Kabete et al_Madagascar

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Journal of African Earth Sciences 45 (2006) 87–122

The geology, SHRIMP U–Pb geochronology and metallogenicsignificance of the Ankisatra-Besakay District,

Andriamena belt, northern Madagascar

J. Kabete a,*, D. Groves b, N. McNaughton b, J. Dunphy b

a Ashanti Exploration (Tanzania) Ltd., P.O. Box 744, Mwanza, Tanzaniab Centre for Exploration Targeting (formerly Centre for Global Metallogeny), University of Western Australia, Crawley 6009, Australia

Received 15 February 2005; received in revised form 11 January 2006; accepted 18 January 2006Available online 15 March 2006

Abstract

The Ankisatra-Besakay District (A-BD), located about 200 km north of Antananarivo and 75 km east of Maevatanana in central-northern Madagascar, hosts two historical mines, the Ankisatra Pb–Zn ± Au and Besakay Pb–Ag mines. These shear-hosted en echelon

quartz veins at Besakay and deformed tensional quartz veins at Ankisatra produced a total of 4446 t of lead and 156 t of zinc in the early1940s. In addition, there is Pb–Zn–Cu mineralisation in both quartz-feldspar leucosome veins/bands and metasomatised granulite-faciesmafic orthogneiss, Cu–Zn and associated Fe–Mn mineralisation in magnetite–pyrite enderbite breccia, and Cu–Zn mineralisation in ret-rograde shear zones in granulite-facies paragneiss in the A-BD. The country rocks in the A-BD consist of amphibolite to granulite-faciesmafic and granitoid orthogneisses and paragneisses with horizons of silicate-facies BIF. The paragneissic rocks in the district tectonicallyoverlie the biotite–granitoid hornfels and sub-volcanic mangeritic rocks, which are separated from amphibolite-facies alkali-feldspargranitoid and enderbitic rocks by a major structure. The A-BD is structurally characterised by: (1) E–W-trending tensional fractures,quartz veins and dolerite dykes; (2) buckling-related axial-planar fractures; and (3) N–S, NE–SW and NNE–SSW trending dextralstrike-slip shear zones, dolerite sills and quartz veins in transpressional extensional zones.

Uranium–Pb SHRIMP II geochronology of zircon constrains the peak of magmatic, metamorphic, deformational and metasomaticevents in the A-BD. An important constraint is whether hosting terranes contain signatures of the ca. 1690–1590 Ma critical age windowfor world-class BHT Pb–Zn–Ag deposits elsewhere in the world. At least two magmatic events are recorded from the A-BD. An earlymagmatic event is recorded by a 2725 ± 12 Ma single xenocrystic magmatic zircon in the >2676 ± 6 Ma precursor to the granulite-faciesmafic orthogneiss. A ca. 2503–2460 Ma event is recorded by a 2465 ± 6 Ma minimum age of magmatism for the precursor to metaso-matised granulite-facies mafic orthogneiss and 2483 ± 20 Ma for the precursor to biotite–granitoid hornfels. Zircons extracted from boththe metasomatised and unaltered granulite-facies mafic orthogneisses record peak metamorphic ages of 2465 ± 12 and 2390 ± 10 Ma,probably representing compressional deformation, partial melting, and associated local magmatic events within the ca. 2475–2380 Maperiod. Inherited zircons from the quartzo-feldspathic granulite-facies paragneisses return ages of protolithic supracrustal rocks rangingfrom ca. 2870 to 1700 Ma. A widespread period of rifting, anatectic magmatism, basic-ultrabasic and mangerite magmatism, and relatedgranulite-facies metamorphism occurred between ca. 820 and 780 Ma. The possible exhalative units (silicate-facies BIF and metasomatic,garnet–quartz–plagioclase rock) are of late-Archaean to early-Palaeoproterozoic, rather than Mesoproterozoic age. The terrane lacks thecritical evolution age window of ca. 1770–1550 Ma, characteristic of well-documented BHT Provinces in the Broken Hill Block and Mt.Isa Block, Australia and ca. 1959–1135 Ma from the Bushmanland Ore District, South Africa. This suggests that either such an event didnot occur in the crust now forming the A-BD or that the equivalent supracrustal rocks containing these age signatures were eroded dur-ing Proterozoic times. It is less likely that the intense 780–820 Ma event destroyed evidence for their prior existence.

The galenas from the Ankisatra and Besakay deposits have signatures characteristic of small-scale mineralised systems which derivedmost of their lead from local crustal rocks older than ca. >2.7 Ga. They are thus atypical of BHT deposits and associated vein-style min-eralisation from well-endowed terranes.

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doi:10.1016/j.jafrearsci.2006.01.008

* Corresponding author. Tel.: +255 28 2500323; fax: +255 28 2500322.E-mail address: [email protected] (J. Kabete).

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88 J. Kabete et al. / Journal of African Earth Sciences 45 (2006) 87–122

It is concluded that there are neither direct signs nor indirect temporal signals of giant stratiform/stratabound BHT Pb–Zn–Ag min-eralisation, nor clear evidence for the presence of characteristic transitional sequences and alteration styles associated with BHT miner-alisation in the A-BD, thus downgrading its prospectivity.� 2006 Elsevier Ltd. All rights reserved.

Keywords: Madagascar; Ankisatra-Besakay; Zircon geochronology; BHT deposit; Tectonic evolution

1. Introduction

There have been dramatic advances in geologicalresearch and mineral exploration in strategic areas of Mad-agascar since 1992 (Ashwal, 1997; Tucker et al., 1999;Kroner et al., 2000; Collins and Windley, 2002). Theseadvances have so far upgraded: (1) understanding of thecrustal growth history of the Malagasy Shield (Collinset al., 2001; Collins et al., 2003; De Wit, 2003); (2) its posi-tion as part of the Meso- and Neoproterozoic Rodinia andGondwana continents, respectively (e.g. Windley et al.,1994; Yoshida and Santosh, 1996; Ashwal, 1997; Handkeet al., 1999; Collins and Pisarevsky, 2005); and (3) impor-tantly an understanding of its metallogenic potential (e.g.Auge and Legendre, 1992; Windley et al., 1994). Reliablegeochronological data have been acquired, mostly by con-ventional U–Pb and Pb–Pb evaporation methods and,most recently, by U–Pb sensitive high-resolution ion micro-probe (SHRIMP) analytical techniques (Ashwal, 1997;Tucker et al., 1999; Collins et al., 2003). These data havegradually replaced the old and unreliable Rb–Sr and K–Ar ages (Cahen et al., 1984).

Significant mineralisation styles situated in the stronglyreworked high-grade metamorphic belts of the A-BD(Fig. 1) include Ni–Cr (Co) with PGE potential hostedby the ca. 790 Ma Andriamena mafic–ultramafic complexes(Ohnenstetter et al., 1991), the Besakay Pb–Ag deposit inshear-hosted quartz-veins, and the Ankisatra Pb–Zn–Audeposit in deformed quartz-veins (Besairie, 1961). In addi-tion, there is high exploration potential for Zn–Cu in goss-anous magnetite–pyrite enderbite breccia and Pb–Zn–Cuin both metasomatised mafic orthogneisses and quartzo-feldspathic granulite-facies paragneisses. These mineralisa-tion styles are related to the last intense geotectonic cycle inthe region, the ca. 800–520 Ma Neoproterozoic orogeniccycle represented by Pan-African events (Kroner et al.,1997; Goncalves et al., 2003).

Of all deposit styles in high-grade metamorphic belts,BHT (Broken-Hill Type) Pb–Zn–Ag deposits are the mosthighly targeted (e.g. Walters, 1998). They represent a dis-tinct category of Pb–Zn–Ag mineralisation commonlyassociated with unusual exhalative chemical sedimentaryrocks (lode horizons) situated in polydeformed amphibo-lite- to granulite-facies orogenic belts (Kerr, 1994; Reidet al., 1997; Giles et al., 2004), specifically of early Meso-proterozoic age e.g. Groves et al., 2005). BHT provincescommonly display an extensive meta-supracrustal stratig-raphy, comprising a lower sequence of quartzo-feldspathic

rocks and an upper sequence of psammo-pelitic rocks, sep-arated by transitional sequences of hydrothermal exhala-tive origin (Roche, 1994). Thin siliceous units, magneticlode horizons, including BIF, calc-silicate units, gahnitequartzite and other unusual chemical sedimentary rocks,and most importantly BHT stratiform/stratabound Pb–Zn–Ag deposits, typify these transitional sequences. Thereare narrow, small-scale cross-cutting mineralised veins andshear zones in these BHT Provinces (Gulson et al., 1985;Reid et al., 1997). Based on Pb-isotope data, vein-typePb–Zn–Ag deposits in some of these provinces are shownto be genetically related to giant stratiform/strataboundPb–Zn–Ag deposits (e.g. Aggeneys, Gamsberg, Broken Hilland Black Mountain in the Bushmanland Ore District,Namaqua Metamorphic Province, South Africa: Reidet al., 1997). Based on this knowledge, it is possible thatvein-type mineralisation in potential BHT provinces couldbe used as pathfinders for locating giant stratiform/strat-abound deposits, which may have been structurally relo-cated at depth or along strike during polydeformationevents. However, in the world-class BHT Provinces ofBroken Hill Main Lode in New South Wales, and theCannington deposit in Queensland, Australia, stratiform/stratabound Pb–Zn–Ag deposits are not apparentlygenetically linked to the cross-cutting mineralised veinsand shear zones, at least based on Pb-isotope signatures(Gulson et al., 1985; Stevens et al., 1990), so Pb-isotopedata are equivocal. In addition to the Palaeoproterozoicand Mesoproterozoic orogenic belts, some analogous,but much smaller and less economic BHT deposits havebeen recognised in younger mobile belts (e.g. CambrianKanmantoo Belt of southern Australia: Willis, 1996).These occurrences suggest some metallogenic cyclicity forBHT deposits, but the younger examples are not primaryexploration targets.

The timing of formation of BHT deposits and precur-sors to the hosting metamorphic rocks are difficult todetermine due to high-grade metamorphism, intense meta-somatism and polydeformation (Walters, 1998; Nutmanand Ehlers, 1998b). In consequence, dating of the BHTmineralisation has relied on two indirect approaches: (1)Sm–Nd-isotope modelling of primary crustal-formationage together with SHRIMP U–Pb dating of high-gradepeak-metamorphic zircons; and (2) isotope modelling ofPb in pristine galena and sphalerite assumed not to containsignificant U and Th (Reid et al., 1997).

In this paper, key geological factors (i.e. lithostratigra-phy, mineralisation styles, SHRIMP II U–Pb zircon dates

Fig. 1. The main tectonic domains and belts comprising the central and north-eastern Madagascar Province.

J. Kabete et al. / Journal of African Earth Sciences 45 (2006) 87–122 89

and Pb-isotope signatures) of the A-BD are presented thenexamined in terms of its potential to host large BHT Pb–Zn–Ag deposits, thus highlighting the potential mineralbase-metal endowment of the A-BD. It is stressed that itis the temporal evolution of the District that is considered,not its spatial position, relative to BHT provinces else-where in the world. It is nowhere implied that the A-BD

was adjacent to any of the world-class BHT provinces dur-ing BHT mineralisation.

2. Rationale for the study

Criteria for the recognition of mineral districts or prov-inces which are prospective for BHT mineralisation have


been summarised by many workers including AGCRC(1995), Giles and Ehlers (1997) and Reid et al. (1997).These features are used here to determine whether theA-BD has the potential to host BHT deposits. Suchan assessment is warranted, because most giant BHTPb–Zn–Ag deposits occur in extensive Palaeoproterozoicand Mesoproterozoic orogenic belts with various types ofvein-type and young intrusion-related mineralisation styles(Nutman and Ehlers, 1998b; Reid et al., 1997). Some ofthese orogenic belts have an Archaean to Palaeoprotero-zoic geological history that is comparable to that of theA-BD (e.g. ca. 2700–2100 Ma events in both the BrokenHill Block of Australia and A-BD: Nutman and Ehlers,1998a). They also occur in high-grade metamorphic beltswhich were involved in poly-deformational and multiplemagmatic events, comparable to the Pan-African Belts ofMadagascar. Furthermore, galenas extracted from vein-type Pb–Zn–Ag mineralisation in the A-BD indicated Pbmodel ages of ca. 1850 Ma and 1750 Ma (Besairie, 1961),which are within the ca. 1920–1570 Ma terrane evolutionage window for BHT Pb–Zn–Ag deposits in the BrokenHill Block of Australia.

Based on these factors, BHP Minerals InternationalExploration Inc. decided to target BHT deposits in theA-BD, and this research study arose from the ensuingexploration-based work.

Fig. 2. Geological map of central and northern Madagascar Province showingzircon conventional/SHRIMP unless otherwise stated.

3. Regional geological setting

Two-thirds of the island of Madagascar is underlain byPrecambrian rocks of the Malagasy Shield (Besairie, 1961;De Wit, 2003). Windley et al. (1994) and Windley andRazakamanana (1996) originally split this shield, alongthe sinistral Ranostara Shear Zone, into the central-north-ern Madagascar Province (CNMP) of predominantlyArchaean age components and the southern MadagascarProvince of largely Proterozoic cover (Fig. 1). The CNMPhas since been subdivided into several tectonic units basedon available robust geochronology data. Collins et al.(2000) distinguish five tectonic units in the CNMP, whereasDe Wit (2003) split the shield into three major belts to illus-trate the growth history from the Archaean, through Meso-Neoproterozoic to Neoproterozoic belts.

In the Maevatanana–Andriamena Belts (Fig. 2), thegreenstone belts (De Wit, 2003), also referred to as theTsaratanana Thrust Sheet (Collins et al., 2001), are sepa-rated by N–S trending lower-crustal level migmatites,which are conformable with the ca. 600–550 Ma granitoids(Nedelec et al., 1994; Windley et al., 1997; Kroner et al.,1997; Paquette and Nedelec, 1998). In the MaevatananaGreenstone Belt, thin greenstone lithologies are interca-lated within the 2518–2505 Ma (U–Pb zircon: Tuckeret al., 1999) granodiorites and tonalites. All of these are

known geochronology and mineral deposit potential. Age dates are U–Pb


intruded by the ca. 780 Ma arc-magmatic gabbro andgranitoids, part of the ca. 870–740 Ma arc-magmatic gab-bro and granitoids in the ca. 2550–2500 Ma AntananarivoBlock (Fig. 1).

The Andriamena Belt is underlain by mafic–felsic gran-ulites and gneisses including unspecified Archaean base-ment rock (ca. 2.7 Ga: Rb/Sr whole-rock isochron:Windley and Razakamanana, 1996) and the ca. 3.2–2.4 Ga tonalitic gneisses (Collins et al., 2001). These base-ment rocks are unconformably overlain, and in most placesdeformed together with, cordierite–sillimanite–garnet andgraphitic paragneisses and granulites exposed along theAndriamena-Alaotara synclinal structure (Fig. 2). The787 ± 6 Ma (step-wise Pb-evaporation U–Pb in zircon:Guerrot et al., 1993) layered mafic–ultramafic complexesin the belt are interpreted to have intruded in the back-arc intracratonic rift setting related to subduction of theMozambique Ocean. The synformal structure in the Andri-amena Belt is probably a result of thrusting and imbrica-tion of the supracrustal rocks and associated ultramaficcomplexes following the ca. 600–550 Ma collision betweenNeoproterozoic India and the Congo/Tanzania/Bangwe-ulu Block (Collins and Pisarevsky, 2005). This is consistentwith the suggestion by Windley and Razakamanana (1996)that the ca. 2517–2494 Ma (Ashwal, 1997) Andriamena,Masora, Beforona and Androna Greenstone Belts arethrust belts which were part of a thrust nappe that wasprobably a once-continuous sheet. The high-grade meta-morphic rocks underlying the Andriamena Belt were finally

Fig. 3. The geology of the Ankisatra-Besakay District showing areas of mitechnique in this study.

intruded by conformable granitoids and post-tectonic per-alkaline complexes during the late Pan-African (Fig. 2:Nedelec et al., 1995; Goncalves et al., 2003).

4. Lithology and structure of the A-BD

The major lithotectonic units of the A-BD are groupedunder three major sub-terranes: (1) the orthogneissic thrustsub-terrane in the west; (2) the tectonic corridor sub-ter-rane in the east; and (3) the tectonic boundary sub-terranein the centre (Fig. 3). The orthogneissic thrust sub-terranecomprises gneissic basement rocks such as voluminous dio-rite to granodiorite gneiss, orthopyroxene–quartz–feldspargneiss (e.g. tonalitic gneiss), granitoid gneiss, thin units ofinterbanded quartz-magnetite–amphibolite gneiss (cf. sili-cate-facies banded iron formation: BIF) and thin green-stone belts (e.g. Fig. 4N). These rocks are folded andfaulted and intruded by early Neoproterozoic mafic–ultra-mafic complexes (see Guerrot et al., 1993), orthopyroxenealkali-feldspar granitoid and minor late-Neoproterozoicper-alkaline granite (sensu stricto) sheets and Cretaceousmafic dykes (Besairie, 1961).

In the eastern A-BD, the tectonic corridor comprisesthrust-transported quartzo-feldspathic granulite-facies par-agneiss, biotite–granitoid hornfels and mangerite (Fig. 3).North–south-trending mafic–ultramafic intrusive rocks,dolerite dykes and conformable granite (sensu stricto)bodies intrude quartzo-feldspathic paragneiss in the south,whereas N–S trending gabbroic sills and NNW-trending

neral occurrences and locality of samples dated by the U–Pb SHRIMP


pink granites (sensu-stricto) intrude the northern portion(Fig. 3). The tectonic boundary sub-terrane comprises anextensive N–S to NNE-trending structure consisting of

Fig. 4. A. Greenish-grey pyroxene-amphibole-rich melanosome and quartzo-famphibolite to transitional granulite facies mafic orthogneiss. C. Laminatedmagnetite and quartz-magnetite interbands in C. E. Strong metasomatism in debands in mafic orthogneiss. G. Mineralised orthopyroxene-quartzo-feldspatmineral assemblage typical of granulite facies metapelitic rock with a basic cMangerite typified by uralitised and biotite-altered pyroxenes, plagioclase andsyn-deformation mangerite. L. Photomicrograph showing near granoblastic texmagnetite in alkali-feldspar granitoid. M. (i) Strong brittle–ductile shear fabultramylonitic enderbite. N. Photomicrograph of moderate to strongly shearehosting enderbite. P. Amphibolite greenstone lithology intercalated within ma

amphibolite-facies alkali-feldspar granitoid, stronglydeformed and intruded by mangerite, diorite and basalticdykes. Ubiquitous Cretaceous basalt dykes are common

eldspathic leucosome bands in mafic orthogneiss. B. Photomicrograph ofquartz-magnetite amphibole gneiss (cf. silicate facies BIF). D. Silicate–formed mafic orthogneiss. F. Photomicrograph of a magnetite-plagioclasehic gneiss. H. Photomicrograph showing biotite–garnet–orthopyroxeneomponent. I. Sugary and mottled texture of biotite granitoid hornfels. J.minor quartz. K. Field relationship between alkali-feldspar granitoid andture of quartz-feldspar uralitised and biotite-altered green hornblende andric and pervasive hydrothermal alteration in enderbite; (ii) mylonitic tod enderbite. O. Laterite (gossan) expression of the manganiferous Cu–Znfic orthogneiss, after diorite and granodiorite.

Fig. 4 (continued)


at all lithotectonic levels, implying that they are roots toplateau basalts mapped in the west (Fig. 3). The degreeof metamorphism and deformation across the A-BDincreases from east to west (Kabete, 1999).

4.1. Orthogneissic thrust sub-terrane

The orthogneissic thrust sub-terrane comprises region-ally extensive tectonically-juxtaposed late-Archaean dio-ritic to granodioritic gneisses and thin greenstonelithologies of basaltic andesite composition (Kabete,1999). The late-Archaean basement rocks contain composi-tional inter-layering possibly evolved during earlier partialmelting events related to metamorphism. These basementrocks and the ultramafic complexes appear to have beenjuxtaposed before the open to tight isoclinal folding andpenetrative bulk-shearing events responsible for the N–S

and NNE-trending regional structural grain (Figs. 3 and4A). Axial traces to these folds and the shear plane fabricdip at moderate angles to the east. These rocks containlocal, yet extensive, zones of parallel to crosscutting shearand fault zones ranging from 50 cm to up to 40–60 m widecomprising smoky quartz, plagioclase, magnetite, garnetand altered pyroxenes. Petrographic study of these struc-tures indicates that they are zones of strong metasomaticalteration in granulite-facies mafic orthogneiss (Fig. 4Eand F). The intensity of deformation and mineralogy ofthe metasomatised granulite-facies mafic orthogneiss sug-gests that the structures which accessed the hydrothermalfluids were detachment zones related to early thrustingevents and were reactivated during younger events. Themetasomatised granulite-facies mafic orthogneiss typifiesthe nature and fabric of the bulk-shear structure of theA-BD, which is further crosscut by NW-trending dextral


faults (Fig. 6). The relatively younger, Cretaceous basaltdykes are aligned along NE to N–S and NNW–SSW struc-tures (Fig. 3).

4.1.1. Amphibolite–granulite-facies mafic orthogneiss

The granulite-facies mafic orthogneiss constitutes athick unit of dominantly uralite-amphibole and biotite-altered, pyroxene-rich melanosome bands, interlayeredand/or tectonically deformed together with centimetre-to metre-scale quartz-feldspar leucosome veins/bands(Fig. 4A and B). There are local shear zones, developedparticularly in the vicinity of interpreted deep-seated struc-tures, which define zones of intense deformation and meta-somatism. They comprise smoky and blue-quartz anduralitised orthopyroxene together with fracture-filling pyr-rhotite, pyrite after pyrrhotite and magnetite.

Petrographic studies also reveal hornfelsic decussatefeatures shown by green hornblende and subordinate bio-tite, forming as clusters with granular quartz, suggestingprograde amphibolite to granulite facies metamorphism.Orthopyroxene and hornblende generally show preferredorientations consistent with the regional fabric (Fig. 4Aand B). The growth of biotite and hornblende along inter-granular spaces, and amphibole and biotite rimming ofstubby pyroxenes, possibly represents retrograde meta-morphism (e.g. Opx + Pl + Kfs + Vapor! Hbl + Bt +Qtz) following exhumation of granulites. Field and petro-graphic investigation and geochemical analyses (Kabete,1999) show that the granulite-facies mafic orthogneissresulted from partial melting of diorite to granodioriticrocks.

4.1.2. Quartz-magnetite–amphibolite gneiss

Quartz-magnetite–amphibolite gneiss is essentially anamphibolite- to granulite-facies, silicate-facies BIF crop-ping out extensively as thickly bedded units within thegranulite-facies mafic orthogneiss and metasomatisedgranulite-facies mafic orthogneiss (Figs. 3 and 4C andD). The rock comprises millimetre to centimetre-scale,quartz-magnetite, shear-foliated siliceous units that areinter-laminated with centimetre to metre-scale, magnetite-supported silicate breccia. In the magnetite-supportedsilicate breccia bands, coarse-grained uralitised pyroxenecontains plagioclase and quartz inclusions, a feature absentin uralitised pyroxenes in shear-foliated silica-rich bands.Laminations, chemical composition and petrographicfeatures suggest that the quartz-magnetite–silicate gneissis a chemically precipitated silicate-facies BIF. The bulkC-S shear fabric is of contrasting micro-textural–structuralfeatures, possibly related to post- to syn-deformationalmetasomatism and magnetite flooding along bands domi-nated by the S-fabric (see lower portion in Fig. 4D).Inter-layering and the mineralogy suggest an exhalativevolcanic origin, before the rock was subjected to granu-lite-facies metamorphism. Inclusions of elongate quartzand feldspar grains, which are parallel to internal cleavageplanes in uralitised orthopyroxene, provide evidence for the

granulite-facies metamorphism. Intense uralitisation of sil-icate minerals is interpreted to result from retrogrademetamorphism.

4.1.3. Metasomatised granulite-facies mafic orthogneiss

Metasomatised granulite-facies mafic orthogneiss com-prises metre-scale layering of massive and dark-grey gar-net-bearing magnetite–quartz–feldspathic rock (Fig. 4Eand F) cropping out in the southwestern A-BD (Fig. 3).These garnet-bearing magnetite–feldspathic layers andunits are located in parallel to sub-parallel structureswithin granulite-facies mafic orthogneiss. They are stronglyfractured and consist of orthopyroxene, plagioclase, inclu-sion-free euhedral garnet porphyroblasts and intergranularmagnetite containing pyrrhotite exsolutions. Elongate tostubby-prismatic pyroxenes form a weak foliation alternat-ing with quartz–feldspar-rich areas. Uralite, garnet andmagnetite appear to be concentrated in areas of high strainand abundant orthopyroxene. These petrographical fea-tures suggest that at least garnet, and possibly orthopyrox-ene, are the result of metasomatism following interactionbetween hydrothermal fluids and mafic granulite interac-tion along low-angle thrust faults splaying off deep-rooteddetachment zones. In addition, uralite and biotite are tran-sitional, partial replacement products of orthopyroxenefrom granulite to lower amphibolite facies (Miyashiro,1994).

4.2. Tectonic corridor sub-terrane

The tectonic corridor sub-terrane is underlain to thenorth by granulite-facies biotite–granitoid hornfels andgranulite-facies quartz–monzonite (i.e. mangerite), and tothe south by quartzo-feldspathic paragneiss and sporadicultramafic intrusive complexes (Fig. 3). A greater numberof N–S trending basalt dykes and faults crosscut NNE–SSW trending fabrics which are coincident, and/or parallel,with the alignment of mangerite and gabbro outcrops in thenorth. These structures are further crosscut by late NW-trending dextral faults. In the southern part, there arenumerous closely-spaced NNE-trending major shear andfault zones, locally ranging from mylonite to ultra-mylonitezones that traverse the tectonic corridor (Fig. 4H).

4.2.1. Quartzo-feldspathic granulite-facies paragneiss

Quartzo-feldspathic granulite-facies paragneiss is a leuc-ocratic orthopyroxene- and sillimanite-bearing lithotec-tonic unit extending along the N–S trending AndriamenaBelt in the southern part of the tectonic corridor sub-terrane (Figs. 3 and 4G). Petrographic studies show thatsillimanite overgrows garnet and orthopyroxene overgrowsgarnet and biotite (Fig. 4H), consistent with upper-amphibolite to granulite metamorphic facies (Miyashiro,1994). The retrogression of biotite and garnet to chloriteis common along local, but late, dextral shear zones, whichalso contain graphite, rutile and sulphide mineralisation(Fig. 5D).

Fig. 5. A. Photomicrograph illustrating the structural controls and galena textures in quartz from the Besakay Pb–Ag deposit. B. Remobilised galena inclosely-spaced fractures in quartz at the Ankisatra Pb–Zn–Au deposit. C. Photomicrograph showing texture and ore mineralogy of Pb–Cu–Zn sulphides inmetasomatised mafic orthogneiss. D. Photomicrograph showing pyrrhotite and pyrite in late anastomosing shear zones in quartzo-feldspathicparagneisses. E. Sheared enderbite showing coincidence of sulphides in areas of abundant garnet. F. Photomicrograph illustrating the structural controlspertaining to pyrrhotite and pyrite in mineralised enderbite. G. Metre-scale quartzo-feldspathic leucosome bands with ribbons of quartz containing Pb–Agmineralisation. H. Strongly altered quartz-epidote and Pb-bearing leucosome bands in mafic orthogneiss.


The mineral assemblage Bt + Grt + Sil + Kf suggeststhat the main component of the precursor quartzo-feld-spathic granulite-facies paragneiss contained significantpelitic components. Orthopyroxene in this assemblage sug-gests the presence of a basic component in the precursorrock, which further implicates a supracrustal volcano-sed-imentary rock. Both garnet and orthopyroxene showsyn- to post-metamorphic deformation and retrogrademetamorphism by their hydration to biotite and chlorite

(Figs. 4H and 5D). Graphite along shear zones may haveeither precipitated from CO2-rich hydrothermal fluids(Miyashiro, 1994), or be related to organic matter in theprecursor rock (Rajesh-Chandran et al., 1996).

4.2.2. Biotite–granitoid hornfels

Biotite–granitoid hornfels crops out in the northern partof the tectonic corridor sub-terrane as elongate slices ofmoderately to strongly weathered basement rock intruded


or juxtaposed with mangerite (Fig. 4J) and gabbro (Fig. 3).Biotite–granitoid hornfels is a pink to pinkish-grey, med-ium-grained recrystallised rock consisting of granoblasticquartz, plagioclase and feldspars (Fig. 4I). The clusteringof biotite and finely recrystallised granular quartz, the typ-ical decussate texture of local overgrowths of biotite byclinopyroxene, and triple-junction textural features dis-played by granoblastic quartz and feldspars are evidencefor high-temperature low-pressure amphibolite- to granu-lite facies metamorphism (Miyashiro, 1994; Blat andTracy, 1996). This is a contact metamorphic facies, com-monly referred to as pyroxene-hornfels facies (Miyashiro,1994), characterised by serrated margins and recrystallisa-tion of fine granular quartz along quartz grain boundaries.Accessory titanite, zircon and opaque minerals are com-monly located along intergranular spaces, with sub-rounded and sub-angular zircons enclosed in quartz andfeldspars. However, minor zircons and titanites that occuras inclusions in biotite show radiogenic coronas.

Field and petrographical work suggest that biotite–gran-itoid hornfels resulted from high-temperature low-pressurepyroxene-hornfels metamorphism of the granitoid. Theintrusive character of the extensive granulite-facies mange-rite potentially explains the source of heat.

4.2.3. Enderbite

Enderbite is a moderately to strongly deformed, two-pyroxene-bearing quartzo-feldspathic tonalitic to quartz-dioritic sub-volcanic rock cropping out extensively in thesouthern extension of the tectonic boundary (Fig. 4K). Itforms a brittle–ductile structural transition from theamphibolite-facies alkali-feldspar granitoid to the orthog-neissic thrust sub-terrane. The rock is generally coarsegrained and granoblastic, consisting of about 20–40%orthopyroxene and clinopyroxene, 35–55% plagioclase,15–35% quartz, and minor garnet, biotite and alkali-feldspar (Fig. 4L). These major diagnostic textural andmetamorphic features suggest an igneous precursor, over-printed by retrograde metamorphism. A strong foliation,possibly secondary, is shown by pyroxenes that are elon-gate parallel to garnet, biotite and opaque minerals.However, in some places, orthopyroxene and plagioclasepreserve relic igneous textures locally mimicked by uraliti-sation, biotite and sericite alteration (Fig. 4L). Garnet isinclusion-free and lacks chemical zonation, diagnostic evi-dence for a high-grade metamorphic overprint (Miyashiro,1994).

In the field, enderbite shows significant brittle–ductilestructures, parts of which are mineralised (Fig. 5E andF). The age of these structures is not constrained, asthe age of the rocks themselves is not known. Enderbiteexhibits high strain features in which dextral shearingprogressively forms mylonite and ultra-mylonite zonesproducing granulated pyroxenes in silicate-rich bands(Fig. 4K and L). These silicate bands alternate withbands or ribbons of recrystallised granular quartz wrap-ping around rotated feldspars (Fig. 4L). Opaque minerals

form thin layers in these bands, despite their occurrencealong fractures and intergranular spaces in other ender-bite varieties.

It is evident that enderbite in the A-BD represents two-pyroxene bearing tonalitic to dioritic sub-volcanic rocks,which were emplaced, crystallised and deformed at mid-crustal levels.

4.3. Tectonic boundary sub-terrane

The tectonic boundary sub-terrane (Fig. 3) is defined bythe structurally-bound, amphibolite-facies alkali-feldspargranitoid, comprising buckling-related radial axial planesand strong N–S trending faults and shear zones throughwhich subsequent dolerite and basalt dykes intruded(Fig. 6). These structures are evidence for horizontal com-pressional tectonics, which created zones of extension inthe tectonic corridor basement-rocks, through which latersills intruded (Figs. 3 and 4L). Rejuvenation of these exten-sion zones during progressive and subsequent orogenicevents may have allowed further zones of extension in thebasement through which emplacement of ca. 590–560 Maconformable granites (sensu-stricto) and Cretaceous basaltdykes took place (Figs. 2, 3 and 7).

4.3.1. Amphibolite facies alkali-feldspar granitoid

Amphibolite facies alkali-feldspar granitoid forms a N–S to NNE-oriented fold structure defining a possible tec-tonic boundary, which wraps around a tectonic corridorof quartzo-feldspathic granulite-facies paragneisses, bio-tite–granitoid hornfels and quartz monzonite to the east(Figs. 3 and 6). Parallel, N–S-trending shear zones and dol-erite dykes, which further cut and intrude the granitoid(Fig. 4O), are possibly related to the extensive N–S toNNE-trending dextral shear and extension zones in thenearby orthogneissic thrust sub-terrane (Fig. 6).

The alkali-feldspar granitoid is a coarse-grained pinkish-grey and moderately magnetic rock consisting of abundantperthitic-alkali-feldspars, uralitised pyroxenes, and greenand brown hornblende (Fig. 4O). The relatively minormagnetite and other opaque minerals, and other silicates,form thin bands mainly at granoblastic quartz-feldsparintergranular spaces and in anastomosing shears(Fig. 4P). In places, hornblende shows decussate textureindicative of contact metamorphism at amphibolite togranulite facies, whereas strain features in quartz andabundant exsolution lamellae in feldspars are evidencefor deformation. These features, as well as corona texturesshown by biotite around uralite and hornblende, suggestpost-metamorphic deformation with subsequent retrogrademetamorphism. At outcrop and specimen scales, amphibo-lite–granulite facies alkali-feldspar granitoid shows igneousmorphology with weak metamorphic overprints of incipi-ent gneissic banding. However, in contrast, petrographicobservations show clear metamorphic textures (Fig. 4P),thus presenting ambiguities on how granulite-facies

Fig. 6. A. Landsat image band-five image showing lithologic and structural interpretations. B. Proposed lithostructural evolution of the ca. <810–780 MaAnkisatra-Besakay District based on A and field mapping results.


metamorphic and igneous rocks form (see Windley, 1995,for comments).

Possible origins for the protolith of the amphibolite-facies alkali-feldspar granitoid include: (1) intrusion and

crystallisation of precursor magma at deeper crustal levels;(2) intensive metasomatism by CO2-rich fluids emanatingfrom high-grade metamorphism of basement rocks,focussed from deep seated structures; and (3) formation

Fig. 7. Detailed geology of the area around the Besakay Pb–Ag Deposit.


of a residue of a partially melted precursor basement rockthat remained and recrystallised at deeper crustal levels(e.g. Tarney and Weaver, 1987).

5. Common features of BHT deposits

BHT deposits are discussed here as a basis for discussionof the U–Pb geochronology and Pb-isotope data from theA-BD, which follows. BHT districts are commonly charac-terised by their: (1) intracratonic rift basin setting; (2)upper amphibolite- to granulite-facies metamorphism, withstrong metasomatic overprints; (3) multiple episodes ofdeformation and magmatism; and (4) unique transitionallode horizons (Gulson et al., 1985; Kerr, 1994). Debateon the genesis of BHT deposits highlights two main contro-versial models: (1) the syn- to post-metamorphic, mag-matic-hydrothermal origin; and (2) the syn-sedimentaryor diagenetic model (Giles and Ehlers, 1997).

The syn-metamorphic model involves scavenging of Pb–Zn–Ag metals by pervasive partial melting of basementrocks during high temperature metamorphism and theirdeposition in structurally or chemically favourable sites(AGCRC, 1995). The syn-sedimentary model suggests ero-sion and accumulation of base metals from the syn-riftingquartz-feldspar basement rocks (Gulson et al., 1985;Sawkins, 1989; Stevens et al., 1990; Solomon and Groves,1994).

5.1. Hosting lithology characteristics

Critical geologic features of BHT ore deposits and theirhosting terranes in the Broken Hill Block, Mt. Isa Blockand the Bushmanland Ore District are largely outlinedfrom Gulson et al. (1985), Ryan et al. (1986), Joubert(1986), Skrzeczynski (1993), Solomon and Groves (1994),Kerr (1994), AGCRC (1995), Walters (1996), Willis(1996), Giles and Ehlers (1997) and Reid et al. (1997).BHT districts typically display the following features:

(1) They are Palaeo- to Mesoproterozoic intracratonicmobile belts with evidence of extensional tectonicand long-lived orogenic evolution (e.g. it took100 m y for the Willyama Supergroup to evolve from1690 to 1590 Ma: Giles et al., 2004).

(2) These belts are metamorphosed to upper amphiboliteto granulite facies, with evidence of intense deforma-tion and metasomatism, which complicates precursorrock identification.

(3) Regional stratigraphy comprises lower quartzo-feld-spathic sequences and upper metapelitic, psammiticand psammopelitic sequences separated by transi-tional, syn-rift and thin siliceous and magnetic lodehorizons of hydrothermal exhalative origin. Thesehorizons comprise BIF and gahnite quartzite andother unusual chemical sedimentary rocks.


(4) Deposits are strongly to less-deformed, disseminated-sulphide stratiform orebodies, conformable mas-sive-sulphide stratabound deposits, and cross-cuttingquartz and pegmatite vein-type deposits, the majorityof which are commonly located in transitional zones.

(5) There is extensive sillimanite–garnet metasomatism inquartz-feldspar rocks (e.g. Cannington), and complexFe–Mn–Ca silicate haloes, expressed by garnet,pyroxene and pyroxenoids (e.g. Broken Hill Block).

Metamorphosed siliceous Fe–Mn garnet–quartzite andsandstones form a blanket around ore bodies.

5.2. BHT deposit characteristics

World class BHT districts show the following majorBHT deposit-scale characteristics: (1) obvious gross con-cordance between ore bodies and hosting lithotectonicunits; (2) original sedimentary features preserved in orebodies that are restricted in low-strain stratigraphic inter-vals (Reid et al., 1997); (3) extensive regional stratigraphiccorrelation of the hosting terranes (Giles and Ehlers, 1997);and (4) Si–Mn–Fe and K-metasomatic geochemical, miner-alogical and Pb-isotopic haloes (e.g. Bushmanland OreDistrict) associated with ore bodies (Walters, 1996; Willis,1996; Reid et al., 1997).

Apart from stratabound/stratiform deposits, vein-typeand intrusion-related mineralisation occurs in these prov-inces. The most critical features of the widely acceptedsyn-sedimentary genetic model for BHT deposits are: (1)initial crustal extension of the Archaean to Palaeoprotero-zoic basement; (2) older basement and associated basalsupracrustal sequences comprising quartz-feldspar rocks;and (3) long-lived orogenic cycles and evidence of transi-tional syn-rifting depositional environments.

6. SHRIMP U–Pb zircon geochronology of critical A-BD

Rocks

6.1. Sample selection and preparation

A total of five samples comprising amphibolite to gran-ulite-facies mafic orthogneiss, metasomatised amphiboliteto granulite-facies mafic orthogneiss, quartzo-feldspathicamphibolite to granulite-facies paragneiss, amphibolite-facies alkali-feldspar granitoid, and biotite-orthogneissicgranitoid were selected for SHRIMP U–Pb analyses inorder to cover the entire history of the A-BD (Fig. 3). Sam-ples were crushed, milled and sieved to retain the 60# meshfraction, which was cleaned with normal tap water toremove finely clayey material. Retained silty material wasdried and mixed with a 2.85 ± 0.02 g/mil heavy liquid(Lithium-Silica-Tungsten: LST) and allowed to settle, sep-arating heavy from less-dense material. Distilled water fol-lowed by acetone was used to clean the heavy fraction. Amagnetic separator was used on the dry fraction. Zircons

from the non-magnetic fraction were handpicked under ahigh-magnification binocular microscope. These grainswere mounted together with chips of a standard zirconCZ3 (Nelson, 1996), polished and cleaned to expose half-section zircon grains. The mounts were imaged by environ-mental scanning electron microscopy (ESEM) using thecharge contrast technique (Griffin, 1997) to reveal detailedinternal morphological features. The samples were thencleaned and gold coated in preparation for SHRIMP ion-microprobe analysis.

6.2. Nature of U–Pb zircon SHRIMP geochronology

The U–Pb and Th–Pb geochronology technique utilisesminerals with high U/Pb and Th/Pb ratios, such as zircon,titanite, apatite and monazite in order to determine reliableand absolute ages of rock units. The U–Th–Pb isotopein situ analyses of 25–30 lm size spots on sectioned zircongrains or titanite by SHRIMP are among the best high-pre-cision dating techniques capable of resolving ambiguousU–Pb discordance patterns commonly obtained from anal-ysing a mixture of zircon populations by conventional U–Pb and Pb–Pb evaporation techniques. Compston et al.(1984), Nelson (1996) and Smith et al. (1998) describe inmore detail the analytical methods applied to measureU–Th–Pb isotopes on SHRIMP and calculate ages. Agesquoted in the text are at 2-sigma error, whereas those inthe table of SHRIMP analytical results are reported at 1-sigma error.

6.3. Imaging techniques

Transmitted and reflected light photographs providepreliminary information about the zircon morphology, col-our, size, and where appropriate, core–rim relationshipand zoning in the grains. In terms of zircon overgrowthand damage complexities due to multiple deformation,magmatic and metamorphic processes, it is important toselect specific geological targets to be analysed by theion-microprobe SHRIMP technique. Scanning ElectronMicroscope (SEM) imaging techniques, including back-scatter electron (BSE) and cathodoluminescence (CL), pro-vide far greater information about the detailed internalmorphology of the selected zircons. Essentially, these tech-niques employ a finely focused primary beam of electronswhich sweeps across the sample resulting in the generationof several emissions from the sample: secondary electrons,backscattered electrons and CL. At the time of thisSHRIMP study, Griffin (1997) had developed the chargecontrast technique (CCI) to be used within an environmen-tal scanning electron microscope (ESEM). As this tech-nique was stated to produce images similar to thoseproduced by CL, but to be quicker, not to require carboncoating, to have a better signal/noise ratio, and to high-light details not visible in BSE and CL (Griffin, 1997), itwas decided to use this technique to image zircons in thisstudy.


7. Dated rocks and morphology of their zircons

7.1. Granulite-facies mafic orthogneiss

The granulite-facies mafic orthogneiss (sample 67-04/UWA mount 98-19B; Fig. 4A and B) was selected for geo-chronological studies to constrain the age of the major hostrock to the vein-type Pb–Ag mineralisation in the Besakayarea (Figs. 2 and 7). Zircons extracted from the sample arebrown, anhedral and elongate, showing external igneousmorphologic features. Reflected light and ESEM examina-tion of zircons reveals external and internal morphologicfeatures as classified as: (1) least-modified igneous zirconscomprising 100–300 lm long zircons with internal igneousgrowth zoning, some of which are mechanically fragmentedand others have very thin colourless zircon overgrowths(Fig. 8: grains 30 and 6); (2) moderately modified igneouszircons with resorbed edges, lack of internal growth zon-ing, and presence of secondary internal features (Fig. 8:grain 14); and (3) heterogeneously fragmented and feature-less zircons with angular edges and homogeneous internalmorphology (Fig. 8: grain 3). These morphological varia-tions are interpreted to be related to partial recrystallisa-tion or annealing of primary igneous zircons duringsubsequent magmatic and metamorphic events.

7.2. Metasomatised granulite-facies mafic orthogneiss

A dark-grey garnet–magnetite–quartzo–feldspathicgneiss (sample 21349; Fig. 4E and F), interpreted to bea metasomatised granulite-facies mafic orthogneiss, islocated in the southwestern part of the A-BD (Fig. 3). Thisrock was chosen for SHRIMP geochronology in order toconstrain the timing of the metasomatic event which ishypothesised to have occurred along deep-rooted thrust-fault splays from detachment soles developed during a tec-tonic shortening event of unconstrained age. Based on fieldobservations, these faults and shear zones are proximal toboth the layered magnetite–silicate unit (cf. silicate-faciesBIF) in the south and the eastern part of the BesakayPb–Ag deposit in the northern part of the study area(Fig. 3).

About 75% of zircons extracted from sample 21349(UWA mount 98-19C) are brownish, relatively fine-grainedand have a well rounded, sub-spherical external morphol-ogy (�40 lm in diameter), whereas the rest are elongate(up to 250 lm long). Following transmitted light andESEM examination of these zircons, two major groupscan be delimited: (1) oval and elongate zircons (Fig. 8:grain 51), with simple to homogeneous internal structure(Fig. 8: grain 114), locally with fragmented margins; and(2) near-spherical, well-rounded zircons ranging fromhomogeneous, featureless to simple internal morphologies(Fig. 8: grain 58). Some of these zircons show multifacetedexternal features in reflected light images, with distinctcore–rim relationships (Fig. 8: grain 60). Both groups showvariable degrees of irregular external morphologies, possi-

bly related to mechanical abrasion or metamorphicresorption.

7.3. Amphibolite-facies alkali-feldspar granitoid

Sample 21380 represents an amphibolite-facies alkali-feldspar granitoid (Fig. 4O and P), which defines a tectonicbreak between the orthogneissic thrust sub-terrane to thewest and the tectonic corridor sub-terrane to the east(Fig. 3). Its synclinal structural setting is interpreted to bea result of syn-deformational emplacement into a NE–SWtranspressional stress regime (Fig. 6). Whereas field rela-tionships show its igneous granitoid features, it has weaklygneissic textures, and petrographic examinations show thatit has experienced amphibolite- to transitional granulite-facies metamorphism (Fig. 4P). Its low SiO2 content (e.g.53.41%) and high alkalinity (Na2O + K2O = 7.9%) suggestthat the rock is a monzodiorite.

Zircons extracted from this rock are distinct transparentfragmental zircons. Some grains have internal growth zon-ing, as shown by ESEM charge-contrast (e.g. Fig. 9: grain98), whereas others are featureless, with rounded externalmorphologies (e.g. Fig. 9: grains 71 and 80). Both zircontypes have no overgrowth rims, whereas some show relicexternal morphologies of primary igneous zircons with irreg-ular to fragmented edges. There are others with oval externalmorphologies comprising resorbed edges (Fig. 6: grain 80).

7.4. Quartzo-feldspathic granulite-facies paragneiss

The quartzo-feldspathic granulite-facies paragneiss(Fig. 4G) comprises a garnet, sillimanite and orthopyrox-ene-bearing leucocratic rock in the southern tectonic corri-dor sub-terrane in the A-BD (Fig. 3). Petrographic studysuggests that its protolith was a volcano-sedimentary rock.It overlies, or is juxtaposed against, granulite-facies maficorthogneiss and granitoid gneisses, and together they formpart of older basement rocks in the study area. The ortho-pyroxene–quartzo-feldspathic granulite hosts small-scale,locally mylonitic siliceous and ferruginous shear bandsand irregular fracture planes with sulphide rich zones (seeinset in Fig. 4G) bounded by NNW-trending dextral shearzones (Fig. 6). They comprise zones of Mn (400–850 ppm),Ni (85–120 ppm), Cu (80–210 ppm) and Zn (80–210 ppm)-bearing shear zones. Petrographical evidence suggests themineralisation to be syn- to late-orogenic.

The SHRIMP study of sample 21337 was undertaken toconstrain the time of volcano-sedimentary accumulationand/or magmatic emplacement of the precursor rock, aswell as subsequent metamorphic and deformation eventsthat affected this lithotectonic unit. Although most zirconsextracted from this rock are colourless and have well-rounded morphologies, they nonetheless show the follow-ing internal morphologic variations: (1) elongate andprismatic, up to about 250 lm long (Fig. 9: grain 11) andstubby-prismatic, up to 450 lm long crystals with relicinternal growth zoning, which are probably original

Fig. 8. Charge-contrast environmental scanning-electron microscopy images illustrating ages from dating of least-modified igneous zircon morphologies(A and B); moderately-modified (D) and featureless zircons (C) from sample 67-04, a granulite-facies mafic orthogneiss; subhedral and elongate zircons (Eand F); spherical and strongly overgrown metamorphic zircons (G and H), possibly illustrating metasomatic effects from sample 21349, a metasomatisedgranulite-facies mafic orthogneiss.


igneous crystals; and (2) multifaceted grains with well-rounded external morphology (Fig. 9: grain 33).

The zircon morphologies vary from those clearly show-ing an igneous source (e.g. core of grains 33: Fig. 9) tothose with ovoid morphology and well-rounded external

morphology showing prior recycling by either metamor-phic or mechanical abrasion before their deposition. Thezircon population shown by sample 21337 is common involcano-sedimentary rocks which contain detrital grainsfrom a wide range of source rocks.


7.5. Biotite–granitoid hornfels

Biotite–granitoid hornfels (Fig. 4I) occupies the lowerpart of the lithotectonic units underlying the tectonic corri-dor sub-terrane in the A-BD (Fig. 3). It crops out as slicesof weathered brownish-grey granitoid gneisses juxtaposedagainst pyroxene-bearing quartz-monzonitic syenite andamphibolite-facies alkali-feldspar granitoids. Their tectonicjuxtaposition obscures their prior relationships. Approxi-mately 70% of the zircons in biotite–granitoid hornfelsare sub-rounded to oval, occurring as xenocrysts in feld-spars and quartz. The remaining 20–30% subhedral toangular zircons and titanite overgrow secondary brownbiotite located along granoblastic quartz and feldspar grainboundaries. In addition, titanites and zircons show radia-tion-damaged corona rims within biotite. Transmittedand reflected light microscopic studies on zircons showtheir elongate, multifaceted external morphology and inter-nal structure. Further examination by charge contrastESEM techniques reveals detailed internal morphologicalfeatures defining zircons with either external or relic inter-nal igneous features, or both, and up to 120 lm long (e.g.grain 25: Fig. 9), and completely recrystallised featurelesszircons up to 140 lm long (e.g. grains 11 and 28 in Fig. 9).

In addition, there are zircons whose internal structuresindicate partial recrystallisation (e.g. grain 28: Fig. 9).Some zircons in group 2 are extremely recrystallised, form-ing pseudo-rims (e.g. grains 11 and 28: Fig. 9). The U andTh contents of these rims are neither consistent with newzircon growth nor overgrowth during high-grade metamor-phism. The similar ages of the cores and rims (e.g. grain 28:Fig. 9) reflect high-grade metamorphic conditions at aboutthis time. Importantly, zircon does not commonly recordretrograde metamorphic processes (Kroner et al., 1997).

8. U–Pb SHRIMP results and interpretation

8.1. Granulite-facies mafic orthogneiss (67-04)

Table 1 presents corrected U–Th–Pb compositions andradiogenic isotope ratios processed from raw ion-micro-probe SHRIMP analyses of 36 zircon grain-spots. The Bro-ken Hill 204Pb composition is used to correct common-Pbeffects (Wiedenbeck, 1995). The Pb/U calibration errorduring the two sessions when the data were collected was±1.28% on 15 standards and ±1.34% on 14 standards:the more conservative value has been used in the calcula-tion of the ages reported here. Fig. 10A shows the dataon a U–Pb concordia diagram.

The data produce a 207Pb/206Pb age range of 2725–2390 Ma. The spread of ages is consistent with the morpho-logical classification, showing a general decrease of207Pb/206Pb ages with increase in the degree of crystalannealing. For example, the irregular, but most igneous-like faintly-zoned zircon grains give older ages (e.g.2678 ± 8 Ma in grain 3: Fig. 8), whereas their frag-mented equivalents preserve relatively younger ages of

between 2661 ± 6 Ma and 2566 ± 8 Ma (e.g. grain 6:Fig. 8). Similarly, relatively younger ages, ranging between2621 ± 8 Ma and 2542 ± 8 Ma, are recorded in near-prismatic zircons lacking internal igneous zoning. An olderconcordant 207Pb/206Pb age of 2725 ± 12 Ma is preservedin the core of grain 34-1, which is possibly a xenocryst.Three relatively unmodified igneous zircons form a popula-tion with a concordant statistical 207Pb/206Pb age of2676 ± 6 Ma (Fig. 10A), defining a possible minimum ageof protolith magmatism. The 207Pb/206Pb dates, whichrange between 2465 ± 8 Ma and 2390 ± 10 Ma, are con-cordant to disconcordant ages measured from fragmentedand featureless zircons containing concentrations of 64–776 ppm U. The 207Pb/206Pb age of 2390 ± 10 Ma, specifi-cally, is preserved in a featureless and fragmented zirconinterpreted to be of metamorphic affinity (Fig. 10A andTable 1). Therefore, the 2390 ± 10 Ma age (ca. 2400 Ma)could be interpreted as a maximum age of metamorphism(i.e. partial melting) of the precursor granulite-facies maficorthogneiss, if Pb-loss in this sample is related to the pres-ent weathering processes. However, it is equally possiblethat the ca. 2400 Ma represents the minimum age of proto-lithic magmatism which was reset during the ca. 800 Magranulite-facies metamorphic event. This age is consistentwith the age of ca. 2400 Ma preserved by the reset, but par-tially healed zircons (e.g. grain 30, Fig. 8).

In summary, the 2725 ± 12 Ma age is interpreted to beof a xenocrystic zircon which was recycled during the ca.2676 Ma magmatic event responsible for the formation ofthe dioritic–granodioritic rocks. From the data above,these rocks appear to have been subsequently metamor-phosed into granulite-facies mafic orthogneiss between ca.2465 Ma and 2390 Ma. Lead-loss in this sample could berelated to either the present weathering processes or tointrusion of the ca. 800 Ma alkali-feldspar granitoid atamphibolite to granulite-facies conditions. The latter couldexplain why the apparent age of granulite-facies metamor-phism recorded here is slightly younger than the ca.2500 Ma recorded by Paquette et al. (2004) for similarrocks elsewhere in Andriamena.

8.2. Metasomatised granulite-facies mafic orthogneiss

The SHRIMP analyses and data from sample 21349 arepresented in Table 2 and are plotted on a U–Pb concordiadiagram (Fig. 10B). The analyses indicate relatively lowcommon-Pb corrections, high U contents and low Th/Uratios in the least-modified, subhedral to elongate zircons(e.g. grain-spot 51-1: U = 352 ppb and Th/U = 0.436);grain 114: U = 335 ppb and Th/U = 0.405: Table 2), whichhave textures more typical of igneous grains (Fig. 8E). Incontrast, the near-spherical multifaceted zircons recordlow U and higher Th/U ratios (e.g. metamorphic over-growth in grain-spot 58-1: U = 20 ppb and Th/U = 0.666:Table 2, Fig. 8F).

In the concordia plots (Fig. 10B), the data define a largespread along, above and below the concordia. Analyses

Fig. 9. Charge-contrast environmental scanning-electron microscopy images of zircons extracted from an amphibolite-facies alkali-feldspar granitoid,sample 21380. I. Showing internal growth zoning and external morphological features are neither typical of igneous nor metamorphic zircons, but could berelated to crystallisation at granulite-facies. J and K. Featureless, oval to near-spherical homogenous metamorphic zircons typical of high-grademetamorphism. L. Elongate prismatic zircons of igneous origin M. Zoned inherited core with featureless overgrowth and sub-rounded externalmorphologies typical of metamorphic zircons. N. Well-rounded oval-shaped detrital zircons extracted from a quartzo-feldspathic granulite-faciesparagneiss. N. A biotite hornfels with zircons showing relic igneous features, external overgrowth morphologies (O) and recrystallised features (P).


with P5% discordance are not considered in the followingdiscussion, as their isotopic systems have been disturbedby complex Pb-loss, possibly related to either the ca.800 Ma metamorphic resetting and/or recent weathering

processes. The >95% and <105% concordant data for sam-ple 21349 show that most zircons with relic igneous featuresgive older ages, with the 207Pb/206Pb age of 2465 ± 12 Maconsidered to be the minimum age for magmatism that

Table 1Results of U–Th–Pb SHRIMP analyses of part of the mafic-granulite (i.e.sample 67-04, Mount UWA 98-19B) from Ankisatra-Besakay District, Andriamena Belt, central-northern Madagascar

Grain-spot U (ppm) Th (ppm) ThU % comm

207Pb206Pba

208Pb206Pb

206Pb238U

207Pb235U

208Pb232Th

% conc. Age (Ma)207Pb206Pb

Age (Ma)206Pb238U

2-1 458 757 1.654 0.008 0.1783 ± 6 0.4515 ± 16 0.4570 ± 57 11.236 ± 150 0.1248 ± 17 92 2637 ± 6 2426 ± 252-2 371 527 1.422 0.036 0.1787 ± 6 0.3874 ± 15 0.4828 ± 61 11.892 ± 160 0.1316 ± 18 96 2640 ± 6 2539 ± 263-1 1237 142 0.115 0.003 0.1797 ± 3 0.0310 ± 3 0.4853 ± 58 12.021 ± 149 0.1307 ± 20 96 2650 ± 3 2550 ± 253-2 803 275 0.342 0.000 0.1827 ± 4 0.0939 ± 5 0.4874 ± 59 12.279 ± 155 0.1338 ± 18 96 2678 ± 4 2559 ± 264-1 337 307 0.911 0.000 0.1580 ± 6 0.2502 ± 12 0.4291 ± 55 9.348 ± 129 0.1179 ± 17 95 2434 ± 7 2302 ± 255-1 847 118 0.140 0.032 0.1766 ± 5 0.0375 ± 4 0.4922 ± 60 11.986 ± 154 0.1323 ± 23 98 2621 ± 4 2580 ± 266-1 417 113 0.272 0.000 0.1492 ± 6 0.0785 ± 6 0.3426 ± 43 7.048 ± 96 0.0988 ± 15 81 2337 ± 7 1899 ± 216-2 366 335 0.914 0.000 0.1802 ± 6 0.2440 ± 11 0.4920 ± 62 12.223 ± 165 0.1313 ± 18 97 2654 ± 5 2579 ± 276-3 223 72 0.324 0.144 0.1779 ± 10 0.0879 ± 15 0.4915 ± 65 12.058 ± 179 0.1334 ± 30 98 2634 ± 9 2577 ± 287-1 776 202 0.261 0.034 0.1539 ± 4 0.0722 ± 5 0.4200 ± 51 8.913 ± 114 0.1163 ± 17 95 2390 ± 5 2260 ± 237-2 636 306 0.481 0.006 0.1473 ± 5 0.1322 ± 9 0.3344 ± 41 6.792 ± 90 0.0918 ± 13 80 2315 ± 6 1860 ± 207-3 871 568 0.653 0.022 0.1557 ± 4 0.1801 ± 7 0.4066 ± 49 8.728 ± 111 0.1123 ± 15 91 2409 ± 4 2199 ± 238-1 280 502 1.793 0.022 0.1820 ± 7 0.4917 ± 20 0.4978 ± 64 12.492 ± 175 0.1365 ± 19 98 2671 ± 7 2604 ± 288-2 1121 225 0.201 0.022 0.1609 ± 3 0.0621 ± 4 0.4488 ± 54 9.954 ± 124 0.1386 ± 19 97 2465 ± 4 2390 ± 249-1 540 205 0.379 0.000 0.1684 ± 6 0.1067 ± 7 0.3486 ± 43 8.095 ± 107 0.0982 ± 14 76 2542 ± 6 1928 ± 219-2 772 58 0.075 0.026 0.1684 ± 4 0.0186 ± 4 0.4149 ± 51 9.635 ± 123 0.1026 ± 24 88 2542 ± 4 2237 ± 2310-1 1228 600 0.488 0.002 0.1809 ± 3 0.1381 ± 5 0.4988 ± 60 12.444 ± 154 0.1411 ± 18 98 2661 ± 3 2609 ± 2610-2 342 37 0.107 0.058 0.1605 ± 7 0.0249 ± 8 0.4103 ± 52 9.078 ± 127 0.0957 ± 34 90 2461 ± 8 2216 ± 2414-1 196 152 0.773 2.435 0.1461 ± 25 0.2365 ± 56 0.3016 ± 41 6.074 ± 141 0.0923 ± 26 74 2300 ± # 1699 ± 2015-1 637 423 0.665 0.021 0.1769 ± 5 0.1889 ± 8 0.4530 ± 56 11.049 ± 143 0.1288 ± 17 92 2624 ± 5 2409 ± 2516-1 343 182 0.531 0.021 0.1652 ± 8 0.1436 ± 12 0.3681 ± 47 8.387 ± 118 0.0996 ± 16 81 2510 ± 8 2021 ± 2218-1 1578 147 0.093 0.102 0.1720 ± 3 0.0257 ± 4 0.4468 ± 53 10.598 ± 131 0.1234 ± 24 92 2577 ± 3 2381 ± 2419-1 707 223 0.315 0.118 0.1710 ± 6 0.0827 ± 8 0.3996 ± 49 9.424 ± 123 0.1048 ± 17 84 2568 ± 5 2167 ± 2321-1 405 680 1.679 0.013 0.1786 ± 7 0.4650 ± 18 0.4683 ± 59 11.533 ± 158 0.1297 ± 18 94 2640 ± 6 2476 ± 2630-1 64 95 1.485 0.000 0.1447 ± 15 0.4001 ± 43 0.3271 ± 55 6.526 ± 136 0.0881 ± 19 80 2284 ± # 1824 ± 2730-2 225 80 0.357 0.000 0.1589 ± 8 0.0980 ± 9 0.4171 ± 55 9.136 ± 134 0.1146 ± 20 92 2444 ± 8 2247 ± 2534-1 253 179 0.710 0.045 0.1880 ± 7 0.1930 ± 12 0.5285 ± 72 13.698 ± 200 0.1436 ± 23 100 2725 ± 6 2735 ± 3035-1 234 84 0.360 0.000 0.1482 ± 7 0.0979 ± 9 0.3034 ± 41 6.197 ± 93 0.0824 ± 14 73 2325 ± 8 1708 ± 2036-1 872 59 0.067 0.015 0.1563 ± 4 0.0168 ± 3 0.4203 ± 54 9.057 ± 120 0.1052 ± 23 94 2416 ± 4 2262 ± 2437-1 1623 98 0.060 0.000 0.1812 ± 3 0.0164 ± 1 0.4972 ± 63 12.421 ± 160 0.1355 ± 20 98 2664 ± 2 2602 ± 2737-2 867 258 0.298 0.043 0.1708 ± 4 0.0847 ± 5 0.4037 ± 52 9.509 ± 126 0.1147 ± 16 85 2566 ± 4 2186 ± 2438-1 1167 131 0.112 0.029 0.1825 ± 3 0.0309 ± 2 0.4910 ± 62 12.351 ± 161 0.1350 ± 21 96 2675 ± 3 2575 ± 2731-1 320 109 0.341 0.111 0.1806 ± 7 0.0905 ± 9 0.4652 ± 62 11.585 ± 165 0.1236 ± 21 93 2659 ± 6 2462 ± 2732-1 231 235 1.020 0.130 0.1687 ± 9 0.2850 ± 19 0.4200 ± 58 9.768 ± 149 0.1173 ± 19 89 2545 ± 9 2260 ± 2639-1 2114 2277 1.077 0.010 0.1728 ± 2 0.2964 ± 5 0.4492 ± 57 10.701 ± 138 0.1236 ± 16 93 2585 ± 2 2391 ± 2540-1 589 143 0.242 0.086 0.1757 ± 5 0.0659 ± 6 0.4411 ± 57 10.683 ± 145 0.1199 ± 20 90 2612 ± 5 2355 ± 2641-1 686 83 0.121 0.279 0.1473 ± 6 0.0341 ± 9 0.3256 ± 42 6.610 ± 92 0.0913 ± 28 79 2314 ± 7 1817 ± 20

Samples listed in numerical order.% comm. = common Pb.% conc. = measure of concordance (Pb-loss and U-loss).

a Broken Hill 204Pb common-Pb corrected isotopic composition.

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Fig. 10. Concordia plot of SHRIMP data for (A) sample 67-04 illustrating the maximum age of metamorphism (i.e. 2400 and 2300 Ma) from grain-spots30-1 (core) and 30-2 (rim) elaborated by a dashed line, suggesting metamorphic resetting at ca. 800 Ma: (B) sample 21349 illustrating possible majorgeological events. Darker error boxes represent weakly-modified igneous zircons, whereas hatched boxes show metamorphic zircons. Ages are quoted at 2-sigma for single analyses and 1-sigma for average ages.


produced the protolith. The 207Pb/206Pb ages older than ca.2465 Ma (n = 3) are coincidentally preserved in irregular tosub-rounded zircons with possible relic inherited cores.Rims to these pseudo-cores give relatively younger ages,consistent with ages measured from featureless, low-uranium metamorphic zircons (Table 2, and from grain60, Fig. 8). The youngest zircon growth and overgrowthconsisting of low-U and low-Th featureless zircon has206Pb/238U ages of between 625 ± 8 Ma and 879 ± 27 Ma,with a mean age of 785 ± 7 Ma (Table 2). The ca. 785 Maevent either caused metamorphic resetting of the older zir-cons, in which a considerable amount of Pb- and U-lossoccurred, or new zircon overgrowths formed at this time.The ca. 785 ± 2 Ma age recorded in featureless new andovergrowth zircons is interpreted to be due to granulitefacies metamorphism (Fig. 10).

In summary, U–Th–Pb isotopic ratios from zirconsextracted from sample 21349 suggest that: (1) the minimum

age of magmatism was at 2465 ± 6 Ma, defined from a con-cordant, least modified zircon, that incorporated olderxenocrysts which indicate a minimum protolith age of2475 ± 12 Ma (grain 51, Fig. 8); (2) metamorphic resettingcaused a spread in ages between 2465 ± 12 Ma and ca.1900 Ma (Fig. 10B), suggesting intense metamorphism atca. 2465–2390 and possibly to 2262 ± 3 Ma, as measuredfrom featureless, rounded to ovoid zircons interpreted asmetamorphic grains (Fig. 10B); and (3) the ca. 785 Maage of peak granulite facies metamorphism.

8.3. Amphibolite-facies alkali-feldspar granitoid

The U–Th–Pb SHRIMP results for sample 21380 aresummarised in Table 3 and plotted on a U–Pb concordiadiagram in Fig. 11A. Despite morphological differences,all zircons contain consistently low U and Th concentra-tions and high Th/U ratios.. The high common-Pb

Table 2Results of SHRIMP U–Th–Pb analyses of part of the metasomatised mafic granulite (sample 21349, mount UWA98-19B) from Ankisatra-Besakay District, Andriamena Belt, central-northernMadagascar


207Pb206Pba

208Pb206Pb

206Pb238U

207Pb235U

208Pb232Th

% conc. Age (Ma)207Pb206Pb

Age (Ma)206Pb238U

1-1 575 213 0.371 0.014 0.1504 ± 5 0.1044 ± 8 0.3706 ± 48 7.685 ± 107 0.1042 ± 16 86 2350 ± 6 2032 ± 231-2 316 133 0.422 0.000 0.1591 ± 6 0.1152 ± 7 0.4884 ± 66 10.713 ± 154 0.1334 ± 21 105 2446 ± 6 2564 ± 293-1 2033 221 0.109 0.010 0.1522 ± 2 0.0310 ± 2 0.3901 ± 49 8.184 ± 106 0.1115 ± 16 90 2370 ± 3 2123 ± 234-1 454 148 0.327 0.024 0.1622 ± 5 0.0896 ± 6 0.5452 ± 72 12.196 ± 169 0.1495 ± 23 113 2479 ± 5 2805 ± 305-1 658 228 0.347 0.000 0.1549 ± 4 0.0973 ± 5 0.3811 ± 50 8.138 ± 111 0.1069 ± 15 87 2401 ± 5 2081 ± 236-1 653 133 0.204 0.023 0.1450 ± 5 0.0565 ± 6 0.3213 ± 42 6.423 ± 89 0.0891 ± 15 79 2288 ± 6 1796 ± 207-1 1583 290 0.183 0.007 0.1514 ± 3 0.0500 ± 3 0.4549 ± 58 9.493 ± 124 0.1240 ± 17 102 2361 ± 3 2417 ± 268-1 352 132 0.376 0.028 0.1469 ± 6 0.1091 ± 10 0.3595 ± 48 7.280 ± 106 0.1043 ± 17 86 2310 ± 8 1980 ± 239-1 517 31 0.060 0.221 0.0616 ± 9 0.0167 ± 18 0.1018 ± 14 0.865 ± 18 0.0285 ± 30 95 661 ± 32 625 ± 8

10-1 565 44 0.077 0.099 0.1057 ± 6 0.0204 ± 8 0.2328 ± 31 3.393 ± 51 0.0613 ± 26 78 1727 ± 10 1349 ± 1611-1 421 60 0.143 0.002 0.1262 ± 6 0.0442 ± 8 0.3404 ± 45 5.924 ± 86 0.1052 ± 24 92 2046 ± 8 1889 ± 2212-1 211 105 0.498 0.032 0.1364 ± 11 0.1383 ± 21 0.2845 ± 40 5.353 ± 91 0.0791 ± 17 74 2182 ± 14 1614 ± 2013-1 300 105 0.350 0.000 0.1475 ± 6 0.0984 ± 8 0.4052 ± 55 8.244 ± 121 0.1139 ± 19 95 2318 ± 7 2193 ± 2514-1 288 131 0.455 0.000 0.1176 ± 8 0.1334 ± 14 0.2213 ± 31 3.590 ± 59 0.0648 ± 12 67 1921 ± 12 1289 ± 1636-1 404 147 0.363 0.013 0.1642 ± 5 0.0906 ± 7 0.6386 ± 80 14.462 ± 192 0.1593 ± 24 127 2500 ± 5 3184 ± 3136-2 2102 375 0.178 0.006 0.1428 ± 2 0.0519 ± 3 0.4194 ± 50 8.258 ± 101 0.1219 ± 16 100 2262 ± 3 2258 ± 2336-3 413 157 0.381 0.002 0.1578 ± 6 0.1049 ± 9 0.4388 ± 55 9.548 ± 130 0.1209 ± 19 96 2432 ± 7 2345 ± 2553-1 73 41 0.568 1.850 0.0693 ± 70 0.1742 ± 163 0.1326 ± 25 1.267 ± 134 0.0407 ± 39 88 907 ± 210 803 ± 14

52-1 70 24 0.342 0.201 0.1609 ± 16 0.0870 ± 28 0.6712 ± 107 14.893 ± 295 0.1705 ± 64 134 2465 ± 17 3310 ± 4151-1 352 154 0.436 0.000 0.1618 ± 6 0.1240 ± 8 0.4535 ± 58 10.116 ± 139 0.1288 ± 19 97 2475 ± 6 2411 ± 2651-2 730 590 0.808 0.023 0.1577 ± 4 0.2221 ± 9 0.4616 ± 56 10.039 ± 129 0.1268 ± 17 101 2432 ± 5 2447 ± 2553-2 492 188 0.381 0.036 0.1528 ± 6 0.1028 ± 9 0.4209 ± 53 8.870 ± 122 0.1135 ± 18 95 2378 ± 7 2265 ± 2455-1 61 20 0.330 0.090 0.1562 ± 21 0.0909 ± 40 0.4978 ± 85 10.719 ± 247 0.1373 ± 66 108 2415 ± 23 2604 ± 3656-1 39 29 0.727 0.782 0.0642 ± 104 0.2166 ± 244 0.1262 ± 32 1.116 ± 187 0.0376 ± 44 103 747 ± 348 766 ± 18

57-1 22 18 0.838 0.892 0.0595 ± 185 0.2425 ± 432 0.1386 ± 48 1.136 ± 360 0.0401 ± 73 143 584 ± 561 837 ± 27

58-1 20 14 0.666 0.229 0.0779 ± 165 0.1995 ± 381 0.1461 ± 48 1.570 ± 342 0.0438 ± 85 77 1145 ± 430 879 ± 2758-2 519 255 0.491 0.038 0.1579 ± 5 0.1369 ± 9 0.4559 ± 57 9.925 ± 133 0.1271 ± 18 100 2433 ± 6 2421 ± 2558-3 243 53 0.217 0.102 0.1439 ± 11 0.0666 ± 18 0.3449 ± 46 6.845 ± 111 0.1060 ± 33 84 2275 ± 13 1910 ± 2260-1 1139 135 0.118 0.004 0.1552 ± 4 0.0369 ± 4 0.4383 ± 53 9.377 ± 118 0.1368 ± 22 97 2404 ± 4 2343 ± 24101-1 277 129 0.466 0.083 0.1629 ± 9 0.1291 ± 15 0.3921 ± 52 8.805 ± 132 0.1087 ± 20 86 2486 ± 9 2133 ± 24102-1 901 164 0.182 0.023 0.1085 ± 5 0.0526 ± 7 0.2121 ± 26 3.173 ± 43 0.0615 ± 11 70 1774 ± 8 1240 ± 14106-1 434 262 0.603 0.040 0.1609 ± 6 0.1649 ± 11 0.4710 ± 59 10.449 ± 142 0.1288 ± 19 101 2465 ± 6 2488 ± 26114-1 335 136 0.405 0.098 0.1613 ± 7 0.1065 ± 11 0.4865 ± 63 10.821 ± 152 0.1278 ± 22 103 2470 ± 7 2555 ± 27116-1 817 181 0.222 0.042 0.1462 ± 4 0.0612 ± 5 0.3692 ± 48 7.445 ± 101 0.1019 ± 16 88 2303 ± 5 2026 ± 22116-2 820 84 0.103 0.008 0.1140 ± 4 0.0305 ± 4 0.3064 ± 40 4.815 ± 66 0.0908 ± 18 92 1864 ± 6 1723 ± 20116-3 743 134 0.180 0.063 0.1155 ± 5 0.0496 ± 6 0.2652 ± 34 4.222 ± 59 0.0730 ± 13 80 1888 ± 7 1516 ± 18117-1 396 132 0.335 0.032 0.1603 ± 6 0.0921 ± 8 0.4812 ± 64 10.632 ± 151 0.1324 ± 22 103 2458 ± 6 2532 ± 28118-1 34 7 0.199 0.132 0.1466 ± 27 0.0528 ± 46 0.4107 ± 84 8.301 ± 241 0.1088 ± 100 96 2306 ± 31 2218 ± 39119-1 330 211 0.640 0.088 0.1500 ± 7 0.1770 ± 13 0.3721 ± 50 7.696 ± 114 0.1029 ± 16 87 2346 ± 8 2039 ± 24

Samples listed in numerical order.% comm. = common Pb.% conc. = measure of concordance (Pb-loss and U-loss).Errors reported at 1-sigma in the table are quoted at 2-sigma in the text (i.e. 95% confidence).Normal text (not bold) for 207Pb/206Pb age for zircons with negligible common-Pb.


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Table 3Results of ion-microprobe U–Th–Pb SHRIMP analyses of part of the pyroxene alkali-feldspar granitoid (sample 21380, mount UWA98-19A) from Ankisatra-Besakay District, Andriamena Belt,central-northern Madagascar


a 204Pba 206Pb

a 207Pba 206Pb

208Pba 206Pb

206Pb238U

207Pb235U

208Pb232Th

%conc. Age (Ma)207Pb206Pb

Age (Ma)206Pb238U

01-1 13 18 1.409 0.000 0.0715 ± 34 0.4412 ± 160 0.1287 ± 43 1.269 ± 78 0.0403 ± 22 80 972 ± 98 781 ± 25

02-1 36 97 2.734 2.660 0.0475 ± 87 0.8240 ± 267 0.1201 ± 29 0.787 ± 148 0.0362 ± 15 933 78 ± 383 731 ± 17

03-1 33 63 1.873 1.548 0.0582 ± 100 0.5494 ± 267 0.1260 ± 32 1.011 ± 178 0.0370 ± 21 142 538 ± 382 765 ± 18

04-1 30 54 1.799 2.243 0.0515 ± 92 0.5201 ± 251 0.1285 ± 33 0.913 ± 168 0.0371 ± 21 295 264 ± 366 779 ± 19

05-1 29 88 3.009 1.392 0.0566 ± 87 0.8736 ± 275 0.1345 ± 34 1.050 ± 167 0.0390 ± 17 171 476 ± 345 814 ± 20

06-1 38 72 1.885 0.000 0.0664 ± 18 0.5571 ± 105 0.1391 ± 31 1.274 ± 48 0.0411 ± 13 102 819 ± 58 840 ± 17

07-1 30 43 1.435 1.519 0.0525 ± 117 0.4197 ± 301 0.1271 ± 35 0.921 ± 210 0.0372 ± 29 249 309 ± 442 771 ± 20

08-1 30 91 3.031 0.670 0.0599 ± 109 0.9162 ± 327 0.1365 ± 36 1.128 ± 211 0.0413 ± 19 137 601 ± 401 825 ± 21

71-1 21 67 3.183 0.000 0.0741 ± 31 0.9773 ± 249 0.1347 ± 41 1.376 ± 75 0.0414 ± 18 78 1045 ± 85 814 ± 23

73-1 10 22 2.184 2.585 0.0427 ± 395 0.6255 ± 1014 0.1354 ± 81 0.796 ± 745 0.0388 ± 68 0 0 ± 130 819 ± 46

74-1 19 61 3.178 0.341 0.0644 ± 160 0.9408 ± 479 0.1399 ± 50 1.243 ± 317 0.0414 ± 27 112 755 ± 543 844 ± 28

75-1 8 14 1.805 0.000 0.0766 ± 51 0.5715 ± 275 0.1331 ± 60 1.405 ± 120 0.0422 ± 30 73 1110 ± 135 806 ± 34

80-1 18 37 2.097 0.000 0.0681 ± 32 0.6378 ± 197 0.1321 ± 42 1.241 ± 74 0.0402 ± 19 92 872 ± 97 800 ± 24

82-1 11 28 2.432 4.407 0.0455 ± 327 0.6004 ± 845 0.1373 ± 72 0.861 ± 626 0.0339 ± 52 0 0 ± 106 829 ± 41

82-2 14 33 2.360 3.701 0.0454 ± 335 0.7240 ± 882 0.1235 ± 64 0.772 ± 577 0.0379 ± 51 0 0 ± 103 751 ± 37

83-1 9 18 1.939 0.000 0.0774 ± 48 0.5789 ± 259 0.1420 ± 60 1.515 ± 121 0.0424 ± 28 76 1131 ± 124 856 ± 34

84-1 20 46 2.295 9.574 0.0134 ± 280 0.6050 ± 725 0.1252 ± 55 0.232 ± 484 0.0330 ± 43 0 0 ± 70 760 ± 31

85-1 8 16 1.896 4.741 0.0328 ± 586 0.5122 ± 1466 0.1261 ± 101 0.571 ± 1026 0.0341 ± 102 0 0 ± 143 766 ± 58

86-1 25 31 1.259 2.172 0.0494 ± 183 0.3478 ± 455 0.1283 ± 46 0.874 ± 328 0.0354 ± 48 467 167 ± 691 778 ± 26

87-1 14 35 2.420 0.000 0.0679 ± 36 0.7287 ± 245 0.1408 ± 50 1.318 ± 89 0.0424 ± 23 98 866 ± 110 849 ± 28

88-1 24 83 3.455 1.744 0.0666 ± 146 1.0555 ± 460 0.1300 ± 44 1.194 ± 270 0.0397 ± 23 95 825 ± 470 788 ± 25

90-1 17 25 1.464 0.000 0.0715 ± 34 0.4164 ± 153 0.1396 ± 47 1.375 ± 84 0.0397 ± 21 87 970 ± 97 842 ± 26

91-1 19 40 2.118 3.974 0.0357 ± 175 0.5749 ± 472 0.1266 ± 47 0.624 ± 310 0.0344 ± 32 0 0 ± 96 769 ± 27

94-1 9 16 1.700 1.254 0.0666 ± 457 0.4945 ± 1138 0.1356 ± 91 1.245 ± 869 0.0394 ± 95 99 825 ± 1012 820 ± 51

95-1 14 20 1.370 4.241 0.0386 ± 290 0.3662 ± 721 0.1251 ± 61 0.666 ± 506 0.0334 ± 68 0 0 ± 105 760 ± 35

97-1 9 17 1.990 0.000 0.0655 ± 46 0.6248 ± 286 0.1291 ± 58 1.167 ± 102 0.0405 ± 28 99 792 ± 147 783 ± 33

98a-1 36 130 3.658 0.000 0.0699 ± 23 1.0848 ± 207 0.1289 ± 33 1.243 ± 55 0.0382 ± 13 84 926 ± 68 781 ± 19

99-1 20 68 3.336 0.000 0.0719 ± 31 0.9643 ± 247 0.1362 ± 42 1.351 ± 75 0.0394 ± 17 84 984 ± 87 823 ± 24

100-1 25 95 3.798 4.381 0.0310 ± 213 1.0948 ± 633 0.1245 ± 47 0.532 ± 370 0.0359 ± 25 0 0 ± 86 757 ± 27

109-1 23 75 3.229 1.566 0.0599 ± 84 0.9364 ± 288 0.1331 ± 36 1.099 ± 161 0.0386 ± 17 134 600 ± 307 806 ± 21

110-1 19 61 3.276 2.010 0.0493 ± 161 0.9556 ± 477 0.1352 ± 46 0.920 ± 306 0.0394 ± 25 498 164 ± 626 817 ± 26

111-1 24 69 2.856 0.000 0.0643 ± 24 0.8503 ± 187 0.1347 ± 35 1.194 ± 57 0.0401 ± 15 108 753 ± 78 814 ± 20

112-1 21 27 1.311 0.000 0.0698 ± 26 0.4225 ± 121 0.1315 ± 37 1.265 ± 63 0.0424 ± 19 86 922 ± 77 796 ± 21

% comm. = common Pb.% conc. = measure of concordance compositions (Pb-loss and U-loss).Errors are reported at 1-sigma in the table.Ages are quoted at 2-sigma in the text (i.e. approximately 95% confidence).Bold% comm. refers to the 11 out of 33 zircons with high common Pb.


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Fig. 11. Concordia plot for ion-microprobe SHRIMP analyses of zircons extracted from (A) sample 21380, showing concordant age of peakmetamorphism and/or emplacement of an alkali-feldspar granitoid; (B) sample 21337 that suggest a range of possible sources of detrital zircons andmetamorphic resetting by the ca. 780 Ma granulite-facies metamorphic event and/or recent weathering. Ages are within 1-sigma error boxes.


contents measured from 12 out of 33 zircons extracted fromsample 21380 appear primary, and such data are not used inthe age calculation. All data cluster at approximately800 Ma, with the 21 analyses grains unaffected by high com-mon-Pb defining a statistically concordant 206Pb/238U ageof 807 ± 13 Ma (Fig. 11A). This age can be interpreted asa robust peak granulite-facies metamorphic age, recordingeither the emplacement of the alkali-feldspar granitoid orthe high-temperature low-pressure metamorphism of theprecursor older basement rock as suggested above.

The zircons extracted from this rock show growth fea-tures which are uncharacteristic of normal igneous zircons,possibly due to crystallisation and annealing in a high-grade amphibolite to granulite facies environment (seezoning in grain 98, Fig. 9B). In contrast, the internal fea-turelessness and oval to near-spherical external zircon mor-phologies (see grains 71 and 80, Fig. 9C, D) are a result ofhigh-grade metamorphic modification of original igneouszircons.

8.4. Quartzo-feldspathic granulite-facies paragneiss

SHRIMP analytical results from 37 zircon grainsextracted from sample 21337 (UWA mount 98-29D) aresummarised in Table 4. Most zircons contain moderate Uconcentrations (e.g. between 137 and 549 ppm) and lowTh/U ratios. In addition, these zircons contain negligiblecommon-Pb, and are mostly strongly discordant(Fig. 11B), possibly due to extensive Pb-loss during youn-ger metamorphic events and/or recent Pb loss.

The varied morphology of the grains, and interpreta-tion that they have been derived from a number of sourcerocks, suggest that their various protoliths have differentages. The SHRIMP data suggest that these lie betweenca. 2870 and 1750 Ma, as defined from 207Pb/206Pb ages.This time window includes the age of the late-Archaeangranulite (e.g. sample 67–04) and the metasomatisedgranulite-facies mafic orthogneiss (e.g. sample 21439),which may have been part of the source rocks. This

Table 4Results of SHRIMP U–Th–Pb analyses of zircons from quartzo-feldspathic granulite facies paragneiss (sample 21337) from Ankisatra-Besakay District,Andriamena Belt, central-northern Madagascar

Grain-spot U(ppm)

Th(ppm)

ThU %

comm

207Pba

206Pba

208Pba

206Pba

206Pba

238U

207Pba

235U

208Pba

232Th% conc.

207Pb206Pb

206Pb238U

1-2 140 103 0.737 0.231 0.1594 ± 13 0.1954 ± 27 0.3256 ± 49 7.159 ± 129 0.0863 ± 18 74 2450 ± 14 1817 ± 241-3 185 142 0.767 0.050 0.1583 ± 10 0.2125 ± 19 0.3348 ± 48 7.308 ± 120 0.0927 ± 17 76 2438 ± 11 1862 ± 232-2 286 131 0.459 0.055 0.0890 ± 11 0.1430 ± 24 0.1547 ± 22 1.898 ± 38 0.0482 ± 11 66 1404 ± 23 927 ± 12

4-1 338 77 0.229 0.000 0.0745 ± 7 0.0674 ± 11 0.1413 ± 21 1.451 ± 27 0.0416 ± 9 81 1054 ± 20 852 ± 12

5a-1 345 441 1.277 0.017 0.1817 ± 5 0.3162 ± 12 0.6543 ± 90 16.390 ± 237 0.1619 ± 24 122 2668 ± 5 3245 ± 355b-1 186 94 0.503 0.064 0.0750 ± 15 0.1511 ± 34 0.1462 ± 22 1.513 ± 39 0.0439 ± 12 82 1069 ± 39 880 ± 12

6-1 184 114 0.620 0.204 0.1438 ± 10 0.1666 ± 20 0.3774 ± 54 7.481 ± 126 0.1014 ± 20 91 2273 ± 12 2064 ± 257-1 218 247 1.135 0.139 0.1147 ± 10 0.3027 ± 25 0.2626 ± 38 4.155 ± 72 0.0700 ± 12 80 1876 ± 15 1503 ± 198-1 182 95 0.522 0.014 0.1012 ± 16 0.1545 ± 34 0.1601 ± 24 2.234 ± 51 0.0474 ± 13 58 1647 ± 29 957 ± 139-1 408 301 0.738 0.029 0.1076 ± 8 0.2059 ± 17 0.1850 ± 25 2.743 ± 45 0.0516 ± 9 62 1758 ± 13 1094 ± 1410-1 167 164 0.984 0.003 0.1522 ± 10 0.2689 ± 22 0.3384 ± 49 7.099 ± 119 0.0924 ± 16 79 2370 ± 11 1879 ± 2411-1 214 225 1.053 0.008 0.1339 ± 9 0.2562 ± 21 0.2883 ± 41 5.321 ± 88 0.0701 ± 12 76 2149 ± 12 1633 ± 2113a-1 147 101 0.684 0.005 0.2050 ± 13 0.2517 ± 24 0.3473 ± 51 9.818 ± 165 0.1278 ± 24 67 2867 ± 10 1922 ± 2513b-1 312 223 0.714 0.017 0.1849 ± 8 0.2207 ± 14 0.3109 ± 43 7.925 ± 119 0.0961 ± 15 65 2697 ± 7 1745 ± 2114-1 118 93 0.790 0.024 0.1260 ± 14 0.2289 ± 32 0.2649 ± 41 4.602 ± 93 0.0768 ± 17 74 2043 ± 20 1515 ± 2115-1 549 250 0.455 0.026 0.1777 ± 5 0.1314 ± 7 0.4012 ± 54 9.829 ± 139 0.1159 ± 17 83 2631 ± 5 2175 ± 2516-1 180 119 0.661 0.276 0.0780 ± 13 0.1948 ± 31 0.1566 ± 23 1.684 ± 39 0.0461 ± 10 82 1147 ± 32 938 ± 1317-1 137 117 0.850 0.048 0.1690 ± 11 0.2267 ± 21 0.4201 ± 63 9.791 ± 166 0.1121 ± 21 89 2548 ± 11 2261 ± 2818-2 181 147 0.815 0.138 0.1375 ± 9 0.2340 ± 20 0.3502 ± 51 6.641 ± 112 0.1006 ± 18 88 2196 ± 12 1936 ± 2418-3 399 20 0.051 0.049 0.1451 ± 6 0.0141 ± 6 0.3657 ± 50 7.315 ± 108 0.1004 ± 48 88 2289 ± 7 2009 ± 2419-1 171 129 0.757 0.207 0.1418 ± 13 0.2324 ± 28 0.2378 ± 35 4.649 ± 85 0.0730 ± 14 61 2249 ± 16 1375 ± 1820-1 407 106 0.260 0.037 0.1506 ± 6 0.0736 ± 9 0.3341 ± 46 6.938 ± 102 0.0946 ± 18 79 2353 ± 7 1858 ± 2221a-1 326 16 0.050 0.098 0.1274 ± 7 0.0127 ± 9 0.3084 ± 43 5.416 ± 84 0.0787 ± 57 84 2062 ± 10 1733 ± 2121a-2 149 93 0.622 0.058 0.1651 ± 11 0.1619 ± 21 0.3902 ± 58 8.882 ± 151 0.1016 ± 21 85 2509 ± 12 2124 ± 2721b-1 212 203 0.961 0.069 0.1594 ± 11 0.2689 ± 25 0.2822 ± 41 6.201 ± 104 0.0790 ± 14 65 2449 ± 12 1603 ± 2022a-1 173 148 0.860 0.066 0.1861 ± 12 0.2510 ± 25 0.3325 ± 49 8.531 ± 143 0.0971 ± 18 68 2708 ± 11 1850 ± 2323-1 213 132 0.620 0.000 0.1119 ± 8 0.1733 ± 17 0.1965 ± 28 3.031 ± 52 0.0550 ± 10 63 1830 ± 14 1157 ± 1524a-1 219 161 0.732 0.086 0.0784 ± 13 0.2458 ± 34 0.1394 ± 20 1.506 ± 36 0.0468 ± 10 73 1156 ± 33 841 ± 12

26-1 143 83 0.578 0.198 0.1492 ± 12 0.1491 ± 23 0.3437 ± 51 7.071 ± 126 0.0887 ± 20 81 2337 ± 14 1904 ± 2532-1 176 119 0.672 0.117 0.1241 ± 10 0.1518 ± 21 0.2915 ± 43 4.985 ± 88 0.0659 ± 14 82 2015 ± 15 1649 ± 2131-1 217 118 0.543 0.094 0.1293 ± 10 0.1515 ± 20 0.2200 ± 32 3.921 ± 68 0.0614 ± 12 61 2088 ± 14 1282 ± 1733-1 209 164 0.784 0.114 0.1442 ± 10 0.2229 ± 21 0.3096 ± 44 6.153 ± 103 0.0881 ± 16 76 2278 ± 12 1739 ± 2233-2 174 106 0.610 0.053 0.1587 ± 12 0.1816 ± 22 0.2931 ± 43 6.412 ± 111 0.0872 ± 17 68 2442 ± 13 1657 ± 2134-1 273 197 0.724 0.000 0.0785 ± 7 0.2172 ± 21 0.1299 ± 19 1.405 ± 25 0.0390 ± 7 68 1158 ± 18 787 ± 11

35-1 177 158 0.896 0.116 0.1706 ± 10 0.2559 ± 20 0.4340 ± 63 10.207 ± 166 0.1239 ± 21 91 2563 ± 10 2324 ± 2836-1 188 155 0.823 0.075 0.1470 ± 10 0.2246 ± 20 0.3122 ± 45 6.329 ± 105 0.0852 ± 15 76 2312 ± 11 1752 ± 2237-1 398 255 0.642 0.000 0.0863 ± 6 0.1959 ± 16 0.1393 ± 19 1.657 ± 27 0.0425 ± 7 63 1345 ± 14 841 ± 11

Data arranged in order of sample number.% comm = % common Pb.% conc. = measure of concordance (Pb-loss and U-loss).Errors are reported at 1-sigma in the table whereas ages quoted in the text are at 2-sigma (i.e. approximately 95% confidence).Normal text (not bold) for 207Pb/206Pb age for zircons with negligible common-Pb.

a 204Pb corrected isotopic compositions.


observation further suggests that sedimentation occurredin a foreland depository basin during or following thetectonic shortening of the basement rocks between 1700and 800 Ma (Fig. 11B). This event brought together theLate Archaean mafic and felsic granulite-facies gneissesbefore the ca. 780 Ma peak metamorphic event, whichcaused partial to near-complete Pb-loss in some of thesegrains.

Overloading of the crust during the tectonic shorteningand sedimentation may have ruptured it and triggeredthe volcanic activity. However, it is likely also that thispackage was a result of syn-rifting volcanism. The interpre-tation of SHRIMP data suggests that either of these eventsoccurred not earlier than ca. 1700 Ma (Fig. 8B).

8.5. Biotite–granitoid hornfels

The processed SHRIMP data for sample 21323, as plot-ted on a concordia diagram (Fig. 12A and B), are high-lighted by: (1) a population of 12 zircon grains thatdefines peak-granulite facies metamorphism at the lowerpart of the concordia; (2) the linear-spread of 20 other datapoints which defines a Pb-loss chord whose upper andlower intercepts with concordia are ages of magmatismand peak-granulite facies metamorphism, respectively(Table 5).

For the first population, two alternative 206Pb/238U sta-tistical ages of 778 ± 7 and 783 ± 8 Ma are deduced by sep-arating the data into two groups, one excluding the two

Fig. 12. Concordia plot of SHRIMP data from zircons extracted from sample 21323 illustrating: (A) the intercept of the discordia line with the upperConcordia, interpreted as the precursor magmatic age, and the intercept on the lower Concordia, interpreted as peak granulite-facies metamorphism; (B)expansion of Concordia plot near 800 Ma, as shown in Fig. 9. Peak metamorphic age possibly defined by relatively older zircons (open boxes) assumed tocontain inherited radiogenic isotope components from their igneous protoliths, as illustrated by their alignment with grains 9-1, 21-1 and 12-1.


older zircons and the other excluding the two youngest zir-cons with low U, respectively (Fig. 12B). For the firstgroup, the two older zircons are assumed to contain inher-ited radiogenic isotope components and, given the commonoccurrence of inheritance, the group excluding these zir-cons is considered the more likely to give the correct age(i.e. 778 ± 7 Ma for metamorphism).

The lower intercept of the chord with concordia mayalso be calculated (Ludwig, 1990). The best-fit line hasMean Squared Weighted Deviates (MSWD) of 12.6, sug-gesting an excessive scatter or higher degree of misfit fromthe regression line (Fig. 12A). The MSWD is much higherthan the 2.5 suggested cut-off MSWD value (Dickin, 1995),and is most likely due to geological rather than statisticalinfluences. However, the 2483 ± 20 Ma and 785 ± 7 Maages from the intercepts are nonetheless considered reliableages for magmatism and peak metamorphism, respectively,despite the high MSWD. The most likely reason for the

high MSWD is that analysed zircons suffered Pb-loss dueto present-day weathering processes. In addition, the timewindow for the magmatic event may have been slightlywider, but still within the same magmatic cycle, whose peakwas at 2483 ± 20 Ma. Similarly, the concordia interceptage for the metamorphic event is within error of the statis-tically determined age for this population at 778 ± 7 Ma(Table 6).

9. Mineral deposits in the A-BD

9.1. Mineralisation styles

From the several small-scale base metal and associatedprecious-metal deposits in the A-BD, only the followingare discussed in this section: (1) the vein-hosted BesakayPb–Ag, and Ankisatra Pb–Zn–Au deposits and leucosomeveins hosted Cu–Zn mineralisation in tectonically

Table 5Results of ion-microprobe U–Th–Pb SHRIMP analyses of zircons from biotite–granitoid hornfels (sample 21323) from Ankisatra-Besakay District,central-northern Madagascar

Grain-spot U(ppm)

Th(ppm)

ThU % comm

207Pba

206Pba

208Pba

206Pba

206Pba

238U

207Pba

235U

208Pba

235Th%conc.

207Pb206Pb

206Pb238U

1-1 86 111 1.293 0.350 0.0600 ± 22 0.3953 ± 63 0.1261 ± 16 1.042 ± 41 0.0386 ± 8 127 602 ± 78 766 ± 9

2-1 165 141 0.850 0.177 0.0637 ± 14 0.2538 ± 37 0.1276 ± 14 1.120 ± 28 0.0381 ± 7 106 730 ± 46 774 ± 8

3-1 248 127 0.512 0.003 0.1444 ± 7 0.1282 ± 12 0.2962 ± 29 5.898 ± 69 0.0741 ± 11 73 2281 ± 8 1673 ± 154-1 152 169 1.110 0.038 0.0644 ± 13 0.3366 ± 39 0.1291 ± 15 1.147 ± 28 0.0391 ± 7 104 755 ± 42 783 ± 8

6-1 442 355 0.802 0.019 0.1349 ± 5 0.2030 ± 9 0.2764 ± 26 5.141 ± 54 0.0699 ± 8 73 2162 ± 6 1573 ± 137-1 315 126 0.398 1.069 0.0628 ± 14 0.1182 ± 33 0.1277 ± 13 1.107 ± 29 0.0379 ± 11 110 703 ± 48 775 ± 7

8-1 531 265 0.499 0.055 0.1433 ± 5 0.1268 ± 7 0.3188 ± 30 6.298 ± 64 0.0811 ± 9 79 2267 ± 5 1784 ± 149-1 370 172 0.466 0.133 0.0784 ± 7 0.1310 ± 15 0.1410 ± 14 1.525 ± 21 0.0396 ± 6 74 1157 ± 17 850 ± 810-1 349 125 0.358 0.039 0.0648 ± 7 0.1068 ± 15 0.1302 ± 13 1.163 ± 18 0.0389 ± 7 103 767 ± 22 789 ± 7

11-1 440 209 0.474 0.434 0.0648 ± 8 0.1349 ± 20 0.1275 ± 12 1.139 ± 19 0.0363 ± 6 101 769 ± 27 773 ± 7

12-1 412 136 0.330 0.057 0.0996 ± 6 0.1071 ± 11 0.1714 ± 16 2.353 ± 28 0.0556 ± 8 63 1617 ± 11 1020 ± 913-1 382 122 0.318 0.159 0.0643 ± 7 0.0903 ± 15 0.1297 ± 13 1.150 ± 18 0.0368 ± 7 105 752 ± 22 786 ± 7

14-1 394 107 0.271 0.000 0.1513 ± 5 0.0756 ± 5 0.3436 ± 33 7.169 ± 74 0.0959 ± 11 81 2361 ± 5 1904 ± 1614-2 442 127 0.287 0.017 0.1525 ± 5 0.0777 ± 5 0.3596 ± 34 7.559 ± 77 0.0972 ± 12 83 2374 ± 5 1980 ± 1615-1 554 166 0.299 0.080 0.1440 ± 5 0.0793 ± 7 0.3078 ± 29 6.111 ± 62 0.0816 ± 11 76 2276 ± 5 1730 ± 1416-1 116 118 1.018 0.544 0.0658 ± 22 0.3207 ± 57 0.1314 ± 16 1.191 ± 43 0.0414 ± 9 100 799 ± 69 796 ± 9

17-1 347 253 0.729 0.025 0.1353 ± 6 0.2128 ± 12 0.2749 ± 26 5.129 ± 56 0.0802 ± 9 72 2168 ± 7 1565 ± 1318-1 642 110 0.171 0.025 0.1522 ± 4 0.0478 ± 4 0.3659 ± 34 7.680 ± 75 0.1025 ± 13 85 2371 ± 4 2010 ± 1619-1 361 164 0.455 0.022 0.1541 ± 5 0.1250 ± 7 0.3981 ± 38 8.460 ± 88 0.1095 ± 13 90 2392 ± 6 2160 ± 1720-1 406 236 0.581 0.062 0.1360 ± 5 0.1668 ± 10 0.2691 ± 26 5.044 ± 54 0.0772 ± 9 71 2176 ± 7 1536 ± 1321-1 506 248 0.490 0.002 0.0879 ± 5 0.1628 ± 11 0.1546 ± 15 1.874 ± 22 0.0513 ± 6 67 1381 ± 11 927 ± 822-1 473 489 1.032 0.167 0.1343 ± 6 0.2483 ± 12 0.2593 ± 24 4.800 ± 52 0.0624 ± 7 69 2155 ± 7 1486 ± 1323-1 421 167 0.398 0.063 0.1450 ± 5 0.1125 ± 8 0.3285 ± 31 6.567 ± 69 0.0929 ± 11 80 2288 ± 6 1831 ± 1524-1 174 199 1.145 0.000 0.0659 ± 7 0.3605 ± 30 0.1275 ± 14 1.160 ± 19 0.0401 ± 6 96 804 ± 22 774 ± 8

25-1 451 177 0.391 0.016 0.1519 ± 4 0.1095 ± 6 0.3846 ± 36 8.052 ± 82 0.1077 ± 12 89 2367 ± 5 2098 ± 1726-1 367 134 0.364 0.076 0.0645 ± 6 0.1097 ± 14 0.1300 ± 13 1.156 ± 17 0.0391 ± 6 104 757 ± 20 788 ± 7

27-1 399 484 1.211 0.075 0.1358 ± 5 0.3282 ± 14 0.2701 ± 26 5.058 ± 55 0.0732 ± 8 71 2174 ± 7 1541 ± 1328-1 101 102 1.011 0.213 0.0638 ± 20 0.3168 ± 56 0.1261 ± 16 1.109 ± 39 0.0395 ± 9 104 735 ± 67 765 ± 9

28-2 447 129 0.289 0.256 0.0656 ± 8 0.0888 ± 18 0.1316 ± 13 1.190 ± 20 0.0405 ± 9 101 792 ± 26 797 ± 7

29-1 507 121 0.239 0.001 0.1409 ± 5 0.0630 ± 5 0.2993 ± 28 5.814 ± 60 0.0789 ± 10 75 2238 ± 6 1688 ± 1430-1 673 109 0.161 0.001 0.1440 ± 4 0.0445 ± 4 0.3160 ± 29 6.276 ± 62 0.0871 ± 12 78 2276 ± 5 1770 ± 14

Data arranged in order of sample number.% comm. = common Pb.% conc. = measure of concordance (Pb- and U-loss).Errors are reported at 1-sigma in the table whereas ages quoted in the text are at 2-sigma (i.e. approximately 95% confidence).Normal text (not bold) for 207Pb/206Pb age for zircons with negligible common-Pb.

a 204Pb Broken Hill corrected isotopic compositions.

Table 6Summary of the geochronological framework for the Ankisatra-Besakay District

Sample Rock type Ages (Ma) Interpretation

67-04 UWA 98-19B Mafic granulite 2725 Single xenocryst (from basement)2676 ± 6 Magmatism to generate protolith2465–2390 Older metamorphic event

21349 UWA 98-19C Metasomatised mafic granulite 2465 ± 6 Minimum age for magmatism to generate protolith2465–1900 Metamorphic resetting785 ± 2 Juvenile granulite-facies metamorphism

21380 UWA 98-19A Alkali-feldspar granitoid (granulite) 807 ± 13 Peak granulite-facies metamorphism

21337 UWA 98-19A Quartzo-feldspathic paragneiss/granulite 2870 Maximum age of detrital zircons from protolith1750 Minimum age of detrital zircons from protolith

(also maximum age for sedimentary sequence)780 Granulite-facies metamorphism related to severe Pb- and U-loss

21323 UWA 98-19A Biotite-orthogneiss/granulite 2483 ± 20 Magmatism to generate protolith785 ± 7 Granulite facies metamorphism



reworked ca. 2676 Ma amphibolite to granulite-faciesmafic orthogneiss; (2) significant Cu–Zn and associatedFe–Mn mineralisation in brittle–ductile shear zones in end-erbite of unknown emplacement- and metamorphic-ages;and (3) localised Cu–Zn mineralisation in retrogradeshear-zones with intense silicification and iron oxide stain-ing in quartzo-feldspathic granulite-facies paragneiss(Table 7).

9.1.1. Pb–Ag in shear-hosted quartz veins (Besakay)

This mineralisation style is defined by Pb–Ag-bearingquartz veins 1 cm to 1.5 m wide locally which dip northerlyat 65–80� (Figs. 3 and 7), and contain patches of coarse-grained galena infilling fractures, voids and interstitialspaces (Fig. 5A). On the regional scale, the veins hostingthe Besakay Pb–Ag mineralisation occur in the dilationaljogs situated in zones of convergent linear features sitedcloser to the interpretative post-tectonic sub-volcanic rock(Fig. 6). In light of this structural pattern, it is most likelythat the Besakay quartz veins formed as arrays of en eche-

lon stepping quartz veins within Riedel shears consistentwith shear development in wrench systems (Colvine et al.,1988 and references therein).

The bulk NE-trending shear fold and fault fabric in theorthogneissic sub-terrane is post ca. 800 Ma, as the ca.>2700 Ma granulite-facies mafic orthogneiss and the ca.800 Ma ultramafic complexes are structurally juxtaposedand deformed. It is possible that mineralisation was intro-duced into the system after development of the NE-trend-ing shear zones, which are related to ca. 580–560 Macompressional tectonic events (Windley et al., 1997; Tuckeret al., 1997). This mineralisation and the hosting struc-tures were subsequently re-oriented and concentrated bydeformation and metamorphism during late-orogenic,NNW-trending dextral shear and faulting events in thelate-Neoproterozoic.

9.1.2. Pb–Zn–Au in deformed quartz veins (Ankisatra)The Pb–Zn–Au mineralisation at Ankisatra occurs in

deformed and metamorphosed quartz-galena-chalcopyriteveins (e.g. Fig. 5B, Table 7), up to 2 m thick, croppingout along river-channels and in old trench workings. Thesequartz veins trend NE–SW and are mostly located adjacentto dolerite sills. Both quartz veins and dolerite sills crosscutthe N–S to NE-trending regional fabric, but have, in turn,been crosscut by the NNW-trending dextral shear zones(Fig. 3). The quartz vein has a sugary texture and containsdeformed fine-grained galena and chalcopyrite. Fine-grained galena commonly forms pervasive, dark-greyhaloes in fractured quartz, whereas coarse-grained galenaoccurs locally in fractured massive recrystallised quartzveins. These textural and mineralogical features suggestthat quartz veins and contained galena formed before thelatest significant deformation event recorded in the area(Fig. 5B).

Regional structural data extracted from Landsat images(Fig. 6) suggest that structures hosting the Ankisatra Pb–

Zn–Au mineralisation are extensional fracture-filling veins.These veins and associated dolerite sills were possibly re-oriented during late-orogenic NNW-trending dextralshearing as a result of progressive compressional deforma-tion. In addition, a post-tectonic sub-volcanic rock, inter-preted east of the Ankisatra mineralisation, may havehad an influence on the mineralisation. Although thesemineralised structures have been attributed to late-Neopro-terozoic orogenic events, it is also possible that some ofthese structures were developed during older orogenicevents, such that their present regional setting and associ-ated controls are due to late-Neoproterozoic tectonicreactivation.

9.1.3. Pb–Zn–Cu in mafic orthogneiss

Lead–Zn–Cu mineralisation occurs in 1 mm to 15-metrethick quartz-feldspar leucosome veins/bands in granulite-facies mafic orthogneiss at Besakay (Fig. 7). They occuras closely interbanded, tight to isoclinal folded bands andmylonite to ultramylonite ribbons of recrystallised bluequartz and deformed feldspars in melanosome bands. Leu-cosome bands with Pb mineralisation contain yellowishpatches of epidote-altered amphibole, pyroxene and plagio-clase and possible Pb-bearing yellowish-green plagioclase.These zones contain magnetite and fine-grained fracture-filling pyrrhotite (Fig. 5C). Chalcopyrite is microscopic,occurring as discrete inclusions in pyrrhotite, whereas pyr-rhotite forms fine exsolution lamellae in coarse-grainedmagnetite. Locally, fine-grained galena occurs along micro-fractures in these altered leucosome bands. Based on lim-ited geochemical analyses from Pb-bearing leucosomeveins, this Pb–Zn–Cu mineralisation style is insignificant.

9.1.4. Pb–Zn–Cu in mafic orthogneiss

The dark-green to grey, 20 cm to 30-metre wide bands ofsmokey garnet-magnetite quartzo-feldspathic rock occur ina local, but extensive, NE-trending dextral shear zone(Figs. 4E, F and 6). These shear zones are defined bystrongly deformed, plagioclase-quartz, magnetite, pyrrho-tite and fine-grained pyrite. Secondary pyrite cubes occuralong micro-shear zones and crosscut silicate minerals,fine-grained pyrite and magnetite (Fig. 4E).

Shear zones in deformed and metasomatised granulite-facies mafic orthogneiss commonly dip between 58� and70�E and are proximal to exhalative BIF and Pb–Ag min-eralisation at Besakay (Figs. 3 and 4C). The structuresalong which intense metasomatism occurred are low-anglethrust faults or detachment zones, related to the ca. 2.4 Gaaccretion of the late Archaean granitoid-greenstone belts,through which fluid migration and associated metasoma-tism occurred at ca. 2200 Ma. The steep dip of this struc-ture is probably due to Pan-African collisional tectonics.

Local metasomatic zones are interpreted to result frommineral replacement due to the movement of deeplysourced hydrothermal fluids along zones of structuralweakness. At a district scale, the metasomatised granu-lite-facies mafic orthogneiss may represent a detachment

Table 7A summary table showing assay results of selected mineralised systems sampled from the Ankisatra-Besakay District, northern Andriamena, Madagascar

Sampid Eastings Northings Deposit/host terrane

Area Au(ppm)

Ag(ppm)

As(ppm)

Bi(ppm)

Fe(%)

Cu(ppm)

Zn(ppm)

Pb(ppm)

Ba(ppm)

Cd(ppm)

Co(ppm)

Ni(ppm)

Mg(%)

Ca(%)

Na(%)

Mn(ppm)

P(%)

21315 509680 1013030 Ribbon-typeqtz-feld

Area A <0.01 3.5 15 <0.1 33.2 10 190 325 350 <0.5 5 60 1.2 0.16 0.36 510 0.06

21345 511320 1008150 Brecciatedenderbite

Area C <0.01 2.0 <5 0.2 10.0 195 40 11 410 <0.5 <5 25 0.2 0.13 0.35 215 0.09


Area C <0.01 <0.5 <5 <0.1 12.2 425 100 21 100 <0.5 50 145 1.4 0.49 0.28 790 0.18

21349 509942 1010162 Magnetitequartzite

Area D <0.01 0.5 <5 <0.1 8.0 50 70 48 680 <0.5 15 40 1.6 1.18 1.55 820 0.06

21350 509880 1010420 Magnetitequartzite

Area D <0.01 0.5 <5 <0.1 6.8 180 50 115 185 <0.5 15 120 0.6 0.53 0.68 880 0.08

21351 509900 1010301 Silicate-magnetitegneiss

Area D <0.01 <0.5 <5 <0.1 32.0 25 50 54 105 <0.5 10 35 1.4 1.10 0.50 530 0.10

21356 510460 1010530 Galenaquartz vein

Area B 0.25 83 <5 1.0 0.2 25 80 31500 105 4.5 <5 70 <0.1 0.03 0.44 30 <0.01

21357 510460 1010530 Galenaquartz vein

Area B 0.57 250 <5 2.8 0.4 45 20 118000 15 17.0 <5 50 <0.1 0.02 0.31 60 0.02

21358 510520 1010650 Galena-cpyquartz vein

Area B 0.20 12.0 <5 0.6 1.4 960 720 3350 90 11.0 15 65 0.2 0.08 0.62 175 0.37

21363 510770 1010400 Py-cpy ina breccia

Area B <0.01 1.5 <5 <0.1 3.0 65 50 300 3300 <0.5 10 15 0.6 1.28 2.48 450 0.03

21375 509000 1012720 Wall-rockto 21376

Area A <0.01 1.0 <5 <0.1 2.6 20 50 52 940 <0.5 10 55 0.6 3.81 4.00 580 0.04


Area C <0.01 1.5 <5 0.1 5.8 480 40 22 830 <0.5 35 150 0.4 0.30 0.56 630 0.19


Area C <0.01 1.0 <5 <0.1 24.8 1410 290 17 25 <0.5 135 445 4.4 0.43 0.42 1920 0.55

Areas sampled are shown in Fig. 3.Abbreviations:

Area A, Shear-type Besakay Pb–Ag deposit;Area B, Shear-type Ankisatra Pb–Zn-Au deposit;Area C, Breccia-type enderbite hosted mineralisation;Area D, Breccia-type magnetite-quarzite hosted mineralisation;Area E, Vein-stringered banded paragneissic granulite;qtz, quartz;feld, feldspar.

J.

Ka

bete

eta

l./

Jo

urn

al

of

Africa

nE

arth

Scien

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5(

20

06

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7–

12

2113


splay, related to the detachment zone defined by the base-ment amphibolite–granulite-facies alkali-feldspar granit-oid, which intruded at 807 ± 13 Ma.

9.1.5. Cu–Zn–Co in brecciated enderbite

The brittle–ductile magnetite–pyrite quartz-plagioclaseenderbitic rock, with Cu, Zn, Co, Fe–Mn mineralisation,crops out between the tectonic boundary and the late-Archaean basement rocks underlying the orthogneissicthrust sub-terrane (see area C, Fig. 3). In this area, miner-alisation is exposed sporadically along the edges of theridge and more extensively as a manganiferous–ironstonelaterite cap extending for at least 1 km (Fig. 4M). Handspecimen examination shows that a late anastomosing C-S shear-foliation overprints the earlier pyrite–magnetitebreccia texture in the mineralised enderbite. In lessdeformed bands, sulphides and magnetite are remobilisedin veinlets or occur as infillings along silicate–quartz or pla-gioclase mineral boundaries. In these areas, weakly devel-oped C-S fabric contains disseminated sulphides at grainboundaries and along the S-fabric. In the C-planes, sulp-hides and garnet form thin elongate bands parallel to foli-ated plagioclase and elongate quartz (Fig. 5E and F).

The pyrite–magnetite mineralisation and brecciation ofenderbite are probably related to young intrusions inter-preted as post-tectonic sub-volcanic rocks. In the field, end-erbite enclaves in younger feldspathic pyroxenite intrusionsare very common. The intrusion-related brecciation andassociated pyrite and magnetite mineralisation in the brec-cia matrix is primary, but was overprinted by a secondary,C-S anastomosing shear fabric related to the late-orogenicNNW-trending structures (Figs. 4L, 5E, F and 6). Hydro-thermal fluids responsible for pervasive garnet floodingalong the C-S fabric are possibly responsible for Cu, Zn,Co and Fe–Mn mineralisation. Similarly, these fluids mayhave leached metals related to the magnetite–pyrite brecciaand deposited them into late-orogenic structures (Fig. 4G).

9.1.6. Cu–Zn in quartzo-feldspathic granulite-facies

paragneiss

Quartzo-feldspathic granulite-facies paragneiss in thesouthern portion of the tectonic corridor sub-terrane con-tains structurally controlled Cu–Zn mineralisation(Fig. 3). Significant Cu–Zn mineralisation and elevatedconcentrations of Co–Ni and Ag occur in strongly shear-foliated zones ranging from 5 to 50 m wide. They arelocally mylonitic, very siliceous and ferruginous, andlocally contain sulphide rich zones (Table 7). At a meso-scopic to microscopic scale, these zones contain closelyspaced, irregular fracture planes (Fig. 4G and H). Micro-scopic examination of these structures reveals alternatinggrey bands of blue quartz (ribbon quartz), patches ofdeformed garnet and sulphide mineralisation, retrogradechlorite, graphite and titanite bands and pyrrhotite andmagnetite. They occur in a strongly ductile anastomosingfabric wrapping around peak-metamorphic garnet andfeldspar and strongly deformed orthopyroxene (e.g.

Fig. 4G and H), suggesting that the anastomosing struc-tures are post-peak ca. 785 Ma granulite-facies metamor-phism (Fig. 5D). Thus, the NNW-trending dextral shearzones are late in the orogenic history of the district, andare responsible for re-orienting pre-existing zones of miner-alisation of unknown age to their present sites (Fig. 5D).

10. Lead-isotope studies of mineral deposits

10.1. Introduction

The success of radiogenic Pb isotopes as significant trac-ers of geological source regions and for dating the timing ofmineralisation is related to variations in the abundance ofthe isotopes, 206Pb, 207Pb, 208Pb, which are derived fromradiogenic decay of 238U, 235U and 232Th isotopes, respec-tively, at different rates. The fourth Pb isotope, 204Pb, usedas a normalising isotope, is an indicator of the initial Pb.207Pb evolves faster than 206Pb, causing fractionation inthe Earth with time. Minerals such as galena, pyrite andpotassium-feldspar are Pb-rich and contain insignificantamounts of U and Th, so that their U/Pb and Th/Pb ratiosare effectively zero. Therefore, there is no additional sourceof radiogenic 206Pb, 207Pb, and 208Pb in these minerals fromthe time of their first formation. Their present Pb-isotopiccompositions are thus due to common-Pb which wastrapped at the time of their initial formation. Becausegalena and pyrite are common ore-forming minerals, theyare routinely used to help identify different possible metalsources for ores as well as to provide model ages for oredeposition (e.g. Cumming and Richards, 1975; Staceyand Kramers, 1975; Gulson et al., 1985; McNaughtonet al., 1993; Reid et al., 1997). These isotopic systems havebeen successfully used to examine the source and geneticrelationships between stratabound, stratiform and vein-type Pb–Zn–Ag deposits in BHT districts (e.g. Gulsonet al., 1985; Reid et al., 1997), as well as for differentsources of metals in other mineral deposits (e.g. Archaeanlode-gold deposits: McNaughton et al., 1993).

A reconnaissance Pb-isotope study was carried out onfive galena samples extracted from Pb-bearing depositslocated within 15 km of each other in the A-BD (Fig. 6).The major aim was to trace the nature and origin of metals,to estimate model ages of mineralisation based on assump-tions explained above, and to compare the isotopic compo-sitions of galenas from the A-BD with those of vein-typedeposits in BHT provinces.

10.2. Lead-isotope compositions

The initial Pb-isotope compositions from two vein-typePb deposits of the A-BD are given in Table 8 and presentedin Fig. 13. Galena analysed from Besakay is coarse-grained, occurring along fractures in massive en echelonquartz veins. At Ankisatra, both galena and quartz veinsare fractured into thin cleavage planes cutting recrystallisedsugary quartz and galena, and filled with late-stage chalco-

Table 8Lead-isotope compositions of galena from the Ankisatra-Besakay District, Madagascar and the BHT Gamsberg and Broken Hill deposits of theBushmanland Ore District, South Africa (Reid et al., 1997), used for comparison with the lithospheric growth model of Stacey & Kramers (1975)

Sample Deposit Description 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb

21356 Ankisatra, Madagascar Sugary, fractured quartz vein with 16.316 15.828 37.71821357 Ankisatra, Madagascar galena sheared and overprinted by late, 16.307 15.817 37.67921358 Ankisatra, Madagascar chalcopyrite 16.322 15.841 37.736

21369 Besakay, Madagascar Coarse-grained galena in voids 16.353 15.707 37.43421376 Besakay, Madagascar and fractures in massive quartz vein 16.374 15.729 37.508

GAM91 Gamsberg, S. Africa 16.780 15.571 36.669GAM56 Gamsberg, S. Africa 16.789 15.561 36.648GAM54 Gamsberg, S. Africa 16.776 15.565 36.648

BH3 Broken Hill, S. Africa 16.746 15.581 36.692BH4 Broken Hill, S. Africa 16.736 15.573 36.687BH5 Broken Hill, S. Africa 16.718 15.566 36.647*BH 8/23 Broken Hill, S. Africa 16.756 15.585 36.698*BH 8/29 Broken Hill, S. Africa 16.732 15.572 36.694#BH 8/1 Broken Hill, S. Africa 16.728 15.570 36.657

With the exception of the marked samples (i.e. * for massive ore body and # for bedded ore body) the remainder comprise vein-hosted deposits.

Fig. 13. A plot showing Pb-isotope compositions of bedded, massive and vein-hosted galena from the Broken Hill and Gamsberg BHT Pb–Zn (Cu–Ag)deposits in the Bushmanland Ore District, South Africa (Reid et al., 1997) in comparison with those of Pb–Ag-vein bearing galenas of the Ankisatra-Besakay District, northern Madagascar, with respect to the Stacey and Kramers (1975, S&K) crustal growth evolution model.


pyrite mineralisation. Fragments of galena were hand-picked from each sample and dissolved at the Lead-freeLaboratory in the Department of Geology and Geophysics(now School of Earth and Geographical Sciences) at theUniversity of Western Australia. Lead-isotope composi-tions of galena were determined using the VG 354 MassSpectrometer at Curtin University. Analytical proceduresfollowed those of Ho et al. (1994).

10.3. Source-region

Galenas from all A-BD deposits fall well above theStacey and Kramers (1975) and Cumming and Richards(1975) growth curves with galenas from Ankisatra havinghigher 207Pb/204Pb ratios than Besakay galenas (Fig. 13).These data suggest that the Pb from both deposits was

derived from older crust with an U/Pb greater than typicallithosphere (i.e. crustal model of Stacey and Kramers,1975), and that the Pb sources for the two deposits weredifferent. In addition, the Pb growth curves for a simpleevolutionary model cannot relate the two deposits to eachother.

Given the complex geological history of the area (e.g.Collins et al., 2001; Collins and Windley, 2002; Goncalveset al., 2003), the complex evolution of the Pb reservoirs isnot surprising. One important conclusion, however, is thatthe heterogeneity of the samples from the A-BD is atypicalof Broken Hill style ores which typically show homoge-neous Pb-isotope ratios (e.g. Bushmanland District, forboth vein-type and bedded ore: Fig. 13). Such heteroge-neous lead data are, instead, indicative of small or Pb-poorsystems (Gulson et al., 1985; McNaughton, 1987). It is

Table 9Model ages based on the Cumming and Richards (1975) lead growthmodel

Sample Deposit Model Age (Ma)

t7/6 t6 t8

21356 Ankisatra 1708 1398 57421357 Ankisatra 1705 1402 59321358 Ankisatra 1714 1394 56521369 Besakay 1590 1378 71221376 Besakay 1595 1367 676

These ages aret7/6: based on the 207Pb/206Pb ratio.t6: based on the 238U! 206Pb decay only.t8: based on the 232Th! 208Pb decay only.


possible that these variations are due to mixing and/or con-tamination on a local scale as indicated by the contrastingPb-isotope signatures of the largely granitoid-gneiss hostedAnkisatra deposit (Fig. 3) and the mafic-orthogneiss hostedBesakay deposit (Fig. 7). Furthermore, fracture or cleavageplanes in galena and quartz, and the recrystallised nature ofquartz veins at the Ankisatra deposit, suggest deformationof pre-existing mineralisation and subsequent late hydro-thermal modification and overprinting mineralisation.

10.4. Lead model ages

Lead-isotope model ages can give an approximate age ofgalena mineralisation, since there is no U or Th decay sincegalena crystallisation. Lead growth models assume that Pbevolved within a large homogeneous source reservoirbefore being trapped when galena formed. Stacey and Kra-mers (1975) and Cumming and Richards (1975) presentsimilar Pb-isotopic models for the lithosphere, which pro-vide generally reliable ages when data fall on the litho-spheric growth curve (e.g. S & K Growth Curve, Fig. 13).

For galena whose Pb is on the growth curves, and allthree Pb model ages are the same, the age can be consid-ered reliable. This is clearly not the case for the samplesfrom the A-BD (Table 9), and these model ages may bemeaningless. Similar arguments can be levelled at the1850 ± 50 Ma and 1750 ± 70 Ma Pb-model ages reportedby Besairie (1961) from galenas sampled from the Ankisa-tra and Besakay deposits, respectively. These are unlikelyto be reliable estimates of the age of formation, particularlysince these ages are not consistent with geochronologicallyconstrained geological events, such as magmatic and peakmetamorphic events, defined by robust SHRIMP geochro-nology as outlined above.

11. Implications for BHT deposits in the A-BD

11.1. Introduction

It is stressed at the outset of this discussion that there isno implication whatsoever that the A-BD was adjacent toany of the BHT provinces with which it is compared at

the time of BHT mineralisation. The comparisons arebased simply on similarities or differences in lithostratigra-phy, mineralisation styles and associated exhalites, robustgeochronology and Pb-isotope signatures of the A-BDcompared to major BHT provinces, which have specificand distinctive characteristics for each of these parameters,as outlined above.

11.2. Lithostratigraphy

There are broad similarities between lithological unitsmapped in the A-BD and some of the major lithologic unitscompiled from selected BHT Provinces (Kerr, 1994). Base-ment rocks in the BHT provinces comprise amphibolite- togranulite-facies mafic and felsic granulite, quartzo-feld-spathic schist and banded leucocratic gneiss (Kerr, 1994).These lithological units are petrographically and mineral-ogically similar to the granulite-facies mafic orthogneiss,tonalitic gneiss and granitoid gneiss in the A-BD. In thelower sequences of the BHT stratigraphy, quartzo-feld-spathic gneiss, derived from an enigmatic precursor rock(e.g. Broken Hill Block, Australia and Bushmanland OreDistrict, South Africa (Kerr, 1994; Roche, 1994; Reidet al., 1997), is comparable to undifferentiated quartzo-feldspathic gneiss interleaved with tectonically reworkedlate-Archaean basement rocks in the A-BD. In addition,quartzo-feldspathic paragneiss in the A-BD could also beanalogous to some of the lower sequence quartzo-feldspathic rocks documented from selected BHT prov-inces (Roche, 1994; Reid et al., 1997). Apart from thesilicate-facies BIF horizon, there are no other analogoustransitional sequences shelf-facies sedimentary rocks in theA-BD, such as are common at this stratigraphic level inmost BHT provinces.

11.3. Mineralisation styles and associated exhalites

Another way to compare the small A-BD Pb–Ag, Pb–Zn–Au, Cu–Zn and Fe–Mn deposits with BHT-mineralisa-tion is to compare and contrast their mineralisation styles.One of the principal characteristics of BHT-provinces istheir world-class Pb–Zn–Ag mineralisation, a critical fea-ture missing in the A-BD. However, quartz-vein hostedgalena has been known in the A-BD since the early 1930s(Besairie, 1961), and giant BHT deposits occur in provincesthat also contain small-scale vein-type Pb–Ag–Zn–Cu min-eralisation such as: (1) silver–Pb–Cu in quartz veins andAg–Pb in quartz-siderite veins in the Broken Hill and Tha-karinga Groups of the Broken Hill Block (Kerr, 1994;Giles and Ehlers, 1997; Stevens and Burton, 1998); (2) sil-ver–Pb–Zn in migmatitic quartz-feldspar pegmatite veins inthe Soldiers Cap Group, Cannington, Mt. Isa Block(Skrzeczynski, 1993; Roche, 1994; Kerr, 1994); and (3)Pb–Zn–Ag in pegmatite and quartz-veins in the Bushman-land Ore District (Reid et al., 1997). Thus a comparisonbetween deposit-scale to district-scale features of the A-BD and BHT provinces is justified.


Bedded units of quartz-magnetite amphibolite gneiss(Fig. 4C) which occur in the A-BD are interpreted to beexhalative silicate-facies BIF. In addition, the garnet–mag-netite–quartzo-feldspathic rocks (i.e. metasomatised granu-lite-facies mafic orthogneiss), which crop out near the BIFand Besakay Pb–Ag deposit represent strong metasomaticzones (Figs. 3 and 4E and F). Both the BIF and metasoma-tised granulite-facies mafic orthogneiss are possible ana-logues of the iron-oxide facies horizons common intransitional exhalative sequences associated with BHT min-eralisation. However, the key features of transitionalsequences, the Zn-spinel (gahnite)-bearing rocks and/orcalc-silicate units, are not known from the A-BD. Thus,the vein-type deposits of the A-BD do show some similar-ities with those of BHT provinces, but some critical compo-nents are absent.

11.4. U–Pb in zircon geochronology

Based on the specific distribution through Precambriantime of preserved BHT deposits, it is appropriate to com-pare the geochronological evolution of the major litholog-ical units that host Pb–Ag and Pb–Zn–Ag mineralisation inthe A-BD (cf. Fig. 14) with that of the BHT province of theBroken Hill Block, Australia (e.g. ca. 1690 and 1590 Maorogenic events). Nutman and Ehlers (1998a,b) haveshown the probable existence of ca. 2670–2550 MaArchaean crust in the Broken Hill Block. They infer theseages from abundant SHRIMP U–Pb analyses of inheritedzircons extracted from the quartzo-feldspathic amphibo-lite-facies gneiss of the Willyama Supergroup in the south-

Fig. 14. A summary diagram illustrating major orogenic events so far constrainthe Ankisatra-Besakay District. At least two magmatic events are recorded frometamorphic event is related to progressive or new compressional tectonics onmetasomatism may have occurred later. These orogens and other basemendevelopment of the basin in which the precursor to the paragneiss was depositand magmatism following rifting of the basement from ca. 820 Ma.

ern portion of the Broken Hill Block. These Archaean agesagree with the terrane evolution model of AGCRC (1995),Willis (1996), Walters (1998) and many other workers,which advocates rifting of the Archaean crust as the initialphase in the evolution of the mobile belts which host theintracratonic BHT deposits. Apart from the Archaeanages, the quartzo-feldspathic amphibolite-facies gneiss con-tains several inherited zircons preserving ages of ca. 2400–2100 Ma, interpreted by Nutman and Ehlers (1998a,b) as arecord of tectonometamorphic events (Fig. 15).

Whereas the Archaean crust in the Broken Hill Block isoverlain by thick Proterozoic sequences, the Archaean crustin the A-BD is well exposed at surface. In the A-BD,Archaean rocks preserve magmatic U–Pb zircon ages at2676 ± 6 Ma and 2483 ± 20 Ma, recording late Archaeansub-volcanic magmatism of intermediate to felsic composi-tion. Furthermore, the granulite-facies mafic orthogneiss(meta-diorite/granodiorite) contains inherited xenocrysticzircon with a concordant U–Pb age of 2725 ± 12 Ma, whilethe overlying meta-supracrustal rocks (i.e. quartzo-feld-spathic granulite-facies paragneiss) contain detrital zirconwith a maximum zircon age of 2870 Ma. These inheritedages show that earlier crust was involved in Late Archaeancrustal growth processes in the region. In addition to indi-cating the presence of precursor magmatic rocks, the quar-tzo-feldspathic granulite-facies paragneisses contain sixzircons preserving ca. 2870–2563 Ma ages and abundantzircons with ca 2450–2200 Ma ages, and very rare zirconspreserving ca. 1800–1650 Ma ages (see Table 4). Thus, boththe Broken Hill Block and the A-BD had somewhat similarArchaean histories.

ed by U–Pb in zircon SHRIMP geochronology of selected rock types fromm the Late Archaean convergent-margin magmatic event. The older peak-the older rocks along whose thrust detachments and associated splay zonest rocks supplied sedimentary/sedimentary-volcanic material during theed. The young peak metamorphism is related to anatectic metamorphism

Fig. 15. A summary diagram illustrating the comparison between Archaean and Proterozoic terrane evolution of the Broken Hill Block, Australia and theAnkisatra-Besakay District of northern Madagascar. In this diagram, timing of the major orogenic events (magmatic, metamorphic and supracrustalaccumulation) are compared. Age dates used are all from U–Pb in zircon by SHRIMP dating techniques.


The greatest difference between the A-BD and the Bro-ken Hill Block, from the present geochronological study,is the absence of the most critically-important ca. 1690–1574 Ma age-component related to long-lived orogenicevents at Broken Hill and commonly recorded in the hostlithostratigraphy for BHT mineralisation (Fig. 15). Asshown in Table 4, only two out of 38 zircons extractedfrom the quartzo-feldspathic granulite-facies paragneisspreserve ca. 1690–1550 Ma U–Pb zircon ages. This indi-cates that either these orogenic events did not occur inthe A-BD or that they occurred but were effectively obliter-ated by the pervasive ca. 780–820 Ma granulite-faciesmetamorphism. The latter appears unlikely as Archaeanzircons are well preserved in all rocks with Archaean pre-cursors despite their subsequent history of reworking.

11.5. Lead-isotope signature

Lead-isotope signatures characteristic of the Ankisatra-Besakay Pb–Zn–Ag deposits are compared with Pb-iso-tope compositions from massive, bedded and quartz-veinhosted Pb–Zn–Ag deposits at Gamsberg and Broken Hill,in the Bushmanland Ore District (Fig. 13). The reason forusing the Bushmanland galena for comparison with theAnkisatra-Besakay galenas is due to their similarly strongradiogenic signatures, but also the fact that both vein andstratabound/stratiform mineralisation in the Bushman-land Ore District give almost the same isotopic signature(Reid et al., 1997). In contrast, Pb from the Thakaringaveins in the Broken Hill Block is more radiogenic com-pared to that in galena sampled from the adjacent, MainLode stratabound mineralisation (Ryan et al., 1986; Ste-vens et al., 1990). Lead-isotope compositions from bothveins and BHT deposits at Broken Hill Block are homo-

geneous but more highly radiogenic than those from theBHT deposits of the Bushmanland Ore District. The dif-ferent Pb-isotope signatures of the Thakaringa veins andBHT deposits suggest that, although these deposits arespatially related, their Pb was derived from differentsource regions. In contrast, the homogeneity of Pb-isotopecompositions in both veins and stratiform deposits in theBushmanland Ore District implies a common sourceregion and suggests that mineralisation in the veins wasremobilised from primary stratiform BHT mineralisationduring the ca. 1135 Ma Grenvillian Orogeny (Reidet al., 1997).

In the A-BD, the Pb is both heterogeneous and veryradiogenic compared to even Pb from the BHT depositsin the Bushmanland Ore District. The low 206Pb/204Pband high 207Pb/204Pb ratios for the Ankisatra-Besakay gal-enas in particular contrast with respective isotopic ratiosfor the BHT deposits in the Bushmanland Ore District(Table 8). As the Pb-isotope compositions of veins andstratiform deposits in the Bushmanland Ore Districtremained homogeneous even after the ca. 1135 Ma Gren-villian Orogeny, it can be inferred that Pb-isotope compo-sitions of deposits in the A-BD reflect their initial ratiosirrespective of the strong overprint of the Pan-AfricanOrogeny. The heterogeneous nature of the Pb-isotope com-position suggests that Pb was derived from a local sourceinvolving older crust with greater U/Pb ratios than typicallithosphere, contrasting with the larger, more homoge-neous Pb reservoir evident in BHT provinces.

The model ages calculated from the isotopic composi-tion of the galenas from A-BD do not fulfil consistencyassumptions for the radiogenic Pb-isotope systems of Sta-cey and Kramers (1975) and Cumming and Richards(1975). Therefore model ages shown in Table 9 are not


robust and previous model ages should be treated with cau-tion both for internal and external comparisons.

12. Summary and conclusions

12.1. Lithostratigraphy

The present structural complexity of the orthogneissicthrust sub-terrane may have resulted from tectonic stack-ing and reworking of early Precambrian basement rocksof different tectonic signatures during Neoproterozoic toearly Palaeozoic collisional orogenic events (e.g. Windleyet al., 1997; Kroner et al., 1997). In this case, precursorsto granulite-facies mafic orthogneiss and biotite–granitoidhornfels may represent arc-magmatic rocks, whereas pre-cursors to quartzo-feldspathic granulite-facies paragneissmay be part of reworked older back-arc basin sequencesjuxtaposed against continental crustal rocks. Unlike theamphibolite to granulite-facies alkali-feldspar granitoid,which defines a structural break, enderbite defines a zoneof structural continuity from the tectonic boundary tothe upper lithotectonic units in the orthogneissic thrustsub-terrane. Enderbite contains brittle–ductile structuresranging from fluid induced breccia to the east, throughzones of mylonite to ultramylonite in ductile shear-zonesto the west.

12.2. Terrane evolution history

The present SHRIMP U–Pb in zircon study of selectedrocks from the A-BD shows:

1. A ca. 2600–2500 Ma period of convergent margin mag-matism responsible for the generation of the precursorsto the granulite-facies mafic orthogneiss and biotitegranitoid hornfels. These magmatic rocks evolved aspart of the global evolution of Late Archaean granit-oid-greenstone belts (Barley et al., 1998), which involvedolder Archaean crust as indicated by a xenocrystic U–Pbzircon age of 2725 ± 12 Ma in mafic orthogneiss.

2. A ca. 2400–2260 Ma period of progressive convergent-margin tectonics involving terrane accretion, partialmelting and metamorphism of diorite-granodiorite, ton-alite, granitoid, greenstone and undifferentiated quar-tzo-feldspathic gneiss (Figs. 9D, E, 10 and 14). Inaddition, ca. 2260 and 1900 Ma metamorphic resettingin the metasomatised mafic orthogneiss is interpretedto have been related to strong hydrothermal activityalong detachment or related splay zones (Fig. 10B).

3. A ca. 1700–800 Ma period in which potentially therewas intracratonic basin development, and associatedorogenic events, but whose record, if present, was over-printed or reset by granulite facies metamorphism anddeformation at ca. 785 Ma.

4. A ca. 785–820 Ma rifting of part of the basement rockscausing the emplacement of mafic–ultramafic complexesand mangerite magmatism. This caused high-tempera-

ture low-pressure granulite-facies metamorphism of theprecursor biotite–granitoid hornfels and partial to com-plete anatectic melting of the precursor amphibolite-facies alkali-feldspar granitoid.

In general, the geochronological data from the A-BDprovide evidence for a tectonically active Precambrianorogen.

12.3. Comparison with BHT Provinces

12.3.1. Hosting lithologic features

Although they were never juxtaposed, there are broadsimilarities between the Archaean to early Palaeoprotero-zoic and Neoproterozoic evolution of the A-BD and theBroken Hill Block (Fig. 15). The A-BD comprises quar-tzo-feldspathic granulite-facies paragneisses with detritalzircon ages of this protolith between 2780 and 1760 Ma.These Archaean ages and the volcanic component in thequartzo-feldspathic granulite-facies paragneisses suggestthat the precursor to the quartzo-feldspathic granulite-facies paragneiss in the A-BD was deposited in an intra-cratonic basin at ca. 1700 Ma. The period over whichthis basin developed is not constrained. The quartzo-feldspathic granulite-facies paragneiss is a possible ana-logue of the lower sequence quartzo-feldspathic rockswhich are common in BHT provinces. Sporadic silicate-facies BIF in the A-BD could be remnants equivalent tothe transitional zones in the BHT provinces. This couldbe further evaluated through the use of chemical discrimi-nation plots, but the chemical data are not available.

12.3.2. Mineralisation styles and lead-isotope compositions

There are no direct indications of the presence of strat-iform or stratabound Pb–Zn–Ag deposits that typify BHTprovinces and the distinctive metasomatic and metamor-phic haloes commonly associated with these provinces inthe A-BD. There are, however, numerous quartz-veinsand quartz-plagioclase leucosome veins/bands with Pb,Ag, Zn, Cu, Au mineralisation in the A-BD. The locationof these mineralisation styles is controlled by late-Neopro-terozoic NNW-trending dextral faults or shear zones.These are similar to the majority of the vein-type miner-alisation styles known in BHT provinces, with quartzvein-hosted Pb–Zn–Cu–Au and magnetite-pyrite ender-bite-hosted Cu–Zn–Fe–Mn–Co occurrences the most sig-nificant mineralisation styles identified in the A-BD todate.

Lead-isotope compositions of galenas from the A-BDvein systems are heterogeneous and very radiogenic. Thisindicates that Pb was derived locally from rocks older thanca. 2700 Ma and that therefore the veins are likely to rep-resent small-scale mineralised systems, atypical of BHTdeposits and their associated vein-style mineralisation,which have more homogeneous Pb-isotope compositionsmore indicative of larger fluid and metal reservoirs. Modelages of 1850 ± 50 Ma and 1750 ± 70 Ma (Besairie, 1961),


previously interpreted as the age of galena and associatedmineralisation in the A-BD, are not reliable.

12.3.3. Geochronological constraints

The available robust geochronological constraints aresummarised in Fig. 14 and are not discussed in detail here.Suffice to state that the derived history of the A-BD is verysimilar in terms of Archaean to early-Palaeoproterozoicand Neoproterozoic events to BHT provinces such as theBroken Hill Block (Fig. 15). However, the most criticalperiod of history, the late Palaeoproterozoic to Mesoprote-rozoic, for BHT provinces such as the Broken Hill Block isapparently absent from the A-BD. Given the abundant evi-dent for the robustness of the Archaean zircons during Pro-terozoic events in the A-BD, it is unreasonable to suggestthat there was a Mesoproterozoic in the A-BD but that evi-dence for it was obliterated by the ca. 800 Ma metamorphicevent.

This evidence, when combined with the Pb-isotope dataand lack of crucial exhalative rocks in the exposedsequences, downgrades the A-BD as a potential BHT prov-ince despite its superficially similar lithostratigraphicsequences to BHT provinces worldwide.

13. Exploration significance and future research

Despite its low ranking as a potential BHT province,there clearly are some exploration targets in the A-BD.These include the magnetite-pyrite breccia-hosted Cu, Zn,Fe–Mn and Co mineralisation, whose host-rock crops outextensively as a manganiferous gossan ridge for about onekilometre (Figs. 3, 4O and 5E and F). The original originof this mineralisation style is unknown, although it is nowapparently remobilised into N–S dextral strike-slip shearzones and NNW-trending dextral faults and shears(Fig. 6). Other mineralisation which is worthy of furtherinvestigation includes the Pb–Zn–Cu occurrences inquartz-feldspar leucosome veins/bands in the granulite-facies mafic orthogneiss, because the host orthogneiss isextensive in the study area, and because similar mineralisa-tion is significant in the Soldiers Cap Group, Cannington, inthe Mt. Isa Block (Roche, 1994; Kerr, 1994; Willis, 1996).

A regional structural study to constrain the structuralevolution synthesised from Landsat image interpretation,coupled with more detailed compilation and field mappingof all known mineral occurrences and deposits in theAndriamena Sheet, is recommended. This should deter-mine the principal controls on late-orogenic NNW-trend-ing dextral shear-zones and the post-tectonic plutons, therelationships between them and their significance in themetallogenic evolution of the A-BD. Detailed field docu-mentation is required of the enderbite and younger plu-tonic rocks in the area, to determine their controls, ifany, on the Cu, Zn, Fe–Mn and Co mineralisation in themagnetite–pyrite enderbite. Study is also required in thelatter to determine whether garnet flooding in anastomo-sing shear zones was related to the mineralising fluids,

and whether the mineralisation was introduced with themagnetite–pyrite phase of brecciation, or, alternatively,was introduced later by infiltration of hydrothermal fluidsalong the anastomosing shear zones. A geochronologicalstudy to better constrain the emplacement age of the end-erbite and the post-tectonic (ca. <780 Ma) plutons is alsorecommended.

Acknowledgements

This paper is part of a M.Sc. study funded by BHP Min-erals International Exploration Inc. and the Governmentof Tanzania and supervised by David Groves. The supportby the senior management of BHP Minerals InternationalExploration Inc., in particular Miles Shaw, Pablo Marcet,Audace Ntungicimpaye, Roger Kuhns and Geoff Woad, isgratefully acknowledged. JK is grateful to Ian Fletcher,Charter Matthison and Brian Krapez for their kind sup-port and advice during interpretation of the SHRIMPdata. Brendan Griffin is thanked for assistance with theEnvironmental Scanning Electron Microscopy (ESEM)and Marion Marshall for zircon separation and mounting.Abdul Mruma is thanked for reading an earlier version ofthe manuscript. We are grateful to Alan Collins, SospeterMuhongo and Lewis D. Ashwal for their incisive refereecomments on earlier drafts of this paper which resultedin major improvement and updating.

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