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Archaean to Proterozoic Crustal Evolution in the Central Zone of the Limpopo Belt (South Africa–Botswana): Constraints from Combined U–Pb and Lu–Hf Isotope Analyses of Zircon
1Mineralogisches Institut Der Universität Würzburg, Am Hubland, D-97074 Würzburg, Germany
2Institut Für Geowissenschaften, Senckenberganlage 28, D-60054 Frankfurt Am Main, Germany
3Department of Geology, University of Fort Hare, Private Bag X1314, Alice, 5700, South Africa
RECEIVED JANUARY 9, 2007; ACCEPTED MAY 24, 2007
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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A combined set of U–Pb and Lu–Hf in situ laser ablation ICP-(MC)-MS zircon analyses were obtained from orthogneisses and granitoids in the Central Zone of the Limpopo Belt, which comprises the Beit Bridge and Mahalapye complexes. The results indicate that by combining the two isotope systems primary magmatic zircon domains can be distinguished from those formed during later metamorphic events, even if the distinct zircon domains underwent multiple Pb loss and the texture–age relationships, as obtained by cathodoluminescence images and U–Pb analyses, are ambiguous. Furthermore, the applied technique allows distinction of zircon grains formed in juvenile magmas from those generated by melting of older continental crust or affected by substantial crustal contamination. The combined U–Pb and Lu–Hf data reveal that the Sand River gneiss suite of the Beit Bridge Complex was emplaced at 3283 ± 8 Ma and formed from melting of an older Archaean crust, which was derived from a depleted mantle source at around 3·65 Ga. The hafnium model age (TDMHf) is significantly older than those obtained from zircons from numerous Neoarchaean granitoids of the Beit Bridge Complex, comprising the Singelele gneiss (2647 ± 12 Ma), the Bulai granite (2612 ± 7 Ma), the Regina gneiss (2649 ± 9 Ma) and two samples of the Zanzibar gneiss (2613 ± 6 Ma). These granitoids show initial
Hf(t) values between + 0·5 and –7·1, which correspond to initial TDMHf between 3·46 and 3·01 Ga. These variable TDMHfinitial and Hf(t)initial values are interpreted to be the result of different mixtures of reworked 3·65 Ga Palaeoarchaean crust with juvenile magmas extracted from the depleted mantle during the Neoarchaean at 
2·65 Ga. This conclusion is supported by results obtained from the Mahalapye Complex, which was affected by migmatization and granite intrusions during the Palaeoproterozoic at 2·02–2·06 Ga. The Mokgware granite (2019 ± 9 Ma) contains zircon xenocrysts with Pb–Pb ages of 2·52–2·65 Ga and 2·93 Ga and hafnium model ages of 3·0–3·4 Ga, indicating that this granite is derived from remelting of Archaean crust. In contrast, uniform TDMHfinitial ages of 2·61–2·67 Ga obtained from a diorite gneiss (2061 ± 6 Ma) of the Mahalapye Complex indicate that its protolith may have been formed from remelting of a Neoarchaean juvenile crust. Variable Hf(t)initial values from –3·7 to +6·3 of zircon cores (2711 ± 11 Ma) in an adjacent leucosome also support a model of mixing of juvenile mantle derived matter with older crust in the Neoarchaean. KEY WORDS: Archaean; Palaeoproterozoic; Limpopo Belt; zircon, U–Pb dating; Lu–Hf isotopes; LA-ICP-MS
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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Isotope methods applied to Archaean magmatic rocks provide important information about the formation and evolution of the continental crust during the early Earth's history. Such methods can provide answers to some of the most controversial questions; for instance, about the timing of Archaean to Proterozoic crust-forming events and the recycling of this crust during subsequent orogenic processes (e.g. Armstrong, 1968
, 1991; O’Nions et al., 1979; O’Nions & Hamilton, 1981; Patchett et al., 1981; Allègre, 1987; Albarède & Brouxel, 1987; Allègre & Lewin, 1989; Stevenson & Patchett, 1990; Amelin et al., 1999; Rino et al., 2004; Condie et al., 2005; Harrison et al., 2005; Iizuka et al., 2005). In detail, such investigations provide valuable information on the timing of magma generation and its source, such as a juvenile depleted mantle, a reworked older crust or a combination of both (e.g. Vervoort & Blichert-Toft, 1999; Davis et al., 2005; Gerdes & Zeh, 2006; Hawkesworth & Kemp, 2006).
For such studies, the mineral zircon has been shown to be an ideal candidate, for the following reasons. (1) Zircon crystallization ages can be obtained very precisely using the U–Pb isotope system even for very old rocks (e.g. Krogh, 1973; Compston et al., 1984; Gerdes & Zeh, 2006; Iizuka et al., 2006). (2) Zircon can incorporate high concentrations of Hf (>1 wt %) but only minor Lu (1 ppm) in its lattice and, thus, the initial Hf isotope compositions of the magma at the time of zircon crystallization can be obtained with high precision, in particular by means of inductively coupled plasma mass spectrometry (ICP-MS) techniques (e.g. Patchett et al., 1981; Vervoort & Blichert-Toft, 1999; Griffin et al., 2004; Gerdes & Zeh, 2006). (3) In situ analytical methods such as ion probes (e.g. sensitive high-resolution ion microprobe; SHRIMP) and laser ablation (LA)-ICP-MS allow the analyses of different parts of individual zircon grains. Consequently, it is possible to recognize inherited zircon domains as well as younger overgrowth and/or alteration zones (e.g. Gerdes & Zeh, 2006, and therein). Furthermore, the combined consideration of the U–Pb and Lu–Hf system for zircon grains in orthogneisses may help to constrain whether distinct zircon grains or domains were affected by single or multiple Pb loss, and/or if distinct zircon zones were formed at different times (e.g. Amelin et al., 2000, and therein).
In this study we investigate zircon from orthogneisses and granitoids in the Central Zone of the Limpopo Belt, which represents a mobile belt squeezed between the Zimbabwe Craton to the north and the Kaapvaal Craton to the south (Fig. 1). As shown by previous studies, the (meta)magmatic rocks of the Limpopo Belt were formed during at least three distinct magmatic periods, during the Palaeoarchaean at 3·25 Ga, the Neoarchaean at 2·5–2·7 Ga and the Palaeoproterozoic at c. 2·0 Ga (e.g. Barton et al., 1994; Barton & Sergeev, 1997; Jaeckel et al., 1997; Kröner et al., 1998, 1999; Chavagnac et al., 2001). In addition, a few detrital zircon grains and xenocrysts point to magmatic events as old as 3·8 Ga (Armstrong et al., 1988; Kröner et al., 1998). The available geochronological data show that the Limpopo Central Zone represents a unique orogenic belt where Archaean to Proterozoic magmatism, which took place over a period of more than 1300 Myr, can be studied in a very restricted area. Furthermore, the position of the Limpopo Belt between the Kaapvaal and Zimbabwe Cratons makes it important for obtaining information about the crust–mantle processes that took place during the amalgamation of these two cratons.
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In contrast to the abundant geochronological data mentioned above, little is known about the source(s) of the magmatic rocks that were emplaced during the three distinct events. A few Sm–Nd isotope whole-rock analyses indicate that the orthogneisses with ages of 3·25–3·3 Ga and 2·6 Ga were derived from a source with an average crustal residence age (TDMNd) between 2·8 and 3·4 Ga (Harris et al., 1987
; Barton, 1996; Kröner et al., 1998; 1999), whereas some 2·0 Ga migmatitic–magmatic rocks yielded younger crustal residence ages between 2·4 and 2·9 Ga (Chavagnac et al., 2001). To clarify the crustal evolution of the Limpopo Belt's Central Zone in more detail, we present new in situ U–Pb LA-ICP-MS age data from magmatic zircons and combine this information with multicollector (MC-)ICP-MS Lu–Hf isotope analyses from the same zircon domains. |
GEOLOGICAL SETTING, CURRENT INTERPRETATIONS AND SAMPLES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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The Limpopo Belt is a high-grade metamorphic province bounded by the Zimbabwe and Kaapvaal Cratons, comprising Archaean to Palaeoproterozoic rocks (Fig. 1) (e.g. Roering et al., 1992
; Berger et al., 1995; Kamber et al., 1995a, 1995b; Rollinson & Blenkinsop, 1995; Berger & Rollinson, 1997; Holzer et al., 1998, 1999). Lithologically and structurally, it can be subdivided into three distinct domains, which are separated by major shear zones (McCourt & Vearncombe, 1987, 1992; Holzer et al., 1999; Schaller et al., 1999). The Central Zone is structurally bounded by the >10 km wide Palala–Zoetfontain and Triangle shear zones against the Southern and Northern Marginal zones, respectively. These steeply dipping shear zones were formed and/or activated under high-grade metamorphic conditions at about 2·0 Ga (Holzer et al., 1998, 1999; Chavagnac et al., 2001; Kreissig et al., 2001). The Southern Marginal Zone consists of high-grade, intensely deformed and metamorphosed, tonalitic, trondjemitic and granodioritic rocks (TTGs) with minor greenstones, interpreted as reworked counterparts of the Kaapvaal Craton (e.g. Coward et al., 1976; van Reenen et al., 1987, 1992), whereas the Northern Marginal Zone is dominated by enderbite and charnoenderbitic orthogneisses with some greenstones and granitoid rocks (e.g. Mason, 1973; Berger et al., 1995; Kamber & Biino, 1995; Berger & Rollinson, 1997; Jelsma & Dirks, 2002). The rocks of both marginal zones were thrust onto their adjacent cratons along >10 km wide, steeply dipping shear zones, the Hout River and Umlali shear zones, respectively (Fig. 1). According to Kreissig et al. (2001), magmatic intrusions and metamorphism of the Southern Marginal Zone occurred between 2·69 and 2·62 Ga, coeval with or slightly younger than the magmatic–metamorphic evolution in the Northern Marginal zone between 2·72 and 2·58–2·62 Ga (Berger et al., 1995; Kamber & Biino, 1995; Mkweli et al., 1995).
What was traditionally called the Central Zone proper was subdivided by Aldiss (1991) into three complexes, which were interpreted by Barton et al. (2006) as terranes: (1) the Mahalapye Complex, which is dominated by c. 2·02–2·0 Ga granites with minor high-grade sedimentary rocks (Hisada & Miyano, 1996; McCourt & Armstrong, 1998; Holzer et al., 1999; Chavagnac et al., 2001); (2) the Phikwe Complex, which is dominated by Archaean, hornblende-bearing, tonalitic and trondjemitic gneisses and igneous rocks (Brandl, 1992); (3) the Beit Bridge Complex. The Beit Bridge Complex hosts the c. 3·2–3·3 Ga Sand River TTG suite and Messina layered intrusion (Barton, 1983; Barton & Sergeev, 1997; Kröner et al., 1999) and numerous Neoarchaean granitic to granodioritic orthogneisses with ages between 2·73 and 2·60 Ga, such as the Alldays, Singelele, Bulai, Zanzibar and Zoetfontain gneisses (e.g. Jaeckel et al., 1997; Kröner et al., 1999). Based on detrital zircon ages and lithological–structural differences, the supracrustal paragneisses of the Beit Bridge Complex can be divided into three unconformity-bounded successions, with ages >3·1 Ga, 3·1–2·6 Ga and < 2·6 Ga (e.g. Erikson et al., 1988; Brandl, 1992; Kröner et al., 1998, 1999; Barton et al., 2003; Buick et al., 2003). Quartzites and metapelites from within the oldest unconformity-bounded succession contain zircons with U–Pb ages of 3·2–3·8 Ga, suggesting that a yet unrecognized, very old protolith must exist within the Beit Bridge Complex (Armstrong et al., 1988; Barton & Sergeev, 1997; Kröner et al., 1998).
Further subdivision of the Limpopo Belt comes from Pb isotopic data, indicating that the Zimbabwe Craton, Northern Marginal Zone, and Beit Bridge and Phikwe terranes were all derived from a source with a long-lived, high µ value (>11–12), whereas the rocks in the Southern Marginal Zone and the Kaapvaal Craton were derived from a source with lower µ values close to Bulk Earth (Barton et al., 1983; Taylor et al., 1991; Barton, 1996; Berger & Rollinson, 1997; Kreissig et al., 2000). Hence, the rocks of the Zimbabwe Craton, Northern Marginal Zone, and Beit Bridge and Phikwe terranes may stem from the same mantle source, but they cannot be genetically related in a straightforward way to those of the Kaapvaal Craton and Southern Marginal Zone (Barton et al., 2006). Based on geochemical and Sm–Nd–Pb isotope data and petrogenetic modelling, Berger & Rollinson (1997) suggested that the 2·6–2·7 Ga enderbitic to charnoenderbitic rocks of the Northern Marginal Zone represent a mixture of a pre-existing Archaean TTG–crust, which was partially remelted and mixed with a more juvenile magma derived from the depleted mantle at 2·6–2·7 Ga.
In the present study zircon grains from nine granitoid samples were investigated, six from the Beit Bridge Complex (Sand River gneiss, Singelele gneiss, Bulai granite, Regina gneiss, grey and leucocratic Zanzibar gneiss) and three from the Mahalapye Complex (Mogkware granite, a garnet–biotite gneiss and a leucosome from the Lose quarry). The sample localities and co-ordinates are shown in Fig. 1 and Table 1, respectively.
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ANALYTICAL TECHNIQUES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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U–Pb and Lu–Hf LA-ICP(MC)-MS analyses
Zircon concentrates were prepared at Würzburg, Giessen and Frankfurt Universities using standard crushing and heavy mineral separation techniques. Selected grains were separated under alcohol and set in epoxy resin to form discs. Subsequently, the zircons were polished to expose their centres. Prior to analysis all mounts were photographed and zircon grains were imaged by SEM cathodoluminescence (CL) to identify homogeneous growth domains, using the JEOL JSM-6400 electron microprobe at the Institute of Geosciences, Frankfurt University. Selected zircon domains were analysed, following the procedure as outlined in detail by Gerdes & Zeh (2006
). First, U–Pb analyses were carried out by LA-ICP-MS at Frankfurt University using a Thermo-Finnigan Element II sector field ICP-MS system coupled to a New Wave UP213 ultraviolet laser system. A teardrop-shaped, low-volume laser cell was used to allow the precise detection of heterogeneous material (e.g. inclusion or different growth zones) during time-resolved data acquisition (see Janousek et al., 2006). Each analysis consisted of c. 20 s background acquisition followed by 35 s data acquisition, using a laser spot-size of 30 µm. A common-Pb correction based on the interference- and background-corrected 204Pb signal and a model Pb composition (Stacey & Kramers, 1975) was carried out if necessary. The necessity of the correction is judged on whether the corrected 207Pb/206Pb lies outside the internal errors of the measured ratios. Discordant analyses were generally interpreted with care. Reported uncertainties were propagated by quadratic addition of the external reproducibility (standard deviation) obtained from the standard zircon GJ-1 (n = 12; 0·5% and 0·5–0·9% for 207Pb/206Pb and 206Pb/238U, respectively) during individual analytical sessions (33 analyses of unknowns and 12 standards) and the within-run precision of each analysis (standard error). Concordia diagrams (2
error ellipses), concordia ages and upper intercept ages (95% confidence level) were calculated using Isoplot/Ex 2.49 (Ludwig, 2001). Subsequently, the same zircon domains were studied using the same laser system with a 40 µm spot-size, and the Lu, Hf and Yb isotopes were measured with a Thermo-Finnigan Neptune multi-collector ICP-MS system. The isotopes 172Yb, 173Yb and 175Lu were simultaneously monitored during each analysis step to allow the correction of isobaric interferences between Lu and Yb isotopes on mass 176. 176Yb and 176Lu were calculated using 176Yb/173Yb = 0·796218 (Chu et al., 2002) and 176Lu/175Lu = 0·02658 (see Gerdes & Zeh, 2006), and by taking the instrumental mass fractionation of each individual analysis into account. For instrumental mass bias correction Yb isotope ratios were normalized to 172Yb/173Yb = 1·35274 (Chu et al., 2002) and Hf isotope ratios to 179Hf/177Hf = 0·7325 using an exponential law. The protocol was tested by replicate analysis of a 50 ppb JMC 475 solution that was doped with variable amounts of pure Yb and Lu (see Gerdes & Zeh, 2006). The accuracy of the interference correction could be verified even for a 176Yb/177Hf ratio of 0·241 and a 176Lu/177Hf ratio of 0·021, which clearly exceed the highest observed Yb and Lu amounts in the studied zircon grains. Analyses of a pure 50 ppb JMC 475 solution yielded 176Hf/177Hf = 0·282146 ± 8 (2, n = 22). For better comparison all analyses are reported relative to the JMC 475 176Hf/177Hf value of 0·282160. Ten and 16 LA-MC-ICP-MS analyses of the GJ-1 zircon (c. 9600 ppm Hf) during the two analytical session gave a 176Hf/177Hf ratio of 0·282007 ± 15 (2) and 0·282006 ± 15 (2), respectively (Table 3). This is identical, within error, to results obtained by solution MC-ICP-MS analyses of the Lu- and Yb-free Hf fraction (0·281999 ± 8, 2, n = 10). Multiple LA-MC-ICP-MS analyses of the 91500 zircon standard over a period of 9 months yielded 176Hf/177Hf = 0·282297 ± 22 (2, n = 112). Our results are shown in Tables 2 and 3, and in Figs 2–8. All uncertainties in this paper are quoted at the ±2 level.
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, discordant analyses.
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Constants and calculation parameters
Depleted mantle Hf model ages (TDMHf) were calculated using a decay constant of 1 · 867 x 10–11 (Scherer et al., 2001
; Söderlund et al., 2004), 176Lu/177Hf = 0·0384 and 176Hf/177Hf = 0·28325 for the depleted mantle (Chauvel & Blichert-Toft, 2001), a mean 176Lu/177Hf value of 0·0113 for the Palaeoproterozoic–Archaean crust [mean of average continental crust as suggested by Taylor & McLennan (1985) and Wedepohl (1995)], and the measured 176Lu/177Hf value of each spot for the time since zircon crystallization. The applied decay constant is, according to various workers (e.g. Davis et al., 2005; Albarède et al., 2006; Scherer et al., 2007), more appropriate to derive information about Archaean crustal growth processes than that estimated by the analyses of meteorites [(1·958 ± 0·044) x 10–11; Albarède et al., 2006, and references herein). The use of the ‘meteoritic’ constant would increase the calculated Hf(t) for our samples on average by about 2·5 epsilon units, and would give slightly younger (0·10–0·003 Ga) hafnium model ages. A more significant impact on the calculated hafnium model ages results from the assumed 176Lu/177Hf value for the Palaeoproterozoic–Archaean crust. A lower 176Lu/177Hf value of 0·007, which is typical for present-day upper crust (Taylor & McLennan, 1985; Wedepohl, 1995), would decrease the TDMHf of our zircon analyses by about 0·01–0·25 Ga (mean for all analyses = 0·10 Ga). A higher 176Lu/177Hf value of 0·015, as proposed by Rudnick & Gao (2003) for the bulk continental crust, would increase our calculated TDMHf by about 0·12 Ga (average). Additional uncertainties with similar influence on the calculated TDMHf result from the 176Lu/177Hf heterogeneity of depleted mantle melts (e.g. Chauvel & Blichert-Toft, 2001).
For calculation of Hf(t) we used the chondritic uniform reservoir (CHUR) recommended by Blichert-Toft & Albarède (1997; 176Lu/177Hf and 176Hf/177Hf values of 0·0332 and 0·282772, respectively), and the apparent Pb–Pb ages obtained for the respective zircon domains. By means of this procedure many zircon analyses from individual granitoid samples plot on linear arrays in the Hf(t) vs apparent Pb–Pb age diagram, in the initial 176Hf/177Hf vs apparent Pb–Pb age diagram, and in the TDMHf vs apparent Pb–Pb age diagram (Figs 4 and 8). Given the case that the initial 176Hf/177Hf values of all of these zircon analyses are identical, within error, such an array suggests that all zircon domains were formed at the same time, but that some of them underwent multiple Pb loss afterwards, while maintaining their initial 176Hf/177Hf (see Amelin et al., 2000). Thus, for all zircon analyses that plot on such an array initial Hf(t) values [=Hf(t)initial] were recalculated using the respective magma crystallization ages (e.g. Fig. 9a and b). The same criterion was used to calculate initial TDMHf values (=TDMHfinitial), which are younger than the TDMHf obtained by using the respective apparent Pb–Pb age (Table 3). The depleted mantle curve, as shown in Fig. 9, is extrapolated from average modern-day values of mid-ocean ridge basalts (MORB) (176Lu/177Hf = 0·0384, 176Hf/177Hf = 0·28325; Chauvel & Blichert-Toft, 2001) assuming a linear behaviour.
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RESULTS AND INTERPRETATION OF U–PB AND LU–HF ZIRCON ISOTOPE ANALYSES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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Beit Bridge Complex
Sand River gneiss (sample SR)
This tonalitic gneiss sample, which mainly consists of plagioclase, quartz, biotite, minor hornblende, and orthopyroxene, was taken from the bed of the Sand River near the Causeway locality south of Messina (Fig. 1, Table 1). It represents a grey banded gneiss that is truncated and cut by numerous later melt batches. CL images of most zircon grains show cores with oscillatory zoning patterns, which are in some grains overgrown by structureless rims (Fig. 2a). In addition, a few completely structureless zircon crystals are observed (Fig. 2a). During this study 27 U–Pb spot analyses were carried out on cores and rims of 21 zircon grains (Table 2). In addition, Lu–Hf isotope spot analyses of nine zircon cores were obtained (Table 3). As shown in Fig. 3a, four groups of concordant to near-concordant U–Pb ages can be distinguished. Zircon cores yield concordant ages of about 3·28 Ga and one of 3·14 Ga, whereas zircon overgrowths and structureless zircon grains yield concordant ages of about 2·02 Ga and one of c. 2·61 Ga (Figs 2a and 3a). All four concordant age groups are also reflected in the Pb–Pb intercept ages of discordant zircon analyses (Table 2, Fig. 3a).
We suggest that the oldest concordant zircon grains, which yield a mean age of 3283 ± 8 Ma, identical to an upper intercept age of 3280 ± 7 Ma (Fig. 3a), reflect the time of the Sand River gneiss protolith intrusion, whereas the younger concordant age of 3143 ± 13 Ma dates the time of zircon growth during a Palaeoarchaean anatectic event. Our estimated intrusion age is identical to or slightly younger than that obtained by SHRIMP U–Pb analyses from the Sand River granitoid suite at the Causeway locality (3314 ± 5 Ma, 3240 ± 5 Ma), the Macloutse farm (3290 ± 17 Ma), Bellvue farm (3297 ± 7 Ma), Verbaard farm (3296 ± 4 Ma) and Vrouenroom farm (3292 ± 4 Ma), but older than zircon SHRIMP ages obtained from the Esmefour farm (3205 + 69/–62 Ma) and the Sand River bed (3240 ± 5 Ma; 3197 ± 30 Ma) (Kröner et al., 1998, 1999). The suggested Palaeoarchaean anatectic event at 3143 ± 13 Ma conforms with field relationships and geochronological data; the highly deformed and migmatitized Sand River gneiss is locally cut by basalt dykes, which yielded Rb–Sr and Pb–Pb ages of 3060 Ma and 2922 Ma, respectively (Barton et al., 1977; Barton et al., 1990).
The upper intercept age of 2614 ± 11 Ma is interpreted to reflect zircon formation in the Sand River gneiss during a Neoarchaean metamorphic–magmatic event, whereas the upper intercept age of 2014 ± 10 Ma (five spots, MSWD = 1· 0) dates zircon growth during the Palaeoproterozic granulite-facies metamorphic overprint. The latter age is slightly younger than U–Pb isotope dilution thermal ionization mass spectrometry (ID-TIMS) and Pb–Pb zircon evaporation ages of 2031 ± 6 and 2026 ± 1 Ma, respectively, which were obtained on metamorphic zircon grains from pelitic gneisses of the Causeway locality (Jaeckel et al., 1997), but identical within error to the upper intercept ages of strongly discordant U–Pb ID-TIMS zircon analyses of 2006 ± 8 Ma and 2003 ± 11 Ma, which were obtained from zircon grains from melt patches in the Sand River gneisses and from a pegmatitic granite (Jaeckel et al., 1997). The 2614 Ma age is identical within error to the intrusion age of several orthogneisses found throughout the Central Zone (Jaeckel et al., 1997; Kröner et al., 1999; this study). However, it is significantly younger than the two concordant U–Pb SHRIMP analyses of about 2734 ± 4 Ma in a c. 3·29 Ga quartz monzonitic augen gneiss, which were interpreted to represent a local melting event during the Neoarchaean (Kröner et al., 1999).
The Lu–Hf analyses of eight zircon cores yielded relative homogeneous initial 176Hf/177Hf values between 0·28063 and 0·28066, which correspond to Hf(t)initial between –0·1 and –1· 6, and provide evidence that the zircon cores crystallized in a magma, which was derived from a source with an average crustal residence age of TDMHfinitial 3·64 ± 0·04 Ga. A slightly higher initial 176Hf/177Hf value (0·28071) was obtained from the 3·14 Ga zircon domain (spot 32c/zrc19) interpreted to be formed during Palaeoarchaean anatexis about 140 Myr after the intrusion of the Sand River gneiss protolith (Figs 2a and 3a). The increased initial 176Hf/177Hf is consistent with the incorporation of additional radiogenic hafnium during anatectic zircon crystallization. This additional 176Hf was formed by 176Lu decay in the Sand River gneiss matrix between 3·28 and 3·14 Ga.
The presented hafnium model ages for the 3·28 Ga zircon cores are in line with the U–Pb age data, but older than the TDM values of 3·1–3·4 Ga, which were derived by Sm–Nd whole-rock isotope analysis of the Sand River gneisses (Kröner et al., 1999). The younger neodymium model ages may result from Palaeoarchaean open-system behaviour, which led to a resetting of the Sm–Nd system (see Moorbath et al., 1997). In combination with the U–Pb ages, our TDMHfinitial values indicate that the crustal source was extracted from a depleted mantle reservoir about 300 Myr prior to the formation of the Sand River granitoid suite.
Bulai granite (sample Bu)
Sample Bu is a porphyric, weakly deformed granite containing K-feldspar, plagioclase, biotite and quartz, and was taken from the type locality of the Bulai pluton (Table 1). From this sample 27 U–Pb spot analyses on 24 zircon grains and 19 Lu–Hf analyses were carried out (Tables 2 and 3; Fig. 4d–f). CL images of most zircon crystals reveal an oscillatory magmatic zoning and two bright zones (cores and rims), which are separated by a dark band of weak luminescence (Fig. 2b). Some zircon grains show additional overgrowths with a diffuse luminescence (Fig. 2c). U–Pb analyses obtained from all of these distinct zircon domains yielded within error identical ages (Fig. 2b and c), indicating that they were formed during the same magmatic event. Fourteen concordant U–Pb zircon analyses (99–101% concordance level) gave an age of 2612 ± 7 Ma, which is within error identical to an upper intercept age of 2607 ± 7 Ma obtained from five discordant analyses plus the 14 concordant zircon analyses (Fig. 3b). The lower intercept age of 790 ± 67 Ma is interpreted as geologically meaningless and may result from multiple Pb loss. Furthermore, there are seven U–Pb analyses that group together with the 14 concordant analyses on a discordia with a lower intercept at 1943 ± 150 Ma (Fig. 3b). This suggests that some zircon domains were affected by Pb loss during the Palaeoproterozoic, which is in agreement with previous age data of Barton & Sergeev (1997), Jaeckel et al. (1997) and Holzer et al. (1998).
The concordant 2612 ± 7 Ma U–Pb zircon age of the Bulai pluton is interpreted to reflect the time of emplacement. This age is identical within error to a U–Pb zircon SHRIMP upper intercept age of 2620 ± 8 Ma from a granitic phase (Kröner et al., 1998) and an ID-TIMS U–Pb zircon upper intercept age of 2605 ± 2 Ma from a deformed enderbitic phase of the Bulai pluton (Barton et al., 1994). However, it is significantly older than a U–Pb ID-TIMS zircon age of 2572 ± 4 Ma from a granitic phase (Barton et al., 1994) and a Pb–Pb zircon evaporation age of 2587 ± 1 Ma from a deformed enderbitic phase (Kröner et al., 1999). It should be noted that the younger zircon age reported by Barton et al. (1994) is an upper intercept U–Pb age, which results from the fitting of only two highly discordant zircon analyses. Thus, it cannot be excluded that these zircon grains were subjected to multiple Pb loss, an effect that is well documented by our in situ LA-ICP-MS analyses (Fig. 3b). Multiple Pb loss may also account for the young Pb–Pb evaporation age (2587 ± 1 Ma) obtained by Kröner et al. (1999).
Nineteen Lu–Hf analyses of concordant and discordant zircons were carried out. As shown in Fig. 4d, the zircons have initial 176Hf/177Hf ratios of about 0·28102 ± 0·00003 (2 SD = standard deviation), which correspond to Hf(t)initial between –2·3 and –4·3 when the 2612 Ma Bulai granite crystallization age is applied to all Hf analyses. This procedure is justified, as the initial 176Hf/177Hf values of all analysed zircon domains show relatively minor variations. Nevertheless, there seems to be a small increase of the initial 176Hf/177Hf with decreasing Pb–Pb age (Fig. 4d), a feature that can be explained in different ways. Taking the errors into account, the initial 176Hf/177Hf of the apparently younger domains is identical to that obtained from most of the apparently older domains. Thus, the observed array can be interpreted (in the most simple way) to reflect multiple Pb loss in younger zircon domains, which preserved their initial 176Hf/177Hf incorporated during magma crystallization. Alternatively, the small increase of the initial 176Hf/177Hf could reflect some minor incorporation of more radiogenic Hf in the younger zircon domains, perhaps during partial recrystallization of metamict zircon domains caused by fluid–zircon–matrix interactions. Initial hafnium model ages of 3·22 ± 0·06 Ga indicate that the Bulai granite was derived from a crustal source, which was extracted on average from a depleted mantle reservoir 600–700 Myr prior to the Bulai granite intrusion.
Singelele granodiorite gneiss (sample Sin)
Sample Sin is a high-grade granite gneiss, which was collected from the type locality at the Singelelekop near Messina (see Kröner et al., 1999). It is characterized by the mineral assemblage plagioclase–K-feldspar–biotite–quartz–garnet. A total of 53 spots on 43 zircon grains from this sample were analysed for U–Pb and 45 spots for Lu–Hf (Tables 2 and 3; Figs 3c and 4g–i). The calculated 207Pb/206Pb ages show large variations from 1·54 to 2·66 Ga with 38 of 53 analyses yielding discordant results (below or above 95–105% concordance; Table 2, Fig. 3c). For many zircon grains core and rim domains were distinguished by their CL images (Fig. 2d–f). However, this textural relationship is not always reflected by the geochronological data. Based on the texture–(Pb–Pb)-age relationships three zircon groups can be distinguished: (1) zircon grains with significantly younger cores than rims (Fig. 2d); (2) zircon grains with nearly identical Pb–Pb ages of cores and rims (Fig. 2d); (3) zircon grains with older cores and younger rims (Fig. 3f). These different texture–age relationships are interpreted to reflect that the zircon cores and rims grew either at the same time or at different times [textures (2) and (3)], and/or that the distinct zircon domains were affected by alteration processes causing non-zero or multiple Pb loss, effects that have been described in detail by several researchers (e.g. Vavra et al., 1996; Geisler et al., 2001).
Whereas the texture–age relationships are difficult to interpret, the additional Hf isotope data reveal a relatively simple story. In fact, they allow us to distinguish between zircon zones that either are affected by alteration or are the result of new zircon growth. Most zircon domains (Fig. 4g) have, despite their different apparent Pb–Pb ages, which range from 1539 to 2658 Ma, very similar initial 176Hf/177Hf values of about 0·28099 ± 0·00005 (2 SD). Therefore, they display a simple trend in the Hf(t) vs age diagram, which is characterized by a positive correlation between apparent zircon Pb–Pb ages and Hf(t) (Fig. 4h). This trend is similar to that observed for the Bulai zircons (Fig. 4e), although more pronounced. It suggests that all of these zircon domains or grains (including most zircon cores and rims of sample Sin1; Tables 2 and 3) were formed during the same magmatic event, but were subjected to Pb loss with a different intensity. It should be noted that the calculated TDMHf of these zircons become apparently older with decreasing apparent Pb–Pb age (Fig. 4i). This effect results from the two-stage model used. Thus, for the geological interpretation only initial hafnium model ages (TDMHfinitial) can be used. These are between 3·2 and 3·5 Ga for the Singelele gneiss.
The Hf isotope analyses indicate that nearly all zircon grains or domains in sample Sin crystallized during a single magmatic event, and that the zircon grains are free of older inheritance. Thus, the oldest Pb–Pb and concordant ages obtained from these zircon domains are nearest to the crystallization age. Eleven spot analyses define a discordia with an upper intercept at 2646 ± 10 Ma, and nine of those define an identical concordia age of 2647 ± 12 Ma, which is interpreted as the intrusion age of the Singelele granite gneiss (Fig. 3c). In contrast, all younger concordant and discordant zircons, which have similar initial 176Hf/177Hf values (0·28099 ± 0·00005, 2 SD) are assumed to have undergone Pb loss. Among them are seven that fall on a discordia with a lower intercept age of 1989 ± 160 Ma (Fig. 3c), indicating a Palaeoproterozoic Pb loss similar to that obtained for sample Bu (Fig. 3b). However, the large scatter of the discordant analyses indicates that most zircon domains have undergone multiple Pb loss. In this context, it should be noted that the Pb–Pb evaporation ages and discordant U–Pb ID-TIMS ages of 2568–2582 Ga for the Singelele gneiss (Jaeckel et al., 1997) could date the time of zircon alteration (Pb loss) rather than zircon formation.
In contrast to the zircon analyses discussed above, six zircon rim analyses, which show significantly higher initial 176Hf/177Hf values of up to 0·28166, do not follow the general Hf(t) trend and show significantly younger TDMHf ages extending to 2·2 Ga (Fig. 4g–i). A concordant U–Pb analysis of one of these rims (spot 27r/zrc3) indicates that it was formed at 2042 ± 19 Ma. This age is slightly younger than two Pb–Pb upper intercept ages of 2081 ± 20 Ma and 2091 ± 13 Ma, which were obtained for two other overgrowths (spots 15r/zrc19 and 20r/zrc23), but significantly older than three discordant Pb–Pb zircon ages of about 2196, 2444 and 2538 Ma (spots 54r/zrc38, 47r/zrc35 and 55r/zrc39). The three younger zircon ages suggest that these zircon overgrowths were formed during the Palaeoproterozoic, perhaps during the high-grade metamorphic event that was dated at 2·02–2·05 Ga (e.g. Barton & Sergeev, 1997; Jaeckel et al., 1997; Holzer et al., 1999). During this event the Singelele orthogneiss underwent partial anatexis, as evident from a few melt patches observed in the outcrops at the Singelelekop. The three older Pb–Pb ages could reflect mixing ages, which result from the analyses of older (2·65 Ga) and younger (2·04 Ga) zircon domains. These ages could not be resolved by the technique used. Consequently, their 176Hf/177Hf initial ratios should also represent mixtures.
The distinctly higher initial 176Hf/177Hf values of the six rims are thought to be due to partial melting of the Singelele orthogneiss at c. 2·04 Ga. During that melting event most of the c. 2·65 Ga zircon grains remained undissolved. Consequently, only minor amounts of Hf, incorporated by these zircon crystals (= zircon hafnium), were released into the melt from which the new zircon overgrowths were formed. Thus, the zircon overgrowths incorporated abundant radiogenic 176Hf formed by 176Lu decay in the granite gneiss matrix (= matrix hafnium) between 2·65 and 2·04 Ga. This assumption is best supported by spot 27r/zrc3, which shows nearly identical zircon crystallization and TDMHf ages of 2·04 and 2·2 Ga, respectively (Table 2 and 3; Fig. 4i). In contrast, the other five spot analyses yield less radiogenic Hf, thereby indicating that these zircon domains incorporated a mixture of highly radiogenic ‘matrix hafnium’, and weakly radiogenic ‘zircon hafnium’, released by partial dissolution of 2·65 Ga zircons during anatexis.
In summary, the combined U–Pb and Lu–Hf analyses of the zircon grains indicate that the protolith of the Singelele orthogneiss intruded at about 2646 ± 10 Ma and that the magma contained considerable amounts of an older crust, as reflected by average crustal residence ages between 3·2 and 3·5 Ga. The Singelele gneiss underwent partial anatexis during the Palaeoproterozoic at about 2·0 Ga, which caused Pb loss in many zircon domains and the formation of a few zircon overgrowths incorporating abundant highly radiogenic ‘matrix hafnium’.
Regina granite gneiss (Reg)
This gneiss sample was collected from the Regina farm c. 30 km north of Alldays, and contains mainly plagioclase, K-feldspar, quartz, amphibole, garnet, and magnetite. The Regina gneiss is part of the Krone metamorphic terrane forming the basement below the Venetia klippen complex, which underwent an amphibolite-facies overprint at about 2·06–2·0 Ga (Barton et al., 2003; Klemd et al., 2003; Zeh et al., 2005a, 2005b). So far, nothing is known about the timing of the orthogneiss protolith intrusions in this area. From sample Reg, 22 U–Pb spot analyses on 13 zircon grains and 11 Lu–Hf spots were analysed (Tables 2 and 3; Figs 3d and 4j–l).
The CL images indicate that most zircon grains have oscillatory zoning and some of them distinct cores and rims (Fig. 5a). Twelve U–Pb spot analyses yield similar 207Pb/206Pb ages of 2580–2660 Ma for the different domains, indicating that the grains formed during a Neoarchaean magmatic event. In detail, however, more than one age population seems to exist. Five grains yielded a concordia age of 2649 ± 9 Ma, which is identical to an upper intercept age of 2651 ± 6 Ma as defined by eight analyses (Fig. 3d). Three other zircon analyses, however, yielded a significantly younger concordia age of 2600 ± 11 Ma and one analysis a concordia age of 2530 ± 16 Ma. The remaining 10 spots gave 20–79% discordant results with 207Pb/206Pb ages that vary from 1684 to 2487 Ma. We interpret the older concordant age to represent the intrusion age of the Regina gneiss protolith, whereas the younger ages result from partial Pb loss either at about 2·0 Ga or during multiple events (inset in Fig. 3f).
All zircon grains have within error identical initial 176Hf/177Hf values of 0·28096 ± 0·00004 (2 SD), which correspond to Hf(t)initial of –3·4 to –5·5 and TDMHfinitial of 3·3–3·4 Ga. It is worth noting that these values are similar to that of the Bulai granite and to most values obtained from the Singelele gneiss.
Zanzibar granodiorite gneiss (sample ZAG)
The investigated grey gneiss sample was taken from the Seoka river bed near the Botswana–South Africa boundary (see Barton & Key, 1983). From this sample 23 U–Pb and Lu–Hf spot analyses were performed on 17 zircon grains (Tables 2 and 3; Figs 3e and 4m–o). CL images indicate complexly zoned zircon grains and most of them reveal a core and rim structure (Fig. 5b and c). In addition, there are abundant zircon grains that show parallel banding (Fig. 5b). U–Pb spot analyses provided a wide scatter of apparent Pb–Pb ages for all domains ranging between 2227 and 2629 Ma (Table 2). Some zircon grains yield within error identical ages for their core and rim (e.g. spots 9 and 10, Fig. 5c) whereas other zircon grains show apparently younger cores and older rims (e.g. spots 24 and 25, Fig. 5b). The latter indicates that the analysed core domains underwent a much stronger Pb loss than the rim domains of the same zircon.
Eight U–Pb analyses (99–101% concordance) yielded a concordia age of 2613 ± 7 Ma (Fig. 3e), which is interpreted as the time of granodiorite intrusion. Five apparently concordant zircon analyses (96–98% concordance) gave ages of around 2540 and 2596 Ma (Table 2). These ages are interpreted to result from partial Pb loss during a Palaeoproterozoic (c. 2·0 Ga) metamorphic event. The remaining eight spot analyses are 77–94% discordant, reflecting either partial Pb loss to zero or multiple Pb loss. It should be noted that the intrusion age of 2613 ± 7 Ma is significantly younger than the Rb–Sr whole-rock isochron age of 3227 ± 40 Ma from the same gneiss (Barton & Key, 1983). However, it is identical to or slightly younger than two SHRIMP U–Pb zircon ages of 2614 ± 13 Ma and 2659 ± 10 Ma of a granodiorite and monzonitic gneiss, respectively, which were collected from the nearby Tapalaphala river bed (Kröner et al., 1999).
Lu–Hf analyses reveal that all 23 zircon spots have within error identical initial 176Hf/177Hf values of 0·28108 ± 0·00005 (2 SD; Fig. 4m). This supports the conclusion that all investigated zircon domains were formed during the same magmatic event and that all of them underwent partial Pb loss during multiple events. This conclusion is also reflected by the array of the Hf(t) values, which shows a straight line between –1·2 at 2·62 Ga and –9·4 at 2·2 Ga (Fig. 4n). Initial hafnium model ages between 3·01 and 3·18 Ga reveal that the Zanzibar gneiss results from remelting of substantial amounts of an older Archaean crust, which was formed on average c. 400–500 Myr prior to the intrusion.
Zanzibar granite gneiss (sample ZAL)
The Zanzibar granite gneiss sample was also collected in the Seoka river bed. Twenty U–Pb and 15 Lu–Hf spot analyses were obtained on 16 zircon grains from this sample (Tables 2 and 3; Figs 3f and 4p–r). Many zircon grains have cores, which are separated from their rims by dark, U-rich bands (Fig. 5d). Most of the cores show oscillatory zoning, whereas some of the rims are structureless.
U–Pb spot analyses of all zircon domains yield discordant results (97–48% discordance) with apparent Pb–Pb ages between 1785 and 2619 Ma (Table 2). However, the initial 176Hf/177Hf values of all spot analyses are identical within error (0·28108 ± 0·00005; 2 SD). This indicates that all analysed zircon domains, comprising the different core and rim areas, must have been formed during the same magmatic event, which is also supported by the Hf(t) and TDMHf vs Pb–Pb age trends (Fig. 4q and r). The TDMHfinitial values indicate that the granite gneiss protolith stems from a crustal source with an average hafnium model age of 3·02–3·18 Ga, which is identical to that of the Zanzibar granodiorite.
Nine U–Pb spot analyses fall on a discordia (MSWD = 1· 9) with intercepts at 2614 ± 9 Ma and 716 ± 46 Ma (Fig. 3f). The upper intercept age of 2614 ± 9 Ma is interpreted to date the time of the ZAL granite intrusion. In fact, this age is identical to that obtained from the nearby sample ZAG, which shows an identical Hf isotope composition (see above). The scatter of various data points in the concordia diagram of sample ZAL indicates than many zircon domains underwent post-crystallization multiple Pb loss, perhaps during Palaeoproterozoic and recent time (see inset in Fig. 3f).
Mahalapye Complex
Mokgware granite (sample Mo)
The investigated sample represents a medium-grained granite variety, which contains mainly K-feldspar, plagioclase, biotite and quartz. Seventeen U–Pb spot analyses on 14 zircon grains and an 14 Lu–Hf spot analyses were carried out (Tables 2 and 3). CL images indicate different zircon zoning patterns (Fig. 6a–e). Type-I zircons show U-poor cores with a bright luminescence and ‘dark’ U-rich overgrowths (Fig. 6c and d). Some of these cores display oscillatory zoning, whereas others are structureless–diffuse (Fig. 6d). The latter are similar to recrystallized metamict zircons as described by Vavra et al. (1996). Type-II zircons show a balanced luminescence throughout the entire crystal with a weak oscillatory zoning, although some of them have distinct cores and rims (Fig. 6a, b and e).
U–Pb spot analyses of most zircon domains from type-I zircon and from the oscillatory cores of type-II zircon define a discordia with intercepts at 10 ± 12 and 2019 ± 8 Ma (Fig. 7a). The upper intercept age is within error identical to a concordia age of 2026 ± 10 Ma, which was calculated from three zircon analyses (Fig. 7a). It is interpreted to reflect the time of granite emplacement. In contrast, two younger, strongly discordant Pb–Pb analyses, which have apparent Pb–Pb ages of about 1463 and 1721 Ma, are interpreted to result from multiple Pb loss after 2·03 Ga. These two analyses show the highest U contents (1432 and 2832 ppm, respectively). In addition, a few zircon cores give significantly older ages. One zircon core yields a concordant age of 2929 ± 14 Ma (spot 8c/zrc6), and three others have discordant Pb–Pb ages between 2520 and 2645 Ma (Fig. 7a). The age data indicate that the Mokgware granite intruded at 2019 ± 8 Ma and contains some Archaean crustal components. This is in agreement with the results obtained from the Lu–Hf spot analyses (Fig. 8a–c), which reveal Archaean hafnium model ages between 3·0 and 3·4 Ga for the different zircon domains (Fig. 8c).
The 10 spots that define the 2·02 Ga Mokgware discordia age show slightly variable initial 176Hf/177Hf of 0·28120 ± 0·00005 (2 SD). This suggests that magma homogenization during zircon crystallization was incomplete (TDMHfinitial = 3·1–3·2 Ga; Hf(t)initial = –10·5 ± 1· 9). The inherited components with variable age (>2·02 Ga) and Hf isotope composition (Hf(at 2019 Ma) = –11· 3 to –20·1) clearly point to a heterogeneous crustal source.
It should be noted that the assimilation–melting of Archaean material during the Palaeoproterozoic granite formation in the Mahalapye Complex is not restricted to the Mokgware granite, but has also been reported from the nearby Mahalapye granite (Fig. 1). This granite intruded at 2023 ± 11 Ma and contains abundant zircon xenocrysts with Pb–Pb ages between 2·45 and 3·15 Ga (McCourt & Armstrong, 1998). The age of the oldest xenocryst is well within the range of the hafnium model ages, which are estimated in this study for the Mogkware granite.
Lose quarry samples
The Lose quarry is situated c. 50 km south of Palapye (Fig. 1) and exposes ortho- and paragneisses that are transected by garnet-bearing leucosomes. The leucosomes form veins of centimetre to several metres width, which occupy c. 25% of the outcrop volume. A detailed description of this outcroup, lithologies and metamorphic conditions has been given by Chavagnac et al. (2001) and Hisada et al. (2005). For this study two samples were collected, a dark garnet–biotite gneiss and a leucosome.
Garnet–biotite gneiss (sample Ma1h). This sample shows melanocratic and leucocratic domains on the centimetre scale, both of which contain garnet, biotite, plagioclase, quartz and gem quality zircon grains. CL images show typical magmatic zoning patterns such as oscillatory zoning, sector zoning and a similarly bright luminescence for all zircon grains (Fig. 6f). Twenty U–Pb spot analyses on 20 zircon grains and 15 Lu–Hf spot analyses were carried out (Table 2 and 3). All U–Pb analyses fall on a discordia with intercepts at zero and 2058 ± 5 Ma, which is identical to the concordia age of 2061 ± 6 Ma as defined by 17 analyses (Fig. 7b). The Lu–Hf analyses of all zircon grains yielded identical initial 176Hf/177Hf of 0·28145 ± 0·00002 (Fig. 8d), which corresponds to Hf(t)initial between 0·1 and –1·2 (Fig. 8e) and TDMHfinitial of 2·61–2·67 Ga (Fig. 8f).
Similar characteristic CL patterns, Hf isotope data and identical U–Pb ages of all zircon grains indicate that the protolith of the garnet–biotite gneiss was a magmatic rock, perhaps a diorite, which was emplaced at 2061 ± 6 Ma, and syn- or post-intrusively deformed and metamorphosed. Our LA-ICP-MS U–Pb age is identical to, but more precise than, the zircon SHRIMP U–Pb upper intercept age of 2053 ± 21 Ma, which was obtained by McCourt & Armstrong (1998) from a granodiorite dyke of the same outcrop. In contrast to McCourt & Armstrong (1998), we did not find zircon xenocrysts with ages between 2·6 and 3·19 Ga in our sample.
The proposed magmatic origin for the garnet–biotite gneiss contradicts the findings of Chavagnac et al. (2001), who designated texturally similar rocks as migmatitic paragneisses, which were interpreted to be due to in situ anatexis of metagreywackes. However, in such a scenario one would expect to find a more heterogeneous zircon population with different inherited cores (see Zeh et al., 2003). Nevertheless, some sedimentary rocks must have been present in the Lose quarry, as documented by sillimanite- and andalusite-bearing metapelites (Hisada et al., 2005).
Garnet-bearing leucosome (Ma1b). This sample was taken from a 5 m thick leucosome dyke, adjacent to sample Ma1h. It contains minor garnet in a leucocratic K-feldspar–plagioclase–quartz matrix. CL images of zircon grains reveal internal structures that are completely different from those from sample Ma1h (Fig. 6g and h). Most zircon grains have bright, U-poor cores with oscillatory zoning (co), which are surrounded by dark, U-rich rims (rd) (Fig. 6g and h). In some zircon grains the oscillatory zoning patterns of the cores are partly or completely replaced by patchy zoning patterns (cp), which can also be observed in some dark zircon overgrowths, where these patterns commonly start at the rims and grade inward (rp) (Fig. 6g and h). In other zircon grains such patterns are observed to form bands that penetrate both the cores and rims. Judging from these observations, we conclude that the patchy patterns result from recrystallization of metamict zircon domains, perhaps as a result of fluid–zircon interaction (see Rizvanova et al., 2000; Geisler et al., 2001). Most grains have euhedral cores (Fig. 6g and h), indicating that the dark U-rich rims represent zircon overgrowths and are not the result of alteration processes. The latter commonly cause the formation of reaction fronts, which cross-cut older zircon zoning patterns (e.g. Vavra et al., 1996).
Twenty-five U–Pb spot analyses on 20 zircon grains and 22 Lu–Hf spot analyses were performed (Tables 2 and 3). Most core U–Pb analyses, in particular those of oscillatory zoned domains (co), fall on a discordia with intercept ages at around 468 and 2707 ± 12 Ma (Fig. 7c). The latter age is within error identical to a concordia age of 2711 ± 11 Ma obtained from nine analyses. Seven spots, including two analyses of patchy zircon cores, yield somewhat younger Pb–Pb ages of about 2·47–2·62 Ga. They plot on a discordia with a lower intercept age of 2031 ± 120 Ma when forced through 2711 Ma, the age of the nine concordant spots. In contrast to the oscillatory zoned cores (co), no reliable U–Pb ages were obtained from the zircon rims (domains rd, rp), and from most of the patchy zircon cores (cp). These uranium-rich domains (1815–8477 ppm) yield strongly discordant U–Pb analyses, with Pb–Pb intercept ages between 2·35 and 1·1 Ga (Fig. 7c). This indicates that all of these domains underwent Pb loss, a feature commonly observed in uranium-rich zircons (e.g. Vavra et al., 1996; Mezger & Krogstad, 1997). Judging from the textural relationship we believe that the U-rich zircon overgrowths grew during leucosome formation and crystallization. This assumption agrees with the fact that the zircon grains in leucosomes are commonly uranium-rich (e.g. Rubatto et al., 2001; Zeh et al., 2003). Because the leucosome veins observed in this study cross-cut the garnet–biotite (diorite) gneiss, the zircon rims must be younger than 2·06 Ga. This agrees with a U–Pb monazite age of 2002 ± 10 Ma, and Sm–Nd garnet ages between 2023 ± 7 and 1989 ± 38 Ma, which were reported for the leucosomes (Chavagnac et al., 2001). Thus for Hf(t)initial and hafnium model age calculations we use an age of 2·0 Ga for the zircon rim analyses (Table 3).
Initial 176Hf/177Hf of the cores yield a wide scatter from 0·28095 to 0·28124 (Fig. 8g). This scatter is most pronounced for the >95% concordant analyses (mostly oscillatory zircon cores), where the Hf(t) values vary from –3·5 to 3·8 and the TDMHf values are between 2·9 and 3·3 Ga (Fig. 8h and i). In contrast, the Hf(t) and TDMHf data for the discordant, patchy zircon cores (cp) show well-defined positive or negative trends when correlated with the apparent 207Pb/206Pb age data (Fig. 8h and i). This indicates that the ‘discordant’ cores underwent multiple Pb loss, whereas their initial 176Hf/177Hf remained nearly unchanged. Thus, when these zircon analyses are recalculated with the ‘true’ zircon crystallization age of 2711 Ma their resulting Hf(t)initial (–0·6 to 2·6) and TDMHfinitial values (3·1 ± 0·3 Ga) are well within the range of the concordant analyses (Fig. 8h and i).
The Lu–Hf analyses of the uranium-rich overgrowths yield initial 176Hf/177Hf2·0 Ga of 0·28123–0·28139, indicating that they contain more radiogenic hafnium than the zircon cores (Fig. 8g), and that the rims crystallized during a later (anatectic) event than the cores in the same precursor rock, similar to the scenario proposed for sample Sin (see above). It should be noted that most hafnium model ages obtained from the zircon cores and rims are higher (TDMHfinitial = 2·8–3·2 Ga) than those previously estimated by Sm–Nd whole-rock analyses of leucosomes from the Lose quarry (TDMNd = 2·7–2·9 Ga; Chavagnac et al., 2001). This may have the same cause (open-system behaviour; see Moorbath et al., 1997) as previously discussed for the Sand River gneiss.
The combined U–Pb and Lu–Hf data indicate that the zircon cores in sample Ma1b were formed at 2711 ± 11 Ma. The wide scatter of the Hf(t)initial values may indicate that the cores were formed in different magmatic rocks, which contained different amounts of recycled Palaeoarchaean crust [Hf(t)initial = –3·5; TDMHfinitial 
3·3 Ga] and juvenile Neoarchaean crust [Hfinitial = +6·3; TDMHfinitial 
2·8 Ga]. During the Palaeoproterozoic (at c. 2·0 Ga) this heterogeneous crust was remelted and new zircon overgrowths formed.
The isotope dataset and the CL images indicate that the garnet–biotite (diorite) gneiss (sample Ma1h) is neither the palaeosome nor the restite of the abundant leucosomes exposed in the Lose quarry. It seems more likely that the leucosomes were formed by partial anatexis of a deeper crustal source, rather than from in situ anatexis as proposed by Chavagnac et al. (2001). In fact, the thick leucosome veins of the Lose quarry may represent magma conduits, which are genetically unrelated to the surrounding diorite gneiss. Our isotope data in combination with the field relationships from the Lose quarry support a model in which the diorite gneiss protolith results from Palaeoproterozoic (2·06 Ga) melting of a juvenile Neoarchaean lower crust formed at 2·6–2·7 Ga. In contrast, the younger garnet-bearing leucosomes resulted from Palaeoproterozoic melting of the upper crust, which was predominantly made up of pre-Neoarchaean material and affected by magmatism at c. 2·7 Ga.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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Zircon formation vs alteration
From our combined CL, U–Pb and Lu–Hf dataset we are able to discriminate between zircon growth and alteration processes. Zircon domains affected by single or multiple Pb loss yield initial 176Hf/177Hf values that are identical within error to those of unaffected zircon domains formed at the same time. This is generally reflected by zircon spot analyses obtained from the samples Bu, Sin, ZAG, ZAL and Ma1b, as the analyses of co-genetically formed zircon domains form horizontal lines in the 176Hf/177Hfinitial vs apparent Pb–Pb age diagram and show a positively correlated array in the
Hf(t) vs apparent Pb–Pb age diagram (Figs 4 and 8). In contrast, newly formed zircon overgrowths, which may result from partial anatexis of the orthogneisses during later events, always show significantly higher radiogenic 176Hf/177Hfinitial and Hf(t) than the older zircon domains. This is well demonstrated by zircon analyses from samples Sin and Ma1b (Fig. 4g and i), and agrees with results obtained from migmatitic gneisses from the Antarctic Peninsula (Flowerdew et al., 2006). Therefore, our results indicate that the hafnium isotopes once incorporated into a certain zircon domain will not (or only slightly) be fractionated during later metamictization and fluid-driven processes, which generally lead to partial or complete Pb loss in distinct zircon domains (e.g. Vavra et al., 1996; Rizvanova et al., 2000; Geisler et al., 2001). Increased initial 176Hf/177Hf values of zircon overgrowths with respect to zircon cores indicate that the overgrowths were formed during partial anatexis, with most zircon cores not being dissolved.
Timing of magmatism and metamorphism
The combined information from CL images, U–Pb and Lu–Hf isotope spot analyses provides evidence that zircon grains in granites and orthogneisses of the Limpopo Central Zone were formed during distinct magmatic events at about 3·28, 2·71, 2·65, 2·61, 2·06 and 2·02 Ga. In addition, our data indicate that some zircon domains were formed during anatectic events at about 3·14, 2·61 and 2·02 Ga, whereas others were subject to single and/or multiple Pb loss. Our age data combined with those from previous studies (e.g. Jaeckel et al., 1997; Holzer et al., 1998, 1999; Kröner et al., 1998, 1999; McCourt & Armstrong, 1998) indicate that granitoid rocks with ages of 3·28 Ga and 2·61–2·65 Ga are restricted to the Beit Bridge Complex, whereas granitoids with ages of 2·06 and 2·02 Ga occur only in the Mahalapye Complex (Fig. 1). However, zircon xenocrysts reveal that the Mahalapye Complex also comprises large quantities of Archaean magmatic rocks formed at 2·52–2·8 Ga and >2·9–3·15 Ga (McCourt & Armstrong, 1998, and references therein). In fact, most xenocryst ages of the Mahalapye granitoids are identical to orthogneiss protolith ages obtained from the Beit Bridge Complex. This also holds true for the zircon cores with an age of 2711 ± 11 Ma found in the Lose quarry leucosomes. Their ages are identical to U–Pb SHRIMP ages of 2718 ± 5 and 2717 ± 3 Ma obtained from granodiorite or dacite samples from the Beit Bridge Complex (Kröner et al., 1999).
A concordant 3143 ± 13 Ma zircon age and rare 2614 ± 11 Ma zircon overgrowths around 3·28 Ga zircon cores indicate that the Sand River gneiss of the Beit Bridge terrane was subject to partial anatexis during the Palaeo- and Neoarchaean, respectively. So far, the reason for the Palaeoarchaean thermal overprint, which is also suggested by the age data of Barton et al. (1977, 1990) and our hafnium isotope results (see above), is not clear. The Neoarchaean event was obviously induced by the voluminous granite intrusions in the Central Zone, comprising the Singelele gneiss (2647 ± 12 Ma), the Regina gneiss (2649 ± 9 Ma), the Zanzibar gneiss (2613 ± 7 Ma), and the Bulai pluton (2612 ± 7 Ma). In fact, these granitic intrusions occurred contemporaneously with the Razi granite of the Northern Marginal Zone (2627 ± 7 Ma; Mkweli et al., 1995; 2·52–2·67 Ga: mostly discordant ages of Frei et al., 1999), and the Matok granite of the Southern Marginal Zone (2620 ± 75: Barton et al., 1983; 2667–2664 Ma: Barton et al., 1992; 2643 ± 1 Ma: Kreissig et al., 2001).
Numerous U–Pb analyses provide evidence that new zircon overgrowths were formed during partial anatexis of the Sand River and Singelele orthogneisses at 2·02–2·04 Ga. In addition, the U–Pb analyses indicate that many zircon grains and/or domains in orthogneisses from the Beit Bridge Complex underwent Pb loss at the same time. This is well supported by data from the Bulai granite, the Singelele gneiss, the Regina gneiss and the Zanzibar gneiss (Fig. 3b–e). We assume that new zircon growth as well as partial or complete Pb loss at about 2·02 Ga is related to structural–metamorphic processes, which affected the Central Zone during the Palaeoproterozoic, in agreement with data and conclusions suggested by several other workers (e.g. Jaeckel et al., 1997; Holzer et al., 1998, 1999; Kröner et al., 1999; Zeh et al., 2004; Buick et al., 2006). During the Palaeoproterozoic parts of the Central Zone were affected by a high-grade metamorphic event, as confirmed by petrological data from the Messina, Alldays and Mahalapye areas (Van Reenen et al., 2004; Zeh et al., 2004; Hisada et al., 2005; Buick et al., 2006), whereas the area around the Venetia kimberlite pipes underwent amphibolite-facies metamorphism (Barton et al., 2003; Klemd et al., 2003; Zeh et al., 2005a, 2005b).
Crustal evolution of the Central Zone
Palaeoarchaean granitoids
The Lu–Hf isotope data indicate that the 3·28 Ga Sand River TTG suite results from melting of a crust with an average crustal residence age of TDMHfinitial = 3·64 ± 0·04 Ga. This indicates that the oldest crust of the Limpopo Central Zone was formed c. 300–400 Myr prior to the formation of the Sand River TTG suite. It is interesting to note that the c. 3·6–3·7 Ga hafnium model age is identical to a Pb–Pb SHRIMP age of a single detrital zircon grain from a pelitic schist of the Beit Bridge formation (Kröner et al., 1998). This suggests that the oldest crust constituting the Central Zone was recycled immediately after its formation. Nevertheless, a detrital zircon U–Pb SHRIMP age of 3·8 Ga, as reported by Armstrong et al. (1988), indicates that the Limpopo Belt may contain even older crust, which, however, is not confirmed by our data. The TDMHfinitial values obtained from the Limpopo Belt are similar to those calculated for felsic rocks of the Barberton greenstone belt (Amelin et al., 2000; TDMHf = 3·61–3·75 Ga), and to Sm–Nd isotope data for basic and ultrabasic rocks from the Kaapvaal Craton (Wilson & Carlson, 1989) and for felsic rocks of the Tokwe segment of the Zimbabwe Craton (TDMNd = 3·55–3·67 Ga, data from Moorbath et al., 1986; Taylor et al., 1991). Similarly to the Sand River gneiss area, the Tokwe segment also contains a minor 3·75–3·8 Ga zircon population (Dodson et al., 1988). This suggests that the Sand River TTG suite and some of the associated metasedimentary rocks could be part of the same Palaeoarchaean ‘granite–greenstone belt crust’, which now constitutes the cratons south and north of the Limpopo Belt. However, these similarities do not necessarily mean that these ‘granite–greenstone belt crusts’ were formed in a close geographical proximity to each other.
Neoarchaean and Palaeoproterozoic granitoids
The protoliths of all investigated Neoarchaean orthogneisses of the Beit Bridge Complex (Bulai granite, Singelele gneiss, Regina gneiss, Zanzibar gneiss) intruded between 2·65 and 2·61 Ga. Zircon grains from these granitoids show Hf(t)initial values between +0·5 and –7·1 (Fig. 9a and b), and TDMHfinitial of 3·01–3·46 Ga. Similar TDMHfinitial values are also obtained from the 2·5–2·9 Ga zircon xenocrysts of the Mokgware granite from the Mahalapye Complex. In contrast, the 2·71 Ga zircon cores from the Lose quarry leucosome of the Mahalapye Complex (sample Ma1b) show a wide range of Hf(t)initial from –3·7 to +6·3, and TDMHf between 2·76 and 3·32 Ga. The highest Hf(t)initial of these zircon grains is close to that of the depleted mantle (Fig. 9a).
The oldest TDMHfinitial age of 3·46 Ga is based on a zircon grain from the 2·65 Ga Singelele orthogneiss. This model age is only slightly younger than those obtained from the nearby Sand River orthogneiss (3·60–3·68 Ga), indicating that the Singelele gneiss may have formed by partial melting of Sand River gneiss equivalents at about 2·65 Ga. In fact, anatexis of the Sand River gneiss during the Neoarchaean is reflected by zircon overgrowths, which yielded U–Pb ages of about 2·61 Ga (Fig. 3a). The minor age differences between the Singelele granite gneiss intrusion (2·65 Ga) and the Sand River gneiss anatexis (2·61 Ga) may be explained by partial Pb loss of the analysed zircon overgrowths. Apart from the 3·46 Ga hafnium model age, most zircon grains of the Singelele gneiss yield somewhat younger TDMHfinitial values of 3·17–3·40 Ga (Fig. 9a), which correspond to higher initial 176Hf/177Hf and Hf(t)initial (Figs 4g, h, and 9a). It should be noted that the younger hafnium model ages of the Singelele gneiss are identical to those obtained from the Bulai pluton (TDMHfinitial = 3·17–3·28 Ga) and the Regina gneiss (TDMHfinitial = 3·25–3·37 Ga), but are predominantly older than those obtained from the Zanzibar granodiorite (TDMHfinitial = 3·01–3·18 Ga) and Zanzibar granite gneiss (TDMHfinitial = 3·04–3·18 Ga) (Fig. 9a and b).
In general, two models can explain the younger hafnium model ages of the Neoarchaean gneisses with respect to those of the Palaeoarchaean Sand River gneisses. According to the first model, abundant zircon was fractionated during partial melting of the c. 3·65 Ga ‘Sand River gneiss’ crust at 2·65 Ga and thereby causing the ascending magmas to have higher radiogenic hafnium (matrix hafnium) than the re-molten precursor rocks. It should be noted that this model does not require the addition of any juvenile magma at c. 2·65 Ga. If true, however, the Neoarchaean gneisses should have higher 176Lu/177Hf ratios than their Palaeoarchaean precursor rocks. As illustrated in Fig. 9b, the Neoarchaean gneisses should have whole-rock 176Lu/177Hf ratios between 0·018 and 0·030 if they are formed by partial anatexis of Sand River gneiss. These values are higher than the mean Archaean–Palaeoproterozoic crustal composition (176Lu/177Hf = 0·0113) as suggested by Taylor & McLennan (1985) and Wedepohl (1995). Unfortunately, there are no 176Lu/177Hf whole-rock isotope data available, which could prove or disprove this model.
The second model is based on the assumption that all of the Neoarchaean gneisses represent mixtures of remelted c. 3·65 Ga crust (Sand River gneiss equivalents) and a juvenile component derived from the depleted mantle at c. 2·65 Ga. So far, there is only one Rb–Sr isotope dataset, that of Watkeys & Armstrong (1985), which provides evidence for the existence of Neoarchaean juvenile magmas in the Limpopo Belt, whereas all the Lu–Hf and Sm–Nd data are ambiguous (Fig. 9a). The Rb–Sr dataset was derived from alkaline lamprophyre dykes, which transect the 2·61 Ga Bulai granite. These dykes yield an initial 86Sr/87Sr of 0·70228 and Rb–Sr and Pb–Pb ages of 2671 ± 163 Ma and 2641 + 125/–131 Ma, respectively (Watkeys & Armstrong, 1985). These data are interpreted to reflect the formation of the lamprophyres from a Neoarchaean primitive mantle source beneath the Central Zone of the Limpopo Belt, which was ‘fertilized’ by an incompatible trace element-enriched fluid derived from the dehydration of a slab, which subducted beneath the region at about 2·7 Ga.
Further evidence for the existence of Neoarchaean juvenile magmas underneath the Limpopo Belt is provided by our zircon Hf isotope data from the Mahalapye Complex. The first set of evidence comes from zircon grains of the 2·061 Ga garnet–biotite (diorite) gneiss of the Lose quarry, which shows hafnium model ages between 2·61 and 2·67 Ga (Fig. 9a). They indicate that the protolith of the diorite gneiss perhaps results from melting of a lower mafic crust, which was derived from the depleted mantle during the Neoarchaean. Identical initial 176Hf/177Hf ratios obtained from all zircon grains of sample Ma1h (Fig. 8d–f) may reflect that the source region of the diorite melt was homogeneous at 2·06 Ga. Alternatively, the homogeneity of the initial 176Hf/177Hf could also result from intense mixing between a Palaeoproterozoic (2·06 Ga) juvenile magma and a ‘heterogeneous’ Neo- to Palaeoarchaean crust (<2·6 Ga). This possibility, however, is considered to be less likely. In fact, there is no evidence for the formation (existence) of any juvenile Palaeoproterzoic magma throughout the Limpopo Belt and the adjacent Kaapvaal and Zimbabwe Cratons (see Fig. 9c). Even the rocks of the nearby mantle-derived Bushveld layered intrusion (Fig. 1), which has the same intrusion age as the Lose quarry diorite (2·06 Ga, Walraven & Hattingh, 1993; Buick et al., 2001), show Archaean neodymium model ages between 3·0 and 3·7 Ga (Maier et al., 2000), which are much older and more variable than the hafnium model ages obtained from the Lose quarry diorite.
The second set of evidence is provided by Hf(t)initial values of +6·3 to –3·7, as obtained from the 2·7 Ga zircon cores of the Lose quarry leucosome, which itself was formed during the Palaeoproterozoic at c. 2·0 Ga (Fig. 9a). The most positive Hf(t)initial is close to that of the depleted mantle (Fig. 9a) and, thus, suggests that zircon was formed in a juvenile magma at c. 2·7 Ga. The large scatter of the Hf(2·71 Ga)initial values of the zircon cores conforms with an interpretation that the leucosome of the Lose quarry was formed by anatexis of a mixed crust, which contained Palaeoarchaean (TDMHf >3·3 Ga) and juvenile Neoarchaean components (TDMHf = 2·7–2·8 Ga). Recycling of Palaeoarchaean crust during the Palaeoproterozoic is reflected by negative Hf(t)initial values of –7·1 to –12·6 obtained from zircon grains of the Mokgware granite (Fig. 9a).
Geotectonic model for the Neoarchaean–Palaeoproterozoic evolution
The formation of juvenile magmas at 2·6–2·7 Ga, as suggested from the Lose quarry data and the Bulai lamprophyes, is probably linked with the formation of the abundant Neoarchaean TTGs and greenstone belts, which constitute wide parts of the Zimbabwe Craton to the north of the Limpopo Belt (e.g. Hamilton et al., 1977; Moorbath et al., 1986; Taylor et al., 1991; Jelsma et al., 1996; Zhai et al., 2006). As shown in Fig. 9c, most magmatic rocks of these greenstone belts show positive to slightly negative Nd(2·6–2·7 Ga) values between +3·8 to –2·2, indicating the formation of new juvenile crust as well as the recycling of some older material during the Neoarchaean. Figure 9c also shows that the Nd(2·6–2·7 Ga) data from the Zimbabwe Craton overlap with those of the Northern Marginal Zone, which is interpreted to represent an active continental, Andean-type, margin at 2·7 Ga (Berger & Rollinson, 1997). In contrast, rocks from the Neoarchaean greenstone belts of the Kaapvaal Craton and the Southern Marginal Zone generally show lower Nd(2·6–2·7 Ga) values, indicating their different crustal evolution, which is also documented by their different Pb isotope data [see compilation by Barton et al. (2006)].
The occurrence of juvenile and mixed 2·6–2·7 Ga magmatic rocks in the Limpopo Central Zone is explained by a model that starts with the Neoarchaean subduction of juvenile oceanic crust underneath evolved pre-Neoarchaean crust (TDMHf of 3·65 Ga) of the Limpopo Belt. The absence of any juvenile Neoarchaean granitoids (TTGs) throughout the Limpopo Belt, including the marginal zones, may indicate that the exposed granite bodies are not the result of Neoarchaean juvenile slab-melting. Instead, it is more likely that the subducted slab underwent dehydration, which caused melting (and incompatible trace element enrichment by the released fluid) of the overlying subcontinental mantle wedge and led to the formation of juvenile Neoarchaean mafic magmas. Such a scenario conforms with the interpretation of Watkeys & Armstrong (1985) for the formation of the Bulai lamprophyres (see above). The juvenile mafic magmas, because of their high density, ponded at the crust–mantle boundary, where they caused magmatic underplating, which in turn triggered the melting of the overlying Palaeoarchaean crust. In consequence, the juvenile, mantle-derived mafic magma and evolved crustal materials were mixed to different degrees, leading to the formation of different types of magmatic rocks with distinct hafnium isotope composition at 2·6–2·7 Ga. Rocks with the oldest TDMHf values, such as the Singelele granite gneiss (3·3–3·5 Ga), perhaps represent nearly pure crustal melts, whereas rocks with younger TDMHf values, such as the Bulai granite and the Zanzibar gneisses (3·1–3·2 Ga), may contain a higher amount of juvenile components. The crust constituting the Mahalapye Complex was intruded by granitoids with the largest compositional variations, as reflected by the Hf isotope data of sample Ma1b (TDMHf = 2·8–3·3 Ga). In addition to the granitoid formation, some portions of the underplated mantle-derived mafic magma solidified without any significant isotopic modification, and formed a Neoarchaean lower crust. During Palaeoproterozoic heating at 2·06 Ga, this Neoarchaean mafic crust was remelted, leading to the formation of diorite magmas with TDMHf of about 2·65 Ga. Solidified relics of such a magma are now exposed in the Lose quarry in the Mahalapye Complex. The same heating event caused the melting of the isotopically more variable upper crust (TDMHf = 2·8–3·3 Ga), which led to the formation of the garnet-bearing leucosomes exposed in the Lose quarry and the final intrusions of the Mokgware and Mahalapye granites at about 2·02 Ga. So far, the reason for the Palaeoproterozoic heating event, which caused the high-grade metamorphic overprint throughout wide areas of the Limpopo Belt, is not understood. P–T paths, structural investigations and the geochronological data of Holzer et al. (1998, 1999), Kröner et al. (1999) and Zeh et al. (2004, 2005a, 2005b) indicate that the 2·02 Ga event is related to crustal convergence that was caused by the oblique collision between the Zimbabwe and Kaapvaal Cratons.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING, CURRENT...ANALYTICAL TECHNIQUESRESULTS AND INTERPRETATION OF...DISCUSSIONCONCLUSIONSREFERENCES |
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The results indicate that by combining the U–Pb with the Lu–Hf isotope system primary magmatic zircon domains can be distinguished from those formed during later metamorphic events, even if the distinct zircon domains underwent multiple Pb loss and the texture–age relationships, as obtained by CL images and U–Pb analyses, are ambiguous. Furthermore, the applied technique allows us to distinguish zircon grains formed in juvenile magmas from those generated by melting of older continental crust or affected by substantial crustal contamination.
Taking all data into account we conclude that most Palaeoarchaean to Palaeoproterozoic granitoids of the Central Zone of the Limpopo Belt result from major recycling of an old Archaean crust, which was derived from the depleted mantle at c. 3·65 Ga, and is equivalent to some of the oldest rocks constituting the Kaapvaal and Zimbabwe Cratons. At 3·283 Ga this old crust was at least partially remelted, leading to the formation of the Sand River TTG suite, which underwent a first anatectic overprint at 3·14 Ga. Subsequently, during the Neoproterozoic at 2·7–2·6 Ga, large quantities of the Palaeoarchaean crust were recycled again and mixed with new juvenile magma to different degrees, causing variable hafnium model ages between 3·0 and 3·45 Ga. The influence and amount of juvenile magmas seems to increase from east to west, as supported by increasing Hf(2·6–2·7 Ga)initial from the Sand River area of the Beit Bridge Complex (–7·1 to –1·8), to the Zanzibar area (–2·3 to +0·6) and to the Mahalapye Complex (–3·7 to +6·3) (Fig. 9a and b).
The coexistence of juvenile and mixed 2·6–2·7 Ga magmatic rocks in the Limpopo Belt is considered to result from the subduction of juvenile Neoarchaean oceanic crust beneath pre-Neoarchaean continental crust constituting the Limpopo Belt. This subducted crust underwent dehydration, triggering mantle wedge melting, and caused the underplating of mantle-derived mafic magma below the pre-Neoarchaean continental crust. This underplated juvenile magma triggered the formation of crustal melts and of mixed crust–mantle-derived melts, and parts of it formed a Neoarchaean mafic lower crust. Finally, the juvenile Neoarchaean lower crust and the mixed upper crust were remelted during the Palaeoproterozoic at 2·02–2·06 Ga.
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ACKNOWLEDGEMENTS |
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A.Z. thanks Hartwig Frimmel for stimulating discussions, Wolfgang Dörr for help with the sample preparation, and the German Science Foundation (DFG) for financial support (grants KL 692/15-2 and GE 1152/2-2). Furthermore, we thank Kent C. Condie, Jan Kramers and Hugh Rollinson for their constructive critical reviews of an earlier version of the manuscript.
*Corresponding author. Telphone: +49-931-888-5415. E-mail: armin.zeh{at}mail.uni-wuerzburg.de
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