji orogeny and suture. 3.3. Chronology of the Hout River Shear Zone

Two dates that can be regarded as giving depositional ages are the 3203 ± 2 and 2949 ± 2 Ma ages of zircons from a meta-andesite in the NW ultramafic-mafic-BIF domain of the Giyani GB and a quartz porphyry in the simatic basement of the Pietersburg GB, respectively (Fig. 8; Kröner et al., 2000). In both cases multiple grains yield consistent results and the zircons are small and not rounded, thus do not appear to be xenocrysts. The dates clearly preclude a correlation between these two units of the Giyani and Pietersburg GB’s, respectively. However, as discussed above, the NW Giyani GB domain is regarded by us as being probably a separate tectonic unit to the central-SE domain of the belt. Conclusions regarding the NW domain therefore cannot necessarily be extended to the central-SE domain. The youngest concordant detrital zircons found within the Uitkyk Formation of the Pietersburg GB are 2854 ± 9 and 2879 ± 9 Ma old (Fig. 8; Zeh and Gerdes, 2012), providing an oldest age limit for this Formation. For the mafic and clastic sedimentary central-SE domain of the Giyani GB, there is no such upper limit. However, youngest age limits for both, the Uitkyk Formation and the central-SE domain supracrustals of the Giyani GB, are given by the zircon dates on three pre- to syntectonic felsic intrusive rocks (Fig. 8) which, for the Giyani and Pietersburg GB’s, have ages that are identical, with error limits overlapping at 2871 Ma. This age is also just within the error limits of the disturbed Pb/Pb step-leach garnet age of sample HL 108 in the Giyani GB, and close to the youngest detrital zircon ages for the Uitkyk Formation listed above, confirming the syntectonic character of that formation.

Along the NW edge of the Giyani GB, 3 ages for peak metamorphism and pre-retrogression tectonism, determined by Pb/Pb step-leaching of pre-to syntectonic garnet and staurolite (Fig. 8; this work and Kreissig et al., 2001), agree within error limits and yield a weighted mean of 2713 ± 49 Ma (95% confidence). They also agree with a Pb/Pb step-leaching age of 2729 ± 19 Ma on fabricaligned titanite from an amphibolite in the footwall of the HRSZ in the Rhenosterkoppies GB (Fig. 8; Passeraub et al., 1999); the weighted mean for these four HRSZ results is 2717 ± 28 Ma (95% confidence). Thus an age of 2717 ± 28 Ma for the peak metamorphism in the immediate footwall of the HRSZ is reasonably well established. Since this metamorphism is seen as a hot-iron effect of the overthrust nappe (Roering et al., 1992a,b; Perchuk et al., 1996; van Reenen et al., 2011), this time of metamorphism is a minimum age for the thrusting itself. In a basement window within the Rhenosterkoppies GB, at 2 km distance from the HRSZ, a banded tonalitic biotite augengneiss is intruded by a finer grained granodiorite which also carries a foliation. Syn- to posttectonic titanite from the augengneiss yielded a Pb/Pb step-leaching age of 2754 ± 4 Ma, while (retrograde) epidote plus titanite from the granodiorite gave 2743 ± 21 Ma (Fig. 8; Passeraub et al., 1999). Both ages are slightly older than the peak metamorphism in the HRSZ discussed above, the first one significantly so. The U–Pb systematics in titanite are retentive at high temperature, but the mineral readily recrystallizes under stress, even at low grade metamorphic conditions. Our interpretation is

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36

40

32

Ca/K

30

26 20

24 20

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107

3200 3000 2800 2600 2400 2200 2000 1800 1600 1400 1200 1000 800 600 400

16 0

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1

HL 72: no plateau 2 largest release steps: 2593 ± 4 Ma (31%) and 2540 ± 6 Ma (11%)

0

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0.8

1

12 3200 3000 2800 2600 2400 2200 2000 1800 1600 1400 1200 1000 800 600 400

0

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HL 80: no plateau. Integrated age over arrowed range (78% of 39Ar): 2566 ± 20 Ma

0

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cumulative 39Ar fraction

0.4

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0.8

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cumulative 39Ar fraction

120 30

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100 80

20

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HL 89: no plateau. Largest gas release steps: 2964 ± 8 Ma (23%) 2843 ± 8 Ma (30%) 2695 ± 18 Ma (8%) 2918 ± 16 Ma (9%)

0

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0.8

cumulative 39Ar fraction Fig. 7.

40

1

0 3200 3000 2800 2600 2400 2200 2000 1800 1600 1400 1200 1000 800 600 400

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HL 105, almost a plateau. Integrated age over arrowed range (70% of 39Ar): 2654 ± 15 Ma

0

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cumulative 39Ar fraction

Ar/39 Ar age spectra and associated Ca/K ratio data, obtained via 37 Ar/39 Ar ratios. Boxes show 1 SE uncertainty limits.

that these dates may record early deformation in the vicinity of the HRSZ. Dates younger than these pre- to synmetamorphic ages are also encountered in the HRSZ, yielded by Pb/Pb stepwise leaching date on kyanite, and four Ar/Ar dates on hornblendes (Fig. 8). The kyanites, dated at 2672 ± 51 (Kreissig et al., 2001) and 2634 ± 49 Ma (Passeraub et al., 1999; the large errors are due to a low spread in Pb isotope ratios) in the Giyani and Rhenosterkoppies sections of the

HRSZ, respectively, occur in quartz-fuchsite-kyanite-sillimanitetourmaline schists within the HRSZ. The above authors argued that these are metasomatized zones of gneiss, from which fluids leached Ca and Na, transforming feldspars into alumosilicates, while introducing B. Kyanite is statically grown. The hornblendes dated by Ar/Ar are all fabric forming. Those dated by Kreissig et al. (2001) along the HRSZ are ordinary hornblendes with Ca/K ratios between 12 and 20, while the one analyzed

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

Matlala Moletsi

(11) (8) (6) (7) (6) (5) (2) (11) (2)

Late Archean undeformed granitoids

Uitloop granite, PGB Matok pluton Bandelierkop Pegmatites intruding Hout River shear zone

(4)

Fabric forming amphiboles along Hout River shear zone

(2) (1)

zircon, metamorphic, Groot Letaba gn. monazite, Bandelierkop Quarry, SMZ

(11) (2)

garnet, staurolite, kyanite and titanite in footwall of Hout River shear zone, Giyani and Rhenosterkoppies GB garnet, SE edge Giyani GB (xenolith) metamorph. zirc. Kl. Letaba (11) and Hout River gneisses porphyry dyke,Giyani GB (9) foliated granites, Pietersburg GB (8) xenocrysts

(3) (shear zone)

(1)

amphibolites in Klein Letaba Gneiss SE Giyani GB

meta-andesite Giyani GB (9) (8)

quartz porphyry,Ysterberg Fm., Pietersburg GB detrital zircons, Uitkyk Formation Pietersburg GB

Terrane north of Pietersburg & Giyani GB

Terrane south of Pietersburg & Giyani GB

(7) (6) (11)

(10)

Turfloop granite intrusion age intrusion ages of younger units within Groot Letaba gneiss

(6) xenocrysts and older units in Groot Letaba gneiss (11) (6)

(9)

Constraints on Constraints on ages of deformation sedimentation ages and metamorphism

108

(11) (9)

younger Hout River gneiss and granulites 20 km N Giyani xenocrysts older Hout River gn. (11) xenocrysts (6) Meriri granite & xenocrysts (9) Gneis 20 km N Giyani

2400 2500 2600 2700 2800 2900 3000 3100

3200 3300 3400

Age (Ma) Fig. 8. Compilation of geochronological data from greenstone belts south of the Hout River Shear Zone, the HRSZ itself and adjacent areas. Black: U–Pb zircon and monazite. Dark grey: 207 Pb/206 Pb stepwise leaching dates on metamorphic minerals. Middle grey: 40 Ar/39 Ar on amphiboles. Light grey: Rb–Sr on pegmatitic muscovites. Sources: (1) this work; (2) Kreissig et al. (2001); (3) Passeraub et al. (1999); (4) Barton and van Reenen (1992); (5) Barton et al. (1992); (6) Zeh et al. (2009); (7) Henderson et al. (2000); (8) de Wit et al. (1993); (9) Kröner et al. (2000); (10) Zeh and Gerdes (2012); (11) Laurent et al. (2013).

by us (HL 80) from the Klein Letaba River in the footwall of the HRSZ is a magnesiohornblende (Table 3, Fig. 7) with a lower K content as seen by the Ca/K ratios between 25 and 30. None of the HRSZ amphiboles yield plateaus, and all apparent ages obtained, ranging from about 2550 to 2640 Ma (Fig. 8), are significantly younger than the results from Pb/Pb stepwise leaching in the HRSZ, except for the young kyanite ages discussed above. The age difference is not an artefact of inconsistencies of the 40 K decay constant, as the results were recalculated using the value of Renne et al. (2010), which

yields Ar/Ar ages consistent with U–Pb dates in undisturbed systems. While early heating steps yielding young apparent ages can be discounted as resulting from alteration (as shown by high Cl/K ratios), age variations in the higher temperature steps cannot be explained in this way. Detailed microprobe work by Kreissig et al. (2001) revealed strong heterogeneity within and between amphibole grains of the dated populations, with Ca/Al and Ca/K ratios varying by up to 50% within a sample. Using plots of Ca/Al, Cl/K and age vs. Ca/K., these authors could show that their results reflected

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

109

Fig. 9. Overview tectonic map showing the major structures and their apparent ages. MGB: Murchison greenstone belt; PGB: Polokwane (Pietersburg) greenstone belt; GGB: Giyani greenstone belt; RGB: Rhenosterkoppies greenstone belt. Murchison Suture and its age from Zeh et al. (2013).

mixtures of different generations of hornblende. This is similar to the heterogeneities seen in samples analyzed in this work (Table 3) where step ages correlate with Ca/K (Fig. 7). We agree with the conclusion of Kreissig et al. (2001) that the young and variable amphibole age patterns reflect mixtures of heterogeneous, distinctive populations rather than slow cooling. This is in accord with the finding that the rate of diffusion of large-ion (or atom) incompatible elements in minerals is generally lower than that of the major elements (Villa, 2010). Numerous 40 Ar/39 Ar amphibole dates with accompanying microprobe data, recently reported from samples within the SMZ itself (Belyanin et al., 2014) support this view and yield a cluster of ages in the 2650–2670 Ma range. For the HRSZ, these 40 Ar/39 Ar age spectra, along with the young kyanite ages, thus indicate that the zone remained a conduit for fluid flow under amphibolite facies conditions for >100 Ma after peak metamorphism, and may also have been locally reactivated during this period. Rb–Sr dates between 2630 and 2680 Ma on pegmatitic muscovites that crosscut the HRSZ (Barton and van Reenen, 1992; Fig. 8) yield minimum ages for the local shearing, except for one, aged 2661 ± 14 Ma, which is brecciated and thus documents relatively late movement on the shear zone. When comparing the chronology of events along the HRSZ with that of metamorphism in the SMZ proper (as far as established), it appears at first sight contradictory that the peak of hot-iron metamorphism in the HRSZ footwall is older than the monazite age for the metamorphism in the granulite zone (2679 ± 7 Ma, Kreissig et al., 2001; Fig. 8). However, monazite participates in metamorphic exchange reactions during both prograde and retrograde stages

(Berger et al., 2005; Kelly et al., 2006; Harlov et al., 2011) and can also be reacted out and regrow: Pyle and Spear (2003) could show three stages of monazite growth interspersed with two stages of monazite resorption, through the prograde and retrograde paths within a single high-temperature, medium-pressure metamorphic loop. Thus in spite of the retentivity of monazite for U–Pb systematics being extremely high (Cherniak et al., 2004), monazite U–Pb dates can record retrograde reactions. The zircon U–Pb ages of 2667 ± 4 and 2679 ± 7 Ma obtained for the intrusion of the mainly posttectonic Matok Pluton into the SMZ (Barton et al., 1992; Zeh et al., 2009) do not prohibit the notion that the 2679 ± 7 Ma monazite age of Kreissig et al. (2001) could be a time marker on the retrograde path of SMZ metamorphism. Of three hornblendes analyzed in this study that are from samples near the SE edge of the Giyani GB (where Pb/Pb stepwise leaching on garnet yielded a date of 2833 ± 39 Ma for peak metamorphism, see above), one (HL 105) yields a near-plateau age of 2654 ± 15 Ma, one (HL 72) shows a disturbed spectrum below 2600 Ma with some mid-temperature steps ranging from apparently 2200 to 2500 Ma (Figs. 7 and 8). Meanwhile the apparent step age range up to 3000 Ma of the third sample (HL 89) serves to demonstrate again that the young ages of the other dated amphiboles from the same region are not cooling ages. Samples HL 105 and 72 were specifically taken because of the presence of aligned amphiboles, and thus represent a sample bias towards shear zones. It now appears that late reactivation and fluid activity along shear zones is not confined to the HRSZ itself, but extends at least 15 km into the Craton margin.

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4. The PGB-Lwaji suture and the HRSZ in perspective 4.1. Can there be a connection between the PGB-Lwaji and HRSZ events? A broad synthesis of structural data and related geochronology of region of the Giyani, Pietersburg and Rhenosterkoppies GB’s (Fig. 9) shows the proposed PGB-Lwaji suture as well as the frontal and lateral ramps of the HRSZ, areas that were overridden by the SMZ nappe, and the region where this nappe is currently exposed as a horizontal sheet. The relatively high pressure metamorphism documented in areas behind frontal ramps (Khavagari arm and Rhenosterkoppies GB) either indicates a great thickness (about 15 km) of the overriding nappe, or an upwarp of the ramp region (doming) following the thrust event. Doming was suggested for the Rhenosterkoppies GB, as discussed above (Passeraub et al., 1999). A scenario in which the high grade metamorphism of the SMZ resulted from crustal thickening related to N-vergent structures such as the PGB-Lwaji suture has been proposed by Smit et al. (2014). Since the HRSZ is not considered to be a terrane boundary, but rather a metamorphic step in the Craton (Kreissig et al., 2000; Zeh et al., 2009; Kramers and Zeh, 2011), it is not easy to fit this phenomenon into common models of orogenic belts. In contrast, the proposed Pietersburg GB-Lwaji suture can probably be regarded as the product of a terrane collision following subduction of the northern plate underneath the southern one. For the Pietersburg GB this has been argued by de Wit et al. (1992a), and it is in accord with Hf isotope data from the Kaapvaal Craton south of the suture (tonalitic and granodioritic Groot Letaba Gneiss and Turfloop Granite, Zeh et al., 2009) which show a mixture of older crustal and juvenile mantle derived components between 2750 and 2850 Ma, indicating arc to postcollisional magmatism in the foreland (Kramers and Zeh, 2011). The age constraints on the PGB-Lwaji suture preclude this structure from being related to the Limpopo Orogeny as reviewed by Roering et al. (1992a,b) and indeed, the great age difference between the tectonism and metamorphism along the suture (ca. 2870 Ma) and the HRSZ (ca. 2710 Ma and younger) at first sight appears to rule out any connection between the two. Nevertheless we have explored the possibility of such a link by first-order thermal modelling.

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4.2. Heating rates for the SMZ

5 10

A fundamental question of the metamorphism in the SMZ is how the crust could have become so hot. The U, Th and K contents of both the SMZ and the adjoining Kaapvaal Craton are unusually low (U: 0.1–2 ppm; Th: 0.1–6 ppm; K2 O 0.2–4%, averaging 2%), resulting in average crustal heat production rates of 0.8 and 1.3 ␮W/m3 , respectively, calculated back to 2800 Ma (Kramers et al., 2011; see also Andreoli et al., 2011). With these values for the lower and upper half of the crust, respectively, and a conservative estimate of the basal heat flux (20 mW/m2 ), a steady-state geotherm going through the peak metamorphic conditions observed in the SMZ (∼900 ◦ C at 9–10 kb) ultimately results if the crust is thickened to 54 km (Kramers et al., 2011). In Fig. 10A the thermal evolution of such a thickened crust is illustrated for the parameters given in the caption. It can be seen that granulite facies conditions are reached in the lowermost crust after about 30 Ma, and 60 Ma are required for high grade metamorphism to occur in most of the lower crust. Invoking a large increase in basal crustal heat flux at the time of crustal thickening would shorten the time scale (e.g., Ashwal et al., 1992), but there is no evidence of any mantle-derived magmatism between 2900 and 2700 Ma in the SMZ, neither from Nd nor from Hf isotope

C

40 0

200

400

T (°C)

600

20 80

800

1000

Fig. 10. Development of geotherms following (A) crustal thickening, (B) and (C) subsequent overthrust of SMZ over KC, to illustrate rates of thermal adjustment. Calculated using a finite difference routine; parameters for all cases: Heat capacity Cp = 1.15 J/◦ Cg, heat production H = 1.3 ␮W/m3 , density 2.8 g/cm3 , basal heat flow into the crust constant at 20 mW/m2 , heat conductivity = 2.5 W/◦ Cm, crustal thickness covered by frames. In (A) thickening of a 35 km thick Kaapvaal Craton-type crust to 54 km is envisaged. “Initial”: conditions immediately after crustal thickening, simplified to two isothermal layers having the average temperatures of the upper and lower half of the steady state geotherm before thickening. Thermal adjustment over time is shown by grey geotherms marked by time (Ma) after the thickening event. (B) and (C) Illustrates development after burial of previously low-grade upper crust below a nappe stack of hot lower crust (“crustal overturn”, an extreme scenario for the Hout River Shear Zone overthrust). “Initial”: temperature profiles immediately after the event, The total crustal thickness is reduced to 44 km (see text). Grey geotherms marked with time (Ma) after lower crustal overthrust show the thermal adjustment.

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113

studies (Kreissig et al., 2001; Zeh et al., 2009; Kramers and Zeh, 2011), which renders a very high basal heat flux unlikely. With values for the basal heat flux significantly lower than 20 mW/m2 lower crustal heating is slower, but then high grade metamorphic conditions are not reached at all. Therefore the 30–60 Ma time interval appears fairly robust. This is much shorter than the age difference between the tectonometamorphism along the PGB-Lwaji suture (2850–2870 Ma) and the apparent peak of hot-iron metamorphism along the HRSZ (2717 ± 28 Ma). Thus there cannot be a direct causal link between the proposed PGB-Lwaji orogeny and the high grade metamorphism in the SMZ. In an alternative scenario, Smit et al. (2014) suggest crustal thickening at a time based on titanite Pb/Pb ages in the Rhenosterkoppies GB. In this case (considering ages of 2754 ± 4 Ma and 2743 ± 21 Ma, Passeraub et al., 1999; see Section 3.3) the time lag is comparable to the constraints from thermal modelling, and this scenario thus merits further investigation. A question that arises in this context is why an overthickened crust, with the resulting topographic relief and isostatic anomalies, could have persisted for the length of time needed to increase lower crustal temperature to high metamorphic grade, without erosion or, more likely, collapse through lower crustal flow. From about 550 ◦ C upwards, the viscosity of granitoids and many metasedimentary rocks dramatically decreases (e.g., Whitney et al., 2004), and this temperature is reached in the lower crust many millions of years before granulite facies conditions (Fig. 10A). Yet in the SMZ, because of the low radioactive heat production, overthickening of the crust is necessary for these high temperatures to be reached. While we are unable to solve this problem, the question of why the SMZ uplift happened could be answered via a related mechanism. In the scenario of a relatively stable thickened crust section that heats up slowly, density inversions are likely to be initiated due to heating. Gerya et al. (2004) have shown that density reductions of the order of 0.1 g/cm3 occur at mid- to lower crustal level in both granodiorites and metapelites, as rocks are heated through the 750–800 ◦ C range, while viscosity is reduced at the same time. It would thus be predicted that the SMZ crust would become more isostatically unstable as peak metamorphic conditions are approached, and uplift could occur along high temperature shear zones such as the HRSZ. The puzzling apparent doming in the Rhenosterkoppies GB (Passeraub et al., 1999) could also be explained as a consequence of such a density decrease. While this scenario has elements of published gravitational redistribution models for the SMZ (e.g., Perchuk and Gerya, 2011), the tectonic framework and mode of heating envisaged here are both different.

5. Post-peak metamorphism and the question of fluid sources The hornblende Ar/Ar ages around 2.6 Ga and even younger, found along the HRSZ by Kreissig et al. (2001) and within the Giyani GB in this study, have been interpreted above as resulting from fluid flow focussed along shear zones, that led to the crystallization of new generations of amphiboles, ultimately producing heterogeneous populations with mixed ages. Clearly no fluids could be produced in the SMZ crust itself after peak metamorphism and uplift along the HRSZ. Apart from the dehydrating effect of the metamorphism, the thinning of the crust following uplift should have stopped further heating of the granulites. This means that the only possible crustal source of fluids in regions around the HRSZ is located in portions of the northern Kaapvaal Craton that were at low metamorphic grade prior to the HRSZ event, and were overridden by the SMZ nappe or buried to mid/lower crustal levels in some other way during the event.

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The Giyani GB in areas that were not overridden by the HRSZ nappe (this work) as well as the Pietersburg GB (de Wit et al., 1992a,b) provide good examples of supracrustal units that could be similar to such fluid sources. In places where they form N vergent nappes at some distance away from the main PGB-Lwaji suture, their metamorphic grade is greenschist facies with evidence for ocean floor metamorphism surviving in each greenstone belt. In the Rhenosterkoppies GB and areas of the Giyani GB that were once covered by the SMZ nappe (Figs. 5 and 9), the hot-iron metamorphism and shearing of the HRSZ is superimposed on this low-grade metamorphism and mostly obliterated its features. From the regional comparison it is clear that the supracrustal units of the three belts, metasediments as well as mafic and ultramafic rocks, were hydrous before they were overridden by the SMZ nappe and ramps, and that prograde hot-iron metamorphism should have produced metamorphic fluids from them. By extrapolation and based on geophysical evidence (de Beer and Stettler, 1992), such originally greenschist-facies mafic, ultramafic and metasedimentary units are expected to exist beneath the flat lying SMZ nappe (Fig. 9). A source for fluids percolating along the HRSZ and its vicinity can therefore be identified with some confidence. Any model invoking this type of fluid source has to account for the observed time lag of up to 100 Ma from the HRSZ tectonism to the apparent timing of the fluid flow. This is a problem because hot-iron metamorphism (heat adjustment over a length scale of only a few km) is a geologically fast process, as discussed below. Increased heat flux from the mantle appears an unlikely alternative, first because there is no evidence of any mantle-derived magmatism as mentioned above, and second, because such a process would manifest itself as a relatively sharp pulse in time, rather than a very protracted effect as observed. A further possibility to be explored is, again, intracrustal radioactive heating, this time affecting greenstone belt material that became buried at greater crustal depth in the course of HRSZ tectonism than it was before, and then heated up slowly. The heating rate of supracrustal rocks, previously at greenschist facies, brought to greater crustal depth in the HRSZ event depends on their distance from the SMZ nappe sole (hot-iron effect), the depth to which they were buried, and the thickness of the crust after the HRSZ event. In Fig. 10B and C models for the crustal thermal evolution after this burial are shown for a total crustal thickness 45 km and two values for the thickness of the top hot layer (see Ashwal et al., 1992, for a similar model). The assumption of the still overthickened crust after the HRSZ event accommodates the observation that kyanite postdating this event is currently at surface (Section 3.3), while there is no evidence of post HRSZ major tectonic events in the region, and the present-day crustal thickness is normal. Close to the SMZ nappe sole the heating of buried supracrustals is seen to be very rapid and of very short duration, in accord with the closely clustered peak metamorphic mineral ages discussed above. The hot top layer cools down rapidly, followed by much slower heating of the whole crust. This heating is slower for models with thinner top layers. The 16 km thickness of the hot layer in Fig. 10C is the minimum value for which kyanite could still be expected to occur in the footwall, and therefore crustal heating slower than shown in this figure is not expected. Further, lower crustal heating is faster for thinner total crust models (not shown). Following the model of Fig. 10C, the main temperature range for important dehydration reactions, between ca. 600 ◦ C and 800 ◦ C, is reached at about 10 Ma after the HRSZ event in the lowermost crust, and the geotherm progressively traverses it there and reaches it at higher crustal levels, up to 25 km depth. At 80 Ma after the event the geotherm is close to steady state, and thus at that time fluid production from dehydration reactions in Kaapvaal Craton crust overridden by the HRSZ is expected to cease. The probably fluidrelated kyanite ages and two amphibole dates from the HRSZ fall in

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this 80 Ma time range, as do a number of amphibole dates reported by Belyanin et al. (2014) from the SMZ itself. While the amphibole dates of around 2600 Ma and younger are not readily explained by this model, it appears that deep crustal burial of low-grade supracrustal rocks is a viable mechanism to provide fluid activity, as observed along the HRSZ and its environs, for an extended period after the event. 6. Summary and concluding remarks The supracrustal sequences of the three greenstone belts discussed (Giyani GB, Pietersburg GB and Rhenosterkoppies GB) allow some limited speculation on correlation. The NW domain of the Giyani GB, dominated by ultramafic schists and rich in BIF, is much older (ca. 3200 Ma) than the units summarized as “Simatic Basement” in the Pietersburg GB (ca. 2950 Ma). On the other hand it is possible that the Rhenosterkoppies GB and the SE Giyani GB domain are equivalent to the Simatic Basement, since all contain abundant metagabbro sheets that show a probable ocean floor metamorphism, as well as various ocean floor type sediments. However there are no age data to confirm this. There are no equivalents of the clastic sedimentary Uitkyk formation of the Pietersburg GB in either the Rhenosterkoppies or the Giyani GB. Data yielding ages for tectonism and peak metamorphism in the Pietersburg GB and the SE domain of the Giyani GB cluster at 2840–2870 Ma, which is also the most likely age for the syntectonic Uitkyk Formation. It is proposed that an orogeny, probably related to terrane accretion onto the Kaapvaal Craton, occurred along the two greenstone belts at that time, with the major PGBLwaji suture located along their SE boundaries. In contrast, uplift of the SMZ along the HRSZ occurred at 2717 ± 28 Ma, as shown by ages of peak “hot-iron” metamorphism in the Rhenosterkoppies and Giyani GB’s. The modelled time scale for lower crustal heating to granulite facies conditions in the SMZ after moderate crustal thickening and in the absence of mantle derived magmatism, is 30–60 Ma, which is much shorter than the time lag between the PGB-Lwaji orogeny and the HRSZ uplift. Therefore a connection between the two events is precluded. The possibility that two titanite Pb/Pb dates of ∼2750 Ma in the Rhenosterkoppies GB signal tectonics related to crustal thickening in the SMZ merits further investigation. Further, metamorphism-induced density inversion may have triggered the SMZ uplift along the HRSZ and associated nappe tectonics. Hornblende 40 Ar/39 Ar dates along the HRSZ and within the Giyani GB are diverse and mostly 50–100 Ma younger than the HRSZ peak metamorphism. The results are found not to reflect cooling ages, but rather mixtures of different generations of amphibole, formed as fluid flow was channelled through the HRSZ and related shear zones long after the main HRSZ uplift event. Thermal modelling indicates that supracrustal rocks similar to those of the Rhenosterkoppies, Pietersburg and Giyani GB’s, at low metamorphic grade before the HRSZ event and buried to lower crustal depths during it, could be a possible source for such fluids up to about 80 Ma after the overthrust. Acknowledgements O.T.R. Mining Company (Klein Letaba Section) is thanked for accommodation at Klein Letaba Mine. We thank Chief Homu and the local population for their hospitality. Andre Smit is thanked for introducing the team to the geology of the Giyani Greenstone Belt. This research was supported by the Swiss National Science Foundation (Grant 20-53865.98 to JDK). Balz Kamber and an anonymous referee made insightful and constructive comments that helped to improve the manuscript significantly.

References Andreoli, M.A.G., Brandl, G., Coetzee, H., Kramers, J.D., Mouri, H., 2011. Intracrustal radioactivity as an important heat source for Neoarchean metamorphism in the Central Zone of the Limpopo Complex. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir Boulder, Colorado, pp. 143–161. Ashwal, L., Morgan, P., Hoisch, T.D., 1992. Tectonics and heat sources for granulite metamorphism of supracrustal-bearing terrains. Precambrian Res. 55, 525–538. Barton, J.M.J., Doig, R., Smith, C.B., Bohlender, F., van Reenen, D.D., 1992. Isotopic and REE characteristics of the intrusive charnoenderbite and enderbite geographically associated with the Matok Pluton, Limpopo Belt, Southern Africa. Precambrian Res. 55, 451–467. Barton, J.M.J., van Reenen, D.D., 1992. When was the Limpopo Orogeny? Precambrian Res. 55, 7–16. Belyanin, G.A., Kramers, J.D., Vorster, C., Knoper, M.W., 2014. The timing of successive fluid events in the Southern Marginal Zone of the Limpopo Complex, South Africa: constraints from 40Ar-39Ar geochronology. Precambrian Res. 254, 169–193. Berger, A., Scherrer, N.C., Bussy, F., 2005. Equilibration and disequilibration between monazite and garnet: indication from phase-composition and quantitative texture analysis. J. Metamorphic Geol. 23, 865–880. Cherniak, D.J., Watson, E.B., Grove, M., Harrison, T.M., 2004. Pb diffusion in monazite: a combined RBS/SIMS study. Geochim. Cosmochim. Acta 68, 829–840. de Beer, J.H., Stettler, E.H., 1992. The deep structure of the Limpopo Belt from geophysical studies. Precambrian Res. 55, 173–186. de Wit, M.J., Armstrong, R.A., Kamo, S.L., Erlank, A.J., 1993. Gold-bearing sediments in the Pietersburg greenstone belt: age equivalents of the Witwatersrand Supergroup, South Africa. Econ. Geol. 88, 1242–1252. de Wit, M.J., Jones, M.G., Buchanan, D.L., 1992a. The geology and tectonic evolution of the Pietersburg greenstone belt, South Africa. Precambrian Res. 55, 123–153. de Wit, M.J., et al., 1992b. Formation of an Archaean continent. Nature 357, 553–562. de Wit, M.J., van Reenen, D.D., Roering, C., 1992c. Geologic observations across a tectono-metamorphic boundary in the Babangu area, Giyani (Sutherland) greenstone belt, South Africa. Precambrian Res. 55, 111–122. Frei, R., Kamber, B.S., 1995. Single-mineral Pb–Pb dating. Earth Planet. Sci. Lett. 129, 261–268. Frei, R., et al., 1997. Single mineral dating by the Pb–Pb step leaching method: assessing the mechanisms. Geochim. Cosmochim. Acta 61, 393–414. Gan, S., van Reenen, D.D., 1996. Geology of gold deposits in the Southern Marginal Zone of the Limpopo belt and adjacent Sutherland greenstone belt, South Africa: Klein Letaba. S. Afr. J. Geol. 100, 73–83. Gerya, T.V., Perchuk, L.L., Maresch, W.V., Willner, A.P., 2004. Inherent gravitational instability of hot continental crust; implications for doming and diapirism in granulite facies terrains. In: Whitney, D.L., Teyssier, C., Siddoway, C.S. (Eds.), Gneiss Domes in Orogeny. Geological Society of America Special Paper, Boulder, Colorado, pp. 97–115. Harlov, D.E., Wirth, R., Hetherington, C.J., 2011. Fluid-mediated partial alteration in monazite; the role of coupled dissolution-reprecipitation in element redistribution and mass transfer. Contrib. Mineral. Petrol. 162, 329–348. Henderson, D.R., Long, L.E., Barton, J.M.J., 2000. Isotopic ages and chemical and isotopic composition of the Archean Turfloop batholith, Pietersburg granitegreenstone terrane, Kaapvaal Craton, South Africa. S. Afr. J. Geol. 103, 38– 46. Jahn, B.-M., Condie, K.C., 1995. Evolution of the Kaapvaal Craton as viewed from geochemical and Sm-Nd isotopic analysis of intracratonic pelites. Geochim. Cosmochim. Acta 59, 2239–2258. Kelly, N.M., Clarke, G.L., Harley, S.L., 2006. Monazite behaviour and age significance in poly-metamorphic high-grade terrains: a case study from the Western Musgrave Block, Central Australia. Lithos 88, 100–134. Kleywegt, R.J., et al., 1987. The structure of the Giyani greenstone belt, as derived from geophysical studies. S. Afr. J. Geol. 90, 282–295. Kramers, J.D., McCourt, S., Roering, C., Smit, C.A., van Reenen, D.D., 2011. Tectonic models proposed for the Limpopo Complex: mutual compatibilities and constraints. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir, Boulder, Colorado, pp. 311–324. Kramers, J.D., Mouri, H., 2011. The geochronology of the Limpopo Complex: a controversy solved. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir, Bulder, Colorado, pp. 85–106. Kramers, J.D., Zeh, A., 2011. A review of Sm-Nd and Lu-Hf isotope studies in the Limpopo Complex and adjoining cratonic areas and their bearing on models of crustal evolution and tectonism. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir, Boulder, Colorado, pp. 163–188. Kreissig, K., et al., 2001. Geochronology of the Hout River Shear Zone and the metamorphism in the Southern Marginal Zone of the Limpopo Belt, Southern Africa. Precambrian Res. 109, 145–173. Kreissig, K., Nägler, T.F., Kramers, J.D., van Reenen, D.D., Smit, C.A., 2000. An isotopic and geochemical study of the northern Kaapvaal Craton and the Southern

J.D. Kramers et al. / Precambrian Research 253 (2014) 96–113 Marginal Zone of the Limpopo Belt: are they juxtaposed terranes? Lithos 50, 1–25. Kröner, A., Jaeckel, P., Brandl, G., 2000. Single zircon ages for felsic to intermediate rocks from the Pietersburg and Giyani greenstone belts and bordering granitoid orthogneisses, northern Kaapvaal Craton, South Africa. J. Afr. Earth Sci. 30, 773–793. Laurent, O., Paquette, J.-L., Martin, H., Doucelance, R., Moyen, J.-F., 2013. LA-ICP-MS dating of zircons from Meso- and Neoarchean granitoids of the Pietersburg block (South Africa): crustal evolution at the northern margin of the Kaapvaal craton. Precambrian Res. 230, 209L 226. Manning, C.E., MacLeod, C.J., Weston, P.E., 2000. Lower-crustal cracking front at fastspreading ridges: evidence from the East Pacific Rise and the Oman ophiolite. In: Dilek, Y., Moores, E.M., Elthon, D., Nicolas, A. (Eds.), Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America Special Paper, Boulder, Colorado, pp. 261–272. Manning, C.E., Weston, P.E., Mahon, K.I., 1996. Rapid high-temperature metamorphism of East Pacific Rise gabbros from Hess Deep. Earth Planet. Sci. Lett. 144, 123–132. McCourt, S., van Reenen, D.D., 1992. Structural geology and tectonic setting of the Sutherland Greenstone Belt, Kaapvaal Craton, South Africa. Precambrian Res. 55, 93–110. Miyano, T., Ogata, H., van Reenen, D.D., van Schalkwyk, J.F., Arakawa, Y., 1992. Peak metamorphism of sapphirinebearing rocks in the Rhenosterkoppies greenstone belt, northern Kaapvaal Craton South Africa. In: Glover, J.E., Ho, S.E. (Eds.), The Archean; Terrains, Processes and Metallogeny. Proceedings of the Third International Archean Symposium. University of Western Australia. Geology Department, Perth, W.A., pp. 73–87. Passeraub, M., Wüst, T., Kreissig, K., Smit, C.A., Kramers, J.D., 1999. Structure, metamorphism and geochronology of the Rhenosterkoppies greenstone belt, South Africa. S. Afr. J. Geol. 102, 323–334. Perchuk, L.L., Gerya, T.V., 2011. Formation and evolution of Precambrian granulite terranes: a gravitational redistribution model. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High-Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir, Boulder, Colorado, pp. 289–310. Perchuk, L.L., Gerya, T.V., van Reenen, D.D., Safonov, O.G., Smit, C.A., 1996. The Limpopo Metamorphic Complex South Africa, 2. Decompression-cooling regimes of granulites and adjacent rocks of the Kaapvaal Craton. Petrology 4, 571–599. Pyle, J.M., Spear, F.S., 2003. Four generations of accessory-phase growth in lowpressure migmatites from SW New Hampshire. Am. Miner. 88, 338–351. Renne, P.R., Mundil, R., Balco, G., Min, K., Ludwig, K.R., 2010. Joint determination of 40 K decay constants and 40 Ar*/40 K for the Fish Canyon sanidine standard, and improved accuracy for 40 Ar/39 Ar geochronology. Geochim. Cosmochim. Acta 74, 5349–5367. Roering, C., et al., 1992a. Tectonic model for the evolution of the Limpopo belt. Precambrian Res. 55, 539–552. Roering, C., et al., 1992b. Structural geological and metamorphic significance of the Kaapvaal Craton/Limpopo belt contact. Precambrian Res. 55, 69–80.

113

S.A.C.S. (South African Commission for Stratigraphy), 1980. Stratigraphy of South Africa, Part I (Comp. L. E. Kent). Lithostratigraphy of the Republic of South Africa, South West Africa/Namibia, and the Republics of Bophuthatswana, Transkei and Venda. Handbook of the Geological Survey of South Africa, 8, Geological Survey of South Africa, Pretoria, pp. 690. Schoene, B., Dudas, F.O.L., Bowring, S., de Wit, M.J., 2009. Sm-Nd isotopic mapping of lithospheric growth and stabilization in the eastern Kaapvaal Craton. Terra Nova 21, 219–228. Smit, C.A., Roering, C., van Reenen, D.D., 1992. The structural framework of the southern margin of the Limpopo belt, South Africa. Precambrian Res. 55, 51–67. Smit, C.A., van Reenen, D.D., Roering, C., 2014. The role of fluids in the exhumation of the Southern Marginal Zone of the Limpopo Complex, South Africa. Precambrian Res. 253, 81–95. Smit, C.A., van Reenen, D.D., Gerya, T.V., Perchuk, L.L., 2001. Conditions of decompression of the Limpopo high grade terrain: record from shear zones. J. Metamorphic Geol. 19, 249–268. van Reenen, D.D., Smit, C.A., Perchuk, L.L., Roering, C., Boshoff, R., 2011. Thrust exhumation of the Neoarchean ultrahigh-temperature Southern Marginal Zone Limpopo Complex: convergence of decompression-cooling paths in the hanging wall and prograde P-T paths in the footwall. In: van Reenen, D.D., Kramers, J.D., McCourt, S., Perchuk, L.L. (Eds.), Origin and Evolution of Precambrian High Grade Gneiss Terranes, with Special Emphasis on the Limpopo Complex of Southern Africa. Geological Society of America Memoir, Boulder, Colorado, pp. 189– 212. van Schalkwyk, J.F., de Wit, M.J., Roering, C., van Reenen, D.D., 1993. Tectonometamorphic evolution of the simatic basement of the Pietersburg greenstone belt relative to the Limpopo Orogeny: evidence from serpentinite. Precambrian Res. 61, 67–88. Vearncombe, J.R., 1988. Structure and metamorphism of the Archean Murchison belt, Kaapvaal Craton, South Africa. Tectonics 7, 761–774. Villa, I.M., 2010. Disequilibrium textures versus equilibrium modelling: geochronology at the crossroads. In: Spalla, M., Marotta, A.M., Gosso, G. (Eds.), Advances in Interpretation of Geological Processes: Refinement of Multi-scale Data and Integration in Numerical Modelling, Special Publications, 332. Geological Society, London, pp. 1–15. Whitney, D.L., Teyssier, C., Vanderhaeghe, O., 2004. Gneiss domes and crustal flow. In: Whitney, D.L., Teyssier, C., Siddoway, C.S. (Eds.), Gneiss Domes in Orogeny. Geological Society of America Special Paper, Boulder, Colorado, pp. 15– 26. Zeh, A., Gerdes, A., 2012. U/Pb and Hf isotope record of detrital zircons from gold-bearing sediments of the Pietersburg greenstone belt (South Africa); is there a common provenance with the Witwatersrand Basin? Precambrian Res. 204–205, 46–56. Zeh, A., Gerdes, A., Barton, J.M.J., 2009. Archean accretion and crustal evolution of the Kalahari Craton – the zircon age and Hf isotope record of granitic rocks from Barberton/Swaziland to the Francistown Arc. J. Petrol. 50, 933–966. Zeh, A., et al., 2013. Juvenile crust formation in the Northeastern Kaapvaal Craton at 2.97 Ga; implications for Archean terrane accretion, and the source of the Pietersburg gold. Precambrian Res. 233, 2–43.