Journal of Petrology 45(6) © Oxford University Press 2004; all rights reserved
Variation in Metamorphic Style along the Northern Margin of the Damara Orogen, Namibia
1 CONTINENTAL EVOLUTION RESEARCH GROUP, SCHOOL OF EARTH AND ENVIRONMENTAL SCIENCES, UNIVERSITY OF ADELAIDE, ADELAIDE, S.A. 5005, AUSTRALIA
2 SCHOOL OF EARTH SCIENCES, UNIVERSITY OF MELBOURNE, PARKVILLE, VIC. 3010, AUSTRALIA
RECEIVED FEBRUARY 4, 2003; ACCEPTED DECEMBER 12, 2003
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONREGIONAL GEOLOGY OF THE...STRUCTURAL EVOLUTION AND...METAMORPHIC EVOLUTIONDISCUSSIONSUPPLEMENTARY DATAREFERENCES |
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The northern margin of the Inland Branch of the Pan-African Damara Orogen in Namibia shows dramatic along-strike variation in metamorphic character during convergence between the Congo and Kalahari Cratons (M3 metamorphic cycle). Low-P contact metamorphism with anticlockwise P–T paths dominates in the western domains (Ugab Zone and western Northern Zone), and high-P Barrovian metamorphism with a clockwise P–T path is documented from the easternmost domain (eastern Northern Zone). The sequence of M3 mineral growth in contact aureoles shows early growth of cordierite porphyroblasts that were pseudomorphed to biotite–chlorite–muscovite at the same time as an andalusite–biotite–muscovite transposed foliation was developed in the matrix. The peak-T metamorphic assemblages and fabrics were overprinted by crenulations and retrograde chlorite–muscovite. The KFMASH P–T pseudosection for metapelites in the Ugab Zone and western Northern Zone contact aureoles indicates tight anticlockwise P–T loops through peak metamorphic conditions of 540–570°C and 2·5–3·2 kbar. These semi-quantitative P–T loops are consistent with average P–T calculations using THERMOCALC, which give a pooled mean of 556 ± 26°C and 3·2 ± 0·6 kbar, indicating a high average thermal gradient of 50°C/km. In contrast, the eastern Northern Zone experienced deep burial, high-P/moderate-T Barrovian M3 metamorphism with an average thermal gradient of 21°C/km and peak metamorphic conditions of c. 635°C and 8·7 kbar. The calculated P–T pseudosection and garnet compositional isopleths in KFMASH, appropriate for the metapelite sample from this region, document a clockwise P–T path. Early plagioclase–kyanite–biotite parageneses evolved by plagioclase consumption and the growth of garnet to increasing XFe, XMg and XCa and decreasing XMn compositions, indicating steep burial with heating. The developed kyanite–garnet–biotite peak metamorphic parageneses were followed by the resorption of garnet and formation of plagioclase moats, indicating decompression, which was followed by retrogressive cooling and chlorite–muscovite growth. The clockwise P–T loop is consistent with the foreland vergent fold–thrust belt geometry in this part of the northern margin. Earlier formed (580–570 Ma) pervasive matrix foliations (M2) were overprinted by contact metamorphic parageneses (M3) in the aureoles of 530 ± 3 Ma granites in the Ugab Zone and 553–514 Ma granites in the western Northern Zone. Available geochronological data suggest that convergence between the Congo and Kalahari Cratons was essentially coeval in all parts of the northern margin, with similar ages of 535–530 Ma for the main phase of deformation in the eastern Northern Zone and Northern Platform and 538–505 Ma high-grade metamorphism of the Central Zone immediately to the south. Consequently, NNE–SSW-directed convergent deformation and associated M3 metamorphism of contrasting styles are interpreted to be broadly contemporaneous along the length of the northern margin of the Inland Branch. In the west heat transfer was dominated by conduction and externally driven by granites, whereas in the east heat transfer was dominated by advection and internally driven radiogenic heat production. The ultimate cause was along-orogen variation in crustal architecture, including thickness of the passive margin lithosphere and thickness of the overlying sedimentary succession.
KEY WORDS: Pan-African Orogeny; P–T paths; pseudosections; low-P metamorphism; contact metamorphism; Barrovian metamorphism
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY OF THE...STRUCTURAL EVOLUTION AND...METAMORPHIC EVOLUTIONDISCUSSIONSUPPLEMENTARY DATAREFERENCES |
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Along-orogen variation in the metamorphic response to orogenesis is almost never investigated, and examples that have been studied are restricted to transpressional orogens (e.g. Goscombe et al., 2003b
), which have also been thermally modelled for variation along orogen length (Thompson et al., 1997). Most investigations of metamorphism in high-angle convergent orogens are restricted to profiles across the grain of the orogen. Although this monoclinic view produces significant results in understanding metamorphism in orogens (e.g. Inger & Harris, 1992; Vannay & Hodges, 1996; Harrison et al., 1997; Goscombe & Hand, 2000; Goscombe et al., 2003b) and the thermo-mechanical modelling of orogens (England & Thompson, 1984; Jamieson et al., 1996, 2002; Huerta et al., 1998, 1999; Beaumont et al., 2001), it is a very limited view that runs the risk of missing significant aspects of orogenesis in general. This recent thermo-mechanical modelling effort has quantified first-order parameters giving insights into how the metamorphic response may vary with changes in crustal architecture along the length of orogenic belts. Orogens are not regular homogeneous entities; first-order, crustal-scale variation in orogen architecture along strike does occur in many, if not all orogens. Along-strike variation in metamorphic response is anticipated where radical variation in orogen architecture exists. This paper forms part of a large project to investigate along-strike variations in metamorphic response to collisional orogenesis from two classic high-angle convergent orogens, the Inland Branch of the Damara Orogen in Namibia and the central Himalayan metamorphic front. The northern margin of the Damara Orogen displays a problematic metamorphic pattern of high-P Barrovian metamorphism in the east, apparently coeval with low-P contact metamorphism in the west, with all regions apparently undergoing NNE–SSW shortening at this time. Laterally juxtaposed terranes that experienced coeval but contrasting style of metamorphism, within the same orogenic front, are rarely drawn attention to but nevertheless may be more common than currently thought. The tectonic interpretation of these adjacent, but contrasting, terranes is problematic. This paper presents supporting evidence to characterize the along-orogen variation in metamorphism that exists and proposes a field-based explanation for the variation in metamorphic character, but it is beyond the scope of this paper to develop causative thermo-mechanical models of first-order processes that could be responsible for such variations.
The northern margin of the Damara Orogen offers a unique opportunity to investigate the variable metamorphic response along the strike of a convergent orogen exhibiting contrasting crustal architecture. In this paper we investigate the metamorphic history of a broad ENE-trending band (450 km long) of metamorphic rocks, which we call the ‘northern margin’ of the Inland Branch, between the high-grade orogen core (the Central Zone) and the Congo Craton and Kaoko Belt elements to the north (Fig. 1). The northern margin is composed of the Northern Zone of the Inland Branch in the east, and the Ugab Zone in the west (Miller, 1983). Both zones share a common deformation and thermal history, which can be correlated with events in the transpressional Kaoko Belt to the north (Goscombe et al., 2003a, 2003b). The western extremity of the investigated northern margin is marked by the steep, 5 km wide Ogden Mylonite Zone, which is interpreted to be the lateral extension of the Purros Mylonite Zone, the crustal-scale median structure controlling the architecture of the Kaoko Belt (Fig. 1; Goscombe et al., 2003a, 2003b). The eastern northern margin contains a thick succession of Palaeoproterozoic Damara Sequence sediments deposited in a deep trough (Porada & Wittig, 1983); this basinal sequence was inverted as a foreland vergent fold–thrust belt during the Pan-African Damara Orogeny (Miller, 1983). In contrast, the western Northern Zone and Ugab Zone comprise a thin, 1·7 km, succession of Damara Sequence sediments deposited on an unexposed basement high (Miller et al., 1983; Swart, 1992; Freyer & Hälbich, 1994; Passchier et al., 2002).
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No metamorphic study has been previously undertaken in any part of the entire northern margin of the Damara Orogen, and the few structural studies published (Miller, 1980
; Coward, 1983; Freyer & Hälbich, 1994; Maloof, 2000; Passchier et al., 2002) are not integrated with studies of evolving mineral parageneses and metamorphism. Consequently, this paper presents the first integrated structural and metamorphic investigation of the northern margin. Near 100% exposure in the western 240 km allowed detailed mapping of the entire Ugab Zone and western Northern Zone (Fig. 2) to be undertaken between 1997 and 2002. Detailed investigation of metamorphic rocks in nine widely spaced regions representing all major structural domains and all contact aureoles within this portion of the northern margin were undertaken. Exposures of suitable metapelite rocks for metamorphic studies are absent from the eastern Northern Zone; only one critical metapelite sample from a drill core was available for investigation from this portion of the northern margin. The rationale of our work was to link our detailed structural analysis with the P–T evolution in all domains containing suitable metapelitic rocks. In this way the structural and metamorphic evolution was quantified across much of the northern margin and correlated with the available geochronology in the literature.
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REGIONAL GEOLOGY OF THE NORTHERN MARGIN OF THE DAMARA OROGEN |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY OF THE...STRUCTURAL EVOLUTION AND...METAMORPHIC EVOLUTIONDISCUSSIONSUPPLEMENTARY DATAREFERENCES |
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The Damara Orogen of Namibia has triple junction geometry between the Kalahari Craton in the south, the Congo Craton in the North and the Rio de la Plata Craton in South America to the west (Fig. 1), and contains a Neoproterozoic succession that was intensely deformed and metamorphosed in the protracted Neoproterozoic to Cambrian Pan-African Orogeny (Miller, 1983
). The Kaoko Belt component is a NNW-trending, sinistral transpressional belt that extends northward from the Ugab Zone into Angola (Dürr & Dingeldey, 1996; Goscombe et al., 2003a, 2003b); the Gariep Belt in the south is a north-trending sinistral transpressional belt (Davies & Coward, 1982; Frimmel, 1995; Hälbich & Alchin, 1995); the Inland Branch trends ENE–WSW into Botswana (Coward, 1981; Miller, 1983; Porada et al., 1983) and links with the Pan-African Orogenic System (Goscombe et al., 2000) across the southern and eastern African continent (Fig. 1).
The northern margin of the Inland Branch has been divided into two distinct zones (Miller, 1983). (1) The Ugab Zone (Fig. 2) is at the junction between the Kaoko Belt and Inland Branch components of the Damara Orogen, and can be considered the southern extension of the Kaoko Belt (Miller, 1983; Passchier et al., 2002; Goscombe et al., 2003a, 2003b). The Damara Sequence in the Ugab Zone comprises a thin (1·7 km), distal turbiditic succession that is of greenschist-facies grade and has been pervasively deformed by a complex evolution of intense folding without apparent involvement of the basement (Fig. 3; Miller et al., 1983; Swart, 1992; Freyer & Hälbich, 1994; Passchier et al., 2002; Goscombe et al., 2003a). The western margin of the Ugab Zone is the Ogden Mylonite Zone (Fig. 2). (2) The Northern Zone abuts the Palaeoproterozic Kamanjab Inlier portion of the Congo Craton, which is overlain by platform carbonates of the Northern Platform (Fig. 1). The Northern Zone is bounded to the north by a steep thrust complex, north of which low-amplitude folding of the carbonate platform sequence occurs (Hedberg, 1979). To the south, the Northern Zone is a strongly deformed fold belt with a complex deformation history dominated by foreland vergent structures (Fig. 3; Miller, 1980, 1983; Coward, 1983). The western portion of the Northern Zone has a thin greenschist-facies grade sedimentary sequence that is similar to the Ugab Zone, and overlies a basement high called the Huab Ridge (Porada et al., 1983) that extends to the SW under the Ugab Zone to the Walvis Ridge offshore. In the eastern portion of the Northern Zone, the Damara Sequence is a markedly thicker succession of distal turbidites deposited in a trough called the Northern Graben (Miller, 1983; Porada, 1983; Porada & Wittig, 1983); metamorphic grade reaches Barrovian series, mid-amphibolite-facies conditions. East of the Otavi Region (Fig. 1) there is no exposure of the Inland Branch, which is covered by Cretaceous to Recent deposits of the Kalahari Basin.
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Neoproterozoic Damara Sequence
The Damara Sequence is a marine sequence deposited on a passive margin, progressing from shelf carbonates on the Northern Platform to slope and distal facies to the south and west (Miller, 1983
; Porada & Wittig, 1983; Swart, 1992; Hoffman et al., 1998). Deposition of the Damara Sequence is interpreted to have spanned the Neoproterozoic between 770 and 600 Ma (Miller, 1983; Hoffmann, 1994; Frimmel, 1996; Prave, 1996). The basal Damara Sequence is restricted to the Northern Zone (Fig. 2) and represented by rift-related siliciclastics of the Nosib Group, comprising quartzites, conglomerates and arenites. Quartz-syenite, alkaline ignimbrite and alkali-rhyolite units in the upper Nosib Group have U–Pb and Pb–Pb zircon ages ranging from 757 ± 1 to 746 ± 2 Ma (Hoffmann, 1994; Hoffman et al., 1996: de Kock et al., 2000), constraining the minimum age of the Nosib Group to be c. 750 Ma (Prave, 1996; Hoffman et al., 1998). The overlying Otavi Group is dominated by a turbiditic greywacke sequence with pelitic schists and psammites. Within this succession are two turbiditic carbonate formations, parts of which are correlated with regional diamictite horizons that are elsewhere interpreted to be of 750–735 Ma and 700 Ma age (Hoffmann, 1994; Frimmel, 1996; Folling et al., 1998; Hoffman et al., 1998). The uppermost Otavi Group is the widespread Kuiseb Formation, which is composed of turbiditic greywackes and pelite schists with thin calc-silicate bands (Fig. 2). The sedimentary succession in the Ogden Mylonite Zone (Fig. 2) is composed of greywackes and arenites that are correlated with the Kuiseb Formation (Miller, 1983; Miller & Grote, 1988; Freyer & Hälbich, 1994), and are identical to sequences immediately to the north within the Coastal Terrane portion of the Kaoko Belt (Goscombe et al., 2003a, 2003b). The uppermost Damara Sequence is the Mulden Group siliciclastic molasse of 620–600 Ma age (Miller, 1983), which is preserved only in the Northern Platform (Fig. 2; Miller, 1980; Miller & Grote, 1988). Deposition of the Damara Sequence was terminated by collision between the Kalahari and Congo Cratons late in the Neoproterozoic and was followed by a protracted period of Pan-African tectonothermal events collectively called the Damara Orogeny, which spanned from the late Neoproterozoic to Cambrian (Miller, 1983; Prave, 1996).
Neoproterozoic and Palaeozoic Pan-African granitoids
The boundary between the Northern Zone and high-grade Central Zone is marked by a broad region that is dominated by numerous Pan-African granitoid bodies (Fig. 1). Two periods of widespread granitoid emplacement are recognized in the northern margin of the Inland Branch (Electronic Appendix A; extended datasets are given as Electronic Appendices A–D, which can be downloaded from the Journal of Petrology website at http://www.petrology.oupjournals.org). (1) Early syn-kinematic granitoids that both pre-date and crosscut the pervasive foliation and main phase deformational structures are c. 590–570 Ma. (2) Younger, less strained granitoids were emplaced during north–south shortening in the latest deformational period of the Damara Orogeny at c. 550–510 Ma. Older U–Pb zircon and monazite ages are reported from a restricted portion of the Coastal Terrane in the western Kaoko Belt; granite orthogneiss crystallization ages are 656 ± 8 Ma (Seth et al., 1998) and high-grade metamorphic parageneses with metamorphic zircon and monazite age are 645 ± 4 Ma (Franz et al., 1999). These ages represent the M1 metamorphic cycle (Goscombe et al., 2003b), which has not been recorded by any other reliable age data from elsewhere in the Damara Orogen (Fig. 4; Electronic Appendix A). Consequently, the earliest stage of mineral growth recorded in the Ugab Zone and Northern Zone cannot be correlated with this M1 metamorphic cycle, which is thus considered beyond the scope of this paper.
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656–645 Ma granitoids only recorded from the Coastal Terrane.Early syn-kinematic granitoids, c. 590–570 Ma in age, occur throughout the Inland Branch and Kaoko Belt. Syn-kinematic granites from the southernmost Coastal Terrane, immediately north of the Ogden Mylonite Zone, give a U–Pb zircon age of 570 ± 20 Ma (Miller & Burger, 1983a
), from the Ugab Zone a 573 ± 33 Ma Rb–Sr whole-rock age (Kröner, 1982) and from the western Northern Zone a 589 ± 40 Ma U–Pb zircon age (Miller & Burger, 1983b). S-type granitic orthogneiss bodies of similar age are common in the western Kaoko Belt and these were emplaced immediately prior to, and possibly also during, the pervasive transpressional deformation (Goscombe et al., 2003b). Concordant U–Pb ages from these syn-kinematic granites range from 580 ± 3 Ma (Seth et al., 1998) to 576 ± 5 Ma (Franz et al., 1999), which is coincident with Sm–Nd peak metamorphic garnet ages, which range from 579 ± 15 Ma to 571 ± 6 Ma (Goscombe et al., 2003b). Post-kinematic granites give minimum ages by Pb–Pb evaporative techniques in the range 567–552 Ma (Seth et al., 1998; Franz et al., 1999), and provide a minimum age constraint for the main phase of transpressional deformation in the Kaoko Belt (Goscombe et al., 2003b).
The younger period of granitoid emplacement, associated with north–south shortening between the Congo and Kalahari Cratons (Goscombe et al., 2003a, 2003b), is not recognized in the Kaoko Belt but is otherwise volumetrically dominant in the Inland Branch. These granitoids are typically composite bodies, some concentrically zoned, with at least three intrusive phases ranging from syenite to biotite granite and late-stage aplite dykes. The protracted intrusive history records a progressive decrease in Shortening Phase strain from weak L–S fabrics in early syenites to essentially undeformed youngest phases. Granitic sills and pegmatites in the contact aureoles can be intensely deformed by isoclinal folding caused by the focusing of strain in these relatively hot ductile zones. The composite Footspore Pluton (Fig. 2) in the Ugab Zone post-dates the pervasive transpressional structures and was emplaced during north–south shortening (Passchier et al., 2002; Goscombe et al., 2003a, 2003b). The oldest syenite phase in this pluton has been dated at 530 ± 3 Ma by the Pb–Pb single zircon evaporative technique (Seth et al., 2000). No other U–Pb zircon age data are available from the Northern Zone granitoids; Rb–Sr whole-rock ages range between 500 and 550 Ma (Hawkesworth et al., 1983). This age range is consistent with the 546–490 Ma range of U–Pb monazite, zircon and titanite ages (Briqueu et al., 1980; Allsopp et al., 1983; Miller & Burger, 1983a; de Kock & Walraven, 1995; Jacob et al., 2000; Jung et al., 2000a, 2000b) from syn-tectonic granites emplaced in the Central Zone during pervasive deformation and metamorphism of the Inland Branch during the NNE–SSW convergent Shortening Phase.
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STRUCTURAL EVOLUTION AND TECTONIC FRAMEWORK |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY OF THE...STRUCTURAL EVOLUTION AND...METAMORPHIC EVOLUTIONDISCUSSIONSUPPLEMENTARY DATAREFERENCES |
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As elsewhere within the Pan-African Orogenic System (Goscombe et al., 2000
), the Pan-African Damara Orogeny is protracted, with tectonothermal events spanning c. 650–500 Ma. Three distinct tectonothermal phases make up the Damara Orogeny: the Thermal, Transpressional and Shortening Phases described below (Goscombe et al., 2003a, 2003b). These tectonothermal phases are recognized in both the northern margin of the Inland Branch (Fig. 4), and in the Kaoko Belt with which it is linked (Goscombe et al., 2003a, 2003b). Throughout this paper specific deformation events (i.e. D1, D2,...), within the framework of tectonothermal phases, are defined by their mutually common textural relationships and are neither similar in orientation in different domains (Fig. 5) nor necessarily time equivalent across the region investigated. Nevertheless, a consistent sequence of texturally similar deformation structures is recognized across the Ugab Zone and western Northern Zone.
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, intersection lineation, crenulation lineation or fold axis. Open stars indicate mean orientation of long axis of boudins. Dotted lines in the West Ugab domain outline the distribution of poles to S2 and shallow north plunge of L2 in the Ogden Mylonite Zone.Thermal Phase (early M2)
The earliest recognizable tectonothermal event, the Thermal Phase, was responsible for the growth of early porphyroblasts of biotite and less commonly ilmenite, chlorite and andalusite (Table 1; Fig. 6a–c) that pre-date pervasive reworking and formation of the main mica foliation (S1 and S2). Thermal Phase porphyroblasts (early M2) are enveloped by the S2 foliation (M2), develop pressure shadows of quartz, chlorite and muscovite, and are typically aligned with each other, defining a relic early foliation (S? in Fig. 6a–c). Samples containing pre-S2 porphyroblasts are rare, but widely scattered throughout the Ugab Zone (Fig. 2) and western Northern Zone (Fig. 3). These samples are not necessarily associated with outcropping granitoid intrusives, but nevertheless have mineral parageneses indicative of low-P, high average thermal gradient metamorphism that is otherwise commonly associated with heat advection by pluton emplacement. Consequently, growth of these porphyroblasts is interpreted to be associated with high heat flow during the earliest episode of granite emplacement at c. 590–570 Ma in the Ugab Zone and western Northern Zone region. Similar relationships are described in the Kaoko Belt, where coarse-grained Thermal Phase mineral parageneses (early M2) and granitoids of 580–571 Ma age are reworked by intense foliation development in the Transpressional Phase (Goscombe et al., 2003b
). Similarly, rocks in the Ogden Mylonite Zone (Figs 1 and 2) had early coarse-grained quartzo-feldspathic gneiss assemblages that were downgraded and reworked by a mylonitic Transpressional Phase foliation (Fig. 6d). The Thermal and Transpressional Phases comprise sequential periods during progressive oblique collision between the Rio de la Plata and Congo Cratons (Goscombe et al., 2003a, 2003b). Thermal Phase porphyroblasts in the Ugab Zone and western Northern Zone are not correlated with the 656–645 Ma episode of granite emplacement (M1), which is restricted to the westernmost Kaoko Belt and is not evident in the Inland Branch.
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Transpressional Phase (M2)
Transpressional Phase structures (D1–D3) of the M2 metamorphic cycle are pervasive, indicating intense deformation and crustal shortening that occurred during oblique SE-directed convergence (Fig. 4) and closure of the Adamastor Ocean between the Congo and Rio de la Plata Cratons (Dürr & Dingeldey, 1996
; Maloof, 2000; Passchier et al., 2002; Goscombe et al., 2003a, 2003b). Three texturally distinct fabrics (S1, S2 and S3) and associated fold structures formed during intense progressive deformation of the Ugab Zone and western Northern Zone during the Transpressional Phase, producing the regional metamorphic (M2) matrix assemblages. Aspects of Transpressional Phase structures have been previously described in the Ugab Zone (Freyer & Hälbich, 1994; Maloof, 2000; Passchier et al., 2002; Goscombe et al., 2003a) and in the western Northern Zone (Miller, 1980; Coward, 1983). A previously unrecognized S1 fabric is preserved in part, throughout the Ugab Zone and western Northern Zone. S1 foliations and rare quartz-aggregate and mineral stretching lineations are the earliest formed fabrics. S1 is always parallel to bedding and developed as a fine-grained schistosity of aligned white micas–chlorite ± biotite, that is best preserved in the lowest grade (sub-biotite grade) greywacke rocks and by grain-shape fabrics in carbonates. Small-scale (centimetre-scale) F1 isoclinal folds with axial planar S1 foliation are rare. Where preserved, S1 foliations are evident in outcrop by being overprinted at low angles (typically <20°) by the regionally pervasive S2 foliation and folding by F2. Pre-S2 flattening across bedding is evident by boudin trains that are shortened by F2 folds, thus constituting polyphase, non-congruent structures (Fig. 5; Goscombe et al., 2004).
S2 is the main foliation developed in all parts of the Ugab Zone and western Northern Zone; it is typically a strong schistosity of aligned biotite, muscovite and chlorite, and is also developed as a very fine crenulation cleavage. S2 is axial planar to tight map-scale chevron folds with inter-limb angles averaging 41° ± 24° and 0·1–6·0 km wavelengths (Figs 2 and 3). F2 folds dominate the Ugab Zone and western Northern Zone (Figs 2 and 3); these had sub-horizontal hinges prior to F5 refolding and trends ranging from NNW to NNE (Fig. 5). Axial surfaces are typically sub-vertical but are moderately inclined SE in the Ugabmond Region and SW in the Khorixas Region (Figs 2, 3 and 5). In the Ogden Mylonite Zone, S2 is an upper greenschist-facies mylonitic foliation developed by the grain-refinement of pre-existing coarse-grained quartzo-feldspathic gneisses (Fig. 6d). This mylonitic foliation trends north–south and is sub-vertical to steeply inclined (70°) to the east (Figs 2 and 5); 
-type mantled porphyroclasts (Passchier & Simpson, 1986) and asymmetric boudin trains (Goscombe & Passchier, 2003) indicate a sinistral shear sense along shallowly north-plunging stretching linations (Fig. 5). The pervasive and intense D2 deformation of the Ugab Zone occurred in a transpressional environment (Passchier et al., 2002) and is correlated with the pervasive deformation of the Kaoko Belt during the Transpressional Phase (Passchier et al., 2002; Goscombe et al., 2003a, 2003b). Transpressional Phase deformation and metamorphism occurred between 580 Ma and 552 Ma in the Kaoko Belt (Goscombe et al., 2003b), which is consistent with the syn- to post-D2 granitoids of 570 ± 20 Ma to 573 ± 33 Ma age in the Ugab Zone (Kröner, 1982; Miller & Burger, 1983a).
S2 is overprinted by a spaced (1–5 mm) S3 crenulation cleavage with axial planar alignment of stubby biotite grains, producing distinctive biotite-rich seams (Figs 7a–c and 8e). D3 crenulations and crenulation lineations are similarly oriented to D2 structures in most domains (Fig. 5), but are locally best developed at high angles to S2, where S2 has rotated into orientations conducive to crenulation. S3 is uncommonly developed but heterogeneously partitioned across all domains (Fig. 5), being most common in the Ugabmond and Twyfelfontein Regions (Figs 2 and 5). In the Ugab Zone, tight to isoclinal map-scale F3 folds are developed only in the Goantagab Region, where they are the dominant map-scale structures developed. Map-scale F3 folds are common in the western Northern Zone and responsible for fold interference patterns between coplanar F2 and F3 structures (Miller, 1980; Coward, 1983).
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Shortening Phase (M3)
The Shortening Phase or M3 metamorphic cycle (D4–D6 in Ugab Zone and western Northern Zone and main fabric in the eastern Northern Zone) encompasses all deformation structures, granites and metamorphic parageneses formed during NNE–SSW shortening across the Inland Branch (Fig. 4) during high-angle convergence of the Congo and Kalahari Cratons after closure of the Khomas Sea (Coward, 1981
, 1983; Miller, 1983; Freyer & Hälbich, 1994; Maloof, 2000; Passchier et al., 2002; Goscombe et al. 2003a). Shortening Phase deformation in the Ugab Zone is interpreted to have involved NNE–SSW shortening with a component of sinistral transpressional flow (Passchier et al., 2002; Goscombe et al., 2003a). In the Ugab Zone and western Northern Zone a sequence of three texturally distinct deformation fabrics (D4, D5 and D6) were generated during the Shortening Phase. Shortening Phase strain was insignificant in the Kaoko Belt and is progressively more intense across the Ugab Zone and Northern Zone towards the east, producing the dominant, main phase structures in the eastern Northern Zone and Central Zone (Fig. 3; Miller, 1983; Coward, 1983; Goscombe et al., 2003a). Throughout the Ugab Zone and western Northern Zone, Shortening Phase structures are restricted to ductile fabrics in contact aureoles and widespread development of upright folds, crenulations and kinkbands. In contrast, the eastern Northern Zone has a strong ductile mica schistosity that developed during the regionally pervasive and tight, north over south vergent fold and thrust nappes. This foliation both pre-dates and envelops garnet porphyroblasts (Fig. 9a–c), indicating progressive evolution of the main foliation during north–south shortening, and is not overprinted by any late-stage foliations or crenulations. Shortening Phase deformation was synchronous with emplacement of the 530 ± 3 Ma Footspore Pluton in the Ugab Zone (Seth et al., 2000) and 530 ± 11 Ma mineralization (Kamona et al., 1999) and 535 ± 13 Ma K–Ar whole-rock ages from phyllites in the Northern Platform (Clauer & Kröner, 1979). Maximum K–Ar whole-rock ages of 490 ± 11 Ma from the Ugab Zone (Ahrendt et al., 1983) indicate cooling through 350°C and cessation of deformation by this time.
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Protracted Shortening Phase deformation accompanied large volumes of granitoid emplacement and deformation continued after the youngest aplite dykes in the Ugab Zone (Miller, 1983
; Goscombe et al., 2003b). Shortening Phase deformation was partitioned into the contact aureoles, the higher temperatures being conducive to relatively intense ductile deformation and the formation of new mineral parageneses (M3). D4 ductile fabrics are developed in contact aureoles and within some granite bodies; including steep schistosity (S4) subparallel to granite margins and sub-vertical stretching lineations (Fig. 5). S4 is a transposed foliation of aligned micas that is typically so intense that it obliterates earlier foliations in all but a few samples. D4 strain in contact aureoles produced mutually orthogonal boudin neck axes, indicative of flattening strains (Casey et al., 1983). D4 strain in contact aureoles is highly variable and folding of late aplite dykes produces interlimb angles ranging from 0° to 130°.
Broadly coeval with the D4 structures within the aureoles, D5 structures are developed beyond the contact aureoles and are expressed as upright, open to close folds and associated crenulation cleavages (Weber & Ahrendt, 1983; Freyer & Hälbich, 1994; Hoffman et al., 1994; Maloof, 2000; Passchier et al., 2002). D5 structures are heterogeneously partitioned into 5–10 km wide zones of relatively high sinistral strain, and are often associated with outcropping, and possibly subsurface, granite plutons. F5 folds have interlimb angles ranging from 35° to 150° (averaging 104° ± 31°), east–west- to NNE–SSW-trending axial surfaces and vertical to moderate plunging hinges (Fig. 5). The scale of F5 folds ranges across many orders from 0·1 m to 20 km wavelengths. S5 crenulation cleavages invariably have sinistral asymmetry vergence and orientations consistent with the macroscopic fold structures (Freyer & Hälbich, 1994; Maloof, 2000; Passchier et al., 2002). S5 crenulation cleavages overprint the S4 foliation in contact aureoles (Fig. 8c and f). Multiple episodes of S5 crenulation cleavages are developed in the Twyfelfontein Region, the latest formed being shallowly inclined to the ESE with down-dip plunging hinges. The latest formed, regionally developed structures are vertical S6 kinkbands, with NE-trending sinistral kinkbands and NW-trending dextral kinkbands indicating north–south 1.
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METAMORPHIC EVOLUTION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY OF THE...STRUCTURAL EVOLUTION AND...METAMORPHIC EVOLUTIONDISCUSSIONSUPPLEMENTARY DATAREFERENCES |
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Petrography
Eastern Northern Zone: metapelite schists
Avdale Namibia Pty. Ltd. supplied drill core of Damara Sequence metapelite (sample NZ1) from Platveld Farm, 50 km SW of Otavi. This is the only amphibolite-facies metapelite described from the eastern Northern Zone of the Inland Branch. Sample NZ1 has a well-developed schistose matrix of quartz–andesine–biotite with kyanite laths syntectonic with this main foliation (Table 1). Sub-idioblastic garnet porphyroblasts overgrow the kyanite–biotite foliation (Fig. 9a and c) and are also enveloped by it, indicating progressive development of the foliation (Fig. 9a–c). Consequently, garnets are considered broadly late syn-kinematic. Vague sigmoidal inclusion trails in garnet are defined by quartz and are continuous with the matrix foliation; matrix biotite and andesine were resorbed during garnet growth and are not represented in the inclusion assemblage (Fig. 9b; Table 1). Matrix andesine occurs as small grains with irregular margins indicating partial resorption. Very thin (0·1 mm) andesine moats corrode garnet margins (Fig. 9c) and represent a second generation of plagioclase growth. Matrix andesine is partially retrogressed to late-stage muscovite ± chlorite, garnet is retrogressed to chlorite and kyanite develops thin muscovite coronas (Table 1). Minor rutile and graphite in the matrix foliation (Table 1) indicate low oxygen fugacity metamorphic conditions.
Western Northern Zone and Ugab Zone: metapelite schists
Metapelite schists in the western Northern Zone and Ugab Zone show a simple metamorphic isograd pattern (Figs 2 and 3). In both zones, metapelites sufficiently distant (typically 5–20 km) from granitoid bodies preserve regional metamorphic (M2) assemblages of biotite to sub-biotite grade. The extent of the M3 contact metamorphic overprint is documented by the incoming of biotite laths across S2 and S3 in metapelites, and the nearly coincident garnet-in isograd in calc-silicate rocks. The M3 cordierite–andalusite isograd is typically within 5 km of the granite margin (Figs 2 and 3), and within this zone the mineral assemblages are almost entirely formed by contact metamorphism. In the Central Region of the Ugab Zone, a zone of cordierite–andalusite parageneses that are distant from any outcropping granites suggests a subsurface granite body (Fig. 2).
Metapelite samples that are distant from the cordierite–andalusite contact aureole (Fig. 3) have a fine-grained phyllitic foliation (S1 and/or S2) with quartz–muscovite ± biotite ± chlorite ± ilmenite ± calcite assemblages. A bedding-parallel schistosity (S1) of aligned micas is overprinted by a very low-angle crenulation cleavage, which in most samples develops as a new transposed ductile schistosity (S2) with similar assemblage to S1 (Fig. 10b). In some samples the S2 foliation envelops early porphyroblasts of andalusite, biotite, chlorite or ilmenite, which develop chlorite ± quartz pressure shadows (Fig. 6a–c). Pre-S2 biotite porphyroblasts have exsolved ilmenite (Fig. 6b), indicating high-Fe and -Ti compositions with respect to S2 biotite compositions. Typically, pre-S2 biotite porphyroblasts are aligned with each other, defining an early foliation (Figs 6a and 10a). These pre-S2 porphyroblasts indicate the existence of a metamorphic mineral growth prior to development of the pervasive, ductile S2 foliation. This relationship is identical to the textural relationships developed during the Thermal Phase (early M2) in the Kaoko Belt (Goscombe et al., 2003a, 2003b). Distant from contact aureoles, post-S2 mineral growth occurs only as discordant laths of biotite and less commonly muscovite or chlorite in the axial surfaces of S3 crenulation cleavages (Figs 7c, 8e and 10c), often forming biotite-rich seams (Fig. 7b); S5 crenulations are devoid of new mineral growth.
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All mineral parageneses developed in contact aureoles grew in association with Shortening Phase deformation (S4–S6) during granite emplacement, and earlier fabrics (S1–S3) were obliterated to varying degrees. In contact aureoles the regionally pervasive S2 foliation is transposed by the growth of coarser-grained biotite and muscovite (Fig. 8a and b), forming a steep S4 foliation as a result of shortening at a high angle to the granite margin, within the contact aureole (Figs 5 and 10e). M3 matrix assemblages in contact aureole metapelites are typically quartz–biotite–muscovite–ilmenite–cordierite–andalusite ± tourmaline ± plagioclase or quartz–biotite– andalusite ± muscovite ± ilmenite ± tourmaline ± plagioclase ± chlorite (Table 1). Less common matrix assemblages are quartz–biotite–muscovite–cordierite ± ilmenite ± tourmaline ± plagioclase and quartz– biotite–garnet ± muscovite ± ilmenite ± plagioclase. No systematic difference has been recognized between the matrix assemblages developed in all nine contact aureoles investigated (Table 1; Figs 2 and 3). Matrix plagioclase is andesine–oligoclase and is present in only a quarter of the investigated contact aureole metapelite samples. Garnet-bearing metapelites are very rare (three samples); the garnets are small (<2 mm), poikiloblastic or inclusion free, and stabilized by very high MnO compositions (see below) as indicated by Tinkham et al. (2001)
. Staurolite occurs in only one sample (U351b), as small relic grains with irregular highly resorbed margins. The staurolite is almost totally resorbed by the growth of contact metamorphic cordierite porphyroblasts and the biotite–ilmenite–plagioclase matrix assemblage. Staurolite grains pre-date the M3 contact metamorphic parageneses and are interpreted to be relics from an earlier (possibly M2) assemblage. No other relic phases are present in this sample, thus the pre-M3 assemblage and its conditions of formation cannot be determined. Furthermore, the relic staurolite contains 2·2% ZnO (see below) and was thus potentially stabilized at lower temperatures and pressures than is typical of the staurolite stability field (Alias et al., 2002). The textural relationships and sequence of contact metamorphic (M3) mineral growth in metapelites within the cordierite–andalusite isograd are consistent throughout all nine contact aureoles investigated across the western Northern Zone and Ugab Zone (Figs 2 and 3; Table 1). The sequence of contact metamorphic mineral growth is typically initiated by the formation of porphyroblasts of cordierite, biotite, ilmenite and rarely plagioclase. These porphyroblasts overgrow the S2 and S3 foliations, and poikiloblastic cordierite often preserves the earlier S2 and S3 relationships as inclusion trails (Figs 8e and 10d). The early M3 cordierite–biotite–ilmenite porphyroblasts are invariably enveloped by a pervasive quartz–biotite–muscovite ± ilmenite ± plagioclase S4 schistosity (Figs 8d and 10e; Table 1). Pre-S4 biotite porphyroblasts are deformed into mica-fish by the newly formed S4 foliation. Cordierite porphyroblasts were flattened and pseudomorphed to varying degrees, forming a foliated aggregate of biotite–muscovite–chlorite ± ilmenite. This pseudomorph assemblage develops a foliation that is aligned with the enveloping S4 foliation, with which it is correlated (Figs 8d, e, and 10e). Andalusite porphyroblasts are typically syn- kinematic with the S4 schistosity (Fig. 8d), or overgrow the S4 foliation (Fig. 8a and f). Pre-S4 andalusite porphyroblasts occur in a few samples, and these are pseudomorphed to muscovite ± biotite during development of the enveloping S4 foliation (Table 1). Fine-grained sillimanite occurs in only two samples, forming as late-stage aggregates that overgrow biotite and occur in cordierite pressure shadows, but typically sillimanite is parallel to the S4 foliation. Post-S4 muscovite and biotite growth is evident as muscovite ± biotite pseudomorphs after andalusite (Fig. 8f), muscovite plates across the cordierite pseudomorph assemblage, and muscovite and biotite plates aligned with S5 crenulation cleavages (Fig. 8c). Chlorite is the latest formed phase, forming both aligned with the S5 crenulation and as random plates across S5 crenulations and also retrogressive after S5 biotite (Fig. 10f). These described relationships constrain the diagnostic sequence of mineral parageneses developed in western Northern Zone and Ugab Zone contact aureoles as: pre-S4 cordierite–biotite–ilmenite ± plagioclase, S4 andalusite–biotite–muscovite–chlorite ± ilmenite, late-S4 sillimanite in few samples, S5 biotite– muscovite–chlorite and retrograde chlorite.
Western Northern Zone and Ugab Zone: calc-silicate rocks and mafic schists
Thin (5–25 cm) felsic calc-silicate inter-beds are a minor component of the Kuiseb Formation and found in all regions investigated. Calc-silicate beds in contact aureoles have a fine-grained (<0·5 mm) polygonal granoblastic matrix assemblage of quartz–anorthite– calcite–titanite–ilmenite with small (1–3 mm) poikiloblastic porphyroblasts of Mn-rich garnet and actinolitic hornblende (Table 2). Minor matrix phases in some samples include K-feldspar, epidote, apatite and ferrosalite–clinopyroxene. Clinopyroxene has epidote–hornblende coronas developed and chlorite is a common retrograde phase replacing garnet and biotite. Calc-silicates that are distant from the cordierite–andalusite contact aureole have similar assemblages but are typically devoid of garnet, anorthite and clinopyroxene (Table 2). Mafic schist has been recorded at only two localities; both are in the easternmost Ugab Zone. Sample U355a is close to the contact aureole and has a medium-grained polygonal granoblastic and aligned assemblage of quartz–albite–magnesio-hornblende–epidote–titanite (Table 2). Sample U170(ii) is distant from the contact aureole and has a fine-grained schistosic assemblage of albite–actinolite–chlorite–epidote– titanite–calcite (Table 2).
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Ogden Mylonite Zone
The western margin of the Ugab Zone is marked by the Ogden Mylonite Zone, a 5 km wide, sub-vertical, sinistral shear zone with shallow north-plunging stretching lineations (Fig. 5). Low-strain domains and isolated porphyroclasts in the mylonites and protomylonites indicate that the original rock types were meta-arenites with coarse-grained Thermal Phase (early M2) assemblages of quartz–plagioclase–K-feldspar–ilmenite ± epidote ± titanite ± muscovite (Table 2). The whole domain was pervasively sheared by ductile to semi-ductile grain-size reduction, resulting in a low-grade mylonitic foliation (Fig. 6d) comprising quartz aggregate ribbons and biotite–chlorite–ilmenite ± muscovite ± calcite (Table 1). Feldspars underwent mechanical grain-size reduction to subgrains during shearing without the growth or recrystallization of new feldspars (Table 1). The mylonitic fabric progresses continuously into and is correlated with the pervasive S2 foliation in the Ugab Zone. The mylonitic assemblages are identical to the biotite-grade regional metamorphic assemblages (M2) developed in the S2 foliation of the Ugab Zone (Table 1). There is no M3 contact metamorphic overprint recorded in the Ogden Mylonite Zone. Thick carbonate units and thin calc-silicate bands form a minor component of the Ogden Mylonite Zone. Calc-silicates are composed of a fine-grained matrix of quartz–anorthite–calcite–ilmenite–titanite–chlorite with grain-shape fabric and post-kinematic poikiloblastic porphyroblasts of Mn-rich garnet and magnesio-hornblende (Table 2).
Mineral chemistry
Methods
Mineral analyses from the Ugab and Northern Zones were obtained using a Cameca SX51 electron microprobe at the University of Adelaide. An operating voltage of 15 kV and current of 20 nA were used for all phases except micas (10 nA) and feldspar (15 nA), and a beam radius of 2 µm for most phases and 4 µm for micas and feldspars. Natural silicates were used as standards and checked periodically. Mineral analyses from metapelite and calc-silicate samples are described in detail below.
Garnet
Garnets in contact aureole metapelites are rare, small and unzoned. The total compositional range of all samples is narrow and is expressed as mole fractions in the octahedral site as: XFe2+ (0·58–0·59), XMg (0·07–0·11), XCa (0·03–0·04) and XMn (0·27–0·33), and Fe/(Fe + Mg) ratios (0·85–0·90). The poikiloblastic garnets in calc-silicates show very weak zoning with Fe-rich and Mn-poor rims. The total compositional range recorded by all samples is: XFe2+ 0·28–0·47, XMg 0·02–0·08, XCa 0·14–0·26, XMn 0·25–0·49 and Fe/(Fe + Mg) ratios 0·85–0·95. Garnets in the Barrovian metapelite sample NZ1 are typically large (3–12 mm) and growth zoning records increase in Fe and Mg, slight increase in Ca and decrease in Mn towards the inner rim of garnets (Fig. 11). A thin outer rim of variable width dependent on the cation component formed by re-equilibration during garnet resorption, and is Mn-rich and Mg-, Fe- and Ca-poor (Fig. 11). Core to inner-rim growth zoning has the following compositional range: XFe2+ 0·72–0·77, XMg 0·11–0·17, XCa 0·08–0·09, XMn 0·09–0·01 and Fe/(Fe + Mg) ratios 0·87–0·82. The thin re-equilibrated rims have compositions with XFe2+ 0·70, XMg 0·13, XCa 0·06, XMn 0·08 and Fe/(Fe + Mg) ratio 0·85.
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Cordierite
Cordierites in contact aureole metapelites are elliptical porphyroblasts (up to 40 mm long). These cordierites are weakly zoned with slightly Mg-enriched rims. Fe/(Fe + Mg) ratios range from 0·30 to 0·39 in all samples. Na cations range from 0·05 to 0·44 in all samples.
Biotite
Biotites in contact aureole metapelites are unzoned. Fe/(Fe + Mg) ratios range from 0·37 to 0·53 and Ti cations from 0·07 to 0·17 in all samples. Matrix biotite in the Barrovian metapelite sample NZ1 is weakly zoned from Fe/(Fe + Mg) values of 0·34–0·37 in cores to 0·35–0·38 in rims and the total range in Ti cations is 0·05–0·11.
Muscovite
Matrix and pseudomorph muscovite in contact aureole metapelites is unzoned. Si cations in all episodes of all samples range from 3·00 to 3·14 and the K/(K + Na + Ca) ratio ranges from 0·82 to 0·94. Late-stage muscovite in the Barrovian sample NZ1 has 3·19 Si cations and a K/(K + Na + Ca) ratio of 0·96.
Chlorite
Matrix and pseudomorph chlorite in contact aureole metapelites is unzoned ripidolite. The Fe/(Fe + Mg) ratios range from 0·41 to 0·46 and Al/(Al + Si) ratios range from 0·52 to 0·54 in all samples. Matrix chlorites in calc-silicates have Fe/(Fe + Mg) ratios ranging from 0·38 to 0·54 and Al/(Al + Si) ratios from 0·47 to 0·50 in all samples.
Feldspar
Matrix plagioclase in contact aureole metapelite is weakly zoned oligoclase and andesine with XAn ranging from 0·18 to 0·42 in cores and from 0·17 to 0·34 in rims. Matrix plagioclase in Barrovian metapelite sample NZ1 is andesine and preserves weak growth zoning in XAn from 0·38 in cores to 0·34 in rims. Thin plagioclase coronas that corrode garnet porphyroblasts (Fig. 9c) are andesine and have XAn = 0·40–0·55. Plagioclase in most calc-silicate samples is unzoned anorthite (XAn = 0·92–0·97) and sample U125 is zoned labradorite with XAn ranging from 0·54 in the core to 0·64 in the rim.
Staurolite
The relic staurolite in contact aureole metapelite sample U351b from the Goantagab Region (Fig. 2) is unzoned and Fe/(Fe + Mg) ratios range from 0·79 to 0·80; the Al/(Al + Si) ratio is 0·70 and weight percent ZnO ranges from 2·17 to 2·18%.
Ilmenite
Ilmenite in both contact aureole metapelites and calc-silicates is Mn-rich, with MnO ranging from 1·0 to 15·8% in all samples.
Amphibole
Matrix amphiboles in calc-silicates range from magnesio-hornblende to actinolite and do not show consistent compositional zoning patterns. In all samples the Fe/(Fe + Mg) ratios range from 0·32 to 0·61, Si cations range from 6·6 to 7·8 and Ca + Na cations range from 1·75 to 2·20.
Epidote
Matrix epidote in calc-silicate sample U355b is weakly zoned with the mole fraction of Fe3+ in the M1 site ranging from 0·58 in cores to 0·51 in rims.
Clinopyroxene
Matrix clinopyroxene in calc-silicate sample U355b is unzoned ferrosalite with Fe/(Fe + Mg) ratios ranging from 0·55 to 0·56.
Aluminosilicates
Andalusites in contact aureole metapelites contain 0·24–0·27% FeO, and kyanite in NZ1 contains 0·18% FeO.
Pressure–temperature calculations
Methods
Samples with typical metapelite bulk compositions, representing all contact aureoles in the Ugab Zone and western Northern Zone and the Barrovian metapelite sample NZ1 from the eastern Northern Zone, were selected for P–T calculations (Electronic Appendices B and C). A representative number of these metapelite samples have also been analysed by whole-rock X-ray fluorescence (XRF) analysis, or bulk composition was estimated from modal mineral proportions and average mineral compositions (Electronic Appendix D). The metapelite samples used in this study have a restricted range of bulk compositions, with FeO/(FeO + MgO) averaging 0·50 ± 0·02 and Al2O3/(Al2O3 + FeO + MgO) averaging 0·47 ± 0·03. Consequently, bulk compositional variations should not exert systematic bias in the calculated P–T results. A number of calc-silicate rocks were also investigated for comparison with the results from metapelites to test for bulk composition effects on P–T calculations.
In all samples from the Ugab Zone and western Northern Zone, mineral rim compositions are used for P–T calculations and the results are interpreted to represent peak metamorphic conditions (Spear, 1993). In the Barrovian metapelite sample (NZ1), sub-peak metamorphic conditions were calculated after the arguments of Spear (1993) and Kohn & Spear (2000). These sub-peak metamorphic conditions were calculated using matrix biotite cores, inner rims of matrix plagioclase and the inner rims of garnet porphyroblasts, being careful to avoid the Mn-rich outer rims. These analyses represent P–T conditions at the latest stage of garnet prograde growth combined with the latest stage of plagioclase growth that are still preserved after dissolution of a portion of the mineral rims. These P–T calculations are interpreted to represent the highest grade metamorphic conditions preserved on the prograde path, below the peak of metamorphism (Spear, 1993; Kohn & Spear, 2000). Mineral end-member activities were calculated by the method of Holland & Powell (1990) using the program AX (Powell et al., 1998). P–T calculations used the average P–T approach of Powell & Holland (1994) and were performed using THERMOCALC v3.1 (Powell & Holland, 1988) with the 1998 thermodynamic dataset (Powell et al., 1998). A fluid composition of XH2O = 1·0 was assumed for sample NZ1 and XH2O = 0·9 was assumed for all cordierite-bearing contact aureole metapelite samples. Calc-silicate samples were calculated assuming a fluid composition of XH2O = 0·5 for calcite-bearing samples and XH2O = 1·0–0·85 for calcite-free samples (Table 3; Electronic Appendix B). All results satisfy the 
2 test and errors from THERMOCALC, incorporating typical uncertainty for each mineral end-member activity and errors in the thermodynamic dataset, average ±35°C and ±1·1 kbar for contact aureole samples and are ±84°C and ±1·6 kbar in sample NZ1 (Electronic Appendix B). For comparison with the average P–T calculations, the depth of crustal residence during crystallization of granitoid bodies was estimated using the Al-in hornblende geobarometer (Schmidt, 1992; Electronic Appendix C), and that for contact aureole metapelites was estimated using the Si-in phengite geobarometer (Massonne & Schreyer, 1987; Electronic Appendix C).
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Results
Peak-metamorphic conditions based on the average P–T calculations are summarized in Table 3 and represented in Fig. 12. Average P–T results (n = 7) from contact aureole metapelites (M3) from across the northern margin form a tight cluster with a pooled mean of 556 ± 26°C and 3·2 ± 0·6 kbar (Fig. 12; Table 3), corresponding to an average thermal gradient of 49·6°C/km. These results are entirely consistent with the phase stability field of the biotite–muscovite–andalusite ± cordierite contact aureole matrix assemblages, in the range of 540–570°C and 2·5–3·2 kbar in the P–T pseudosection for metapelites of the same bulk composition (Tinkham et al., 2001
; Figs 12 and 13) and are thus considered plausible estimates of the equilibration conditions. Such consistent P–T results along a 200 km long sector of the northern margin are remarkable.
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A further 200 km east along the northern margin, M3 metamorphic conditions accompanying main phase deformation in the Shortening Phase are distinctly different, with higher pressures and lower average thermal gradients. Sub-peak, average P–T calculations from sample NZ1, using garnet inner rims, matrix plagioclase inner rims and biotite cores, give pooled results of 570 ± 93°C and 8·0 ± 1·8 kbar (Table 3; Electronic Appendix B). These results are 50°C lower than the phase stability field of the developed peak metamorphic assemblage in NZ1 (Fig. 12). The garnet–biotite– kyanite–quartz matrix assemblage has a phase stability field of >620°C and >7 kbar in KFMASH for average metapelite bulk compositions (Figs 12 and 14; Powell et al., 1998
; Vance & Mahar, 1998; Tinkham et al., 2001; Pattison et al., 2002). This stability field is bounded by the following reactions (Figs 12 and 14);
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The absence of partial melt and K-feldspar constrain an upper limit to peak metamorphic conditions of 710–750°C (Fig. 13; Chatterjee & Froese, 1975
; White et al., 2001). The up-P prograde path documented below (Fig. 12) indicates that the true peak metamorphic conditions must be at higher P–T than the maximum sub-peak P–T calculation of 8·5 kbar (Electronic Appendix B) and constrained within the phase stability field described above. This would suggest a best estimate of peak-T metamorphic conditions of at least 635°C for a pressure of 8·7 kbar (Table 3; Figs 12 and 14). This estimate corresponds to an average thermal gradient of 20·9°C/km (Table 3), which is typical of Barrovian metamorphism. Average P–T calculations from the re-equilibrated outer rim of garnets, biotite rims and andesine coronas give a pooled mean of 513 ± 75°C and 6·3 ± 1·4 kbar (Table 3; Electronic Appendix B). These results are interpreted to represent the conditions of formation of this re-equilibration assemblage, formed during post-peak decompression and andesine corona growth (Fig. 14).
Average P–T calculations from contact aureole (M3) calc-silicate samples have a pooled mean (n = 6) of 501 ± 29°C and 3·4 ± 0·8 kbar, corresponding to an average thermal gradient of 42·1°C/km (Table 3). These results have similar pressures but consistently lower temperatures (
T averages –55°C) compared with the results from associated metapelite samples (Table 3). Calc-silicate assemblages are fine-grained and are interpreted to have continued to re-equilibrate during isobaric cooling from the peak of M3 metamorphism. Re-equilibration may occur without accompanying mineral recrystallization (Goscombe et al., 1998) and the average P–T results represent the cessation of cation exchange (Harley, 1992). The low pressures of M3 contact metamorphism in the Ugab Zone and western Northern Zone indicated by phase stability (Fig. 13) and THERMOCALC results (Table 3) are supported by conventional geobarometry results from both the contact aureole and intruding granitoid. The pooled mean (n = 14) of Si-in phengite geobarometer results from contact aureole white micas is 3·1 ± 0·5 kbar (Table 3; Electronic Appendix C) and identical to pooled THERMOCALC results. The pooled mean (n = 23) of the Al-in hornblende geobarometer results from three Ugab Zone granitoids is 4·1 ± 0·3 kbar (Table 3; Electronic Appendix C). These Al-in hornblende results are consistently higher (P = 1 kbar) than P estimates by all other methods. All hornblende-bearing granitoids represent the earliest intrusive phase in each composite pluton and so the Al-in hornblende P estimates may represent pressures related to hornblende crystallization at deeper crustal levels than final emplacement of the pluton.
A single calc-silicate sample in the Ogden Mylonite Zone at the western edge of the Ugab Zone has an average P–T result of 479 ± 15°C and 4·7 ± 0·8 kbar, corresponding to an average thermal gradient of 29·1°C/km (Table 3). The S2 matrix assemblage in this sample formed during the Transpressional Phase (M2) of deformation that also produced the main foliation during pervasive sinistral reworking in the Kaoko Belt, prior to north–south contraction in the Shortening Phase (M3). Metapelites from the Ogden Mylonite Zone have M2 assemblages of quartz–muscovite–biotite–chlorite– ilmenite, corresponding to temperatures up to 570°C (Powell & Holland, 1990; Vance & Mahar, 1998; Tinkham et al., 2001). Consequently, the average P–T result from the calc-silicate sample may underestimate the metamorphic conditions during M2. This may be the result of re-equilibration of this fine-grained assemblage during cooling, similar to calc-silicate samples in M3 contact aureoles (see above). The Ogden Mylonite Zone average P–T result has pressures comparable with the average P–T calculation of 547 ± 51°C and 4·3 ± 1·0 kbar (Goscombe et al., 2003b), from M2 parageneses in the Coastal Terrane of the Kaoko Belt, with which it is directly along strike (Fig. 15). In contrast to the low pressures during M3 in the Ugab Zone, the higher pressures experienced during M2 in the Ogden Mylonite Zone and Coastal Terrane are also supported by the Al-in hornblende geobarometer from the southernmost Coastal Terrane, which gives a pooled average result of 5·6 ± 0·3 kbar (Table 3; Electronic Appendix C).
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Pressure–temperature paths
Barrovian metamorphism (M3): eastern Northern Zone
Garnet porphyroblasts in sample NZ1 preserve growth zoning compositional patterns with progressive increase in Fe, Mg and Ca and decrease in Mn and Fe/(Fe + Mg) towards inner rims (Fig. 11). These patterns are typical of growth zoning (Tracy, 1982
; Loomis, 1983; Tuccillo et al., 1990; Spear, 1993) and document a portion of the prograde P–T trajectory during garnet growth (Spear et al., 1984; St-Onge, 1987). Calculated modal garnet and garnet compositional isopleths for a range of typical metapelite bulk compositions all have broadly similar patterns of increasing modal garnet and XFe and decreasing XMn towards higher P and T, and decreasing XCa to higher T at mid-amphibolite grade (Spear, 1993; Vance & Mahar, 1998; Vance et al., 1998). Consequently, the growth of garnet in Barrovian sample NZ1 indicates heating with burial. Compositional variation across NZ1 garnet can be used as a semi-quantitative constraint on the prograde P–T path in this sample. Variation in garnet composition is compared with the XFe, XCa and XMn isopleths in the P–T pseudosection (Fig. 14) calculated by Vance & Mahar (1998) for a metapelite bulk composition very similar to that estimated for sample NZ1 (Electronic Appendix D). Decreasing XMn with increasing XFe and XCa in sample NZ1 documents a prograde P–T path that must have involved a steep (P/T) increase in P from c. 4·0 kbar at 560°C to >7·5 kbar at 590°C (Fig. 14). P–T conditions based on intersecting compositional isopleths in garnet inner rims match closely with sub-peak average P–T calculations (Fig. 12; Table 3). These results confirm the maximum sub-peak P–T conditions preserved by the prograde inner-rim mineral compositions in the garnet, with the peak metamorphic rim compositions dissolved by post-peak re-equilibration reactions.
Phase relationships in sample NZ1 can be interpreted within the Vance & Mahar (1998) P–T pseudosection (Fig. 14). Matrix plagioclase has andesine cores preserving prograde growth to lower XAn compositions, indicating up-P trajectories (Spear, 1993; St-Onge & Ijewliw, 1996). Matrix plagioclase occurs as small relic grains, indicating that it was consumed during prograde garnet growth, as is evident by increasing Ca compositions in garnet during growth (Fig. 11). In the Vance & Mahar (1998) P–T pseudosection, the prograde (up-T) path from early kyanite–plagioclase parageneses to matrix garnet–kyanite–biotite parageneses must have involved increasing P, comparable with the trajectory defined by garnet compositional isopleths (Fig. 14). Peak metamorphic conditions constrained by the kyanite–garnet–biotite stability field in KFMASH (Powell et al., 1998; Vance & Mahar, 1998; Tinkham et al., 2001; Pattison et al., 2002) and minimum pressures recorded by the sub-peak P–T calculations (Table 3; Electronic Appendix B), indicate that conditions reached c. 8·7 kbar and 635°C (Fig. 12). Thin andesine coronas corrode peak metamorphic garnet, indicating post-peak decompression back into phase stability fields with plagioclase parageneses (Fig. 14). Retrogressive sillimanite was not formed and decompression was restricted to the kyanite field, followed by cooling and retrogressive chlorite and muscovite growth, documenting a clockwise P–T path (Fig. 14). Calculated isopleths for modal plagioclase and XAn in typical metapelites (Spear, 1993) indicate a decompressive P–T path for the growth of post-peak andesine coronas after garnet. Corona andesine (XAn = 0·35–0·55) has higher Ca compositions than matrix andesine (XAn = 0·34–0·38), supporting post-peak decompression (Spear, 1993). Average P–T calculations based on Mn-enriched garnet outer rims, andesine coronas and biotite rims give pressures at least 2·9 kbar lower than estimated for peak parageneses (Table 3; Electronic Appendix B), supporting a decompressive post-peak evolution. Clockwise P–T paths are typical at mid-crustal levels in collisional orogens, where crustal over-thickening gives rise to burial, radiogenic heat production and metamorphism, which is terminated by isostatic uplift and erosion (England & Thompson, 1984; Spear et al., 1984; Jamieson et al., 1996).
Contact metamorphism (M3): Ugab Zone and western Northern Zone
The general topology of a variety of published P–T pseudosections in KFMASH, for moderately aluminous metapelites (above the garnet–chlorite tieline) at low pressures, are broadly similar over a range of typical metapelite compositions (Xu et al., 1994; Powell et al., 1998; White et al., 2000; Tinkham et al., 2001; Alias et al., 2002). The bulk composition of seven metapelite samples from Ugab Zone and western Northern Zone contact aureoles have been determined directly by whole-rock XRF analysis (Electronic Appendix D). The bulk composition of an additional three samples has been estimated using average mineral compositions and modal proportions (Electronic Appendix D). The bulk composition of all these samples is typical of aluminous metapelites in general, and the range is narrow, centred on means of FeO/(FeO + MgO) = 0·50 ± 0·02 and Al2O3/(Al2O3 + FeO + MgO) = 0·47 ± 0·03. These compositions are comparable with the bulk composition used in the P–T pseudosection calculated by Tinkham et al. (2001) for the KFMASH system with excess quartz and H2O, and with CaO, MnO and Na2O removed (Fig. 13). The T–XFe and T–XAl pseudosections calculated by Tinkham et al. (2001) show that the general topology of the P–T pseudosection (Fig. 13) does not change over the compositional range of the investigated samples. Consequently, the sequence of mineral growth recorded in Ugab Zone and western Northern Zone contact aureole metapelites can be interpreted within the Tinkham et al. (2001) P–T pseudosection (Fig. 13).
The general sequence of M3 mineral growth in cordierite–andalusite metapelite samples is from an M2 muscovite–chlorite foliation overgrown by early M3 cordierite porphyroblasts or rarely andalusite porphyroblasts. Cordierite is enveloped by an andalusite– biotite–muscovite S4 foliation assemblage that rarely also contains late-stage sillimanite. Cordierite is typically pseudomorphed to biotite–muscovite–chlorite assemblages and is in turn overprinted by late-stage chlorite– muscovite assemblages. This sequence documents a tight anticlockwise P–T loop, with up-P trajectories through the peak of metamorphism from cordierite parageneses to andalusite ± sillimanite parageneses followed by cooling to chlorite parageneses (Fig. 13). The sequence from cordierite to andalusite parageneses implies crossing of the reaction muscovite + cordierite 
andalusite + biotite + quartz + H2O, which has either a shallow positive P/T (Tinkham et al., 2001) or shallow negative P/T (Pattison et al., 2002). Consequently, all potential P–T trajectories through the peak of metamorphism must involve either an up-P component or isobaric heating. Decompressive cooling paths and thus clockwise P–T paths are not possible because in all samples andalusite post-dates cordierite and late-stage cordierite has not been documented (Fig. 13). Different samples record varying portions of this mineral growth history, and all document tight anticlockwise P–T loops, together defining a set of self-similar P–T paths centred on 2·5–3·2 kbar with peak temperatures between 540 and 570°C (Fig. 13). The general form of these P–T loops is the same in all contact aureoles investigated (Fig. 13). Both the employed P–T pseudosection (Fig. 13; Tinkham et al., 2001) and the average P–T calculations (Table 3) were calculated using THERMOCALC v3.1 (Powell & Holland, 1988) with the 1998 thermodynamic dataset (Powell et al., 1998), and so both are considered directly comparable. The average P–T calculations from contact aureole metapelites average 556 ± 26°C and 3·2 ± 0·6 kbar during M3 (Table 3; Fig. 12), and span a range identical to peak metamorphic conditions outlined by the phase relationships (Fig. 13). Consequently, the P–T loops defined in Fig. 13 are considered semi-quantitative both in form and location in P–T space. The best estimate of crustal depth for the Ugab Zone and western Northern Zone during M3 is 3·2 kbar, which is approximately equivalent to 11·2 km. Such low pressures of metamorphism and temperatures as high as 570°C indicate geothermal gradients as high as 51°C/km, which, along with the tight anticlockwise P–T loops, is typical of contact metamorphic environments (Spear, 1993).
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M2 tectono-metamorphic evolution
Throughout the Ugab Zone and western Northern Zone, regional metamorphism during the Transpressional Phase (M2) ranged from biotite to sub-biotite grade (Table 1), indicating temperatures in the vicinity of 450–525°C (Powell & Holland, 1990
; Tinkham et al., 2001). Pressures were low, as indicated by rare pre-S2 andalusite porphyroblasts, restricting pressures to below 4·5 kbar (Powell & Holland, 1990) and Al-in hornblende P estimates from early stage granitoids averaging 4·1 ± 0·3 kbar (Electronic Appendix C). The estimates correspond to an average thermal gradient of 29–37°C/km. Average P–T calculations from Transpressional Phase parageneses are 479 ± 15°C and 4·7 ± 0·8 kbar in the Ogden Mylonite Zone and 547 ± 51°C and 4·3 ± 1·0 kbar in the linked southern Coastal Terrane (Fig. 15; Goscombe et al., 2003b). These results are broadly consistent with the above P–T constraints for regional metamorphic conditions in the Ugab Zone and western Northern Zone. M2 average thermal gradients are moderately elevated, but not sufficiently high to be typical of contact metamorphic environments (Spear, 1993). Average thermal gradients may have been even higher during the Thermal Phase, in the early M2 metamorphic period. In the Ogden Mylonite Zone, relict, Thermal Phase migmatitic quartzo-feldspathic parageneses may indicate temperatures as high as 670°C (Powell & Holland, 1990; Fig. 15) and higher thermal gradients than experienced during M2 Transpressional Phase metamorphism.
Nevertheless, both Thermal Phase (early M2) and Transpressional Phase (M2) regional metamorphic conditions are considered anomalously high geothermal gradients and so must have occurred in a high heat flow setting. One plausible high heat flow scenario is lithospheric thinning (Wickham & Oxburgh, 1985; Sandiford & Powell, 1986; Buck et al., 1988). The low pressures and moderately high thermal gradients preclude metamorphism caused by crustal overthickening and radiogenic heat production (England & Thompson, 1984; Jamieson et al., 1996). Similarly, the paucity of 580–570 Ma Transpressional Phase granites in the Ugab Zone and western Northern Zone (Fig. 15) precludes the advection of heat into the upper crust by granites. Lithospheric thinning and associated high heat flow (Buck et al., 1988) may have generated some granitic melts (i.e. the result, not cause, of metamorphism), but these were not pervasive enough to be considered the main mechanism for heat transfer into the upper crust and cannot be considered the cause of regional metamorphism. An alternative scenario is by thermal subsidence following rifting (Sandiford et al., 1998). Thermal subsidence of a radiogenic passive margin basement may produce high thermal gradients in the overlying upper crust, if blanketed beneath an insulating sedimentary sequence (Sandiford et al., 1998; McLaren et al., 1999). Modern terrestrial heat flow in the Ugab Zone and western Northern Zone is anomalously high, being c. 70 mW/m2 (Ballard et al., 1987). Outcropping basement rocks of the Kamanjab Inlier (Fig. 1) are dominated by alkali granites and granitic gneisses that are moderately enriched in radiogenic heat-producing elements (Burger et al., 1976; C. McKenzie, personal communication, 2002). Thus, if basement rocks of the Huab Ridge underlying the Ugab Zone and western Northern Zone are similarly characterized, these could plausibly produce high heat flows (Jaupart et al., 1981) and the observed M2 regional metamorphism in overlying Damara Sequences.
Deformation in the Ugab Zone involved a transpressional flow regime (Passchier et al., 2002) and is correlated with pervasive transpressional deformation in the Kaoko Belt (Maloof, 2000; Passchier et al., 2002; Goscombe et al., 2003a, 2003b). The Ugab Zone and western Northern Zone region spans the triple junction between the Kaoko Belt and Inland Branch. The triple junction geometry with basement promontory resulted in the progressive deflection of the stress field across the region, producing the range in orientation of structures in this orocline bend (Coward, 1981, 1983; Maloof, 2000; Fig. 5). Transpressional Phase deformation, crustal shortening and thickening, burial and associated metamorphism occurred between 590 and 553 Ma within the wider Kaoko Belt, Ugab Zone and western Northern Zone. Pervasive deformation in the Kaoko Belt is constrained by peak metamorphic garnet Sm–Nd ages of 576 ± 11 Ma (Goscombe et al., 2003b) and pre- to syn-kinematic granitoids with zircon ages of 580–570 Ma and bracketed by post-kinematic granites of 553 Ma age (Seth et al., 1998; Franz et al., 1999). Transpressional Phase deformation in the Ugab Zone and western Northern Zone is less well constrained but spans a similar broad range, with zircon ages from syn-kinematic granites ranging from 570 to 589 Ma (Kröner, 1982; Miller & Burger, 1983a, 1983b).
M3 tectono-metamorphic evolution
In the western portion of the northern margin, M3 metamorphism was restricted to growth of new mineral parageneses in the contact aureoles of granitoids emplaced during Shortening Phase deformation (D4–D6), with essentially no new mineral growth beyond these thermal aureoles. NNE–SSW-directed Shortening Phase strain was partitioned into the ductile contact aureoles, developing a strong ductile S4 fabric and rarely also tight to isoclinal folds, obliterating most of the earlier Transpressional Phase fabrics. The sequence of mineral growth recorded in metapelite schists is similar in all contact aureoles investigated, documenting common deformation and metamorphic histories. The thermal effect of granite emplacement was typically first recorded by the growth of cordierite ± biotite ± ilmenite porphyroblasts prior to development of the ductile transposed foliation. This indicates that the contact aureole region was not strongly strained until the yield strength of the rocks was sufficiently reduced by the thermal pulse to deform in a ductile manner. Cordierite porphyroblasts were flattened and pseudomorphed to a foliated chlorite–biotite–muscovite assemblage that is coplanar with the enveloping S4 matrix foliation with muscovite–biotite–ilmenite–andalusite ± plagioclase assemblages. The S4 foliation was overprinted by rare late-S4 sillimanite and S5 crenulation cleavages with axial planar muscovite and late-stage chlorite that crosscuts S5.
The sequence of mineral growth documented in all the investigated aureoles tracks tight anticlockwise P–T loops through 540–570°C and 2·6–3·2 kbar in the KFMASH pseudosection (Tinkham et al., 2001) for metapelites of similar composition (Fig. 13). Metamorphic conditions are verified by average P–T calculations by THERMOCALC v3.1 (Powell et al., 1998), with resulting pooled means of 556 ± 26°C and 3·2 ± 0·6 kbar (Table 3; Fig. 12). The best estimate of crustal depth of the entire western northern margin during M3 is 3·2 ± 0·2 kbar, which is approximately equivalent to 11·2 ± 0·7 km. Such low pressures of metamorphism and temperatures as high as 570°C in contact aureoles indicate average geothermal gradients as high as 51°C/km, which, along with the tight anticlockwise P–T loops, are typical of contact metamorphic environments (Spear, 1993).
M3 Shortening Phase convergence between the Kalahari and Congo Cratons resulted in pervasive deformation and metamorphism in the eastern Northern Zone that contrasts strongly with the low-strain contact metamorphism in the western Northern Zone and Ugab Zone. The eastern Northern Zone experienced moderate-P Barrovian metamorphism of garnet–kyanite–biotite grade, low average thermal gradients and clockwise P–T path. The phase stability field of the garnet– kyanite–biotite matrix assemblage in KFMASH (Powell & Holland, 1990; Powell et al., 1998; Vance & Mahar, 1998) constrains peak metamorphic conditions to be c. 635°C and 8·7 kbar (Fig. 12), corresponding to a low average thermal gradient of only 21°C/km (Table 3). The rims of matrix plagioclase and garnet are dissolved, thus average P–T calculations using THERMOCALC (Powell & Holland, 1994) from inner-rim compositions give sub-peak metamorphic conditions (Kohn & Spear, 2000) with a pooled mean of 8·0 ± 1·8 kbar and 570 ± 93°C (Table 3). Garnet compositional isopleths document a steep (P/T) burial trajectory from 4·0 kbar and 560°C to >7·5 kbar and 590°C in the portion of growth-zoned garnet still preserved (Fig. 14). Peak-T metamorphism was terminated by decompression indicated by the resorption of garnet rims and growth of andesine coronas (Vance & Mahar, 1998). Late-stage sillimanite was not formed and the decompressive P–T path is interpreted to be restricted to the kyanite field (Figs 12 and 14). A post-peak P–T path involving both decompression and cooling is supported by average P–T calculations of 6·3 ± 1·4 kbar and 513 ± 75°C (Table 3; Fig. 12) using the Mn-enriched garnet outer rims, biotite rims and plagioclase coronas that make up the re-equilibration assemblage formed during decompression. Steep burial, low average thermal gradient metamorphism and clockwise P–T path are all consistent with deformation and crustal over-thickening in the fold and thrust nappe belt of the eastern Northern Zone (England & Thompson, 1984; Jamieson et al., 1996).
The Shortening Phase of deformation was associated with NNE–SSW-directed convergence across the Inland Branch between the Congo and Kalahari Cratons, at c. 550–500 Ma (Fig. 4; Electronic Appendix A; Clauer & Kröner, 1979; Kröner & Clauer, 1979; Miller, 1983; Kamona et al., 1999; Jung et al., 2000a, 2000b). M3 assemblages and fabrics all formed in association with NNE–SSW convergence in all regions of the northern margin and Northern Platform and so must be broadly contemporaneous. Peak metamorphism in the Northern Platform (Fig. 15) is constrained by a K/Ar whole-rock age of 535 ± 13 Ma from phyllite (Clauer & Kröner, 1979), white-mica Rb–Sr age of 537 ± 7 Ma (Kröner & Clauer, 1979) and 530 ± 11 Ma Pb–Pb model age for syn-tectonic galena (Kamona et al., 1999). Age of Shortening Phase metamorphism in the western Northern Zone is poorly constrained by granite Rb–Sr whole-rock ages of 548 ± 31 Ma and 517 ± 11 Ma, and a 521 ± 45 Ma Rb–Sr isochron from a metamorphosed felsic volcanic rock (Hawkesworth et al., 1983). The Footspore Granite in the Ugab Zone (Fig. 2) was emplaced during the Shortening Phase (Goscombe et al., 2003a) and the early syenogranite phase has been accurately dated at 530 ± 3 Ma by the Pb–Pb zircon evaporation technique (Seth et al., 2000). Ar–Ar cooling ages are in the range of 510–490 Ma in the Ugab Zone (D. Gray & D. Foster, unpublished data, 2001), indicating that deformation and granite emplacement had terminated before this time. To the south, within the high-grade Central Zone of the Inland Branch, Shortening Phase deformation spanned a protracted period of main phase of deformation. U–Pb monazite and titanite and Sm–Nd garnet ages document two high-grade metamorphic events during shortening of the Central Zone of 538–516 Ma and 511–505 Ma (Jacob et al., 2000; Jung et al., 2000a, 2000b). The inferred deformation ages and cooling ages for the Shortening Phase give a regular T–time distribution throughout the entire northern margin region (see Goscombe et al., 2003b, fig. 14), and no domainal variation is apparent. No direct dating of the M3 mineral assemblages associated with pervasive deformation and NNE–SSW shortening of the eastern Northern Zone is currently available. Nevertheless, given that main phase NNE–SSW-directed shortening occurred between 516 and 537 Ma in all other domains to the north, south and west, M3 metamorphism accompanying shortening of the eastern Northern Zone is interpreted to be broadly coeval with that in these surrounding regions.
Along orogen variation in M3 metamorphic style
Shortening Phase deformation was responsible for closure of the Khomas Sea and NNE–SSW-directed shortening across the Inland Branch of the Damara Orogen (Coward, 1981, 1983; Freyer & Hälbich, 1994; Maloof, 2000; Passchier et al., 2002; Goscombe et al., 2003a). The degree of crustal shortening and resultant crustal thickening increases significantly from west to east along the northern margin of the Inland Branch. Shortening Phase deformation resulted in open to close upright folding in the Ugab Zone (Freyer & Hälbich, 1994; Goscombe et al., 2003a), close to tight folding and a thrust northern margin in the western Northern Zone and intense deformation and pervasive fabric development in the fold and thrust belt of the eastern Northern Zone (Coward, 1981, 1983). This deformation pattern was superimposed on a parallel variation in the thickness of the Damara Sequence overlying the passive margin basement. The Damara Sequence is a very thin, 1·7 km succession in the Ugab Zone (Swart, 1992), and possibly has similar thicknesses in the western Northern Zone (Miller, 1980). In the eastern Northern Zone the Damara Sequence is a thick succession that was deposited in a deep trough centred on the Northern Graben (Miller, 1983; Porada, 1983; Porada & Wittig, 1983). This combination of variations in thickness of the cover sequence and degree of crustal shortening has resulted in very different orogenic processes and metamorphic histories at the two ends of the northern margin.
Tight anticlockwise P–T loops during the Shortening Phase in the western northern margin indicate that burial during deformation was possibly less than 1–2 kbar (Fig. 13). Locally the thermal Peclet number is low (Sandiford, 2002); heat transfer by conduction from the granite is more significant than by advection by deformation. This is in contrast to metamorphism in the eastern Northern Zone, where heat transfer by advection dominates over conduction. Burial in the eastern Northern Zone was a lot more significant (between 4·5 and 8·7 kbar) and also accompanied Shortening Phase deformation. Metamorphism of the buried rocks was the result of radiogenic heat production in the thickened crustal section (England & Thompson, 1984; Jamieson et al., 1996), resulting in a clockwise P–T path (Fig. 15). Although metamorphism in both parts of the northern margin accompanied crustal shortening, the two regions experienced different style of metamorphism and P–T paths. This is due to the cumulative thermal effect of rocks deposited in a thick trough sequence, which then experienced greater crustal shortening and concomitant burial (Fig. 12). Furthermore, the average thermal gradient was significantly lower because the source of heat was internally driven by radiogenic heat production, and not externally driven by the emplacement of granites.
It is often assumed that there is no change in metamorphic style along the length of linear to arcuate convergent orogens. The northern margin of the Inland Branch clearly illustrates the potential for contrasting styles of metamorphism developing at broadly the same time in different sectors along the length of a single orogenic margin. In this case, three processes were important in the disparate thermal histories: (1) initial and evolving differences in crustal architecture, including thickness of the overlying sedimentary succession; (2) variation in degree of shortening and thus crustal thickening and denudation rates; (3) variation in input of heat into different sectors by granite emplacement. Variation in thermal history and the developed metamorphic style, in otherwise apparently simple linear or arcuate margins of high-angle convergent orogens, may be more widespread than is currently acknowledged. Furthermore, along orogen metamorphic variation may be the result of any number of combinations of many other first-order processes active in convergent orogens. For example, conventional wisdom assumes a common, self-similar metamorphic history for the 2000 km long arcuate Central Himalayan Metamorphic Front (e.g. Hodges, 2000). Nevertheless, it has been shown that sectors of the Central Himalayan Metamorphic Front vary considerably in the metamorphic conditions in the High Himal Crystallines of the overriding upper plate (Goscombe & Hand, 2000). The High Himal Crystallines were extruded between the underthrust Indian crust and overriding Tibetan plateau terranes, resulting in vertical and southward mass transfer (Grujic et al., 1996). The variation in metamorphic grade and average thermal gradient recorded in these rocks may be the result of variation in the rates of extrusion along the orogenic front (Beaumont et al., 2001; Grujic et al., 2002).
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Supplementary data for this paper are available on Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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Introduction to the northern margin by Paul Hoffman and discussions in the field with Cees Passchier and Rudolph Trouw are gratefully acknowledged. This research was supported by private funds, Namibian Geological Survey car support awarded to Cees Passchier (1999), ARC Discovery Grant awarded to David Gray (2001–2003) and Adelaide University Small Grant awarded to Ben Goscombe. The Universities of Adelaide and Melbourne are sincerely thanked for their support during the writing up of this work. The work in Namibia was greatly eased by the support and encouragement of Mimi Duneski and Thomas Bekker. Avdale Namibia Pty. Ltd. kindly supplied the critical sample NZ1 from their drill core. Djordje Grujic, Mark St-Onge and Kurt Bucher are gratefully acknowledged for their comments and editorial efforts.
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* Corresponding author. Present address: Geological Survey of Western Australia, PO Box 1664, Kalgoorlie, W.A. 6433, Australia. E-mail: ben.goscombe{at}adelaide.edu.au
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