Journal of Petrology

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Journal of Petrology | Volume 44 | Number 6 | Pages 1031-1054 | 2003
© Oxford University Press 2003

Melt Generation and Fluid Flow in the Thermal Aureole of the Bushveld Complex

NIGEL HARRIS^1,*, ANDY McMILLAN¹, MARIAN HOLNESS², RON UKEN³, MIKE WATKEYS³, NICK ROGERS¹ and ANTHONY FALLICK⁴

¹ DEPARTMENT OF EARTH SCIENCES, OPEN UNIVERSITY, MILTON KEYNES MK7 6AA, UK
² DEPARTMENT OF EARTH SCIENCES, UNIVERSITY OF CAMBRIDGE, CAMBRIDGE CB2 3EQ, UK
³ SCHOOL OF GEOLOGICAL AND COMPUTER SCIENCES, UNIVERSITY OF NATAL, DURBAN 4041, SOUTH AFRICA
⁴ SCOTTISH UNIVERSITIES ENVIRONMENTAL RESEARCH CENTRE, EAST KILBRIDE, GLASGOW G75 0QF, UK

E-mail: n.b.w.harris{at}open.ac.uk

RECEIVED JULY 10, 2002; ACCEPTED DECEMBER 3, 2002

	ABSTRACT

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Granite sheets emplaced into the migmatite zone of the eastern contact aureole of the Bushveld Complex resulted from fluid-enhanced, incongruent biotite melting of the underlying Silverton Formation shales during prograde metamorphism. Ba concentrations are extreme in both the sheets (>1000 ppm) and the hornfels (>800 ppm) into which they have been emplaced. We conclude that a Ba-rich, hydrothermal fluid induced melting in the aureole, and that fluid transport of Ba²⁺, and to a lesser extent, Sr²⁺ and Eu²⁺, persisted in the melt zones under subsolidus conditions. Sr-isotope systematics from high-Ba localities define an errorchron of 2161 ± 106 Ma with an initial (⁸⁷Sr/⁸⁶Sr) ratio of 0·705 ± 0·001. Metasedimentary rocks unaffected by fluid infiltration were homogenized at the same time but with an increased initial ratio, suggesting that whereas isotope homogenization was achieved between outcrops permeated by fluids, there is no evidence of regional homogenization. Oxygen-isotope compositions of psammitic metasediments in the aureole are uncorrelated with distance from the contact, suggesting the infiltrating fluid equilibrated isotopically with the metasediments. Their elevated ¹⁸O values (11·3–12·1) are consistent with a fluid source from devolatilization of the sedimentary lithologies. Textural analysis at both outcrop and thin-section scale of arkose and psammites in the aureole (Lakenvalei and Magaliesberg Formations) shows that the extent of melting was highly heterogeneous, even on the grain scale, and resulted from heterogeneously distributed infiltration of aqueous fluid on dilatant cracks and grain boundaries. Cathodoluminescence imaging of quartz shows a marked difference in the amount of fine structure, with that in the melted rocks having uniform luminescence in contrast to that in rocks containing little or no melt, which preserve textures inherited from the regionally metamorphosed protolith. Simple finite-difference thermal modelling of the aureole suggests that the width of the melt zone (>500 m) is inconsistent with conductive heat transfer, and hence that the thermal structure has been modified by fluid advection.

KEY WORDS: thermal aureole; crustal melting; fluid infiltration; isotope homogenization

	INTRODUCTION

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Thermal aureoles of mafic intrusions provide natural laboratories in which to study the formation and segregation of crustal melts under conditions of rapid prograde heating in the upper crust. In contrast to crustal melts formed during orogenesis, the high rates of prograde heating in thermal aureoles can lead to disequilibrium features between melt and its restite that carry kinetic information on crustal anatexis (Harris et al., 2000) specific to this mechanism of melting. Equally, as thermal aureoles allow both the heat source and the melt protolith to be precisely identified, they are an ideal environment for assessing the relative importance of protolith lithology, temperature and fluid fluxing in crustal melting.

The Rustenburg Layered Suite (RLS) of the Bushveld Complex of southern Africa is the world's largest reported layered mafic intrusion, ranging in thickness between 7 and 9 km and exposed over an area of 65 000 km² (Walraven, 1982). The magmas intruded the sediments of the Transvaal Supergroup (Fig. 1) at 2058·9 ± 0·8 Ma (Buick et al., 2001). Strong thermal metamorphism resulted in crustal melting close to the contact with the floor of the intrusion (Fig. 2). The highest temperatures associated with emplacement of the RLS were generated by the intrusion of the Bushveld Lower Zone magmas, which crystallized to form a 1500 m thick cumulate sequence of pyroxenites, norites and peridotites at temperatures between 1100 and 1300°C (Sharpe & Hulbert, 1985; Wallmach et al., 1995). Evidence for crustal melting is best preserved around the Burgersfort Bulge on the eastern contact of the Lower Zone exposed in the eastern lobe of the Bushveld Complex (Fig. 1). This study focuses on the petrogenesis of these wall-rock melts and assesses the role of fluid infiltration within the aureole.

View larger version (87K): [in this window] [in a new window]   Fig. 1. Geological map of part of the eastern Bushveld Complex and its aureole, with inset showing the location on a map of southern Africa. Long-dashed lines mark approximate isograds at 500°C (andalusite-in) and 600°C (outer limit of coarsely recrystallized quartzites); short-dash indicates 700°C isograd (outer limit of melting) taken from Button (1976), Sharpe & Chadwick (1982), and our own observations. A–B marks line of section (Fig. 2).

 

View larger version (23K): [in this window] [in a new window]   Fig. 2. Schematic SW–NE section across eastern Bushveld aureole. Line A–B from Fig. 1.

 

	THE AUREOLE OF THE BURGERSFORT BULGE

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
The immediate floor of the Lower Zone of the RLS is underlain by quartzites and psammites (Lakenvalei and Magaliesberg formations) and argillaceous metasediments and minor calc-silicates (Vermont and Silverton formations) of the Early Proterozoic Pretoria Group from the Transvaal Supergroup (Fig. 1). The outer margin of the aureole, 3·5 km below the RLS, is marked by the appearance of biotite and andalusite in argillaceous rocks near the lower contact of the basal Pretoria Group, indicating temperatures of 500°C (Sharpe & Chadwick, 1982; Waters & Lovegrove, 2002). As the intrusion is approached, a succession of metamorphic zones have been identified: chiastolite–chloritoid slate, andalusite shale, and andalusite–cordierite hornfels in pelites, and the onset of significant recrystallization of quartzite. The mineral assemblages indicate temperatures of 600°C (Button, 1976; Sharpe & Chadwick, 1982). The innermost zone, between 1 and 2 km wide in map view, is characterized by sillimanite-bearing hornfels, schists and migmatites, indicating temperatures >700°C (Willemse & Viljoen, 1970).

The highest-grade pelites in the aureole are defined by the assemblage alkali feldspar–sillimanite–cordierite, consistent with pressures of equilibration between 2·1 ± 0·4 kbar (Kaneko & Miyano, 1990) and 3·0 ± 0·5 kbar [obtained for garnet-bearing assemblages in the outer aureole by Waters & Lovegrove (2002)]. Recent mineral equilibria studies on a pelitic xenolith enclosed in granite overlying the RLS provide a minimum pressure estimate of 1·5 kbar for the base of the RLS (Pitra & De Waal, 2001). During the emplacement of this massive igneous complex, pressures will have increased progressively at the base of the intrusion, and it is possible that peak temperatures in the migmatite zone were experienced before peak pressure. Indeed, the emplacement of such a large igneous body will have imposed a complex and diachronous stress regime on the underlying rocks. Intense disharmonic folding (Fig. 3e) suggests an early period of compression and shear to the NE, probably associated with loading. Boudins within calc-silicate and quartzite horizons and a system of vertical east–west tension gashes within the pelites (Fig. 3f) mark a subsequent period of extension that may relate to sag of the lower contact (Uken & Watkeys, 1997).

View larger version (146K): [in this window] [in a new window]   Fig. 3. Field photograph of melting phenomena. (a) Anastomosing, interfingering of darker coarse-grained (F) and fine-grained quartzite (SA21); (b) brittle fracture of pelite horizon (dark) under extension with neck infilled by flow of quartzite (SA21); (c) nebulitic migmatite in melt zone in metapelites (SA22); (d) extended granite sheet (G) intruding psammites (SA22); (e) ptygmatic folding of granite vein intruding pelite with pegmatitic patches (P) developed within low-strain zones (SA15); (f) pegmatitic lens formed within tension gash (T) connecting fibrolite layers (V) oblique to primary foliation of pelite (SA20). Size of coin is 25 mm, size of lens cap is 500 mm.

 
Estimates of peak temperatures in the migmatite zone of the aureole are 750 ± 50°C (Nell, 1985; Kaneko & Miyano, 1990). Significantly, there is no evidence for the development of ultra-high temperatures in the aureole, in contrast to granulitic assemblages containing corundum, spinel and sapphirine preserved in pelitic xenoliths within the RLS (Willemse & Viljoen, 1970). Stable-isotope studies of the eastern Bushveld Complex suggest a vigorous hydrothermal system at temperatures of 350–700°C in the aureole (Schiffries & Rye, 1990; Buick et al., 2000). A ¹⁸O traverse of quartzite at the contact (Schiffries & Rye, 1990) is consistent with a metamorphic origin for the fluids, perhaps from the Transvaal Supergroup. ¹⁸O values of carbonates from the aureole (15–25) also suggest infiltration of a fluid derived from devolatilization of metasedimentary country rock (Buick et al., 2000).

Melting within the migmatite zone can be recognized at a range of scales from the grain scale in quartzites and psammites to the presence of metre-scale granite veins and sheets (Fig. 3d) that have segregated from their migmatitic pelitic source. The textures, geochemistry and significance of these melts are described below.

	ANALYTICAL TECHNIQUES

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Bulk-rock samples were collected from the localities indicated in Table 1 by drilling 2 cm diameter cores (length 3–5 cm). Whole-rock major and selected trace elements (Rb, Sr) were analysed on an ARL Fisons wavelength-dispersive X-ray fluorescence (XRF) spectrometer at the Open University. Major elements were determined using glass discs prepared by fusing powdered samples with Spectroflux 105. Trace elements (Rb, Sr) were determined on pressed powder pellets following the procedures described by Potts et al. (1984).

View this table: [in this window] [in a new window]   Table 1: Sample locations and mineral assemblages

 
Remaining trace elements were obtained by inductively coupled plasma mass spectrometry (ICP-MS) with an Agilent 7500 s system at the Open University. Fused glass discs were crushed and digested using standard HF–HNO₃ digestion techniques. Samples were introduced to the ICP-MS system as 2% HNO₃ solutions, with drift monitored by on-line spiking of Be, In, Tm and Re. Calibration was achieved using external and internal standards of known composition.

For Sr isotopic analysis, samples were digested in equal volumes of HF–HNO₃–HCl in sealed beakers using a CEM Mars^®5 microwave. Samples were dissolved at 210°C and 150 psi for a period of 30–60 min. Sr separates were obtained by standard ion exchange procedures and isotopic ratios were determined statically in multi-collector mode on the Nu Instruments^© MC-ICP-MS system at the Open University. Reproducibility of the NBS987 Sr standard over the analysis period was 0·710231 ± 0·000017 (2). Sr fractionation was corrected to ⁸⁶Sr/⁸⁸Sr = 0·1194.

Electron microprobe analyses were obtained with a Cameca SX100, operating at 20 kV accelerating voltage, with a 20 nA beam current and 10 µm beam size.

Samples for stable isotope analysis were crushed and sieved for picking in ethanol under a binocular microscope. Oxygen isotope analyses were carried out at the Scottish Universities Environmental Research Centre by the laser fluorination method of Sharp (1990) as modified by Macaulay et al. (2000). The precision at 1 is ±0·2 or better, and for NBS28 the value measured is +9·60. All data are reported as delta per mil (¹⁸O ) relative to V-SMOW.

Cathodoluminescence (CL) imaging of quartz was carried out using a JEOL JSM-820 scanning electron microscope running at an accelerating voltage of 15 kV, fitted with a retractable MonoCL detector supplied by Gatan UK. A parabolic mirror placed under the pole-piece 3 mm above the sample allows a high proportion of the photon flux to be collected and directed to a photomultiplier tube for panchromatic imaging. Such a system, together with the scanning technique, allows images to be obtained using a current of 3 nA, which minimizes electron beam damage. The minimum magnification for the MonoCL system is 100x and is constrained by the aperture of the paraboloidal mirror. The resultant panchromatic images show variations in luminescence intensity as shades of grey. Black indicates no (or very weak) lumines-cence, and white indicates very strong luminescence.

	CONTACT METAMORPHISM OF QUARTZ-RICH METASEDIMENTS

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Two psammite horizons occur in the high-grade metamorphic aureole in the Burgesfort area: the Lakenvalei Formation, which marks the contact itself at one outcrop (SA19), and the 500 m thick Magaliesberg Formation (Fig. 1), which is coarse grained and recrystallized over a distance of 1–2 km perpendicular to the contact (Button, 1976). Field observations from some outcrops within 100 m of the contact around the Burgesfort Bulge provide evidence for a high degree of extension-related deformation of psammite including flowage into the low-strain necks of extended pelitic horizons (Fig. 3b) and loss of sedimentary structures. However, loss of sedimentary structures such as ripple marks and cross bedding is not ubiquitous within the Magaliesberg Formation, even within a few metres of the contact. Indeed, sedimentary structures coexist with evidence of extreme deformation within a few metres in the same outcrop (SA22), illustrating the localized nature of psammite mobility. A related phenomenon is the presence of coarse fingers of virtually feldspar-free, coarse-grained quartzite within finer-grained psammite (Fig. 3a). These are structurally similar to the pipe-like fingers of coarse peridotite within a generally fine-grained peridotite massif that have been attributed to localized recrystallization related to fluid infiltration (Kostenko et al., 2002).

Samples of Magaliesberg and Lakenvalei Formation psammites and quartzites were collected within the aureole, at the contact with the RLS (SA19, SA23), and at 100–120 m from the contact (SA21, SA18). Magaliesberg Formation psammite was collected 3–4 km from the contact (SA05). Samples within the aureole containing relict sedimentary structures are SA21b and c, and SA23. Samples SA19a and SA21d show outcrop evidence for deformation and contact metamorphism-related destruction of sedimentary structures. Sample SA21d contains the coarse-grained feldspar-poor fingers. Table 2 summarizes these observations.

View this table: [in this window] [in a new window]   Table 2: Stable isotope composition (SMOW) of mineral separates from arkose and psammites

 
The psammites contain up to 20 vol. % feldspar (dominated by alkali feldspar) and muscovite. Significant post-peak grain growth and recrystallization has occurred, with abundant evidence for quartz grain-growth inhibition by feldspar (Fig. 4a). The feldspar grains have turbid, commonly albite-rich, margins, and the quartz contains abundant planar aggregates of fluid inclusions, suggestive of late-stage infiltration of aqueous fluids (e.g. Saigal et al., 1988; Lee & Parsons, 1997).

View larger version (125K): [in this window] [in a new window]   Fig. 4. Photomicrographs of quartzites under crossed polars. The scale bar in each image represents 100 µm. (a) Sample MW21 (collected from the same locality as SA21) showing a small rounded feldspar grain (arrowed, near top), which is pinning the quartz grain boundary and inhibiting grain growth. The high variability of quartz–quartz–feldspar dihedral angle on this length scale should be noted, with some feldspar grains exhibiting high angles, whereas nearby grains have elongate extensions along quartz grain boundaries owing to an extremely low dihedral angle. (b) Sample SA19a. Cuspate and attenuated feldspar aggregate, with many rounded quartz inclusions. This texture is interpreted to have crystallized during the solidification of a melt phase. The two attenuated grain boundary extensions of feldspar (arrowed) with very low dihedral angles should be noted. (c) Sample MW21. Contrast between a rounded feldspar grain (labelled f) included within a quartz grain with two grains of feldspar on the adjacent quartz grain boundary. The latter have very low dihedral angles (arrowed). (d) Sample SA5, collected from outside the contact aureole. The compact form of the feldspar with a high quartz–quartz–feldspar dihedral angle (arrowed) should be noted. This equilibrated in the solid state in the absence of any melt phase.

 
Petrographic observations
Regardless of composition, feldspar grains have two distinct forms within samples collected in the aureole. The first is attenuated and cuspate, with elongate extensions on quartz grain boundaries or into individual quartz grains (Fig. 4a–c). These shapes may be poly- or monocrystalline, and are similar to those observed in partially melted feldspar-bearing quartzites (e.g. Holness & Clemens, 1999).

The second form is more compact and equant, with grains either on quartz grain boundaries (Fig. 4d) or enclosed entirely within single quartz grains. The latter tend to be small and well rounded (Fig. 4a and c). The relative proportion of the two forms varies between samples, with SA19a containing predominantly attenuated grains and SA23 dominated by the more compact form. The significance of these two textural types is illustrated by measurement of the quartz–quartz–feldspar (qqf) dihedral angle. Under conditions of textural equilibrium, the minimization of internal energies results in the development of a characteristic angle subtended at the junction of a grain of one phase against two grains of another. This angle, the dihedral angle, is a function of the relative magnitudes of the two interfacial energies (Smith, 1964). The equilibrium value of the quartz–quartz–plagioclase angle is in the range 105°–110° (Vernon, 1968).

Although a 2-D cross-section through a three-grain junction will not necessarily give the true 3-D value of dihedral angle, the median of a large (>25) population of measured 2-D angles is within 1° of the true 3-D angle (Riegger & Van Vlack, 1960). For a single value of 3-D dihedral angle, a unimodal population of 2-D angles is expected, whereas a bi- or poly-modal distribution indicates the presence of two or more 3-D angles. Optical measurement of the qqf dihedral angle in the unmetamorphosed sample SA05 gives a median value of 105° ± 3° (Fig. 5). Errors are calculated using the method of Wilks (1962) as reported by Stickels & Hucke (1964). The median of the population measured for SA05 is within the range expected for the equilibrium qqf angle (Vernon, 1968), although the distribution of measured angles is slightly skewed from that expected for a rock with only a single 3-D value of dihedral angle (Fig. 5). The qqf angle in a late granitic vein from sample MW15a is shown in Fig. 5 for comparison. The distribution of angles in this sample has a unimodal peak and a median value of 72° ± 3°. As values of dihedral angle much lower than 100° are associated only with fluid-bearing geological systems [see review by Holness (1997)], low values of the qqf angle in this sample must have been inherited from a precursor melt phase (e.g. Laporte, 1994; Laporte & Watson, 1995).

View larger version (22K): [in this window] [in a new window]   Fig. 5. Quartz–quartz–feldspar dihedral angle populations of selected samples. The population sizes are 100, 100, 70 and 150 for samples MW15a, SA19a, SA21c and SA05, respectively.

 
The populations of observed angles in samples SA19a and SA21c show a peak at around 105° but also a significant proportion of low angles (see Fig. 5), with median values of 87° ± 3° and 95° ± 3°, respectively. In these samples, the low angles are associated with the attenuated forms, and the higher angles with the more compact feldspar forms.

In common with earlier studies (Platten, 1982; Hunter, 1987; Holness & Clemens, 1999; Clemens & Holness, 2000; Rosenberg & Riller, 2000), we suggest that the highly cuspate feldspar grains crystallized during the final stages of solidification of partially melted psammite, whereas the more compact forms represent unmelted feldspar. The field evidence for disappearance of sedimentary structures and the evidence for localized extreme deformation in the quartz-rich metasediments are also consistent with partial melting. The coexistence of the two forms of feldspar in any one sample, and the juxtaposition of sedimentary structures with structurally homogeneous rock in a single outcrop are thus indicative of incomplete melting that was heterogeneous on a range of scales from grain to outcrop.

The evidence for partial melting given above, together with the generally low temperatures within the aureole (Nell, 1985; Kaneko & Miyano, 1990; Johnson, 2001), points to the necessity for melting to have occurred in the presence of water, as dry quartz–feldspar rocks require extremely high temperatures to melt. Given the likelihood that the equilibrium fluid–quartz dihedral angle was >60° at the metamorphic peak (Holness, 1993, 1998), we suggest that fluid infiltration occurred along dilatant cracks and grain boundaries.

The amount of melt in the psammites at the metamorphic peak was likely to have been small, with the maximum possible melt volume a function of the amount of feldspar in the rock and the maximum temperature. Given that most of the samples contain <20 vol. % feldspar, we consider the maximum amount of melt was generally <50 vol. %. However, loss of sedimentary structures is likely to occur at low (<10 vol. %) melt fractions during deformation.

Several psammites were analysed for oxygen isotopes to compare unmetamorphosed arkose with partially melted psammite. In contrast to the study of Schiffries & Rye (1990) we found no pattern of decreasing ¹⁸O as the contact is approached. The quartz in all samples has ¹⁸O in the range 11·3–12·3 regardless of location or extent of melting (Table 2). This is consistent with small amounts of infiltrated fluid. Assuming fractionation values for ¹⁸O_qtz-px at 700–800°C from Matthews et al. (1983) and pyroxene data for the RLS from Harris & Chaumba (2001), fluids derived from the RLS would reset quartz values to 8. The ¹⁸O data in this study therefore preclude the RLS as a fluid source and support the interpretation that devolatilization of Transvaal Supergroup sediments is the likely source of the convecting fluids (Buick et al., 2000).

Cathodoluminescence of psammites
Details of recrystallization and fluid infiltration can be inferred from CL images of quartz (e.g. Seyedolali et al., 1997; Holness & Watt, 2001). The causes of luminescence in quartz are not well understood but result from a complex interplay of factors such as lattice defects, trapped electron–hole pairs, oxygen vacancies, and traces of impurities such as Al, Ti, Li, Na and OH (e.g. Marshall, 1988; Waychunas, 1988; Fisher et al., 1990; Fabre et al., 2002). In general, quartz grown at low temperatures (<200°C) has low luminescence (D'Lemos et al., 1997), possibly because the defect density and impurity concentration are lower than those in high-temperature quartz (Watt et al., 2000). The extreme sensitivity of CL to chemical heterogeneities and lattice defects makes it highly superior to conventional optical microscopy for distinguishing different quartz growth generations.

Three samples were chosen for detailed investigation. SA21c was sampled close to the contact but contains well-preserved sedimentary structures and has mainly compact feldspar grains. Sample SA19a has no sedimentary structures and contains attenuated feldspar grains with a low dihedral angle. Both characteristics are indicative of relatively voluminous partial melting. Sample SA21d was collected from near SA21c but has no relict sedimentary features, a significant population of low qqf dihedral angles (Fig. 5), and coarse-grained fingers of virtually feldspar-free quartzite.

Sample SA21c shows a great complexity of quartz textures under CL with at least six distinct generations, which can be assigned an age order based on cross-cutting relationships (Fig. 6). The youngest of these generations (generation 1) is dark and non-luminescent, forming cross-cutting planar features up to 20 µm wide, often with a dominant orientation, associated with fluid inclusion arrays and non-luminescent low-temperature feldspar alteration (Fig. 6a). The same dark, non-luminescent quartz is present at many of the present-day grain boundaries. Such features are commonly observed in both igneous and metamorphic quartz (Valley & Graham, 1996; D'Lemos et al., 1997; Seyedolali et al., 1997; Watt et al., 1997; Holness & Watt, 2001) and are believed to form during late-stage, low-temperature fluid infiltration along fractures and dilatant grain boundaries.

View larger version (203K): [in this window] [in a new window]   Fig. 6. CL images of SA21c. (a) The white patches are unaltered feldspar. Their darker rims should be noted—this is turbid altered material formed during interaction with low-temperature aqueous fluid. Non-luminescent quartz, crystallized during healing of late-stage low-temperature fractures, outlines the present-day grain boundaries and forms irregular planar features that cut across all other textures. The arrows mark regions of very weakly luminescent quartz, parallel to irregularities in the grain boundary, which immediately pre-date the last non-luminescent generation and represent the late stages of grain growth. (b) Oscillatory zoning caused by grain growth during episodic fluid infiltration on grain boundaries. An earlier grain has been totally consumed by growth of the surrounding grains. The evidence for grain boundary pinning by the particle of feldspar should be noted. The left-hand grain contains a network of fine bright planar features thought to be formed during healing of high-temperature microfractures. (c) Pleated grains with a central homogeneous region to the lowermost grain. (d) Homogeneous (h) or mottled centres to grains with pleated margins (p) and outermost oscillatory zoning (o) representing later grain growth. Non-luminescent, late-stage, quartz (n) fills fractures, and outlines the present-day grain boundaries. The scale bar in each image represents 100 µm.

 
The next youngest quartz generation (generation 2) is weakly luminescent and is closely associated with generation 1 quartz on grain boundaries. Generation 2 quartz alternates on either side of irregular grain boundaries and represents quartz recrystallized during the later stages of grain growth (Fig. 6a and b).

These two youngest generations cut sets of oscillatory zoned quartz (generation 3) (Fig. 6b). Such oscillatory zoning is frequently present in phenocrystic magmatic quartz (Seyedolali et al., 1997; Watt et al., 1997; Peppard et al., 2001), in granite plutons (Valley & Graham, 1996; D'Lemos et al., 1997; Seyedolali et al., 1997) and in quartz grown in open vein systems (Wilkinson & Johnston, 1996; Penniston-Dorland, 2001). A similar occurrence found in contact metamorphosed quartzite has been ascribed to grain growth during episodic fluid infiltration on grain boundaries (Yardley et al., 1991; Holness & Watt, 2001).

The fourth type of luminescence (generation 4) has not been previously described in the literature. It resembles a pleated fabric, with a fine network of interconnected pairs of bright and dark lines in two dimensions (Fig. 6c). Quartz of generation 4 is predominantly found at the edges of the present-day grains. Rarely, grains are cut by a fine network of bright features (generation 5) similar to the Bright Vein Networks (BVN) of Watt et al. (2000) (Fig. 6b). These are thought to be healed micro-cracks formed during internally generated overpressure related to reaction.

The oldest generation (6) of quartz is mottled or homogeneous and occurs predominantly at the centres of grains (Fig. 6c and d). Following Lind (1996) and Holness & Watt (2001) we suggest it represents quartz that retained its regional metamorphic signature (Seyedolali et al., 1997), remaining unrecrystallized during contact metamorphism. Adjacent grains show varying levels of luminescence (Fig. 5a)—this is related to the variation of luminescence intensity with crystallographic orientation (Walderhaug & Rykkje, 2000).

In contrast to the detail visible in sample SA 21c, SA19a shows only three distinct quartz generations (Fig. 7). Much of the sample comprises completely homogeneous, weakly luminescent quartz, cut by the non-luminescent late-stage fractures (generation 1) and Bright Vein Networks (generation 5). The latter appear to be slightly more common in the vicinity of feldspar (Fig. 7c). Very rarely the centres of large quartz grains contain a fine tracery of bright irregular planar features (Fig. 7d) that are highly reminiscent of the ‘mesh texture’ of Lind (1996), interpreted as unrecrystallized regional metamorphic relicts (generation 6; Lind, 1996; Holness & Watt, 2001).

View larger version (135K): [in this window] [in a new window]   Fig. 7. CL images of SA19a. (a) Feldspar-rich region showing typical shape inherited from solidification of melt. The quartz shows virtually no internal structure in CL. The two circular features at the centre of the image are due to radiation damage caused by a small accessory phase particle out of the plane of the image. The alteration of the feldspars to non-luminescent material and the tracery of non-luminescent healed fractures in the quartz should be noted. (b) Even though the feldspar is highly attenuated with very low dihedral angles the nearby quartz is featureless apart from late-stage healed fractures. (c) Tracery of fine bright planar features (arrowed) caused by healing of high-temperature fractures. The feldspar grain in the lower central portion has been completely altered at low temperatures and has lost all luminescence. (d) The top right-hand corner shows a tracery of bright lines probably marking a rare patch of unrecrystallized quartz. The common orientation (shown by double-headed arrow) of the late-stage fractures should be noted. (e) and (f) CL images of the coarse-grained fingers in sample 21d showing mottled textures and traces of oscillatory zoning parallel to the present-day grain boundaries [the ends of which are arrowed in (f)] owing to grain growth during episodic fluid infiltration. Unusually, in (f) these present-day grain boundaries are not marked by late-stage non-luminescent quartz. The scale bar in each image represents 100 µm.

 
Sample SA21d shows similarities to both SA21c and SA19a. The regions between the coarse-grained quartz fingers are indistinguishable from SA21c, whereas the fingers themselves show a reduced degree of detail reminiscent of SA19a. The fingers display the dark non-luminescent fracture network of generation 1, occasionally associated with the growth of weakly luminescent generation 2 quartz on grain boundaries (Fig. 7e). Pleating is rare and poorly developed, although much of the quartz shows a mottled texture. Faint oscillatory zoning is present, with zones parallel to the present-day grain boundaries (Fig. 7f).

The above interpretation of the CL images is necessarily tentative. It is notable that the sample in which sedimentary features are preserved displays the greatest amount of detail and the greatest number of distinguishable quartz generations, whereas the sample that shows melting-related deformation in outcrop has very little record of a complex recrystallization history.

The significant proportion of low measured dihedral angles in SA21c (Fig. 5) demonstrates that this sample did contain some melt at the metamorphic peak. A similar study of partially melted arkose has shown that solidification results in the formation of uniformly bright quartz at the present-day grain boundaries (Holness & Watt, 2001). This is not apparent in any of the Bushveld samples imaged. It is possible that the previously undescribed pleated texture is related to melt solidification but this is far from certain. It is perhaps more likely that the amount of melt formed in SA21c was small and is not distinguishable using CL. This is consistent with the lack of disruption of sedimentary structures such as cross-bedding and points to spatially highly variable amounts of melting associated with well-focused fluid infiltration.

The loss of fine CL structure in the fingers of SA21d is possibly attributable to recrystallization at high temperatures. Low-temperature growth would probably have resulted in loss of CL intensity (D'Lemos et al., 1997). We suggest the fingers represent the pathways for focused sub-solidus fluid infiltration closely following the metamorphic peak (see Kostenko et al., 2002). This is consistent with a slightly lower oxygen isotope composition for the fingers relative to their matrix (Table 2).

	GRANITE SHEETS AND VEINS

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
The largest granite body recognized in the aureole is a biotite granite sheet, 50 cm wide, intruded into migmatized Silverton Formation shales containing quartz–biotite–alkali feldspar–plagioclase–cordierite assemblages (Fig. 3c and d). The granite was emplaced 100 m from the contact with the RLS, separated from the latter by the Magaliesberg quartzite (Fig. 1). The sheet has been slightly boudinaged. Twenty metres downslope of the sheet, nebulitic leucosomes enclose undisturbed pelitic schlieren suggesting in situ melting (Fig. 3c). The leucosomes have not been analysed because of their marked inhomogeneity resulting from entrainment of restite (biotite, quartz, sillimanite) and peritectic (cordierite and possibly alkali feldspar) phases. In addition, migmatite leusosomes often incorporate cumulate phases from an extracted melt (Sawyer, 1987; Ellis & Obata, 1992; Otamendi & Patiño Douce, 2001) and back-reaction between melt and host rock is probable during crystallization (Fourcade et al., 1992; Kriegsman & Hensen, 1998). These factors conspire to make the interpretation of leucosome compositions equivocal and so this study focuses on the extracted melt, represented by the sheet, which provides a homogeneous mineralogy.

The granite sheet is medium grained (1 mm grain size), comprising quartz, oligoclase (An₁₂), microcline and subsidiary muscovite and biotite with an interlocking texture. The bulk of the feldspar in the adjacent psammitic hornfels has attenuated and cuspate shapes indicative of in situ melting (e.g. Fig. 4c). In addition to this body, examples of small-scale granitic intrusions include early granitic veins (2–10 cm wide) that have been strongly deformed resulting in disharmonic ptygmatic folding (SA13a, Table 3), and an undeformed post-deformation vein (SA22e, Table 3).

View this table: [in this window] [in a new window]   Table 3: Bulk-rock major and trace element compositions

 
Trace-element abundances of both the granite sheet and the veins show broadly similar trends. Compared with a peraluminous orogenic granite from the Himalaya, Rb is depleted but Sr and Ba are enriched (Fig. 8). High field strength elements are similar for both granite types, with the notable exception of Zr and Hf, which may reflect an inherited zircon component in the granite sheet.

View larger version (19K): [in this window] [in a new window]   Fig. 8. Normalized trace-element patterns of the granite sheet (MW15b, SA22a) and veins (SA13am, SA22e) relative to an average Himalayan biotite leucogranite (Ayres & Harris, 1997).

 
Modelling trace-element abundances in the granite melt
At the minimum pressure of 2–3 kbar (Kaneko & Miyano, 1990; Waters & Lovegrove, 2002), melting of pelites in the presence of aqueous fluid can occur at temperatures of 650°C (Le Breton & Thompson, 1988). However, the mineral assemblages of the metasediments in the aureole indicate that the granites formed at higher temperatures (>700°C). Under these conditions, melts may form in pelitic or semi-pelitic assemblages by fluid-absent melting of muscovite and/or biotite, as well as by fluid fluxing (Fig. 9). Major-element compositions, when plotted in the granite ternary system (Fig. 10), suggest that compositions are depleted in albite compared with granites formed by the muscovite-dehydration reaction.

View larger version (31K): [in this window] [in a new window]   Fig. 9. Pressure–temperature diagram showing melting relations for pelitic assemblages at low pressures. 1, H2O-saturated pelite solidus (Le Breton & Thompson, 1988); 2, muscovite dehydration reaction (Patiño Douce & Harris, 1998); 3, biotite, sillimanite, plagioclase, quartz dehydration solidus (Le Breton & Thompson, 1988). Dashed line indicates approximate melt fractions (F) obtained from melting a biotite gneiss with the addition of 2% H2O (Holtz & Johannes, 1991). Box indicates conditions of formation of granite sheet.

 

View larger version (16K): [in this window] [in a new window]   Fig. 10. Qtz–Ab–Or normative plot for granite sheet, veins and pegmatite lenses; •, experimental melt compositions from melting a biotite gneiss at 3 kbar, 800°C and 1–2% added H2O as indicated by numerals (Holtz & Johannes, 1991); *, biotite dehydration of a cordierite gneiss at 5 kbar, 800°C (Koester et al., 2002). Vertically shaded field shows Himalayan granite compositions resulting from muscovite-dehydration melting of a pelitic schist (Ayres & Harris, 1997).

 
Coarse poikiloblastic cordierite is widespread in migmatites, hornfels and schists and is likely to result from subsolidus conditions. However, fine-grained euhedral cordierite is restricted to the anatectic migmatites, suggesting that it is a solid product of the melt reaction. As H₂O-saturated melting does not generate peritectic phases, the migmatite assemblages suggest that biotite breakdown, under fluid-absent conditions, is the primary mechanism for melt formation.

The most fertile assemblage at the onset of meltingwill be biotite–quartz–alkali feldspar–plagioclase–sillimanite, found in metagreywacke or semi-pelitic lithologies, resulting in a melt reaction of the general form

Experimental melting of a cordierite gneiss by biotite dehydration at 5 kbar (Koester et al., 2002) results in melt compositions that are enriched in quartz relative to the granite sheet (Fig. 10), owing in part to the higher pressures of the experimental melts compared with pressures during anatexis in the Bushveld aureole. At 3 kbar, comparable with aureole pressures, the addition of H₂O at constant temperature decreases the alkali feldspar component in the melt and increases the melt fraction (Fig. 10). Experimental data on the anatexis of a biotite gneiss at 3 kbar (Holtz & Johannes, 1991) suggest that the granite sheet may have been generated by biotite melting, triggered by the influx of 2% H₂O. The precise value of added H₂O required to generate these melt compositions will depend on the mineral compositions and modal abundances of biotite and feldspar in the protolith. However, it should be noted that such low percentages of fluid infiltration are insufficient to allow an aqueous fluid and melt to coexist.

To assess whether such a melt reaction can generate the observed trace-element composition of the granite it is necessary to establish the reaction stoichiometry. For typical feldspar, biotite and cordierite compositions in the hornfels (Table 4), and the bulk analyses of the granite sheet (Table 3), the following melt reaction is obtained by mass balance:

(1)

Only minor cordierite can be generated by this reaction, as a result of the high Fe and Mg abundances in the melt. Equilibrium non-modal batch melting was modelled using equation (4) from Harris & Inger (1992):

(2)

where C_l is the concentration of the trace element in the liquid, C₀ the concentration in the source, F is the melt fraction and D the bulk distribution coefficient of the trace element for the restite assemblage. D is calculated from D = _i X_iKd_i, where Kd_i is the partition coefficient between mineral i and melt.

View this table: [in this window] [in a new window]   Table 4: Representative mineral analyses from pelitic sample SA20a

 
Mass-balance constraints for melting of the Timeball Hill Formation shales (basal Pretoria Group) from the Bushveld contact aureole in the Potgietersus region to the NW of the study area suggest a value of F = 0·65 (Nell, 1985). This is probably too high for the metasedimentary rocks and granite compositions observed in this study as there is insufficient alkali feldspar in the unmelted rocks to allow a melt fraction >0·3. This compares with melt fractions of 0·20–0·35 from the experimental melting of a pelitic gneiss at 750°C, 3 kbar and 2% added H₂O (Fig. 9). A larger melt fraction will result in depletion of the source in alkali feldspar, thereby restricting the K₂O of the melt to <4% (Holtz & Johannes, 1991) in contrast to the observed values of 5% in the granite sheet. However, this may explain the origin of the late vein SA22e, which has low K₂O but high Zr, possibly indicative of higher temperatures than the granite sheet.

Rb, Sr and Ba geochemistry
The geochemistry of Rb, Sr and Ba is highly influenced by the behaviour of protolith feldspar and mica, into which they are strongly partitioned. Mica and feldspar are the main reactant phases during melting, and if we assume a melt fraction of 0·3, the metasediments will generate a modal restite assemblage similar to that in Table 5.

View this table: [in this window] [in a new window]   Table 5: Outputs from melt modelling

 
For quantitative modelling of trace-element behaviour during equilibrium melting, the appropriate partition coefficients (Kd) must be known. The uncertainties in the equilibrium partition coefficients of Rb, Sr and Ba between micas or feldspars and granite melts propagate significant uncertainties in modelled melt compositions. Published data have been derived from experimental studies (Blundy & Wood, 1991) and empirical studies of volcanic glasses (Mahood & Hildreth, 1983; Nash & Crecraft, 1985; Ewart & Griffin, 1994; Icenhower & London, 1995, 1996) and migmatites (Bea et al., 1994). The latter are the most problematical in that diffusion of Rb, Sr and Ba is likely during both crystallization and subsolidus cooling of the melt, leading to lower-temperature re-equilibration (Nabelek, 1999). Variation in published Kd values is partly a result of the sensitivity of partitioning to volatile content, to the degree of polymerization in the melt (Ewart & Griffin, 1994) and to the crystal chemistry of the restite phases (Blundy & Wood, 1991). Ideally, partition coefficients should be determined for a melt of similar composition to the granite, that contains phenocrysts of the mineral of interest and that has been rapidly cooled from temperatures equivalent to the solidus temperature of the granite. Also important is the role of H₂O, which can act as a network modifier or behave as a charge-balancing cation (Kohn, 2000).

For modelling melting under disequilibrium conditions, such that the melt inherits the trace-element characteristics of the reactant minerals, we have assumed the equation of Harris et al. (1993):

(3)

where x_i is the proportion of mineral i entering the melt, X_i is the proportion of mineral i in the source rock and D_j/i is the distribution of the trace elements between phases i and j. D values have been estimated from a typical pelitic schist (Ayres & Harris, 1997).

For F = 0·3 and the range of partition coefficients given in Table 6, the values for C_l/C₀ under both equilibrium and disequilibrium conditions are given in Table 5. To calculate abundances of trace elements in the melt a protolith composition (C₀) must be assumed. For a model protolith we have assumed the composition of MW16 with Rb, Sr and Ba abundances characteristic of metasedimentary values throughout the aureole where melting is not observed.

View this table: [in this window] [in a new window]   Table 6: Partition coefficients used in equilibrium melt model

 
The resulting concentrations of Rb, Sr and Ba in the modelled melt, assuming a melt fraction of 0·3, are given in Table 5. For the pelites and psammites sampled within 10 cm of the granite sheet, or in the proximal migmatites that source similar granites, anomalous concentrations of both Ba and Sr are observed, exceeding 800 and 200 ppm, respectively (Fig. 11). As there are no indications that these lithologies have unusual abundances of detrital feldspar, this suggests metasomatic enrichment of these elements.

View larger version (13K): [in this window] [in a new window]   Fig. 11. Sr vs Ba from high-grade metasedimentary rocks within the aureole. MW15/SA22 is the granite sheet locality.

 
Although modelled trace-element concentrations are relatively insensitive to variations in melt fractions between 0·2 and 0·4, the uncertainties in the relevant partition coefficients (Table 6) generate large variations. Nevertheless, the following conclusions can be drawn. For Rb, trace-element behaviour is consistent with equilibrium melting, but the abundances of Sr and Ba are inconsistent with either equilibrium or disequilibrium melting, suggesting open-system behaviour for these elements. In other thermal aureoles, Ba anomalies mimic oxygen-isotope patterns (Nabelek & Labotka, 1993), indicating that Ba is particularly mobile in hydrothermal fluids and thus providing evidence for high fluid/rock ratios. Under subsolidus conditions, Ba will partition strongly into alkali feldspars from a hydrothermal fluid at 2 kbar and 700–800°C, similar to conditions that prevailed in the Bushveld aureole (Carron & Lagache, 1980). However, it is unlikely that the limited fluid required for melting (2–3%) was solely responsible for the observed Ba enrichment, as this suggests unfeasibly high Ba concentrations in the fluid (3–4% Ba). It is, therefore, probable that metasomatism persisted under sub-solidus conditions during cooling.

Sr has similar distribution coefficients to Ba between fluid, magma and feldspars (Carron & Lagache, 1980). Sr abundances in the metasedimentary rocks are partially determined by the abundance of plagioclase, but a plot of Sr vs Ba suggests that Sr is also enriched in schists close to the granites (Fig. 11). Anomalously high levels of Ba and Sr in the melt are suggested by their strong positive anomalies relative to orogenic granites formed under fluid-absent conditions (Fig. 8). We suggest that advection of Ba- and Sr-bearing fluids promoted melting. The country rock was metasomatized both during crystallization of the Ba-rich granite by late-stage fluid exsolution and during subsolidus metasomatism. The system was therefore open to both elements, overprinting any disequilibrium signature that may result from rapid melting.

Rare earth element geochemistry
Chondrite-normalized rare earth element (REE) plots of Silverton shales from outside the aureole (Fig. 12a) display a negative Eu anomaly, typical of post-Archaean shales, and are indistinguishable from pelites in the aureole (Fig. 12b), suggesting that solid-state metamorphism has no appreciable effect on the REE concentration. However, those metasediments sampled from the locality of the granite sheet display a reduced negative Eu anomaly (Fig. 12c). REE concentrations in the granite are lower than in their protoliths, and are characterized by a positive Eu anomaly in contrast to the marked negative anomaly of Himalayan granites that result from dehydration of metapelitic lithologies (Fig. 12d).

View larger version (24K): [in this window] [in a new window]   Fig. 12. Chondrite-normalized plots for REE: (a) shales, (b) metapsammites, (c) metapsammites near sheet (MW15/SA22); (d) granite sheet and veins with shaded field shown for Himalayan leucogranites (from Ayres & Harris, 1997).

 
Although REE concentrations in basaltic melts are controlled by partitioning between major phases and melt, for felsic magmas, accessory phase dissolution is the dominant control (Ayres & Harris, 1997). This can be illustrated by extending the modelling of Rb, Sr and Ba to the REE to generate REE abundances in the melt from a semi-pelitic source, assuming a melt fraction of 0·3 and the partition coefficients given in Table 6. The results (Fig. 13) demonstrate that if REE concentrations result from major-phase reaction and subsequent equilibration, then, compared with the granites in the aureole, the predicted REE patterns (1) have much higher REE abundances than are observed and (2) have a negative, rather than a positive Eu anomaly. The underlying cause for the discrepancy between modelled and observed data is that a significant proportion of the REE in the metasedimentary source is sequestered in refractory accessory phases that retain these elements throughout melting and melt extraction. Thus only a small proportion of the total REE budget is available to the melt through dissolution of these phases.

View larger version (14K): [in this window] [in a new window]   Fig. 13. Modelled chondrite-normalized REE plots and profile for granite SA22a. Equilibrium melting assumes REE concentrations are entirely controlled by equilibration with major phases in the restite. Disequilibrium melting assumes that 25% of the Eu in the source resides in feldspars (Ayres & Harris, 1997) and that there is no dissolution of accessory phases.

 
The positive Eu anomaly in the granite sheet might be explained by rapid melting and melt extraction, as the REE profile resulting from disequilibrium melting of a plagioclase-bearing source has a strong positive Eu anomaly derived from the reacting plagioclase (Prince et al., 2001). The extent of the anomaly depends on the proportion of the Eu budget in the source that is in accessory phases, but for a typical pelitic schist 25% of the Eu occurs in major phases (principally feldspars). Disequilibrium melting of such a source would yield the REE profile illustrated in Fig. 13. Hence the REE abundances in the granite could result from the combination of disequilibrium melting of major phases and accessory phase dissolution. However, it is also possible that melting did not occur in a closed system for REE. We have shown that the system was open to fluids transporting divalent ions such as Sr²⁺ and Ba²⁺. Given the geochemical similarities between Sr²⁺ and Eu²⁺, it is likely that these fluids also transported Eu²⁺. Evidence for Eu mobility may be derived by comparing the Eu/Eu^* ratio for metasediments that have experienced Ba and Sr metasomatism (0·62–1·03, Fig. 12c) with the ratio from other aureole metasediments (0·54–0·74, Fig. 12b).

	PEGMATITIC LENSES

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
The most widespread evidence for possible melt formation in the aureole is provided by coarse-grained quartz–fibrolite lenses that develop within low-strain areas of the deformed schists (Fig. 3e), or fill tension gashes connecting anastomosing veins of fibrolite (Fig. 3f). The lenses are made up of quartz and fibrolite, tourmaline, muscovite ± biotite, cordierite, alkali feldspar or plagioclase. Although previously described as tourmaline-rich granites representing melt-fractions expelled from residual stromatic leucosomes (Johnson, 2001), the absence of either alkali feldspar or plagioclase in many lenses indicates that they are not granites. The lenses are undeformed, and frequently zoned with tourmaline-rich cores. Quartz and tourmaline display eutectic intergrowths in some lenses. Fibrolite is abundant in all lenses, concentrated along the grain boundaries of the quartz. Although their grain size (1 cm) is insufficient to conform to many definitions of a pegmatite sensu stricto these assemblages share many of the textures and composition of pegmatites and so are termed ‘pegmatitic’ in this study.

The obvious chemical features of these lenses are their highly siliceous nature and strong variability in major elements inconsistent with being a minimum equilibrium melt in the granite system (Table 3, Fig. 10). The alumina saturation indices (ASI) vary from 1·5 to >20, comparable with the normal range of 1–1·4 for peraluminous granites. The coarse grain size, wide range in major-element chemistry and mineral modal zoning suggest that lens formation may be similar to that of granitic pegmatites. Trace-element concentrations in the pegmatites are highly variable (Table 2), and both strongly negative and positive Eu anomalies are observed in the analysed samples. Positive Eu anomalies and low overall REE abundances are found in feldspar-bearing pegmatites, and negative Eu anomalies and higher overall REE abundances are found in tourmaline-rich samples.

In part, these lenses probably owe their origin to a high F activity in the granitic melt which promoted crystallization of sillimanite over alkali feldspar (Nabelek, 1999). The quartz–fibrolite seams (Fig. 3f) therefore represent melt pathways where early crystallized quartz and sillimanite have been left behind. Recent experimental work has indicated that many pegmatite features can be explained by the disequilibrium crystallization of a volatile-rich peraluminous melt of minimum-melt composition (London, 1992). Volatiles lower the solidus temperature of peraluminous melts, and also decrease nucleation rates, resulting in undercooling. Crystallization is then determined by nucleation kinetics (London, 1996). In tourmaline-rich systems, B forms stable, soluble borosilicate complexes with alkalis and Al, inhibiting nucleation of feldspars. During quartz-dominated crystallization from the undercooled melt, an aqueous fluid exsolves and transports these elements out of the system. The pegmatitic lenses therefore may result from the crystallization of peraluminous granitic melts, as do the granite sheets. However, a high volatile activity, as indicated by the presence of F-bearing tourmaline, allowed the magma to persist to lower temperatures before crystallizing under disequilibrium conditions. Fibrolite, muscovite and quartz coexist in many lenses, suggesting temperatures of <680°C at 3 kbar (Fig. 9). The absence of a Ba anomaly in the host rocks indicates that the lenses result from a phase of fluid fluxing distinct from that associated with the formation of the granite sheet. The siting of the pegmatites in vertical east–west tension gashes suggests that the host melt may have accumulated within sag fractures in the intrusion floor.

	TIME SCALE OF FLUID CIRCULATION

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Partial melting within a thermal aureole provides potential for constraining the rates of melting of crustal rocks from determinations of the degree of isotopic disequilibrium between melt and source at the time of melting (Knesel & Davidson, 1999). For the granite sheet, an initial Sr-isotope ratio (⁸⁷Sr/⁸⁶Sr)₀ of 0·7053–0·7058 is obtained at 2059 Ma. This is within error of calculated values for the pelitic protolith (Table 7), indicating that no discernible disequilibrium has been preserved. The Sr-isotope compositions of five samples from Ba-rich localities within the migmatite zone, including granites, migmatites and hornfels, define an errorchron of 2161 ± 106 Ma, with an initial ratio of 0·705 ± 0·001 Ma (Fig. 14), suggesting that Sr isotopes have approached equilibrium between these diverse rock types both at outcrop scale and between outcrops affected by the high-Ba fluids.

View this table: [in this window] [in a new window]   Table 7: Bulk rock Sr-isotope compositions

 

View larger version (12K): [in this window] [in a new window]   Fig. 14. 87Rb/86Sr vs 87Sr/86Sr for metasediments, granites and pegmatite from the Bushveld aureole. Continuous line is errorchron for Ba-poor localities; dashed line is errorchron for Ba-rich localities.

 
For locality SA14, which has not been infiltrated by these fluids but where pegmatitic lenses are well developed, the Sr isotope composition approached homogenization at 2090 ± 51 Ma (within error of the previous age) at an (⁸⁷Sr/⁸⁶Sr)₀ value of 0·715 ± 0·001. This represents a significant increase in the initial ratio, which suggests that whereas both sample sets underwent isotope homogenization at outcrop scale, there was no regional homogenization between rocks that had been affected by Ba metasomatism and those that had not.

For homogenization to occur over a length scale of several metres, volume diffusion within individual grains must first homogenize contiguous grains. Volume diffusion of Sr at 750°C in oligoclase would require 30–80 kyr to homogenize the Sr isotope composition across 1 mm, assuming the diffusivities of Giletti & Casserly (1994). Once homogenization had occurred over this length scale, homogenization at outcrop scale would be accomplished by grain boundary diffusion, accelerated by fluid advection. We conclude that the migmatite zone was held at temperatures >750°C for at least 30 kyr.

Published thermal models of the Bushveld Complex have simulated the cooling history of the RLS rather than the aureole itself. Such modelling is complicated by the fact that the RLS was not emplaced as a single sheet of magma, but as a succession of pulses. Walraven (1982) modelled three pulses, the first being 2500 m thick, intruded at 1400°C. He concluded that 400 kyr was required for complete solidification of the RLS. More recently, Cawthorne & Walraven (1998) modelled the Bushveld as a chamber filled in 14 stages, the first stage (part of the Lower Zone) being 1 km thick (leaving cumulates 0·25 km thick) at 1300°C, emplaced 3 km below the surface. Their model predicts that this layer would have taken 0·25 kyr to crystallize; by their calculations the entire complex was emplaced over 75 kyr and took 200 kyr to crystallize.

We constructed a simple finite-difference model to predict heat flow, assuming conductive transfer, from the Lower Zone of the RLS into the aureole, based on the model of Wartho et al. (2001), using MATHCAD 2001 software. We assumed that a mafic magma sheet, 1500 m thick, was intruded into country rock initially at 125°C over a period of 1 Myr, and that, during this time, turbulent flow was maintained within the intrusion. Estimates of the temperature of emplacement of the RLS range from a lower value of 1100°C inferred from a xenolith in the Upper Zone (Wallmach et al., 1995), to an upper value of 1200°C determined from a pyroxenite of the Lower Zone (Sharpe & Hulbert, 1985). Initial magma temperature (T₀) was assumed to be 1200°C, and to allow for a maximum possible heating of the country rocks, the magma was superheated to 1680°C (T₀^*) to account for the latent heat of crystallization calculated from the equation of Jaeger (1964)

(4)

where latent heat L is 3 x 10⁵ J/kg, and specific heat capacity c is 1300 J/kg per K (Huppert & Sparks, 1985, 1988).

Initially, temperatures close to the contact are extremely high (>1000°C) but decrease rapidly with distance from the contact (Fig. 15a). Our model shows that temperatures >700°C that are recorded throughout the migmatite zone can be attained only within <200 m of the contact (Fig. 15b). At distances of 500 m, comparable with the stratigraphic thickness of the migmatites, modelled temperatures do not exceed 450°C (Fig. 15c and d). Conductive heat flow therefore fails to account for melting at such distances from the contact even though the period of magma emplacement is taken as 1 Myr, almost certainly an overestimate. However, if the system was open to fluid circulation, we believe that advection would have accelerated the heating rate, moderated ultra-high temperatures (>1000°C) at the contact, and maintained near-solidus temperatures (>700°C) over significant distances in the aureole. In contrast, at distances >4 km from the contact, disequilibrium porphyroblast growth suggests rather slow heating rates of 1°C per 10 kyr (Waters & Lovegrove, 2002) perhaps indicative of less vigorous fluid circulation in the outer regions of the aureole.

View larger version (15K): [in this window] [in a new window]   Fig. 15. Results of finite-element thermal modelling at distances of (a) 0 m (on the contact), (b) 200 m, (c) 500 m and (d) 1000 m from the contact, based on Wartho et al. (2001). Emplacement is assumed to have been continuous over a period of 1 Myr.

 

	CONCLUSIONS

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
This combined geochemical and textural study of the migmatite zone in the Bushveld aureole has identified several lines of evidence for a sustained hydrothermal system operating in the innermost zone of the aureole that played a critical role in determining the location and geochemistry of the anatectic melts. These include:

1. Ba and Sr metasomatism in the melt zones;
   
2. homogenization of Sr-isotope ratios in melt zones across the aureole;
   
3. the absence of very high temperature assemblages (>1000°C) adjacent to the contact;
   
4. the sustaining of elevated temperatures (750°C) at a stratigraphic distance of up to 500 m from the contact;
   
5. localized fluid infiltration-driven melting of predominantly unmodified psammites that still preserve sedimentary structures;
   
6. a record of sustained, episodic, post-peak fluid infiltration in the psammites resulting in grain growth and recrystallization forming finger structures on the outcrop scale, and oscillatory zoning in quartz on the grain scale at high temperatures; as the temperature dropped, grain growth ceased but fluid infiltration continued along isolated fractures and dilatant grain boundaries.
   

In situ migmatization, and melt segregation from metamorphosed pelites and psammites at temperatures of >700°C has resulted in the formation of granite veins and sheets. Melt sites were determined by bulk composition (the presence of biotite and feldspar), by fluid pathways, and by the amount of fluid (2% H₂O is required to generate melt fractions >0·3). Fluid fluxing resulted in open-system behaviour of Ba and Sr, and probably Eu, both during and after melting.

Subsequent sagging of the floor of the RLS caused stretching of the granite sheets, boudinaging of calc-silicate bands and extensional fracturing of pelites, synchronous with fluid-fluxed grain-boundary melting of psammites. Prograde infiltration of fluids sourced from devolatilization of the sediments on the intrusion floor was localized, as primary sedimentary structures and evidence for partial melting of psammite are preserved over distances <5 m.

During cooling, small batches of peraluminous magma, enriched in B and F, persisted to temperatures below the solidus of H₂O-dominated melts (<680°C). The high F activities modified the crystallization sequence and promoted significant undercooling, resulting in post-deformation pegmatitic lenses both within vertical sag fractures in the intrusion floor and exploiting earlier fabrics. The Sr isotope data suggest that during the cooling of the aureole isotopic homogenization was approached over distances of <10 m, but not over larger length scales.

Modification of the thermal structure by fluid advection has been recognized in other aureoles, albeit on a smaller scale (Kitchen, 1984). Combining the available data, we conclude that fluid circulation was not pervasive but focused, resulting in isotopic homogenization only within and between zones of high fluid flow, as indicated by the Ba-enriched metamorphosed pelites and psammites. This focused infiltration was highly heterogeneous and localized along dilatant cracks and grain boundaries but sufficiently vigorous to modify the thermal structure of the aureole.

As crustal melting in thermal aureoles will be enhanced, both by the high temperatures of the intrusion and by metamorphism under conditions of high fluid activity, it is probable that large-scale fluid circulation will characterize many, if not all, examples of contact metamorphism in which anatexis is observed. Such fluids can transport diagnostic trace elements and homogenize isotopic ratios, thus overprinting evidence for the rates of melt formation and segregation. In such circumstances, thermal models that assume heat transfer by conductive flow will greatly overestimate the thermal time scales involved.

	ACKNOWLEDGEMENTS

 
This work was funded by NERC Small Grant GR9/3750. Craig Storey is thanked for mineral separation, and Stephen Reed was very helpful during the CL imaging. Simon Kelley is thanked for his guidance in the thermal modelling section. We thank Ian Buick, Peter Nabelek, Kurt Knesel and Gary Stevens for constructive comments that greatly improved the manuscript.

	REFERENCES

TOP ABSTRACT INTRODUCTION THE AUREOLE OF THE... ANALYTICAL TECHNIQUES CONTACT METAMORPHISM OF QUARTZ... GRANITE SHEETS AND VEINS PEGMATITIC LENSES TIME SCALE OF FLUID... CONCLUSIONS REFERENCES

 
Ayres, M. & Harris, N. (1997). REE fractionation and Nd-isotope disequilibrium during crustal anatexis: constraints from Himalayan leucogranites. Chemical Geology 139, 249–269.[CrossRef][Web of Science]

Bea, F., Pereira, M. D. & Stroh, A. (1994). Mineral/leucosome trace-element partitioning in a peraluminous migmatite (a laser ablation-ICP-MS study). Chemical Geology 117, 291–312.[CrossRef][Web of Science]

Blundy, J. D. & Wood, B. J. (1991). Crystal-chemical controls on the partitioning of Sr and Ba between plagioclase feldspar, silicate melts, and hydrothermal solutions. Geochimica et Cosmochimica Acta 55, 193–209.[CrossRef][Web of Science]

Buick, I. S., Gibson, R. L., Cartwright, I., Maas, R., Wallmach, T. & Uken, R. (2000). Fluid flow in metacarbonates associated with emplacement of the Bushveld Complex, South Africa. Journal of Geochemical Exploration 69, 391–395.[CrossRef][Web of Science]

Buick, I. S., Maas, R. & Gibson, R. (2001). Precise U–Pb titanite age constraints on the emplacement of the Bushveld Complex, South Africa. Journal of the Geological Society, London 158, 3–6.[Abstract/Free Full Text]

Button, A. (1976). Stratigraphy and relations of the Bushveld floor in the Eastern Transvaal. Transactions of the Geological Society of South Africa 79, 3–12.

Carron, J.-P. & Lagache, M. (1980). Étude expérimentale du fractionnement des éléments Rb, Cs, Sr and Ba entre feldspaths alcalins, solutions hydrothermales et liquides silicates dans le système Q–Ab–Or–H₂O à 2 kbar entre 700 et 800°C. Bulletin de Mineralogie 103, 571–578.

Cawthorne, R. G. & Walraven, F. (1998). Emplacement and crystallisation time for the Bushveld Complex. Journal of Petrology 39, 1668–1687.

Clemens, J. D. & Holness, M. B. (2000). Textural evolution and partial melting of arkose in a contact aureole: a case study and implications. Electronic Geosciences 5, 4.

D'Lemos, R. S., Kearsley, A. T., Pembroke, J. W., Watt, G. R. & Wright, P. (1997). Complex growth histories in granite revealed by scanning cathodoluminescence techniques. Geological Magazine 134, 549–552.[Abstract]

Ellis, D. J. & Obata, M. (1992). Migmatite and melt segregation at Cooma, New South Wales. Transactions of the Royal Society of Edinburgh, Earth Sciences 83, 95–106.[Web of Science]

Ewart, A. & Griffin, W. L. (1994). Application of proton-microprobe data to trace-element partitioning in volcanic rocks. Chemical Geology 117, 251–284.[CrossRef][Web of Science]

Fabre, C., Boiron, M.-C., Dubessy, J., Chabiron, A., Charoy, B. & Crespo, T. M. (2002). Advances in lithium analysis in solids by means of laser-induced breakdown spectroscopy: an exploratory study. Geochimica et Cosmochimica Acta 66, 1401–1407.[CrossRef][Web of Science]

Fisher, A. J., Hayes, W. & Stoneham, A. M. (1990). Structure of the self-trapped exciton in quartz. Physical Review Letters 64, 2667–2670.[CrossRef][Web of Science][Medline]

Fourcade, S., Martin, H. & De Bremond d'Ars, J. (1992). Chemical exchange in migmatites during cooling. Lithos 28, 43–53.[CrossRef][Web of Science]

Giletti, B. J. & Casserly, J. E. D. (1994). Strontium diffusion kinetics in plagioclase feldspars. Geochimica et Cosmochimica Acta 58, 3785–3793.[CrossRef][Web of Science]

Harris, C. & Chaumba, J. B. (2001). Crustal contamination and fluid–rock interaction during the formation of the Platreef, Northern Limb of the Bushveld Complex, South Africa. Journal of Petrology 42, 1321–1347.[Abstract/Free Full Text]

Harris, N. B. W. & Inger, S. (1992). Trace element modelling of pelite-derived granites. Contributions to Mineralogy and Petrology 110, 46–56.[CrossRef]

Harris, N. B. W., Inger, S. & Massey, J. A. (1993). The rôle of fluids in the formation of High Himalayan leucogranites. In: Treloar, P. J. & Searle, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications 74, 391–400.

Harris, N. B. W., Vance, D. & Ayres, M. (2000). From sediment to granite: timescales of anatexis in the upper crust. Chemical Geology 162, 155–167.[CrossRef][Web of Science]

Holness, M. B. (1993). Temperature and pressure dependence of quartz–aqueous fluid dihedral angles: the control of adsorbed H₂O on the permeability of quartzites. Earth and Planetary Science Letters 117, 363–377.[CrossRef][Web of Science]

Holness, M. B. (1997). Surface chemical controls on pore-fluid connectivity in texturally equilibrated materials. In: Jamveit, B. & Yardley, B. W. D. (eds) Fluid Flow and Transport in Rocks: Mechanisms and Effects. London: Chapman & Hall, pp. 149–170.

Holness, M. B. (1998). Contrasting rock permeability in the aureole of the Ballachulish igneous complex, Scottish Highlands: the influence of surface energy? Contributions to Mineralogy and Petrology 131, 86–94.[CrossRef][Web of Science]

Holness, M. B. & Clemens, J. D. (1999). Partial melting of the Appin quartzite driven by fracture-controlled H₂O infiltration in the aureole of the Ballachulish Igneous Complex, Scottish Highlands. Contributions to Mineralogy and Petrology 136, 154–168.[CrossRef][Web of Science]

Holness, M. B. & Watt, G. R. (2001). Quartz recrystallisation and fluid flow during contact metamorphism: a cathodoluminescence study. Geofluids 1, 215–228.[CrossRef]

Holtz, F. & Johannes, W. (1991). Genesis of peraluminous granites I. Experimental investigation of melt compositions at 3 and 5 kb and various H₂O activities. Journal of Petrology 32, 935–958.[Abstract/Free Full Text]

Hunter, R. H. (1987). Textural equilibrium in layered igneous rocks. In: Parsons, I. (ed.) Origins of Igneous Layering. Dordrecht: D. Reidel, pp. 473–503.

Huppert, H. E. & Sparks, R. S. J. (1985). Cooling and contamination of mafic and ultramafic magmas during ascent through continental crust. Earth and Planetary Science Letters 74, 371–386.[Web of Science]

Huppert, H. E. & Sparks, R. S. J. (1988). The generation of granitic magmas by intrusion of basalt into continental crust. Journal of Petrology 29, 599–624.[Abstract/Free Full Text]

Icenhower, J. & London, D. (1995). An experimental study of element partitioning among biotite, muscovite and coexisting peraluminous silicic melt at 200 MPa (H₂O). American Mineralogist 80, 1229–1251.[Abstract]

Icenhower, J. & London, D. (1996). Experimental partitioning of Rb, Cs, Sr and Ba between alkali feldspar and peraluminous melt. American Mineralogist 81, 719–734.[Abstract]

Jaeger, J. C. (1964). Thermal effects of intrusions. Reviews of Geophysics 2, 443–466.[CrossRef][Web of Science]

Johnson, T. E. (2001). Controls on melt segregation and migration: evidence from migmatites. Geological Society of America, Abstracts with Program 33, 2.

Kaneko, Y. & Miyano, T. (1990). Contact metamorphism of the Bushveld Complex in the northeastern Transvaal, South Africa. Journal of Mineralogy, Petrology and Economic Geology 85, 66–81.

Kitchen, D. (1984). Pyrometamorphism and the contamination of basaltic magma at Tieveragh, Co. Antrim. Journal of the Geological Society, London 141, 733–745.[Abstract/Free Full Text]

Knesel, K. M. & Davidson, J. P. (1999). Sr isotope systematics during melt generation by intrusion of basalt into continental crust. Contributions to Mineralogy and Petrology 136, 285–295.[CrossRef][Web of Science]

Koester, E., Pawley, A. R., Fernandes, L. A. D., Porcher, C. C. & Soliani, E., Jr (2002). Experimental melting of cordierite gneiss and the petrogenesis of syntranscurrent peraluminous granites in Southern Brazil. Journal of Petrology 43, 1595–1616.[Abstract/Free Full Text]

Kohn, S. C. (2000). The dissolution mechanisms of water in silicate melts; a synthesis of recent data. Mineralogical Magazine 64, 389–408.[Abstract/Free Full Text]

Kostenko, O., Jamtveit, B., Austrheim, H., Pollok, K. & Putnis, C. (2002). The mechanism of fluid infiltration in peridotites at Almklovdalen, western Norway. Geofluids 2, 1–13.

Kretz, R. (1983). Symbols for rock-forming minerals. American Mineralogist 68, 277–279.[Abstract]

Kriegsman, L. M. & Hensen, B. J. (1998). Back reaction between restite and melt: implications for geothermobarometry and pressure–temperature paths. Geology 26, 1111–1114.[Abstract/Free Full Text]

Laporte, D. (1994). Wetting behaviour of partial melts during crustal anatexis: the distribution of hydrous silicic melts in polycrystalline aggregates of quartz. Contributions to Mineralogy and Petrology 116, 486–499.

Laporte, D. & Watson, E. B. (1995). Experimental and theoretical constraints on melt distribution in crustal sources: the effect of crystalline anisotropy on melt interconnectivity. Chemical Geology 124, 161–184.[CrossRef][Web of Science]

Le Breton, N. & Thompson, A. B. (1988). Fluid-absent (dehydration) melting of biotite in metapelites in the early stages of crustal anatexis. Contributions to Mineralogy and Petrology 99, 226–237.[CrossRef][Web of Science]

Lee, M. R. & Parsons, I. (1997). Dislocation formation and albitisation in alkali feldspars from the Shap granite. American Mineralogist 82, 557–570.[Abstract]

Lind, A. (1996). Microstructural stability and the kinetics of textural evolution. Ph.D. thesis, University of Leeds.

London, D. (1992). The application of experimental petrology to the genesis and crystallisation of granitic pegmatites. Canadian Mineralogist 30, 499–540.[Web of Science]

London, D. (1996). Granitic pegmatites. Transactions of the Royal Society, Edinburgh: Earth Sciences 87, 305–319.

Macaulay, C. I., Fallick, A. E., Haseldine, R. S. & Graham, C. M. (2000). Methods of laser-based stable isotope measurement applied to diagenetic cements and hydrocarbon reservoir quality. Clay Minerals 35, 313–322.[Abstract/Free Full Text]

Mahood, G. & Hildreth, W. (1983). Large partition coefficients for trace elements in high-silica rhyolites. Geochimica et Cosmochimica Acta 47, 11–30.[CrossRef][Web of Science]

Marshall, D. J. (1988) Cathodoluminescence of Geological Materials. Boston, MA: Unwin Hyman, 146 pp.

Matthews, A., Goldsmith, J. R. & Clayton, R. N. (1983). Oxygen isotope fractionation involving pyroxenes: the calibration of mineral-pair geothermometers. Geochimica et Cosmochimica Acta 47, 631–644.[CrossRef][Web of Science]

Nabelek, P. I. (1999). Trace-element distribution among rock-forming minerals in Black Hills migmatites, South Dakota. American Mineralogist 84, 1256–1269.[Abstract]

Nabelek, P. I. & Labotka, T. C. (1993). Implications of geochemical fronts in the Notch Peak contact metamorphic aureole, Utah, U.S.A. Earth and Planetary Science Letters 119, 539–559.[CrossRef][Web of Science]

Nash, W. P. & Crecraft, H. R. (1985). Partition coefficients for trace elements in silicic magmas. Geochimica et Cosmochimica Acta 49, 2309–2322.[CrossRef][Web of Science]

Nell, J. (1985). The Bushveld metamorphic aureole in the Potgietersus area: evidence for a two-stage metamorphic event. Economic Geology 80, 1129–1152.[Abstract/Free Full Text]

Otamendi, J. E. & Patiño Douce, A. E. (2001). Partial melting of aluminous metagreywackes in the Northern Sierra de Comechingones. Journal of Petrology 42, 1751–1772.[Abstract/Free Full Text]

Patiño-Douce, A. E. & Harris, N. (1998). Experimental constraints on Himalayan anatexis. Journal of Petrology 39, 689–710.[CrossRef][Web of Science]

Penniston-Dorland, S. C. (2001). Illumination of vein quartz textures in a porphyry copper ore deposit using scanned cathodoluminescence: Grasberg Igneous Complex, Irian Jaya, Indonesia. American Mineralogist 86, 652–666.[Abstract/Free Full Text]

Peppard, B. T., Steele, I. M., Davis, A. M., Wallace, P. J. & Anderson, A. T. (2001). Zoned quartz phenocrysts from the rhyolitic Bishop Tuff. American Mineralogist 86, 1034–1052.[Abstract/Free Full Text]

Pitra, P. & De Waal, S. A. (2001). High-temperature, low-pressure metamorphism and development of prograde symplectites, Marble Hall Fragment, Bushveld Complex (South Africa). Journal of Metamorphic Geology 19, 311–325.[Web of Science]

Platten, I. M. (1982). Partial melting of feldspathic quartzites around late Caledonian minor intrusions in Appin, Scotland. Geological Magazine 119, 413–419.[Abstract]

Potts, P. J., Webb, P. C. & Watson, J. S. (1984). Energy dispersive X-ray fluorescence analysis of silicate rocks for major and trace elements. X-Ray Spectrometry 13, 2–15.[CrossRef]

Prince, C., Harris, N. & Vance, D. (2001). Fluid-enhanced melting during prograde metamorphism. Journal of the Geological Society, London 158, 233–241.[Abstract/Free Full Text]

Riegger, O. K. & Van Vlack, L. H. (1960). Dihedral angle measurement. Transactions of the Metallurgical Society of AIME 218, 933–935.

Rosenberg, C. L. & Riller, U. (2000). Partial-melt topology in statically and dynamically recrystallised granite. Geology 28, 7–10.[Abstract/Free Full Text]

Saigal, G. C., Morad, S., Bjorlykke, K., Egeberg, P. K. & Aagaard, P. (1988). Diagenetic albitisation of detrital K-feldspar in Jurassic, Lower Cretaceous, and Tertiary clastic reservoir rocks from offshore Norway, I. Textures and origin. Journal of Sedimentary Petrology 58, 1003–1013.[Abstract/Free Full Text]

Sawyer, E. W. (1987). The role of partial melting and fractional crystallization in determining discordant migmatite leucosome compositions. Journal of Petrology 28, 445–473.[Abstract/Free Full Text]

Schriffries, C. M. & Rye, D. M. (1990). Stable isotope systematics of the Bushveld Complex: ii. Constraints on hydrothermal processes in layered intrusions. American Journal of Science 290, 209–245.[Abstract/Free Full Text]

Seyedolali, A., Krinsley, D. H., Boggs, S., Jr, O'Hara, P. F., Dypvik, H. & Goles, G. G. (1997). Provenance interpretation of quartz by scanning electron microscope–cathodoluminescence fabric analysis. Geology 25, 787–790.[Abstract/Free Full Text]

Sharp, Z. D. (1990). A laser-based microanalytical method for the in-situ determination of oxygen isotope ratios in silicates and oxides. Geochimica et Cosmochimica Acta 54, 1353–1357.[CrossRef][Web of Science]

Sharpe, M. R. & Chadwick, B. (1982). Structures in Transvaal sequence rocks within and adjacent to the eastern Bushveld Complex. Transactions of the Geological Society of South Africa 85, 29–41.

Sharpe, M. R. & Hulbert, L. J. (1985). Mobilized suspensions of early Lower Zone cumulates in a parental magma with boninitic affinities. Economic Geology 80, 849–871.[Abstract/Free Full Text]

Smith, C. S. (1964). Some elementary principles of polycrystalline microstructure. Metallurgical Reviews 9, 1–48.

Stickels, C. A. & Hucke, E. E. (1964). Measurement of dihedral angles. Transactions of the Metallurgical Society of the AIME 230, 795–801.

Uken, R. & Watkeys, M. K. (1997). Diapirism initiated by the Bushveld Complex, South Africa. Geology 25, 723–726.[Abstract/Free Full Text]

Valley, J. W. & Graham, C. M. (1996). Ion microprobe analysis of oxygen isotope ratios in quartz from Skye granite: healed micro-cracks, fluid flow and hydrothermal exchange. Contributions to Mineralogy and Petrology 124, 225–234.[CrossRef][Web of Science]

Vernon, R. H. (1968). Microstructures of high-grade metamorphic rocks at Broken Hill, Australia. Journal of Petrology 9, 1–22.[Abstract/Free Full Text]

Walderhaug, O. & Rykkje, J. (2000). Some examples of the effect of crystallographic orientation on the cathodoluminescence colours of quartz. Journal of Sedimentary Research 70, 545–548.[Abstract/Free Full Text]

Wallmach, T., Hatton, C. J., de Waal, S. A. & Gibson, R. L. (1995). Retrogressive hydration of calc-silicate xenoliths in the eastern Bushveld Complex; evidence for late magmatic fluid movement. Journal of African Earth Sciences 21, 533–646.

Walraven, F. (1982). Textural, geochemical and genetical aspects of the granophyric rocks of the Bushveld Complex. Ph.D. thesis, University of Witswatersrand.

Wartho, J.-A., Kelley, S. P. & Blake, S. (2001). Magma flow regimes in sills deduced from Ar isotope systematics of host rocks. Journal of Geophysical Research 106, 4017–4035.[CrossRef]

Waters, D. J. & Lovegrove, D. P. (2002). Assessing the extent of disequilibrium and overstepping of prograde metamorphic reactions in metapelites from the Bushveld Complex aureole, South Africa. Journal of Metamorphic Geology 20, 135–149.[CrossRef][Web of Science]

Watt, G. R., Wright, P., Galloway, S. & McLean, C. (1997). Cathodoluminescence and trace element zoning in quartz phenocrysts and xenocrysts. Geochimica et Cosmochimica Acta 61, 4337–4348.[CrossRef][Web of Science]

Watt, G. R., Oliver, N. H. S. & Griffin, B. J. (2000). Evidence for reaction-induced microfracturing in granulite facies migmatites. Geology 28, 327–330.[Abstract/Free Full Text]

Waychunas, G. A. (1988). Luminescence, X-ray emission and new spectroscopies. Mineralogical Society of America, Reviews in Mineralogy 18, 639–698.[Abstract]

Wilkinson, J. J. & Johnston, J. D. (1996). Pressure fluctuations, phase separation and gold precipitation during seismic fracture propagation. Geology 24, 395–398.[Abstract/Free Full Text]

Wilks, S. S. (1962). Mathematical Statistics. New York: John Wiley.

Willemse, J. & Viljoen, E. A. (1970). The fate of argillaceous material in the gabbroic magma of the Bushveld Complex. Special Publication of the Geological Society of South Africa 1, 336–366.

Yardley, B. W. D., Rochelle, C. A., Barnicoat, A. C. & Lloyd, G. E. (1991). Oscillatory zoning in metamorphic minerals—an indicator of infiltration metasomatism. Mineralogical Magazine 55, 357–365.[Web of Science]

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