Structural evolution of the Chewore inliers, Zambezi mobile belt, Zimbabwe

Journal of African Earth Sciences, Vol. 19, No. 3, pp. 199-224, 1994 Elsevier Science Ltd Printed in Great Britain 0899-5362/94 $7.00 + 0.00

Pergamon

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Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe B. GOSCOMBE, P. FEY and F. BOTH Geological Survey of Zimbabwe, P.O. Box CY 210, Causeway, Harare

(Received 24 January 1994 : accepted 8 August 1994) Abstract - Detailed mapping of the Chewore Inliers, Zambezi valley, Zimbabwe has recognized three distinct geological terranes. The areally most significant "gneissic terrane" is dominated by amphibolite facies quartzofeldspathic gneisses with minor migmatitic, sillimanite/kyanite-gamet-biotite metapelites, kyanite-staurolitegarnet schists, calc-silicate rocks, amphibolite and quartzite. The "quartzite terrane" is comprised almost entirely of quartzite with minor kyanite-bearing metapelitic gneisses and the "granulite terrane" of anhydrous quartzofeldspathic gneisses, intermediate two-pyroxene gneisses and sillimanite-bearing metapelitic gneisses. The boundary between the quartzite and granulite terranes is a steep mylonite zone that involved oblique overthrusting of the granulites to the SW. The boundary between the quartzite and gneissic terranes is also discordant and possibly a thrust surface, although it is not exposed anywhere. The granulite terrane involved two phases of isoclinal folding (Fg2 and Fg3) by S to N tzansport and with accompanying high grade metamorphism. The granulites have well-annealed granoblastic textures. No strong planar foliation has developed, although two mineral lineations are recognized. The quartzite and gneissic terranes display a strong layer-parallel $2-L2 fabric that immediately post-dates the peak of the major metamorphic event (M1). The orientations of D2 structural elements are different in the two terranes. Furthermore, these terranes display different orientations of the major tight to isoclinal folding (D~) and the gneissic terrane displays a second tight folding event (D4). Both the gneissic and quartzite terranes developed conjugate crenulation cleavages (Ss) and all terranes have N-S trending open warps (F4band F~t). SW over NE tectonic transport during D2 and S over N transport during Dg3 cannot be correlated with any known orogenic periods in central eastern Africa. However, SE over NW transport during D3 is tentatively correlated with the Irumide orogeny (1100-1300 Ma) and the NE over SW transport, during D4-Ds, with the Zambezi/Mozambique orogeny (830i30 Ma). Post-Ds events involved over-thrusting of the granulite terrane and possibly also the quartzite terrane to the SW, followed by the intrusion of the Chewore Ultramafic Complex, dolerite dykes and pegmatites. Late-stage pegmatites may have occurred at Pan-African times (450-650 Ma). The igneous and ductile structural evolution of the Chewore Inliers ceased before Karoo times (150-285 Ma). Metamorphic parageneses indicate that a total of 21 km of crust has been removed at some stage subsequent to M1 metamorphism and prior to the deposition of the Karoo Supergroup sediments on the Chewore gneisses. A protracted period of faulting and uplift of the inl/ers during and after Karoo times resulted in the Chewore Inliers being a significant topographic high within the Zambezi Rift valley. R~sum~ - La cartographie d~taill~ de la fen~tre de Chewore, dans la vall~,e du Zambezi au Zimbabwe, a permis de mettre en ~vidence trois blocs crustaux distincts. Le "bloc gneissique", le plus important en surface, est principalement form~ par des gneiss quartzo-feldspathiques du facies amphibolite; il comprend ~galement en quantit~s subordonn~zs des m/gmatites, des m~tap~lites t s'dliman/te/cyanite-grenat-biotite, des micaschistes a cyanite-staurolite-grenat, des roches calco-silicat~es, des amphibolites et des quartzites. Le "bloc quartzitique" est presque enti~rement compos~ de quartzites avec des quantitC,s mineures de gneiss l~litiques a cyanite. Le "bloc granulitique" est quant a lui form~ de gneiss quartzo-feldspathiques anhydres, de gneiss/~ deux pyrox~mes de composition interm~diaire et de gneiss p~litiques a silllmartite. Le contact entre les blocs quartzitique et granulitique est form~ par une zone mylonitique redress~ qui indique le surcharriage oblique des granulites vers le SW. Le contact entre les blocs quartzitique et gneissique est ~galement discordant et probablement aussi de type charriage, mats il n'affleure nulle part. Le bloc granulitique comprend deux phases de plis isoclinaux (Fg2et Fg~) marquant un transport du sud vers le nord et accompagn~es d'un m~tamorphisme de haut degr~. Les granulites poss~ent de belles textures granoblastiques indiquant une complete recristallisation; la foliation planaire est peu d~velopp~ mais deux lin~ations min~rales sont bien visibles. Les blocs gneissique et quartzitique d~tiennent une fabrique S'z-L2 en ~l~ments parall~les tr~s bien marquee qui suit immc~diatement le pic m~tamorphique principal 0Vii). Les ~16ments structuraux D2 sont orient~ ~ m m e n t dans les deux blocs. De plus, ces blocs poss~dent des I~ (plissement majeur resserr~ A isoclinal)/~ orientation ~ t e et seul le bloc gneissique montre une deuxiCSmep h a ~ de plis serr~s (D4). Par contre, les deux blocs d6veloppent un m&me clivage de ~ u l a t i o n conjugu~ (Ss) et tous les blocs sont affect~ par des courbures ouvertes de direction N-S (F¢oet Fs4). Lea transports tectoniques du SW au hiE lors de D2 et du S au N durant D3 ne peuvent ~tre corral,s/l aucune p~riode orog~nique connue en Afrique centro-orientale. Par contre, nous proposons de rapporter le transport du SE au NW de Eb a l'orog/me irumide (1100-1300 Ma) et le transport du NE au SW de D4-Ds /~ l'orog~me du Mozambique/Zambezi (830-~30 Ma). Les ~v(mements post-Ds comprerment le surcharriage du bloc granulitique et probablement du bloc quartzitique vers le SW suivi de l'intrusion du complexe ultramafique de Chewore, de filons

199

200

B. GOSCOMBE, P. FEY and F. BOTH dol~ritiques et de pegmatites. Les pegmatites tardives pourraient ~tre pan-africaines (450-650 Ma). L'6volution magmatique et structurale des fen~tres du Chewore s'est arr~t~e avant la p~riode Karoo (150-285 Ma). Les paragen~ses m~tamorphiques indiquent qu'une ~paisseur totale de 21 km de croflte a ~t~ enlev~e entre le m~tamorphisme M1 et le d~pOt des s~diments du Supergroupe du Karoo sur les gneiss de Chewore. Une longue p~,riode de faiUage et de remontc~e des fen~tres pendant et aprons la p~riode Karoo est & l'origine de l'anomalie topographique positive des fen~tres du Chewore dans la vall~e du rift du Zambezi. INTRODUCTION

The 1600 kln 2 C h e w o r e Inliers c o m p r i s e 20 horsts of a m p h i b o l i t e a n d g r a n u l i t e facies s u p r a c r u s t a l gneisses located w i t h i n the Z a m b e z i Rift valley, Z i m b a b w e (Figs. 1 a n d 2). T h e p r e s e n t s t u d y b e g a n in

1990 a n d c o n s t i t u t e s the first d e t a i l e d analysis of the region. Earlier references to the inliers i n c l u d e a r e p o r t b y Wiles (1956) a n d m o r e recently a brief mention byMunyanyiwa(1993). The Chewore Inliers are important for u n d e r s t a n d i n g central e a s t e r n A f r i c a n P r o t e r o z o i c

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Tectonic Transport Vectors. Pan-African (/+50-650 Ha). ~

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Prototith Orogeny Ha. Ha. .... Pan-African overprint, south Limit -/+50-650 ~1 Zambezi Belt: deformed, undeformed. 820-880 830-+30 Mozambique Belt: granulites, gneisses. 960-1100 830-+30 B Rushinga & Umkondo Groups. 1075-1376 830+-30 Irumide Belts: deformed, undeformed. 1355-1500 1100-1300 ProfoLith Age Unknown. ? ? Ubendian Belts. 2000-2200 1660-1850 R Archean Pateoproterozoic BeLts. > 2360 ;~2000 ~1 Reworked Archean Craton. [ I Archean Cratons.

Figure 1. Simplified tectonic map illustrating the orogenic framework of central eastern Africa. Some of the thrusts presented in the literature are highly interpretive. The dashed line in the Zimbabwe Craton is the southern limit of 450-650 Ma isotopic ages that dominate the region to the north. Chewore Inliers are indicated by (C). Compiled from: Pinna et al. (1993), Choubert et al. (1968), Johns et al. (1989), Daly (1986a, 1988), Barton et al. (1993), Piper (1989), Kr6ner (1977), Coward and Daly (1984), W i l s o n et al. (1993) and Daly et al. (1984).

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

201

Figure 2. Simplified geological map of the region around the Chewore Inliers, compiled from: Pinna et al. (1993),Johns et al. (1989), Daly (1986a), Barton et al. (1993), Coward and Daly (1984), Daly et al. (1984), Fey and Broderick (1990), Treloar (1988), Thieme and Johnson (1977), Broderick (1976), Wiles (1961), Bache et al. (1990a, 1990b), Kirkpatrick (1976), Hahn et al. (1990b),Barton et al. (1991), Stocklmayer (1980), Treloar and Kramers (1989) and the geological maps of Zimbabwe (1:1 000 000) and Mozambique (1:2 000 000). (a) Late Proterozoic, Karoo Supergroup and younger sedimentary cover, Co) Pan-African (450-650 Ma) granites, (c) Katangan Group, (d) Rushinga Groups and Umkondo Group, (e) high-grade gneisses and granulites of the Mozambique Belt, (f) quartzites of all ages including the Muva Group, (g) Muva Group gneisses not including quartzites, (h) foliation trace in basement gneisses of largely unknown age, (i) Proterozoic granite, (j) Magondi Mobile Belt, (k) possibly Archaean escarpment gneisses and other paragneisses, (1) Archaean Zimbabwe Craton, solid black is ultramaflc igneous bodies, (m) mineral elongation lineation, (n) second elongation lineation at a locality, (o) thrusts - largely interpretive, (p) ductile shear zone, (q) initial translational sense, (r) later reactivation, translational shear sense, (s) tectonic transport vectors in the gneissic terrane of the Chewore Inliers, other arrows as labelled in Fig. 1, (t) location and label of structural domains referred to in Fig. 12, (u) southern limit of 450-650 Ma isotopic disturbance, (v) eastern limit of isotopic disturbance due to the Magondi Mobile Belt (approximately 1700 Ma) and (w) international boundaries. geology because of their location at the junction of the Irumide, Zambezi and M o z a m b i q u e orogenic belts (Figs. 1 and 2). Furthermore, this p a p e r offers structural constraints on the evolution of a complex, but poorly understood, region of central eastern Africa. The identification of three distinct geological terranes (Howell 1989) within the C h e w o r e Inliers implies that the published tectonic f r a m e w o r k of central eastern Africa (e.g. Fig. 1) is too simplistic and that in reality the region comprises a complex mosaic of geological provinces juxtaposed late in their history. These terranes have been labelled the "granullte","gneissic" and "quartzite" terranes (Fig. 3). Within the gneissic terrane at least three tectonic events each with different vectors of tectonic transport are recognized. These have been tentatively correlated with existing Proterozoic age orogenic belts in central eastern Africa. At present there are no

geochronological data from the Chewore Inliers. Thus, all w i d e r tectonic correlations are poorly constrained and based entirely on correlations of structural elements. REGIONAL GEOLOGY Central eastern Africa comprises a mosaic of cratonized blocks that stabilized at approximately 2500 Ma (Zimbabwe and Tanzania Cratons, Cahen et al., 1984) and 1800 Ma (Bangwelu Block, Andersen and U n r u g 1984) and enclosed the Proterozoic mobile belts (Fig. 1). The mobile belts have a variety of ages of formation for the protoliths, but can be broadly classified into five major orogenic periods (Table 1). Each successive orogenic period reactivated (to some degree) earlier-formed mobile belts. In addition, paragneisses with calc-alkaline protoliths (Hahn e t al., 1990a) on the northern and western margins of the

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Lithotocjicat contact.

Chlorite- amphibole, anfigorife - talc ultra- marie rocks.

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Quartzite & quartz-mica schist.

Mafic amphibolife & mafic granulite.

Biotite & tourmaline-rich feldspathic cjneiss.

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Structural evolution of the Chewore Inliers, Zambezi MobileBelt, Zimbabwe

203

Table1. Simplifiedtable of tectonic events in the region between the Tanzania and ZimbabweCratons. Age Range Deposition/Emplacement of Protoliths 150-285 450-650

MajorTectonic Periods

Karoo sediments. Pan-African reworking and coolin ages through most of the region Lufilian-Zambezi Belt& Mozambiqu Belt.

830±30

Katangan Group (Zam.). Makuti Group (Zim.). 960-1100 • Mozambiquian gneisses and granulite formation. 820-880

1075-1376 Umkondo Group (Zim.) ( Rushinga Groups (c)). 1100-1300

Earliest Mozambiquian deformatio and metamorphism.

Irumide Belt.

1355-15007 Irumide Belt protoliths.

Ubendian Belt (Tanz.). Magondi Belt (Zim.).

1660-1850

2000-2200 Ubendian Belt protoliths an early deformation (Tanz.). Magondi Belt protoliths (Zim.).

Western Paragneisses (Zim.) Escarpment Gneisses (Zim.) Zimbabwe Craton

Earliest deformation Earliest deformation

Age(Ma). 150-280 538 450-650 480 820±7 830±30 843 800-860 820-880 3800-850 960-1100 850-1100 31000 1075-1376 1080[+140/-25] 1200-1230 1100-1355 1300±58 970-1100 >1355 <1500 <1610,1625 1725-1864 1780±70 1660-1850 1950-2100 2060 2000-2160 2000-2500 >2000 a2290-2500 a2500-3000 2600-2800

Source.

(a) (b,c)

(v) (d) (c) (0 (e,s,h)

(g) (s) (a) (h)

(w) 0) (d) (e,k) (1) (x,y)

(k) (P) (m) (n) (o) (p,u)

(n) (q) (r) (p) (k) (s) (c,z) (t)

Age ranges are only general maximum and minimum limits from the literature; errors have been mostly ignored. Sources of age data are not all from the original published geochronologicaldata. (a) Pinna et al. 1993; 0a) Barton et al. 1993; (c) Barton et al. 1991; (d) Hanson et al. 1988; (e) N'gambi et al. 1986; (f) Barr et al. 1978; (g) Wilson et al. 1993; (h) Andreoli 1984; (i) Allsopp et al. 1973; ~) Allsopp et al. 1989; (k) Cahen et al. 1984; (1) DrysdaU et al. 1972; (In) SneUing et al. 1964; (n) Lenoir 1993; (o) Treloar and Kramer 1989; (p) Vail et al. 1968; (co Kramers et al. 1989; (r) Master 1991; (s) Loney 1969; (t) Hawkesworthet al. 1975; (u) Hahn et al. 1990a; (v) Pires and Ferreira 1993; (w) Sacchi et al. 1984; (x) Daly 1986; (y) Ring et al. 1993; (z) Vail and Doclson1969. Zimbabwe Craton (escarpment and western paragneisses) have poorly constrained ages of 25003000 Ma (Vail and Dodson 1969; Barton e t al., 1991; L o n e y 1969) a n d t h e r e f o r e m a y h a v e A r c h a e a n protolith and tectonic ages. This is supported by their being intruded by the Great Dyke (2461+16 Ma, Hamilton 1977). The oldest Proterozoic mobile belts are the Ubendian and Magondi Mobile Belts of 2000-2200 Ma age (Fig. 1). The Ubendian Belt was tectonized at 1950-2100 Ma (Lenoir 1993) and both belts at approximately 1700-1800 Ma (Table 1). The Ubendian

Belt is comprised of at least eight terranes, each with different transport vectors, including NW-SE, E-W and NE-SW trending lineations (Daly 1988; Lenoir 1993). The Magondi Mobile Belt involved transport from NW over SE onto the Zimbabwe Craton (Treloar 1988). The rocks constituting the Irumide and Kibaran Belts formed in rift environments. The ages of these rocks are not well constrained but are possibly older than 1355 Ma (Cahen e t al., 1984; Table 1). These rocks, including the Muva Group metasediments (Fig. 2), were deformed at approximately 1100-1300 Ma by SE over N W displacement (tectonic

204

B. GOSCOMBE,P. FEY and F. BOTH

t r a n s p o r t ) r e l a t e d to o v e r - t h r u s t i n g in an intracontinental environment (Daly 1986a; Daly et al., 1984; N'gambi et al., 1986; Cahen et al., 1984; Ring et al., 1993). At 1000-1100 Ma, central eastern Africa under went a period of considerable crustal growth by plate convergence processes, resulting in the formation of the Mozambique Belt protoliths (Pinna et al., 1993). Coincident with this accretion was subduction, burial high-grade metamorphism and E-W collisional shortening (Pinna et al., 1993; Andreoli 1984; Table 1). Daly (1986b) considered the crustal accretion and early orogenesis of the Mozambique Belt to be coincident with the main phase of the Irumide Belt orogenesis. The "Pan-African" event has been used in the literature to incorporate the entire period (and all tectonism) subsequent to the Irumide orogenesis and the formation of the Mozambique Belt protoliths through to Karoo times (Pinna et al., 1993; Hahn et al., 1990a; Munyanyiwa and Blenkinsop 1993; Barton et al., 1991). This period, however, involved two distinct tectono-thermal cycles, thus a distinction must be made between the earlier and major tectonism at 830+30 Ma (N'gambi et al., 1986; Barton et al., 1991; Wilson et al., 1993) and the later wide-spread thermal event at 450-650 Ma (Table 1). The 830+30 Ma event is here named the Zambezi/Mozambique orogeny (Barton et al., 1991), which is equivalent to the Lufilian-Zambezi event of Daly (1986a) and the Mozambiquian orogeny of Andreoli (1984). This name is applied because deformation was restricted largely to the E-W trending Zambezi Mobile Belt at the northern margin of the Zimbabwe Craton, the Katangan Group and older rocks in Zambia and the N-S trending Mozambique Belt. Tectonism of the Mozambique Belt involved SSE to NNW transport onto the Irumide and possibly older basement in the Malawi Province (Cahen et al., 1984) and onto the Tanzania Craton in the north (Fig. 1; Pinna et al., 1993). The Zambezi Belt involved fold repetition and over-thrusting from NE over SW onto the northern, north-eastern and eastern margin of the Zimbabwe Craton (Barton et al., 1991, 1993; Daly 1986a) and onto the Irumide basement to the south in Zambia (Wilson et al., 1993; Daly 1986a; Coward and Daly 1984), as well as SW over NE transport in the Lufilian Arc (Daly 1986a; Fig. 1). To be consistent with the literature, the 450-650 Ma event is called the Pan-African (Cahen et al., 1984; Daly 1986a). Although wide-spread it does not appear to have given rise to regionally extensive penetrative tectonism; deformation was largely restricted to discrete thrust planes and shear zones in central eastern Africa (Pinna et al., 1993; Lenoir 1993; Sacchi et al., 1984). The event is widespread throughout much of the Gondwana fragments and

leads to thermal perturbation, shear zones, charnockite formation, post-tectonic granites and 450650 Ma cooling ages (Cahen et al., 1984; Santosh et al., 1992; Yoshida and Vitange 1993; Stuwe and Sandiford 1993). Pan-African thrusts in the Mozambique belt involved E over W transport (Fig. 1; Pinna et al., 1993; Barton et al., 1993). In the Lufilian Arc and on the northern margin of the Zimbabwe Craton, a second set of elongation lineations indicate N-S shortening (Daly 1986a, 1988); these are possibly post-830 Ma deformation and may therefore be PanAfrican. Subsequent to the Pan-African events, the mobile belts of central eastern Africa have been uplifted, presumably by isostatic rebound in response to crustal over-thickening experienced during the earlier orogenies, and the high-grade metamorphic gneisses were exposed by erosion. In Karoo times, the southern portion of central eastern Africa (Fig. 1) underwent updoming and crustal extension giving rise to the failed triple junction of the Luangwa and upper and lower Zambezi valley rifts (Oesterlen and Blenkinsop 1994). The Chewore Inliers are located near the centre of this failed triple junction. Rifting occurred later, during the Tertiary, to the north in the East African rift system. The orientations of these rift systems are largely governed by the structural grain of pre-existing Proterozoic mobile belts (Piper 1989). The Chewore Inliers are located in a poorly understood region of undated metamorphic tectonites between the Luangwa Rift and the Zimbabwe Craton where the Irumide, Zambezi and Mozambique Mobile Belts coincide. Consequently, it is anticipated that the orogenies that formed these belts could potentially be represented by deformational structures within the Chewore Inliers. The inliers are surrounded by Karoo cover and so their relationship with gneisses to the north and south of the Zambezi valley rift is unknown. MAJOR DIVISIONS OF THE CHEWORE INLIERS Three lithologically, metamorphically and structurally distinct terranes have been recognised in the Chewore Inliers (Fig. 3). The granulite terrane consists of granulite facies gneisses that, as a whole, comprise an entirely different suite of rocks with no compositional equivalents in the other terranes. Similarly, the quartzite terrane, being comprised almost entirely of quartzite, is entirely different to both of the other terranes, which are dominated by quartzo-feldspathic gneisses. Structurally the granulite terrane bears no resemblance to the others and so is described separately and events labelled differently (i.e. Dgl, Dg2 etc; Table 2). The quartzite and gneissic terranes appear to have a very similar tectonic history of

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

AQe (Me)

Protolith

Metamorphism

Structures

Up3

< 150 150 Upper Karoo. 245 285

Lower Karoo.

450-650? Late pegmatitesJ and quartz ve nsI

I

205

Post-Karoo uplift and faulting of Chewore Inliers.

~r Up2 D7 Up~ M.

Multiple faulting events and further uplift of Chewore Inliers. NW-SE open, upright folding of Lower Karoo. Culmination of denudation events since M1, exposing mid-crustal level rocks.

Thermal pulse. I

I Dolerite dykes. I I Chewore Ultra-

mafic Comp ex.

?[~D6

Sinsistral, oblique, southward overthrusting of terranes [NE-SW shortening].

oJT

E N 0

O

+

D4b

o

~,..

>,

c3 ~

_

pegmatltes.

Granulite Terrane.

Steep crenulation cleavages, NNE-SSW and ESE-WNW trends. [NE-SW shortening]

Garnet pegmatite.

I

N-S and NW-SE trending warps and E plunging open folds (steep in Q.T.).

Dolerite Dykes.

I

D4

Tight folds, plunge E to SE. [NE-SW shortening]

O3b

Open to close folds,

D3

M3

Gneissic & Quartzite Terranes.

Minor pegmatltes and sillimanite growth.

plunge SW. Tight to isoclinal folds, plunge SW (vertical in Q.T.). weak axial Sa fabric. [SE over NW transport]

I e'o"e0 I mafic dykes.

D2 Tectonic foliation assembla~les.

M2 I

Pegmatite,

MI

migmatite and orth~ff~i~.

Metamorphic peak, partial melting.

'7

Burial.

[i,

.I

Rare small isoclines, pervasive S2-L=fabric, Lz plunges SW (NE in QT). [SW over NE transport] $1 gneissic layering, granoblastic textures, porphyroblasts and partial melt segregations.

D~

Open steep folds I i and steep N-S cleavage. I 1 [E-W shortening]

Do3

Tight to isoclinal folds, plunge shallowly W. LQ3plunge shallowly W. [S to N transport]

M~2 DQ2

Peak of granullte facies I ._ _ _m_etamo_rph.i..sm_ -I Small intrafolial isoclines, I weak axial Sgzfoliation.| I_~ plunge steeply NNE. I [NNE-SSW shortening]/ S01gneissic layering.

. . . . . . . . . . . . . . . . . . . . . . . . . . .

Pre-D= fabrics.

Lithological layering of supra-crustals. Table 2. Simplified tectonothermal framework of events rccognised in the Chewor¢ Inliers. Correlation of structural events between terranes and correlation of events with absolute ages is discussed in text. No corre|ation between Granulite and other Terranes is implied.

206

B. GOSCOMBE,P. FEY and F. BOTH

events of similar style (Table 2), though structures in each are of different orientations, the quartzite terrane being generally very steep. Consequently, the lithologies of these two terranes are described separately but their tectono-thermal history is analysed together and similar style events labelled the same, though there is no conclusive evidence that they were coeval. Lithological layering within the terranes is discordant on both sides of the terrane boundaries, suggesting structural breaks and not unconformities. These boundaries are sharp and, for the most part, straight and thus presumed steep. There is no evidence of brittle faulting, such as brecciation and silicification, anywhere along these boundaries. The boundary between the granulite and quartzite terranes is a steeply NW dipping mylonite zone that involves oblique sinistral over-thrusting of the granulite terrane towards the SW. The nature of the quartzite/gneissic terrane boundary is not known, but, being of similar orientation to the granulite/quartzite terrane boundary thrust, is also thought to be a thrust (see later discussion). It is presumed that these terranes evolved separately during their ductile deformations and were tectonically juxtaposed late in their history (D6; Table 2). LITHOLOGIES OF THE CHEWORE INLIERS Gneissic Terrane

The gneissic terrane constitutes the south-eastern part of the study area (Fig. 3) and comprises a wide variety of quartzo-feldspathic gneisses (QFG). Most contain biotite, muscovite and garnet, though either muscovite or biotite can dominate as the major mica (Fig. 3). The majority are well-layered (S1) by the variation in mineral proportions on the mm- to dmscale, and well-foliated ($2) by aligned micas and quartz and feldspar aggregate ribbons. The proportion of mafic minerals varies considerably ranging from almost white, leucocratic muscovite+garnet gneiss to intermediate, biotite-rich garnet gneisses. The latter often have mm-dm wide biotite-rich seams alternating with more leucocratic layers. Hornblende-biotite-rich, homogeneous, intermediate quartzo-feldspathic gneisses are rare. This has only vague layering in parts and may have a meta-igneous protolith. Staurolite-garnet-biotite QFG, although rare, is interpreted to have metasedimentary protoliths. The following outlines the distinct types of quartzo-feldspathic gneisses. i) Heterogeneous medium-grained QFG that typically have gneissic layering and lack a granitoid appearance. These comprise the vast bulk of the gneissic terrane. The wide compositional variety, well

developed layering and intimate inter-layering with calc-silicate rocks and quartzites suggest sedimentary protoliths. This layered QFG can be dominated by either muscovite or biotite (Fig. 3). ii) Ten percent of the gneissic terrane is comprised of homogeneous, coarse-grained, biotite~garnet granitic orthogneiss that contain rare intermediate xenoliths. They have large K-feldspar augen with euhedral plagioclase inclusions and are largely devoid of compositional layering. They carry the regional $2-L2 fabric, which varies from nearly absent to very intense. The core of some orthogneiss units are less deformed with D2 deformation having been partitioned into the margins of these bodies. Granitic orthogneiss occurs as elongate, laterally persistent lenses within the typical, layered QFG sequences (Fig. 3). Orthogneiss units range from 1-10 m up to 1-2 km in thickness. They appear to be concordant and their margins are often transitional from more intermediate, layered QFG to progressively coarser grained and more augen-rich gneisses and, finally, into felsic orthogneisses. No intrusive contacts have been seen; the orthogneiss appear to be in situ partial melts of the host QFG. Calc-silicate rock lenses up to 500 m long form linear trains of large xenoliths within orthogneiss units, thus preserving pre-melfing lithological layering and implying passive in situ partial melting of the pre-existing host QFG. iii) Units of QFG with blocky K-feldspar megacrysts differ from typical granitic orthogneiss by having a medium-grained matrix and fewer megacrysts. The granitoid appearance is lacking and the rocks rarely display gneissic layering. These gneisses are often intimately interlayered with granitic orthogneiss and possibly represent a rock unit consisting of an intimate mixture of partially melted gneisses and gneisses that were on the verge of melting. iv) There is one unit of biotite-rich, well layered, intermediate QFG with tourmaline-bearing leucocratic segregations (Fig. 3). v) Inter-layered with the layered QFG are intermediate, homogeneous, massive garnet- and biotite-rich QFG. Some outcrops contain mafic xenoliths, suggesting that these gneisses may have in~usive protoliths. M e t a p e l i t i c biotite- s i l l i m a n i t e / k y a n i t e quartz- plagioclase- K-feldspar+garnet schists and gneisses form a minor component of the gneissic terrane. Sillimanite is the most common aluminosilicate, though kyanite is in textural equilibrium with sillimanite in many samples. The rocks are typically well layered on the cm-scale with alternating leucocratic segregations and biotitesillimanite/kyanite-garnet layers. In one 800 m section the proportion of leucocratic material progressively increased from typically layered

Structuralevolutionof the CheworeInliers,ZambeziMobileBelt,Zimbabwe metapelite to a coarse-grained homogeneous, leucocratic granitic gneiss with thin, randomly dispersed schlieren of sillimanite-biotite-garnet metapelite. This displays progressively larger proportions of partial melt in the host gneiss and implies a genetic connection between partial melting and the development of migmatitic layering in the metapelites. Feldspars in the leucocratic segregations in metapelites are often euhedral, which also suggests formation in a partial melt environment. The boudinage and grain refinement of feldspar augen within the leucocratic segregations suggest that partial melts had solidified prior to the pervasive D2 deformation. Within metapelite units, quartzites and pelitic quartzites are common, foliated amphibolites appear to be more abundant and quartzo-feldspathic gneisses are conspicuously uncommon. Laterally persistent and conformable kyanitebiotite-quartz+staurolite+garnet+tourmaline schists of only 5-50 m thickness are rare, being in only the central and southern portions of the terrane. Biotite (>80%)-quartz schists of cm- to dm-scale concordant seams in QFG may be selvedges formed during metamorphism. Rare amphibolites occur as laterally continuous and concordant layers of typically <1 m width. Most are foliated ($2-L2) and often have a vague gneissic layering defined by relatively plagioclase-rich layers. Their concordant nature may indicate supracrustal (volcanic) origins and their limited thickness (locally 2 cm) suggests that they are not sills. Most amphibolites consist of green hornblendeplagioclase- quartz+biotite+K-feldspar with occasional garnet or epidote. A coarse-grained hornblendite with minor quartz and plagioclase and abundant coarse (up to 6 cm) idioblastic garnet porphyroblasts is a distinctive type of mafic gneiss. Calc-silicate rocks are very rare, very competent lithologies and, as a result, commonly boudinaged into isolated lenses. Typical samples comprise a polygonal granoblastic matrix of quartz and plagioclase w i t h g r o s s u l a r , h o r n b l e n d e , clinopyroxene, epidote, clinozoisite, sphene and minor opaque minerals as subordinate components. A weak gneissic layering is developed. Quartzites constitute approximately <10% of the terrane. They are typically coarse-grained and have a well developed $2-L2 fabric that has a rodded appearance. Most contain muscovite and other phases including biotite, sillimanite, opaques, garnet, plagioclase, K-feldspar and tourmaline. Banded magnetite- and garnet-bearing quartzites are rare. Aluminous quartzites, consisting of sillimanitebiotite-J:garnet, are layered on mm- to cm-scale as distinct metapelitic bands within pure quartzite.

207

These metapelitic quartzites presumably had a metasedimentary protolith of alternating sandstone and siltstone-mudstone. A conformable, laterally continuous unit (<25 m wide) of dark green, homogeneous, ultramafic olivine websterite, with margins of chlorite schist, was mapped in the north. In the central region of the gneissic terrane is a discontinuous train of conformable pods of hornblende-chlorite schist and serpentinite-talc rock (Fig. 3). The concordant nature of these pods suggests that they were ultramafic lava flows or sills that were subsequently metamorphosed and boudinaged. Quartzite Terrane The quartzite terrane consists of a variety of quartzites (60-70 %), pelitic-quartzites and metapelitic gneisses. Quartzites are medium- to coarse-grained, of granoblastic texture and contain minor muscovite, biotite, feldspars and opaque minerals defining a weak compositional layering. Very coarse-grained (510 mm) granular, pure quartzite forms prominent ridges. Quartzites grade progressively into pelitic quartzites, quartz-rich metapelites and then into true metapelites. Metapelites consist of garnet, biotite, kyanite, siilimanite, rare cordierite, feldspars and opaque minerals. All display pronounced gneissic layering defined by the alternating layers of quartzfeldspar and biotite-kyanite/sillimanite-garnet. The metapelitic gneisses rarely develop migmatitic leucocratic segregations. The only quartzo-feldspathic gneiss in this terrane is one highly elongate, apparently discordant, unit of granitic orthogneiss which nevertheless follows the general NE-SW grain of the terrane (locality Z on Fig. 3). Concordant, thin (1-10 m) amphibolites are equally rare and consist of hornblende-plagioclasequartz+biotite and very occasional garnet. Three small 2-5 m wide units of magnetite-hematitequartz+garnet banded iron formation are recognized. No calc-silicate rocks and layered quartzo-feldspathic gneiss have been found in this terrane. Granulite Terrane. The granulite terrane is dominated by nearly anhydrous, l e u c o c r a t i c , l a y e r e d garnetsillimanite- biotite+ orthopyroxene QFG. These grade into layered metapelites giving rise to pelitic QFG rich in garnet, sillimanite and biotite. In contrast to the other terranes si|limanite is the dominant alumino-silicate and kyanite is very rare. Leucocratic coarse-grained migmatitic segregations are developed in garnet-bearing felsic granulite; these are boudinaged and deformed. They are presumably syn- to pre-Mg2 partial melts (Table 2). Metapelites display gneissic layering but migmatitic partial melt segregations are not common. Quartzites and layered

208

B. GOSCOMBE, P. FEY and F. BOTH

pehtic-quartzites are c o m m o n (Fig. 3). Quartzites are generally coarse-grained, comprise of 5-100 m wide units and are intimately inter-layered on the 5-50 cm scale with QFG. The abundance of quartzite and metapelites and close association with QFG suggests that the latter may have supracrustal (volcanic or sedimentary) protoliths. There are numerous concordant and laterally persistent, 2-20 m wide, intermediate two-pyroxene-hornblende-plagioclasequartz gneisses. These have a low-K quartz-dacite composition and their concordant nature suggests either former lava flows or sills. Garnet-bearing banded iron formations and calc-silicate rocks are absent. There is one concordant unit of garnet-biotitebearing homogeneous augen QFG. There is only one lenticular body of homogeneous garnet-biotite granitic orthogneiss with the penetrative Sg3 fohation developed; th'is body is discordant at the map scale (Fig. 3). Intrusive Rocks

Pegmatites There are at least t h r e e generations of coarsegrained K-feldspar- plagioclase- quartz- muscovite pegmatites. The most common are lenticular to tabular bodies (<1 m wide); largely concordant, though locally they cut across gneissic layering ($1) and occur preferentially in QFG and only within the gneissic terrane. These are folded, boudinaged by stretching along the L2 lineation, develop the $2-L2 fabric in their margins and, less commonly, are grain refined throughout and develop a penetrative $2-L2 fabric of quartz-feldspar aggregate ribbons. D2 recrystallized pegmatites have garnet- muscovite± biotite- bearing assemblages. These concordant pegmatites were emplaced prior to D2 deformation and subsequent to the formation of the gneissic layering (Table 2). These tabular pegmatites are discordant to leucocratic partial melt segregations that developed in association with $1 gneissic layering in the host gneiss. Partial melt sweats are widespread and formed at the peak of M1 metamorphism, thus the tabular pegmatites are a late-stage M1 feature. A group of 0.1-2 m wide discordant (by 10-20 °) pegmatites were emplaced immediately prior to D3 but after D2 deformation. These cross-cut the $2-L2 fabric, were emplaced parallel to the weak $3 fabric, generally have ESE-WNW trends and dip shallowly, though consistently steeper than $1, to the SSW. Some of these pegmatites are asymmetrically folded by F3 folds with a weak axial planar $3 fabric developed in the hinge within the host gneiss. Late-stage ductile displacement and asymmetrical boudinage of this pegmatite indicates west over east ductile overthrusting along a N-S trending zone of shear inclined 30 ° to the west. It is not k n o w n w h a t deformation

Milky Quarfz Veins. 12 10

8 6; /,

g2 ¢.-

~

Pegmafifes.

, me,~ I~te-p~,,~+te.

"6

~ Dykes.- Dolerife.

(a) lb)

|I

2 0

20

L,O

60

80 100 120 140 160 180 Trend (degrees).

Figure 4. Frequency distribution of the strike of late-stage, steep planar elements in the Chewore Inliers. Black is dolerite dykes in the granulite terrane only and pre-D2 pegmati~es. Diagonal shading are post-D2, pre~D3 pegmatites (with shallow 20-30° dips to the SSW) and pre-Ds deformed amphibolite dykes. Crosshatching is late-stage pegmatites in the granulite terrane. (a) Trend of [ate-stage nltramafic pods, (b) overall trend of the Chewore Ultramafic Complex.

is responsible for the asymmetrical boudinage. These pegmatites rarely contain garnet as well as muscovite and their formation is indicative of a metamorphic event ( M 3 ) closely associated with the D3 deformation. Swarms of late-stage, undeformed, very coarsegrained (typically <10 cm) muscovite+garnet~ tourmaline-bearing pegmatites occur in the eastern portion of the gneissic terrane. These trend on average 132°+25 ° (Fig. 4), are 1-5 m wide, vertical to steeply inclined (75-90 °) north and laterally continuous for up to 4 km. They are often zoned with finer-grained (0.5-2 cm) margins comprised of quartz and muscovite only; along strike they uncommonly grade into milky quartz veins. All quartz-veins crosscut $2-L2, are sub-vertical and a large proportion have similar trends to the late pegmatites (Fig. 4), suggesting that they may be coeval (Table 2). Dry garnet- quartz- K-feldspar- plagioclase peglnatite veins of only 10-20 cm width and 80-120 ° trends (Fig. 4) intrude gneisses of the granulite terrane. These are often zoned with either finergrained aplitic cores or margins and garnet porphyroblasts concentrated in the coarse-grained zones. They are u n d e f o r m e d and so post-date Dg4

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

ineation Strong 5 2 foliation.

2 transport

3 strain I[ipsoid

2 strain [lipsoid

NW Figure 5. Cartoon illustrating the relationship between D2 and D3 structural elements in the gneissic terrane, their tectonic transport vectors and the respective strain ellipses. Overall regional D3 transport is of the opposite sense to that presented here (see text).

folding. These pegmatites are not recognized in any other terrane, suggesting that they pre-date the overthrusting of the granulite terrane (D6; Table 2). Dolerite Dykes At least two distinct generations of mafic dyke emplacement are recognised in the gneissic and quartzite terranes. Weakly foliated, folded and boudinaged amphibolite mafic dykes, cross-cut the $2-L2 fabric. They have a weak foliation parallel to the $3 foliation and their boudinage geometry is consistent with that of F3 folding (Fig. 5), thus they were emplaced subsequent to D2 and before D3 (Table 2). Early amphibolitic dykes have an annealed, granoblastic metamorphic texture consisting of hornblende-plagioclase-opaque-quartz. No igneous textures are preserved. Late-stage dykes are fine- to medium-grained, homogeneous undeformed dolerite with ophitic texture. They consist of interlocking elongate plagioclase laths with relict clinopyroxene and hornblende phenocrysts; quartz and some plagioclase are interstitial. Secondary hornblende aggregates replace primary clinopyroxene and hornblende. Garnet occurs as coronal rims around all primary phases and secondary hornblende. These dykes have a wide spread in orientations (Fig. 4) and a variety of appearances in the field; most are homogeneous and featureless, others contain plagioclase or clinopyroxene phenocrysts. They range from 0.5 to 10 m in width and have been traced for up to 8 km. Late-stage undeformed dolerite dykes in the granulite terrane have two trends: N-S and NW-SE.

209

All dykes in this terrane are distinct from those in the gneissic terrane by containing primary olivine and not secondary garnet or hornblende. Some N-S trending dykes intrude into the quartzite terrane across the thrust between these terranes. However, all other dykes in the quartzite terrane are petrologically similar to those in the gneissic terrane. E-W trending dolerite dykes with neither olivine nor secondary garnet intrude the Chewore Ultramafic Complex. There is a large range in ages of post-deformation dykes in central eastern Africa; NE-SW dykes in the Magondi Mobile Belt at 1000-1100 Ma (Hahn et al., 1990a), 804-864 Ma age dykes in the Irumide Belt (N'gambi et al., 1986) and Karoo and younger dykes from Mozambique and north Malawi (Pires and Ferreira 1993; Fitches 1968). The large range in the orientation of dykes in the gneissic terrane (Fig. 4) precludes correlation with any of the known dyke events in the region. Dyke trends in the granulite terrane are not reported from the immediate region.

Chewore Ultramafic Complex The Chewore Ultramafic Complex (Wiles 1956) comprises three discrete bodies, separated by Karoo Supergroup sedimentary cover, aligned on a 045° trend (Fig. 3) and considered to be coeval. The most southerly consists of massive dolerite, gabbro and norite. The other two bodies are predominantly serpentinite with a steep southerly inclined layered sequence at their southern margins consisting of harzburgite, orthopyroxenite, plagioclaseorthopyroxenite, chromitite and norite. A 028°trending train of 50-150 m wide elongate pods of dunite, olivine-norite and harzburgite intrudes both the north-eastern ultramafic body and the granulite gneisses to the north (Fig. 3). These presumably exploited much the same crustal weakness as the three larger bodies and may represent late-stage intrusions of the same magmatic event. The Chewore Ultramafic Complex is undeformed with no tectonic fabrics and no grain refinement or strain features in the mineral grains. Thus the complex was emplaced subsequent to D4 and possibly Ds (Table 2). Master (1990) correlates the Chewore Ultramafic Complex and the Atchiza Complex in Mozambique with the Great Dyke (of 2461+16 Ma age; Hamilton 1977) on the basis of their similar rock types. There is, however, no geochronological evidence to support this interpretation. TECTONO-THERMAL EVOLUTION OF THE GRANULITE TERRANE

Ds2 Isoclinal Folding The first deformation recognized is rare smallscale (5-35 cm wavelength) intrafohal isoclinal folds

210

B. GOSCOMBE,P. FEYand F. BOTH

(Fg2) that are also commonly rootless. These fold gneissic layering (Sgl) and a very weak axial planar alignment of biotite in the hinges of these folds are developed. This foliation is not reflected by quartz and feldspar grains, which have an unaligned polygonal granoblastic texture. This layer parallel alignment of biotite may be widespread in quartzofeldspathic gneisses, although it is so weak it is not generally recognized where Fg2 folds are not developed. Axial planes trend E-W and dip steeply (76-90°) to the north and south. These folds plunge steeply and are near parallel to a mineral lineation (Lg2) that is strongly developed in exposures with these folds. Lg2 typically plunges 65° to the NNE (Fig. 6). The colinearity of Lg2 and Fg2 suggests high shear strain, which is s u p p o r t e d by t h e huge amplitude/wavelength ratios of all these folds. Furthermore, some horizontal sections have fold patterns of two connected isoclines of opposite vergence, suggesting tongue shaped folds in three dimensions. These features all indicate non-coaxial shear of high shear strains with movement along the Lg2 vector. Fg2 folds are envisaged to have formed by locally intense layer parallel shear with no large-scale Structures being developed, or recognized.

assemblages and features are present only in the granulite terrane and indicate metamorphic conditions of formation distinct from the rest of the Chewore Inliers, which are essentially anhydrous conditions at >5 kbar and >725°C (Holdaway and Lee 1977; Bohlen et al., 1983).

D~ Isoclinal Folding The granulite terrane was folded throughout by isoclinal to very tight (interlimb angles <25°) folds on outcrop and map scale. These plunge shallowly (15°) to the west and are mostly inclined steeply to the south (Fig. 6). These folds appear to fold Lg2 lineations (Fig. 6) and their larger scale of development suggests that they post-date the smallscale layer-parallel shear deformation during Dg2. A mineral lineation (Lg3) is developed that parallels Fg3 fold axes (Fig. 6). Asymmetrical pairs of Fg3 folds consistently indicate S over N vergence, suggesting that tectonic transport was orthogonal to the fold axes. N-S shortening during both Dg2 and Dg3 suggest that there may be different phases in the one orogenic period, with the vector of tectonic transport being within the plane perpendicular to Fg3 and containing Lg2. The Dg2-Dg3 tectonic period progressed up in scale from locally intense layer-parallel shear during Dg2 to map-scale isoclinal overfolds in Dg3.

M~ Peak Metamorphism The peak of metamorphism in the granulite terrane was responsible for high-grade and near anhydrous assemblages in a well annealed mediumto coarse-grained polygonal granoblastic texture of untectonized phases. Peak metamorphic annealing of this terrane post-dates Fg2 folds. This is indicated by idioblastic garnet porphyroblast growth across both small-scale Fg2 fold hinges and axial planar biotite. Biotite in felsic gneisses is weakly aligned though a fabric is not reflected by the quartz and feldspar grains. Thus the granulites are interpreted to have been deformed giving rise to the Sg2 biotite foliation and the rock was subsequently annealed at the peak of the Mg2 metamorphism, giving rise to the polygonal granoblastic texture. The aligned biotite is an integral part of the granoblastic matrix and so the peak of metamorphism is thought to immediately post-date Fg2. Fg2 is most likely causative of, and closely associated with, Mg2metamorphism. Peak metamorphic assemblages include garnetbiotite- quartz- plagioclase- K-feldspar+ sillimanite+ rutile in felsic gneisses and metapelites with orthopyroxene in some gneisses and orthopyroxeneclinopyroxene- hornblende- plagioclase- quartzruffle in mafic gneisses. Muscovite is absent from all rocks, many felsic gneisses are anhydrous and spinel inclusions in garnet are common. Sillimanite is the dominant aluminosilicate, though rare metapelites contain coexisting kyanite and siUimanite. These

Dg4 Open Folds Only one phase of open folding (Fg4) is recognized in the granulite terrane and it is rare and only recognized at map scale (Fig. 3). They are upright, trend N-S and are sub-vertical owing to the steepness of the terrane. A N-S trending, steeply (80-90°) east dipping planar fracture cleavage (Sg4) is developed solely in quartzites (Fig. 6). No displacements are recognized across these cleavage planes, which are thought to have formed by shortening orthogonal to the cleavage plane. Such E-W directed shortening is attributable to Fg4 open folding in the granulite terrane. The Sg4 cleavage is only recognized in the granulite terrane, but upright, open, N-S trending folds are recognized in all terranes. TECTONO-THERMAL EVOLUTION OF THE GNEISSIC AND QUARTZITE TERRANES

M1 Peak Metamorphism Most lithologies develop a mm-cm scale gneissic layering (S0 based on different mineral proportions, or sometimes on distinct layers of different assemblages. Calc-silicate rocks have alternating diopside layers and garnet layers, while metapelites have sillimanite/kyanite-biotite-garnet selvages and quar tz-plagioclase-K-feldspar leucosomes. More commonly, gneissic layering is due to differences in the proportions of minerals in a rock consisting of

211

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

Granul,ite Terrane (I)

Ouarfzite

5neissic

Terrane (II)

Terrane

13neissic Layering. pol,e to S l 7

parallel $2]

Structural Elements. pole to $2] 2 02

D3 Tight Fol,ding

JFg

l~~--~

i-o= fold axis

L",,=axial plane]I ~ DL, Tight Folding.

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Crenul.ation Cl,eavages °=kcren~ '~,,=Scren]

[

Figure 6. Lower hemisphere, equal area stereoplots of structural data from each of the three terranes comprising the Chewore Inliers. Major and minor folds are included. Contour intervals are at 1, 2, 4, 6, 8, 10, 12%/1% area. The dashed [me is the best fit profile plane and the black square is the pole to that plane. The n u m b e r of readings is in the bottom right hand comer.

only the one assemblage. For example, in QFG gneissic layering is defined by variations in the proportion of biotite and/or muscovite. The correlation of gneissic layering with prograde metamorphism is best illustrated by the metapelitic gneisses. These are stromatic migmatites with distinct quartz-feldspar leucosomes alternating with darker sillimanite/kyanite-garnet-biotite-rich layers. The layered nature of the rock is due to partial melting during prograde metamorphism. Thus, gneissic layering in adjacent gneisses is also thought to have formed during M1 metamorphism, though may have been influenced by pre-existing (sedimentary/volcanic) compositional inhomogeneities in the rock. Also formed at the peak of metamorphism are polygonal granoblastic textures

in most rock types and porphyroblasts, such as garnets in QFG, metapelites and some amphibolites, and K-feldspar in QFG. The peak of metamorphism immediately predates the first ductile deformation (D2) as indicated by the presence of $2-L2 fabrics in partial melt segregations in migmatites, early pegmatites and orthogneisses, and by porphyroblasts enclosed by this fabric. The M1 metamorphic cycle persisted to D2 deformation, as evidenced by coexisting sillimanite and kyanite laths aligned parallel to L2 within $2 and syntectonic, poikiloblastic staurolite. In most samples, sillimanite is a late-stage phase that either grows intimately, and in textural equilibrium, with peak kyanite or replaces kyanite. This suggests that sillimanite growth occurred at the peak of metamorphism and continued

212

B. GOSCOMBE,P. FEY and F. BOTH

to grow during decompression after peak-pressure conditions in the kyanite stability field were experienced. Sillimanite growth continued after D2 deformation as indicated by random sillimanite grains overgrowing the $2 fabric and fine sillimanite also growing within the axial planar $3 fabric during D3. M1 metamorphic assemblages, such as coexisting kyanite and sillimanite- garnet- biotite- melt, kyanitecordierite-biotite-melt and kyanite-staurolite-garnetbiotite in the gneissic and quartzite terranes, indicate pressure conditions of at least 6.0 kbar at temperatures of 650-700°C (Schreyer and Seiferts 1969; Holdaway 1971; Winkler 1979; Bickle and Archibald 1984; Powell and Holland 1990). Thus, the predominantly supracrustal (sedimentary/volcanics) protoliths of the Chewore Inliers were buried to at least 21 km depth prior to attaining peak temperature conditions. This burial may have occurred during the same tectono-thermal cycle as the M1 metamorphism.

D2 Deformation ($2-L2 Fabric) Gneisses in both terranes often carry an intense fabric that is strongly linear as well as planar. The foliation ($2) is typically defined by biotite and by flattened ribbon-like quartz-feldspar aggregates formed by grain-refinement in the more deformed rocks. L2 is a mineral elongation lineation defined by biotite trains, quartz±feldspar aggregate, ribbons, boudinaged feldspar augen, fine-grained recrystallization tails on feldspar augen and hornblende laths. In some samples in the gneissic terrane this fabric is almost mylonitic and S-C fabrics are common. The $2-L2 fabric is the pervasive tectonic fabric throughout both these terranes. The $2 foliation is always parallel to the gneissic layering, except in F2 fold hinges. However, $1 and $2 are temporally distinct as indicated by sub-concordant late-M1 pegmatites that crosscut $1 but themselves develop the $2-L2 fabric. Assemblages developed within the $2 foliation are indicative of the M2 metamorphic event that accompanied the D2 deformation. Small-scale (cm-scale wavelength), intrafolial isoclinal folds (F2) have a strong axial planar foliation parallel to the regional $2 foliation, and thus F2 and $2 are considered coeval. F2 folds are rare but distinct from F3 folds, which do not develop a strong axial planar foliation. F2 fold axes are sub-parallel to the L2 lineation. The scarcity and small scale of F2 folds suggest that D2 deformation involved layer parallel shear with only small-scale folding and many of the folds developed may have been obliterated by continued shearing along L2 within the $2 plane. The near co-lmearity of F2 axes with the L2 transport vector is suggestive of non-coaxial shear at high shear

strains, as is the presence of asymmetrically enclosed s-type porphyroblasts and S-C fabrics (Goscombe 1991). The orientation of L2 throughout the gneissic terrane is remarkably consistent, plunging 5-40 ° towards 200-240 ° (typically 220°). Some lineations plunge shallowly (<10 °) towards 040 ° (Fig. 6), possibly indicating the presence of late-stage, upright, NW-SE trending folds (D4b). Both $2 and gneissic layering have a large range in strike but most commonly dip shallowly (5-40 °) towards the SW to SE (Fig. 6). In the quartzite terrane L2 plunges moderately to the NE and $2 dips steeply (>70 °) to both the NW or SE (Fig. 6). M1 peglnatites are commonly boudinaged by stretch along an axis exactly parallel to the L2 lineation (i.e. the long axis of boudins is perpendicular to L2). This indicates ductile stretching during D2 and that the maximum principal strain axis (X) during D2 was parallel to L2 (Nicolas and Pokier 1976). The sense of tectonic transport during D2 is almost always from the SW over NE along the L2 vector in the gneissic and quartzite terranes (Fig. 7). In one small region in the north of Domain III (Fig. 8) D2 shear sense is from the NE over SW. Shear sense has been derived from S-C t y p e fabrics, the asymmetry of s-type feldspar augen, asymmetrical boudinage of pegmatite layers and K-feldspar

Tectonic transport vectors.

D2 transport paraUe[ to L2 in Quartzite g Gneissic lerrane. D3 transport orthogona[ to asymmetrical F3 folds, same terranes. []g3transport in 13ranu[ite Terrane, Figure 7. Lower hemisphere,equal area stereoplot of D2 tectonic transport along the L2 lineation and D3 transport sense as indicated by the asymmetricalF3 folds. All arrows point towards the directionof tectonictransport of the "upperplate".

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

porphyroblasts (Simpson and Schmid 1983) and deflection of the $2 foliation through tabular early pegmatites. The intensity of the $2-L2 fabric is heterogeneous throughout these terranes and between different lithologies. Within the gneissic terrane the layered QFG develop $2-L2 best with this fabric being virtually absent in calc-silicate rocks. Domains IV, V and VI develop strong, often sub- mylonitic $2-L2 fabrics in contrast to weak development in Domain III. In both terranes, quartzite develops a more linear, rodded fabric in preference to a planar foliation. Pre-D2 tectonic fabrics have been recognised in small (10-30 cm diameter) feldspar-rich xenoliths in some granitic orthogneisses. These xenoliths contain a well aligned biotite foliation at a high angle to the $2 foliation that encloses the xenolith. Pre-D2 foliatlons are also preserved in some garnet porphyroblasts as sygmoidal quartz inclusion trails orientated at high angles and unrelated to the $2 foliation that encloses the porphyroblasts. These early foliations are possibly expressions of deformation associated with the burial and prograde portion of the M1 metamorphism and are only rarely preserved because of the pervasive and penetrative nature of $2-L2 fabric development.

D3Tight Folding F3 folds are the most dominant in the area and occur on all scales from cm to km wavelengths (Fig. 8) and are tight to isoclinal folds that most often occur as pairs of asymmetrical hinges. The penetrative $2-L2 fabric is folded by F3 folds and at best only a weak axial planar foliation ($3) is developed. In outcrop, $3 is defined by aligned biotite without the grain refinement of pre-existing mineral phases (i.e. quartz and feldspars). Rare examples of a low-angle crenulation cleavage with $3 orientations are developed in biotite-rich QFG, biotite schist and chlorite schists, are closely associated with mesoscopic F3 folds and so are correlated with $3. In thin section, very thin (~1 mm), widely spaced, obliquely cross-cutting slip planes of fine-grained biotite and sillimanite are rarely developed in the hinges of F3 folds in metapelites. F3 axes in the gneissic terrane plunge shallowly to the SW with the same distribution as L2 (Fig. 6). F3 axes in the quartzite terrane are essentially vertical and thus distinct from the L2 orientation. Because of the tight to isoclinal nature of the F3 folding, poles to their axial planes fall on a similar great circle girdle as poles to S1 and poles to $2 in all domains (Figs. 6 and 9). In the gneissic terrane, F3 folds are reclined with axial planes dipping shallowly and mostly to the SW with a few inclined to either the W or SE (Fig. 6). In contrast, axial planes are mostly vertical in the quartzite terrane. The fold axial traces of major folds are interpreted in Fig. 8 with most of the tight to

213

isoclinal folds being considered as F3 folds. In Fig. 8 the large variation in the orientations of fold axial traces in the gneissic terrane is a function of the shallow inclination of the axial plane and a deeply dissected topography. Stereoplots of $2 and $1 data display reasonably tightly constrained singular distributions in most domains, but also define a girdle consistent with the tight F3 folding in any one domain (Fig. 9) indicating that F3 was the dominant folding event in the gneissic and quartzite terranes. General F3 fold orientation, as indicated by the pole to the $1 and $2 great circle, is remarkably consistent (SW plunging) between all but three (VII, IX and X) of the ten domains constituting the gneissic terrane. Domain VII has shallowly, west plunging folding that is consistent with folding in the granulite terrane. Granulites have not been recognized in this domain. Domain IX is steeply inclined with shallowly, east plunging folding. Domain X is a very shallowly inclined domain with shallowly, ENE plunging folds. The orientation of layering and folding in this domain is interpreted as resulting from refolding by NW-SE trending folds, possibly F4 folds (Fig. 8). An intersection lineation of $3 biotite intersecting the folded gneissic layering and $2 fabric, is rarely developed in the hinges of F3 folds. In the southern part of the gneissic terrane (Domain VIII) there are some SE plunging mineral lineations (Fig. 9) which may represent the D3 transport vector. However, no mineral lineation has been conclusively established to be genetically associated with F3 folding. In the absence of D3 mineral elongation lineations, the D3 tectonic transport vector cannot be confidently established. D3 boudinage, with long axes parallel to F3 axes, suggest that there was no extension parallel to these fold axes (Fig. 5). Thus, the maximum extension axis (X-principal strain axis) during D3 is orthogonal to the F3 a x e s and contained within the $3 plane, which is orthogonal to the axis of maximum shortening (Z, Fig. 5) (Nicolas and Poirier 1976). The tectonic transport vector is contained within the plane perpendicular to $3 and containing the X-axis, but will be moderately inclined to $3 and so possibly near horizontal. Since F3 folds invariably occur as asymmetrical pairs, they imply formation by noncoaxial shear and so are employed as sense of movement indicators. The sense of shear along the D3 transport vector defined above is given by the sense of movement that gave rise to asymmetrical folds in outcrop (Fig. 5). Asymmetrical folds in Domains IV, V and VI in the gneissic terrane consistently display NW vergence. Domain III in the gneissic terrane displays NW over SE transport, and Domain X displays both of these senses (Fig. 7). The reversal in transport sense is interpreted as being due to small-scale F3 folds

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Figure 10. Schematic cross-sectionillustrating the spatial relationship between the terranes in the north of the area. The section line is marked on Fig. 3, vertical:horizontal = 10:1. The contact between the quartzite and gneissic terrane has not been observed, but is inferred to also be h thrust (see text). The sense of vergenceof small-scaleparasitic folds across a diagrammatic large-scale F3 antiform (axial plane is a dashed line) is indicated (see text). being parasitic on a very large-scale, NE-SW trending anticlinal F3 fold, the axial trace of which separates these two regions with Domain X constituting the hinge zone of this fold (Fig. 8). The domains with NW over SE transport sense are also the steepest zones and so constitute the short limb of this major reclined structure (Fig. 10). Consequently, the regional sense of tectonic transport for the gneissic terrane is equivalent to that displayed by F3 folds on the long (shallow) limbs, i.e. SE over NW. F3 folding is co-linear to L2 and poles to F3 axial planes fall on the same great circle girdle as poles to the $2 foliation (Fig. 6). Despite such coincident orientation, the two are distinct structural events formed at different times. This is evidenced by the following: i) L2 lineation is folded around F3 fold closures in outcrop, but at such a low angle that the L2 orientation is only slightly changed (Fig. 5). ii) At every outcrop of F3 folds, the strong $2 foliation is folded around the F3 hinge. Furthermore, the axial planar fabric developed in F3 folds ($3) cuts the $2-L2 fabric and is not as intense as the folded $2L2 fabric iS3 typically involves the alignment of biotite grains but not the grain refinement of feldspars and quartz). iii) If F3 fold axes and the D2 tectonic transport vector (L2) were coeval, their co-linearity would imply high shear strains and extreme rotation of fold axes into sheath folds (Goscombe 1991). No such folds are developed and no curvature of F3 fold hinges is evident, neither are extremely high shear

strain features such as mylonites and highly attenuated and boudinaged porphyroblasts. 'iv) The long axes of boudins closely associated with F3 folds, are approximately parallel to the axes of these folds and both are assumed to have formed at the same time. These D3 boudins indicate that the maximum principal strain axis iX) during D3 was orthogonal to the F3 axes. The D3 X-axis is at a high angle (90 °) to the X-axis during D2 (Fig. 5), where X was parallel to L2, thus the two events cannot be related. Open folds with shallow SW plunges and upright to moderate SE inclined axial planes are recognized in the gneissic terrane (Fig. 6). They are of intermediate scale with wavelengths typically >10 m. No fabric is associated with these folds and they refold the L2 lineation. The co-linearity and coplanarity of these folds with tight to isoclinal F3 structures suggest that they may be related and so have been labelled as D3b structures (Table 2). These may have formed in the waning stages of progressive SE over NW transport during D3. Dl T i g h t to O p e n Folds In the gneissic terrane, at least two generations of major tight folding have been preserved. The second generation (D4) is not tightly constrained in orientation and may have involved more than one episode of folding. F4 folds are tight but not isoclinal and plunge shallowly to the E to SE. Fold axial planes are inclined moderately to steeply to the SE to S (Fig. 6). A weak axial planar foliation ($4) is defined by poorly aligned biotite and muscovite. F4 folds are

Structuralevolutionof the Cheworelnliers, ZambeziMobileBelt, Zimbabwe recognized in outcrop only in the central and southern domains of the gneissic terrane. It is inferred that some of the large-scale NW-SE and E-W trending folds recognized in both the quartzite and gneissic terranes (Fig. 8) may be F4 folds, the orientation of which is broadly consistent with NE-SW to N-S shortening. The sense of tectonic transport during D4 is not confidently known. Open folds of moderate to shallow easterly plunge and SE dipping axial planes are recognized in the gneissic terrane (Fig. 6) and are sufficiently co-planar and co-linear to be considered genetically related to F4 folds. These open F4 folds are labelled F4b (Table 2) and were possibly formed in the waning stages of D4. N-S trending open folds are recognized in all terranes and also labelled F4b (Fg4 in the granulite terrane) folds because their timing with respect to other open folds is unknown. Open folds in the granulite terrane differ only slightly in orientation by being sub-vertical, whereas those in the quartzite and gneissic terranes are inclined moderately to the west. The fold axes are sub-horizontal in the gneissic terrane and sub-vertical in the quartzite and granulite terranes (Fig. 6). These folds formed by E-W shortening and so are distinct from the E to SE plunging F4b open folds in the gneissic terrane. Open upright warps with NW-SE trends are recognized in the gneissic terrane by the reversals in the plunge of L2 lineations, the refolding of F4 axes and the refolding of the $3 fabric (Figs. 6 and 9).

Ds Crenulations In both the gneissic and quartzite terranes two orientations of steep (80-90 ° dips) crenulation cleavages are rarely developed in metapelitic gneisses (Fig. 6). Crenulations are asymmetrical kinks, the axial plane of which defines a sharp plane. Hinge zones are <1-2 m m wide and spaced 0.5-10 cm apart. These crenulate muscovite and biotite grains with no grain refinement or deformation of quartz and feldspars. No new or retrogressive minerals were developed. NNE-SSW trending crenulations have dextral movements with a reverse component, ESEWNW crenulations had sinistral movements. These crenulations are interpreted as a conjugate set that formed by NE-SW trending shortening. Rare NNESSW trending shear zones with dextral shear sense have been recognized and maybe associated with the crenulations of the same orientation and shear sense. These crenulations are consistent with shortening during D4 and D4b in the gneissic terrane. Thus D4 and Ds structures in the gneissic terrane are interpreted as having formed in one orogenic period (D4-D5).

217

LATE-STAGE EVENTS Juxtaposition of Terranes (D6) The granulite terrane is separated from the quartzite terrane (Fig. 3) by two parallel 045 °trending mylonite zones of 10-15 m width, one of which is 100 m north of the mylonite at the contact. The mylonite zones dip 80* to the NW and mineral elongation lineations, defined by biotite trains and quartz-feldspar aggregates, plunge 30 ° to the NE. Quartzites immediately south of the contact have been reoriented by drag along this zone. The sense of shear, indicated by s-type asymmetrically enclosed garnet porphyroblasts and asymmetrical hollows on the foliation surface, indicate oblique south-westward over-thrusting of the granulite terrane onto the quartzite terrane. This over-thrusting pre-dates emplacement of the Chewore Ultramafic Complex, which is intruded into both terranes across this boundary (Fig. 3). Similarly, some N-S trending olivine-bearing dolerite dykes cross-cut this boundary. Over-thrusting pre-dates both late-stage pegrnatites and brittle faulting (Table 2). The contact between the quartzite and gneissic terrane is not exposed, therefore its true nature is unknown. The south margin of the major area of quartzite terrane is a straight, sharp linear feature (Fig. 3) and is presumed to be a steeply NW inclined surface as indicated by its mapped outline with respect to topography. This contact is sub-parallel to the mylonite zones at the south boundary of the granulite terrane. Consequently, the contact between the quartzite and gneissic terranes is inferred to also be a steeply NW dipping thrust zone (Fig. 10). There are oufliers of quartzite terrane rocks on the hill tops in the far east of the Chewore Inliers. Layering within these outliers is discordant to layering in the gneissic terrane below and the contact between the two dips shallowly to the west. These outliers are inferred to be klippen of quartzite terrane thrust onto the gneissic terrane, the root of this thrust being the steep boundary between these terranes in the west (Figs. 3 and 10). This inferred thrust has been displaced downwards to the east by post-Karoo normal faulting. $5 crenulation cleavages have similar orientations in both the gneissic and quartzite terranes (Fig. 6) and so these terranes may have been juxtaposed prior to Ds. This is, however, inconclusive and no other timing relations are available other than juxtaposition post-dated D3 and possibly D4 deformations, which are of distinctly different orientations in the two terranes (Fig. 6). The difference in orientation of D1 to D4 structural elements in these terranes may be the result of rotation of the quartzite terrane during its inferred over-thrusting onto the gneissic terrane. The orientation of L2, Fa and F4b axes is consistent with a clockwise rotation of nearly 90 ° around a NW-SE axis of the quartzite terrane from orientations originally similar to the gneissic terrane to their present orientations. This supports the hypothesis that the

218

B. GOSCOMBE,P. FEY and F. BOTH

quartzite terrane was rotated by upramping onto the gneissic terrane during D6. Identical rotation of the, presumably, originally horizontally layered ultramafic bodies, would have resulted in the vertical NW-SE trending layering present in these bodies. If SW directed oblique over-thrusting of the quartzite terrane over the gneissic terrane was responsible for reorientation of the ultramafic complex, thrusting at this boundary must post-date thrusting at the granulite/quartzite terrane boundary, across which the ultramafic complex intrudes. Pre-Karoo Uplift of the Chewore Inliers (Ups) The Karoo Supergroup sediments were unconformably deposited on the gneisses of the Chewore Inliers (Fig. 3). Metamorphic assemblages indicate peak M1 conditions of the quartzite and gneissic terranes involved >_21 km depth of burial. Consequently, since M1 and prior to Karoo times, at least 21 km of crust was denuded, with accompanying uplift, exposing the gneisses at the unconformity surface. It is unlikely that this uplift (Up1) was as one continuous event because the period between M~ and Karoo times saw the occurrence of multiple tectono-thermal cycles, including DrD2, D3 and D4-D5 and D6 over-thrusting. The amount of burial, crustal thickening and consequent isostatic uplift, that can be attributed to each of these deformational periods is unknown. Despite this potentially complex evolution of the vertical crustal profile during and subsequent to M1, it is known that the sum result is the exposure of rocks previously buried at >_21 km depth by Karoo times. This sum effect is the most significant crustal event a n d / o r events to affect the Chewore Inliers. However, no structural feature can be correlated with pre-Karoo uplift and denudation of the Chewore Inliers. All brittle faults and breccia zone trends in the Chewore Inliers are also recognised in the Karoo sediments. Furthermore, palaeostress analysis of faults in the Chewore Inliers is consistent with results from fault sets in the Karoo sequences. Thus all brittle faulting recognized is thought to have occurred during and subsequent to Karoo times (Up2and Up3) in response to development of the lower Zambezi valley rift (see later discussion). Ductile shear zones that could potentially be associated with pre-Karoo uplift are essentially non-existent except for the mylonites at, and hypothesized to be at, the contact between the terranes. D6 over-thrusting may well have contributed a significant amount of both direct up-thrusting of the Chewore Inlier rocks as well as crustal over-thickening and consequent isostatic uplift. Secondary sillimanite growth after peak M1 metamorphic kyanite suggests decompression

through the peak of M1 metamorphism. Such decompression is typical of late-stage isostatic rebound experienced in collisional belts (England and Thompson 1984). Nothing is known of vertical crustal movements during D2 to Ds deformations. However, it is envisaged that pre-Karoo uplift was by multiple passive isostatic uplift events in response to crustal over-thickening subsequent to each major compressional and crustal shortening phase experienced. The sum total result of these thickenings (with potential burial) and isostatic rebound events is >_21 km of crustal denudation since M1. The cooling history of the Chewore Inliers is presently being investigated by generating T-t curves. This may in part constrain Up1 events. Syn- to Post-Karoo Faulting and Deformation

Faulting and Uplift (Up2 and Up3) Millimetre to metre wide, laterally persistent (up to 10 km long), planar to anastomosing breccia zones are very common in some portions of the Chewore Inliers. A minimum of three episodes of deformation in these breccia zones are recognized by: i) pseudotachylites ii) multiple events of dry gouge and brecciation or 4 less commonly, sharp fault planes iii) further brecciation with fluid influx and silica (+fluorite+calcite) veining. The pseudotachylites occur as dark, planar zones and veinlets of only 1-10 mm width, containing glass and very fine-grained mineral fragments, typically n e a r t h e m a r g i n s of the b r e c c i a zone. Pseudotachylites have themselves been brecciated and occur as isolated breccia blocks. They are preferentially intruded along their length by the latestage silica veinlets. Zones of brecciation may attain widths of 1-2 m and consist of a network of planar to curviplanar breccia zones of cm-scale widths. These produce isolated angular blocks (0.1-10 cm wide) of host gneisses enclosed in a gouge matrix of grain-size up to 1-2 mm. There are some 1-2 cm wide zones consisting wholly of fine gouge without larger blocks. The brittle nature of these zones suggests that the breccias formed only at the last stage of uplift within the upper crust at <10 km depth (Tullis and Yund 1985). However, the presence of pseudotachylites is typically indicative of high strain rates in either a ductile or brittle regime (Sibson 1977). The sequence in structures may be due to rapid uplift of the terrane from ductile to brittle crustal levels (Chen and Molnar 1983). The last stage of reactivation of these zones involved the influx of silica-bearing fluids. Most zones contain anastomosing mm- to cm-scale, white, micro-crystalline and often laminated silica veinlets

Structural evolutionof the CheworeInliers,ZambeziMobileBelt,Zimbabwe Breccia & Fau(t in (3neisses. Breccia & Fault in Karoo.

FauJt in North In(ier.

Breccia in North

In[ier.

Percent of fautts satisfying a solution for 0-3. O~axis 0

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Figure 11. Lower hemisphere projection of the results of stress analysis from breccia zones and fault sets from the Chewore Inliers. Contoursrepresent the percentage of faults in the data set that satisfya solution for cr3.Thus 100% represents a solutionfor the (~3axis and 0% represents the al axis. Stress analysisis by the Angelier and Mechler (1977) method and solutions were generatedby the program EQUALAREAJAN89(Sorlein1989). with rare carbonate, fluorite and kaolin. Silica and carbonate filled breccia zones occur in the Karoo Supergroup sediments and the major zones (>5 m wide) that bound the Chewore Inliers. All breccia and fault zones in the Inliers are steep (mostly >70 ° dips) with a multitude of trends, the most pronounced being; NW-SE, NNW-SSE, WNWESE and NE-SW. Slickenlines on fault planes and in breccia zones are rare, those measured all pitching steeply (70-90 °) in the fault plane. Where the shear sense has been recognized by displaced gneissic layering, it is invariably normal. Most margins of the inliers are fault-bounded or closely associated with faults (Figs. 3 and 8) and as such the Chewore Inliers represent horst blocks. The Karoo Supergroup sediments are extensively displaced by major breccia zones; the main trends of which are 0 °, 25°, 60°, 70 °, 110 °, 130 ° and 150 ° (Oesterlen and Blenkinsop 1994), all of which are recognized in the Chewore Inliers. These faults were active during Karoo times and accompanied the rifl~g of the Zambezi valley (Up2). Syn-Karoo faulting and uplift is evidenced by unconformities

219

within the Karoo sedimentary sequence, particularly between the Upper and Lower Karoo Groups (Oesterlen in prep.). The Chewore Inliers remained a highland and sediment source through much of Karoo times (Oesterlen in prep.) and so were presumably uplifted with coeval subsidence of the mid-Zambezi (to the west) and the lower Zambezi (to the east) basins. All fault trends in the Lower Karoo are present up into the youngest Karoo rocks, indicating that these faults were also active in post-Karoo times (Up3, Table 2). Post-Karoo faulting was responsible for the elevation of the previously peneplained Chewore Inliers to their present position at altitudes of approximately 600-800 m above the Karoo sediments forming the Zambezi Valley floor. Palaeo-stress analysis of faults and breccia zones in both the Chewore Inliers and Karoo rocks, by the method of Angelier and Mechler (1977), was undertaken using the computer program EQUAL AREAJAN89 (Sorlein 1989). Because of the paucity of movement vector and sense data, all planes lacking this data were assumed to be normal faults with d o w n dip movements. These assumptions are validated by the data presently available and result in the principal c o m p r e s s i v e stress (oh) being constrained to near vertical and ~3 to be horizontal (Fig. 11). Solutions for c~3 that satisfy 100% of the faults in the Chewore Inlier datasets have trends of 20° and 160 ° (Fig. 11). All data sets have good solutions for a3 trending 20°. Thus these faults and breccias are tentatively thought to have formed by mainly NNESSW trending extension. The solutions for ~3 from faults in Karoo rocks are very similar to those from the Inliers (Fig. 11). Thus it is thought that the fault and breccia zones in the Inliers were initiated in Karoo times (Up2 and Up3) and not earlier. These solutions are compatible with the NNE-SSW trending pure extension of the lower Zambezi rift with no translational movements as reported by Oesterlen and Blenkinsop (1994).

Post-Lower KarooFlexure (DT) Intermediate-scale (wave length >_20 m), NW-SE trending open, upright warps are recognized in the Lower Karoo Group sediments. These are rare and have inter-limb angles of 140-160 °, thus D7 shortening was extremely small. These folds plunge _<5° to the NW and develop vertical NW-SE trending fracture cleavages. Such warps must have occurred subsequent to Lower Karoo times and so are of <245 Ma age (Table 2). Drape folds of 120 ° trend are developed through the entire Karoo succession. These are the latest formed structures recognized south of the Chewore

220

B. GOSCOMBE,P. FEY and F. BOTH

Inliers by Oesterlen (in prep.). These drape folds are possibly due to sub-surface faulting and differential movement of basement blocks and may thus have occurred in the final uplift events (Up3, Table 2) in which the Chewore Inliers were uplifted and further exposed as a highland area after deposition of the Karoo Supergroup sediments. DISCUSSION AND POSSIBLE REGIONAL CORRELATIONS This study has identified three lithologically and structurally distinct terranes within the Chewore Inliers. At least three tectonic periods are recognized in the gneissic and quartzite terranes; D1-D2, D3 and D4-D5 and two in the granulite terrane; Dgl-Dg3 and Dg4. However, in the absence of any geochronological data from this region, the relative ages of the three terranes and their tectono-thermal histories are unconstrained and the correlation of structural events with the known major orogenic periods in central eastern Africa can only be tentative. Such correlations are, however, attempted on the basis of the tectonic transport vectors operating at different times in the evolution of the Chewore Inliers. The presence of three distinct terranes in such a small area as the Chewore Inliers suggests that the Proterozoic mobile belts comprising central eastern Africa may be heterogeneous and consist of a mosaic of small terranes. This is in contrast to the concept of mobile belts in central eastern Africa being broad belts of essentially homogeneous age and tectonic histories throughout as superficially implied in Fig. 1 and by Kr6ner (1977). The Ubendian Belt for example has been shown to be composite in nature and a mosaic of many distinct terranes (Daly 1988; Lenoir 1993). The granulite terrane, preserving distinctly higher peak metamorphic temperatures and being more anhydrous, is assumed to be older than the other terranes, though there is no direct evidence of this. The apparently simple structural history of this terrane relative to the others (Table 2) may not be a true reflection of its structural history. The anhydrous nature of these rocks leaves them less conducive to ductile reworking, mineral growth and fabric development in later orogenic periods and so tectonothermal events subsequent to high-grade Dgl-Dg3 events may not be readily expressed. The transport sense and orientation of Dg2 and Dg3 structures cannot be correlated with any structural events in the other two terranes. Accordingly, the granulite terrane is interpreted to have developed in isolation to the other terranes and juxtaposed by over-thrusting at a latestage. The granulite terrane cannot be correlated with any province and the deformations within it to any major tectonic event in central eastern Africa.

The quartzite and gneissic terranes, though lithologically distinct, have very similar structural histories (Table 2). Consequently, these two terranes are thought to have evolved together, even if they were not juxtaposed until a later stage. The relatively steep orientation of the quartzite terrane may be a result of tilting during D6 over-thrusting and juxtapositioning of the terranes. Three structurally distinct periods are recognized in the quartzite and gneissic terranes (Eh-D2, D3 and D4-Ds) with tectonic transport vectors essentially orthogonal to each other during each of these periods (Fig. 2). The D3 sense oftransport (SE over NW) is similar to that during pervasive deformation in the Irumide (Fig. 12) and associated belts (Daly 1986a; Daly et al., 1984; Ring et al., 1993) at a p p r o x i m a t e l y 1100-1300 Ma t i m e s (Table 1). Furthermore, the D4-D5 sense of transport (NE over SW) is similar to that experienced in the Zambezi Mobile Belt (Fig. 12) at 830+30 Ma times (Barton et al., 1991; Wilson et al., 1993; N'gambi et al., 1986). Thus D3 and D4-D5 are correlated with these major orogenic cycles. Furthermore, NE over SW transport of the granulite terrane during D6 over-thrusting onto the quartzite terrane is also consistent with Zambezi Mobile Belt transport in the region (Fig. 2). Thus the juxtaposition of terranes (D6) may have occurred at a late-stage of D4-D5 tectonism. The D1-D2 tectono-thermal cycle with SW over NE transport cannot be correlated with any orogenic event in central eastern Africa (Fig. 12). However, based on the above assertions, D2 deformation, as well as the protoliths of the quartzite and gneissic terranes, are possibly older than the Irumide deformation age (i.e. >1100-1300 Ma). The NW over SE transport of the Magondi Belt is not recognized in the Chewore Inliers and the two regions are lithologically different. However, there are structural similarities viz.: early folds in the Magondi Mobile Belt trend NE-SW, broadly similar to F3 in the Chewore Inliers but with different vergence and an associated elongation lineation (Fig. 12), although a late-stage NW-SE trending folding (equivalent to F4s) of Pan-African age is recognized in both regions (Fig. 12; Munyanyiwa and Blenkinsop 1993). The Archaean escarpment gneisses on the northern margin of the Zimbabwe Craton are lithologically distinct from the Chewore Inliers (Fig. 12), being comprised of calcareous (hornblende, epidote and clinopyroxene-bearing) QFG (Fey and Broderick 1990) and are considered to represent the metamorphosed equivalents of the granitegreenstone lithologies of the Zimbabwe Craton (Hahn et al., 1990a). Similarly, the possibly Archaean (Loney 1969) western paragneisses on the north-west margin the Magondi Mobile Belt contain hornblende- and epidote-bearing QFG (Fey and Broderick 1990;

Structural evolution of the Chewore Inliers, Zambezi Mobile Belt, Zimbabwe

Broderick 1976). The Chewore Inliers do not have metamorphosed calc- alkaline granite-greenstone type protoliths and hornblende-, epidote-, clinopyroxene-bearing QFG are absent. Thus, the Chewore Inliers cannot beconfidently correlated with any rock units to the south, suggesting a major geological break along the length of the Zambezi valley, below the Karoo Supergroup sediments. From our observations, the gneissic terrane of the Chewore Inliers is both structurally and lithologically very similar to the basement gneisses directly to the north and north-east in Zambia and Mozambique (the Malawi Province, Cahen et al., 1984). These have similar elongation lineation directions, D3-D4 fold style and gneiss compositions and assemblages (Fig. 12; Johns et al., 1989; Goscombe unpubl, data). Thus, the Chewore Inliers are correlated with this region of basement gneisses in Zambia, which is also poorly constrained geochronologically, but considered to be of Irumide age and possibly older (N'gambi et al., 1986; Cahen et al., 1984; Johns et al., 1989; Haslam et al., 1983). The younger Pan-African (450-650 Ma) ages in central eastern Africa represent, for the most part, cooling ages after a thermal perturbation and posttectonic granites and pegmatites. The widespread and common late-stage pegmatites and quartz veins in the Chewore Inliers are tentatively correlated with this Pan-African thermal perturbation (Table 2) and with 550 Ma pegmatites in the Magondi Mobile Belt (Master 1991). Pinna et al. (1993) and Ring et al. (1993) recognized E-W shortening at this time in the Mozambique Belt and this may be represented in the Chewore Inliers by the N-S trending open folds and cleavage in the granulite terrane (Dg4) and F4b open folds in the quartzite terrane, but not recognized in the gneissic terrane.

5neissic Layering~ Mineral L~eation Tight Folding.

® Escarpment Gneisses.

® Magondi Belt

'1

Irumide Be~t. (Masofu).

Zambian Basement (Chipata).

"~ Zambian Basement ~R (Wyimba) oa

Irumide ?

--~ Basement.

Zambezi A[[ochthonous Terrain,

Marginal (~neissS Terrain (Rushinga)

Acknowledgements Edward and Fanta Gwera, Numeri Kaundura and Edimore Mupapudzi are sincerely thanked for their great company, keen interest and eye-sight and professional help with the mapping. Peter Zizhou and Mr Lunga are acknowledged for their mapping in the extreme south. The staff of the Geological Survey of Zimbabwe are thanked for their generous help and c-ooperation and the Directors, Dr. Orpen and Mr Ncube for making this project possible. The D e p a r t m e n t of N a t i o n a l Parks and Wildlife of Figure 12. Lower hemisphere, equal area stereoplots of structural data in the literature from the region around the Chewore Inliers. Only the fold axes of tight folds are presented, except the poles to the $4 cleavage in domain Ill. The dashed line is the best fit profile plane and the black square is the pole to that plane. Arrows indicate the sense of movement of the upper plate. See Fig. 6 for the domain localities. Data are sourced from: Goscombe (unpubl.), Brown (1966, 1967), Munyanyiwa and Blenkinsop (199"3), Hahn et al. (1990a), StiUman (1965), Wilson et al. (1993), Barton et al. (1991) and Stocklmayer (1978).

Umkondo Groups.

6uruve Metamorphic , Complex

)--

Makuti Group.

Katangan Group.

221

222

B. GOSCOMBE,P. FEYand F. BOTH

Zimbabwe are acknowledged for permission to map in the Chewore Safari area. Hubert Munyanyiwa, Tom Blenkinsop and Phillip Oesterlen are thanked for their helpful discussions. This is published with the permission of the acting Director, Mr Ncube. REFERENCES Allmendinger, R.W. 1989. STEREONET v4.0; a ploVdng programme for orientation data for the Macintosh TM Plus, SE and II computers. Cornell University, New York, U.S.A. Allsopp, H.L., Erlank, A.J. and Hornung, G. 1973 RbSr age measurements on the Umkondo dolerites of Rhodesia. In: Symposium on granites, gneisses and related rocks. (Edited by Lister,L.E.) Geological Society of South Africa, Special Publication 3, 217. Allsopp, H.L., Kramers, J.D., Jones, D.L. and Erlank, A.J. 1989. The age of the Umkondo Group, eastern Zimbabwe, and implications for palaeomagnetic correlations. South African Journal Geology 92, 11-19. Andersen, L.S. and Unrug, R. 1984. Geodynamic Evolution of the Bangweulu Block, northern Zambia. Precambrian Research 25, 187-212. Andreoli, M.A.G. 1984. Petrochemistry, tectonic evolution and metasomatic mineralisations of Mozambique belt granulites from S Malawi and Tete (Mozambique). Precambrian Research 25, 161186. Angelier, J. and Mechler, P. 1977. Sur une methode graphique de recherche des constraintes principales egalement utilisable en tectonique et en seismologie: la methode des diedres droits. Bulletin Geology Society France19, 1309-1318. Barr, M.W.C., Cahen, L. and Ledent, D. 1978. Geochronology of syntectonic granites from Central Zambia: Lusaka Granite and granite, NE of Rufunsa. Annals Society Geology Belgium 100, 47-54. Barton, C.M., Carney, J.N., Crow, M.J., Dunkley, P.N. and Simango, S. 1991. The geology of the country around Rushinga and Nyamapanda. Bulletin Geological Survey Zimbabwe 92, 220. Barton, C.M., Carney, J.N., Crow, M.J., Evans, J.A. and Simango, S. 1993. Geology and structural framework of the Zambezi belt, north-eastern Zimbabwe. In: Gondwana eight (Edited by Findlay,R.H. Unrug, R. Banks,M.R. and Veevers,J.J.) pp55-67. Balkema, Rotterdam. Bickle, M.J. and Archibald, N.J. 1984. Chloritoid and staurolite stability: implications for metamorphism in the Archaean Yilgarn Block, Western Australia. Journal Metamorphic Geology 2, 179-203. Bohlen, S.R., Bettcher, A.L., Wall, V.J. and Clemens, J.D. 1983. Stability of phlogopite-quartz and sanidine-quartz: a model for melting in the lower crust. Contn'butions Mineralogy Petrology 83, 270-277.

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Geological maps Geological map of Mount Darwin; 1:100 000. 1990a. Compiled by Bache, J.J., Dallas, S., Milian, J.F., Prost, A.E. and Rolin, P. Geological Survey Zimbabwe, Harare. Geological map of Centenary; 1:100 000. 1990b. Compiled by Bache, J.J., Dallas, S., Milian, J.F., Prost, A.E. and Rolin, P. GeologicalSurvey Zimbabwe, Harare. Tectonic map of Africa; 1:15 000 000. 1968. Compiled by Choubert, G., Muret, A.F., Dubertret, L., Sougy, J., Mestraud, J-L., Pallister, J.W., Truter, E. and VanEeden, O.R.C.C.G.M. UNESCO, Paris. Geological map of Guruve-west; 1:100 000. 1990b. Compiled by Hahn, L., Ott, G., Resch, M. and Steiner, L. GeologicalSurvey Zimbabwe, Harare. Geological map of Zambia; 1"1 000 000. 1977. Compiled by Thieme, J.G. and Johnson, R.L. Geological Survey Zambia, Lusaka, Zambia.