Geochronology of the Precambrian crust in the Mozambique belt in NE Mozambique, and implications for Gondwana assembly

Precambrian Research 170 (2009) 231–255

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Geochronology of the Precambrian crust in the Mozambique belt in NE Mozambique, and implications for Gondwana assembly B. Bingen a,∗ , J. Jacobs b , G. Viola a , I.H.C. Henderson a , Ø. Skår a , R. Boyd a , R.J. Thomas c , A. Solli a , R.M. Key c , E.X.F. Daudi d a

Geological Survey of Norway, Trondheim, Norway University of Bergen, Bergen, Norway British Geological Survey, Keyworth, UK d National Directorate for Geology, Maputo, Mozambique b c

a r t i c l e

i n f o

Article history: Received 21 July 2008 Received in revised form 12 January 2009 Accepted 26 January 2009 Keywords: Pan-African orogeny Irumidian orogeny Gondwana Mozambique East Africa U/Pb

a b s t r a c t Zircon and monazite U–Pb data document the geochronology of the felsic crust in the Mozambique Belt in NE Mozambique. Immediately E of Lake Niassa and NW of the Karoo-aged Maniamba Graben, the Ponta Messuli Complex preserves Paleoproterozoic gneisses with granulite-facies metamorphism dated at 1950 ± 15 Ma, and intruded by granite at 1056 ± 11 Ma. This complex has only weak evidence for a PanAfrican metamorphism. Between the Maniamba Graben and the WSW–ENE-trending Lurio (shear) Belt, the Unango and Marrupa Complexes consist mainly of felsic orthogneisses dated between 1062 ± 13 and 946 ± 11 Ma, and interlayered with minor paragneisses. In these complexes, an amphibolite- to granulitefacies metamorphism is dated at 953 ± 8 Ma and a nepheline syenite pluton is dated at 799 ± 8 Ma. Pan-African deformation and high-grade metamorphism are more intense and penetrative southwards, towards the Lurio Belt. Amphibolite-facies metamorphism is dated at 555 ± 11 Ma in the Marrupa Complex and amphibolite- to granulite-facies metamorphism between 569 ± 9 and 527 ± 8 Ma in the Unango Complex. Post-collisional felsic plutonism, dated between 549 ± 13 and 486 ± 27 Ma, is uncommon in the Marrupa Complex but common in the Unango Complex. To the south of the Lurio Belt, the Nampula Complex consists of felsic orthogneisses which gave ages ranging from 1123 ± 9 to 1042 ± 9 Ma, interlayered with paragneisses. The Nampula Complex underwent amphibolite-facies metamorphism in the period between 543 ± 23 to 493 ± 8 Ma, and was intruded by voluminous post-collisional granitoid plutons between 511 ± 12 and 508 ± 3 Ma. In a larger context, the Ponta Messuli Complex is regarded as part of the Palaeoproterozoic, Usagaran, Congo-Tanzania Craton foreland of the Pan-African orogen. The Unango, Marrupa and Nampula Complexes were probably formed in an active margin setting during the Mesoproterozoic. The Unango and Marrupa Complexes were assembled on the margin of the Congo-Tanzania Craton during the Irumidian orogeny (ca. 1020–950 Ma), together with terranes in the Southern Irumide Belt. The distinctly older Nampula Complex was more probably linked to the Maud Belt of Antarctica, and peripheral to the Kalahari Craton during the Neoproterozoic. During the Pan-African orogeny, the Marrupa Complex was overlain by NW-directed nappes of the Cabo Delgado Nappe Complex before peak metamorphism at ca. 555 Ma. The nappes include evidence for early Pan-African orogenic events older than 610 Ma, typical for the Eastern Granulites in Tanzania. Crustal thickening at 555 ± 11 Ma is coeval with high-pressure granulite-facies metamorphism along the Lurio Belt at 557 ± 16 Ma. Crustal thickening in NE Mozambique is part of the main Pan-African, Kuunga, orogeny peaking between 570 and 530 Ma, during which the Congo-Tanzania, Kalahari, East Antarctica and India Cratons welded to form Gondwana. Voluminous post-collisional magmatism and metamorphism younger than 530 Ma in the Lurio Belt and the Nampula Complex are taken as evidence of gravitational collapse of the extensive orogenic domain south of the Lurio Belt after ca. 530 Ma. The Lurio Belt may represent a Pan-African suture zone between the Kalahari and Congo-Tanzania Craton. © 2009 Elsevier B.V. All rights reserved.

∗ Corresponding author. Tel.: +47 73 90 4240; fax: +47 73 92 1620. E-mail address: [email protected] (B. Bingen). 0301-9268/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2009.01.005

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1. Introduction Gondwana was assembled along a network of Pan-African– Brasiliano orogenic belts during the Ediacaran (635–542 Ma) and Cambrian (542–488 Ma; Fig. 1a; Stern, 1994; de Wit et al., 2001; Meert, 2003; Boger and Miller, 2004; Jacobs and Thomas, 2004; Collins and Pisarevsky, 2005). Geochronological data attributed to Pan-African high-grade metamorphism in eastern and southern Africa, India, Madagascar, Sri Lanka, Arabia, and Antarctica range from ca. 800 to 470 Ma (compilation in Meert, 2003), clearly indicating that assembly of Gondwana was a polyphase process. Three main frequency maxima are apparent in the data, at ca. 760–730, 660–610 and 570–530 Ma, respectively. The two first maxima are typical of a segment of the N–S trending East African Orogen (Stern, 1994) extending from the Arabian–Nubian Shield to Tanzania (Fig. 1a). They are attributed to progressive closure of the Mozambique Ocean, including collisions between volcanic arcs in the Mozambique Ocean and possibly a first collision between the Congo-Tanzania Craton and a minor continent, including Madagascar (“Azania” in Collins and Pisarevsky, 2005). These events are referred to as the East African Orogeny by Meert (2003). The third event (570–530 Ma) is more widespread and is attributed to final assembly of Gondwana. It is referred to as the Kuunga orogeny by Meert (2003), and involved almost simultaneous welding of the main austral cratons, namely the Indian, East Antarctica, Kalahari, and Congo-Tanzania Cratons (Jacobs et al., 1998, 2003a; Hanson, 2003; Meert, 2003; Boger and Miller, 2004; Collins and Pisarevsky, 2005; Johnson et al., 2005; Jöns and Schenk, 2008; Grantham et al., 2008). The Kuunga orogeny probably includes closure of an oceanic basin between the Kalahari and Congo-Tanzania Cratons along the E–W trending Damara-Lufilian-Zambezi Orogen (Fig. 1a; John et al., 2003; Johnson et al., 2005; Gray et al., 2008). Improved understanding of the dynamics of Gondwana assembly requires improved characterization of individual transects through Pan-African orogenic belts. NE Mozambique is situated at an interesting location in a wide Pan-African orogenic domain situated at the intersection between the East African-Antarctic Orogen and the Damara-Lufilian-Zambezi Orogen (Fig. 1). This domain

is commonly regarded as part of the Mozambique Belt, the central segment of the East African-Antarctic Orogen (Holmes, 1951). The geology of Mozambique was remapped between 2002 and 2006 in a Nordic Development Fund/World Bank funded project, which integrated structural, geochemical, petrological, geophysical and geochronological data. The project results for the NE part of country (ca. 260,000 km2 , area of Fig. 2) are summarized in two unpublished reports (Norconsult Consortium, 2007a,b) and are partly reported in thematic publications (Melezhik et al., 2006, 2007, 2008; Engvik et al., 2007; Key et al., 2007, 2008; Jacobs et al., 2008a; Viola et al., 2008; Bjerkgård et al., 2009) and this paper. In this paper, the U–Pb zircon and monazite data are reported on the felsic basement, which underlies allochthonous Pan-African nappes. The dataset includes data collected by secondary ion mass spectrometry (SIMS) and laser ablation inductively coupled plasma mass spectrometry (LA-ICPMS). The data contribute to characterization of the Pan-African and pre-Pan-African geologic events in NE Mozambique. Especially widespread are the Grenville-aged, i.e. Stenian (1200–1000 Ma) to Tonian (1000–850 Ma), magmatic and metamorphic events, variably referred to as Kibaran, Irumidian and Namaquan in various parts of Africa. The objective of this paper is to (1) link and correlate lithotectonic units in NE Mozambique with adjacent lithotectonic units in the Gondwana context, and speculate on their possible ancestries; (2) map the distribution of East African and Kuunga-aged high-grade metamorphism at the intersection between the East African and the Damara-Lufilian-Zambezi Orogens; (3) evaluate late-Pan-African gravitational collapse; and (4) assess the WSW-ENE-trending Lurio (shear) Belt in NE Mozambique as a potential Pan-African suture between the Kalahari and Congo-Tanzania Cratons. 2. Geological context The Mozambique Belt is exposed in NE Mozambique between the Indian Ocean coastal plain, where it is overlain by Jurassic to Neogene strata of the Rovuma basin (Key et al., 2008), and Lake Niassa, along the Neogene Malawi rift. It is segmented by two major structures, the Mesozoic Maniamba Graben and the Pan-African

Fig. 1. (a) Cambrian reconstruction of Gondwana following Meert (2003), with the location of the study area in NE Mozambique. (b) Sketch map of NE Mozambique and adjacent areas showing the location of the main lithotectonic complexes in their Gondwana context. Gondwana reconstruction following Torsvik et al. (2008). Note the distinct Africa-India fit between the two figures.

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Fig. 2. Simplified geological map of NE Mozambique (Norconsult Consortium, 2007a,b) with a summary of U–Pb geochronological data. Lines inside the complexes represent boundaries of lithological units or faults.

Lurio Belt (Fig. 2). The Maniamba Graben is a SW–NE trending Karoo-age rift filled with Permian-Jurassic sediments (Melezhik et al., 2007; Key et al., 2007). The southern shoulder of the graben exposes a SW–NE-trending shear zone array, the Macaloge shear zone (Fig. 2). The Karoo-age graben may conceal an important preKaroo tectonic zone. The Lurio Belt is a prominent, remarkably linear, ca. 600 km long, WSW–ENE-trending, NNW-dipping, Pan-African tectonic zone (Pinna et al., 1993; Sacchi et al., 2000; Engvik et al., 2007; Viola et al., 2008). It is cored by an attenuated and dismembered complex of granulites, highly sheared gneisses and mylonites, less than 25 km wide, referred to as the Ocua Complex. In the east, the Lurio Belt is well defined by extreme flattening of lithological units. Towards the west, it becomes increasingly structurally diffuse but marked by a linear zone of widespread Pan-African felsic magmatism. The Ocua Com-

plex becomes increasingly discontinuous and lenticular westwards (Viola et al., 2008). The Precambrian bedrock has been divided into a number of lithotectonic complexes (Fig. 2; Norconsult Consortium, 2007a,b), partly overlapping in definition and extent with the stratigraphic entities previously defined by Pinna et al. (1993). These complexes can be grouped into four distinct mega-units each with a distinct geological evolution. (1) The Palaeoproterozoic Ponta Messuli Complex is exposed to the NW of the Maniamba Graben. (2) The Unango and Marrupa Complexes represent Stenian–Tonian felsic crust lying between the Maniamba Graben and the Lurio Belt. In the south, they are progressively reworked towards and into the Lurio Belt and their southern boundary is defined by slivers of rocks of the Ocua Complex. (3) The Nampula Complex represents Stenian felsic crust south of the Lurio Belt. (4) The Stenian–Tonian basement (units 2 and 3 above) is overlain by a set of Pan-African nappes. To

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the north of the Lurio Belt, the Nairoto, Meluco, Muaquia, M’Sawize, Xixano, Lalamo and Montepuez Complexes, collectively referred to as the Cabo Delgado Nappe Complex, overlie the Marrupa Complex (Viola et al., 2008). To the south of the Lurio Belt, the Mugeba and Monapo klippen, overlie the Nampula Complex (Sacchi et al., 1984, 2000; Grantham et al., 2008). In this publication, U–Pb data are reported on the felsic basement complexes (units 1–3). Data on the Cabo Delgado Nappe Complex (unit 4) will be reported elsewhere. 3. Methods Samples for U–Pb geochronology were collected from the typical and the less common lithologies of each complexes (Table 1). Data were collected on zircon and monazite, separated from crushed samples by classical methods of mineral separation, including water table, heavy liquids and magnetic separation. A selection of crystals were mounted in epoxy and polished. Cathodoluminescence (CL) images of zircon and backscattered electron (BSE) images of monazite were obtained with a scanning electron microscope (SEM), to characterize zoning. SIMS U–Th–Pb data were collected at the Museum of Natural History, Stockholm, using a CAMECA 1270 instrument and at the Curtin University of Technology, Perth, using a SHRIMP II instrument. Data are reported in Appendix A. Analytical protocols and quality control are summarized in Whitehouse et al. (1999) and Whitehouse and Kamber (2005) for the CAMECA instrument and in Williams (1998) and Nemchin and Pidgeon (1997) for the SHRIMP instrument. The analyses were calibrated relative to the Geostandard 91,500 reference zircon (1065 Ma; Wiedenbeck et al., 1995) for the CAMECA instrument and relative to the TEMORA reference zircon (417 Ma; Black et al., 2003) for the SHRIMP instrument. SIMS analyses were performed on crystals embedded in epoxy and polished (1 ␮m diamond polish). Concordia diagrams and age calculations were performed with the ISOPLOT program (Ludwig, 2001), using the following approach. For most samples, data were corrected for common Pb using the 204 Pb signal. For concordant data, a concordia age (Ludwig, 1998) was calculated whenever possible. For discordant data, an intercept age was calculated. For some of the samples, the data were not corrected for common Pb. These are samples for which SIMS analyses yield negligible 204 Pb signal (four samples analysed with SHRIMP) and one sample with lowU zircon (sample 31997). In this last case, uncorrected data were regressed with a common Pb anchor point (modern isotopic composition) to calculate an intercept age (207 Pb correction method). LA-ICPMS analyses were collected at the Geological Survey of Norway using a Finnigan Element I instrument fed by a 266 nm laser microsampler (Appendix A). The procedure is summarized in Bingen et al. (2005). LA-ICPMS analyses were done by rastering the laser beam over an area of ca. 40 ␮m × 60 ␮m at the surface of the sample, after a gentle surface cleaning ablation. Data were corrected for common Pb using the 204 Pb analysis corrected for Hg interference. The measured isotope ratios are corrected for element- and mass-bias effects using the 91,500 reference zircon (10–20 analyses at the start of each session and at regular intervals between unknowns). Ablation-related elemental fractionation of up to several % is difficult to avoid with the instrumentation used in this study. As a result, analyses of a zircon or monazite population generally define a linear segment in the concordia diagram, regressing through the origin. The slope of this regression or the weighted average 207 Pb/206 Pb age is statistically robust above ca. 10 analyses. As a consequence, the weighted average 207 Pb/206 Pb age is regarded as the most reliable estimate of the crystallization a population. To monitor accuracy, analyses of at least two independent reference samples (10–20 analyses/sample) encompassing the age of the unknowns were performed at the start of each session. These include reference zircon GJ1 (608.5 ± 1.5 Ma; Jackson et al., 2004),

for which an accurate 207 Pb/206 Pb age of 605 ± 10 Ma was obtained in the course of this study (Fig. 3a), and monazite B0109 (Bamble, Norway, 1137 ± 1 Ma) for which an age of 1137 ± 7 Ma was obtained (Bingen et al., 2008a). These also include reference zircons from Seiland (Finnmark, Norway, 515 ± 4 Ma; Pedersen et al., 1989), Sjona (Norland, Norway, ØS9914, 1797 ± 3 Ma; Skår, 2002), and Ontario (Z6412, 1160 ± 2 Ma, unpubl. analyses GSC Ottawa). Analyses of most samples were distributed over more than one analytical session. LA-ICPMS analyses were performed either on crystals embedded in epoxy and polished (6 ␮m diamond polish for optimum laser ablation), or crystals fixed on double sided tape. For part of the samples, both methods were used and results are indistinguishable. The second method proves efficient in the analysis of zircons with a core–rim structure. Analyses collected sequentially (generally two or three) in the same ablation pit penetrate the rim and then the core. The analyses generally show a spread in a concordia diagram (e.g. samples 33296, 33508 and 33502) and define a bimodal distribution, one mode for analyses of the rim (first ablation) and another mode for analyses of the core (subsequent ablations). Analyses sampling the rim–core interface spread between the two modes. In the following section, the U–Pb geochronological data are reported in a concise fashion in their geological context. Samples are located in Fig. 2, and concordia diagrams are presented in Figs. 3–11, together with information on the samples and data. As a general rule, analyses of prismatic oscillatory-zoned zircon are attributed to magmatic events. For zircon populations showing a core–rim structure, as observed on CL images, analyses of the core are generally attributed to the protolith, either magmatic zircon in an orthogneiss or detrital zircon in a paragneiss, while analyses of the rim are attributed to a high-grade metamorphic event. In all samples, a few outlying individual analyses (driving MSWD significantly above 1.0) were not selected for weighted age calculation (white ellipses in Figs. 3–11). The most common reasons for outliers are sampling of inclusions or fractured domains during laser ablation. In the following section, all valid analyses on all analysed samples are reported. These include analyses on samples with generally low-U zircon, providing imprecise dates only interpretable as reconnaissance values. 4. Ponta Messuli Complex The Ponta Messuli Complex is exposed over a small area along the shores of Lake Niassa NW of the Maniamba Graben (Fig. 2). It consists of amphibolite- to granulite-facies migmatitic gneisses and orthogneisses, commonly augen gneisses, with minor amphibolites. A migmatitic gneiss of pelitic composition, sample 31874, contains a granulite-facies assemblage of cordierite + garnet + sillimanite + spinel + biotite, replacing a first metamorphic assemblage including garnet porphyroblasts. The sample contains zircon and monazite. The simplest zircon crystals consist of a core, presumably detrital, surrounded by a metamorphic rim. Crystals with several zones are common. Three SIMS analyses in cores give 207 Pb/206 Pb ages ranging from 2199 ± 21 to 2074 ± 6 Ma (Fig. 3b). Five analyses from composite crystals range from 2012 ± 80 to 1977 ± 16 Ma, and could date detrital or metamorphic zircon crystallization. A conservative estimate for the maximum age of deposition of the sedimentary precursor is thus provided by the youngest detected zircon core of unambiguous detrital origin at 2074 ± 11 Ma (Fig. 3b). Analyses from weakly concentrically zoned zircons at 1950 ± 15 Ma, and analyses of monazites at 1946 ± 11 Ma (Fig. 3b and c) date highgrade metamorphism, presumably corresponding to development of the granulite-facies assemblage and conspicuous migmatitization observed in this sample. A weakly deformed circular granite pluton, known as the Cobué granite, intrudes the Palaeoprotero-

Table 1 Summary of sampling and U–Pb geochronological data in NE Mozambique including published data. Sample

Map sheet

2

31818 31779 31874 26859 26892 26802 33205 33406 33508 33503 33502 33498 33405 33252 33253 34284 33206 33591 31225 33509 WB295 31260 33499 33586 33598 31997 33572 33510 31965 22782 31863 31229 34273 34288 31261 33507 31870 31971 40403 22780 33512 33312 26878 40744 33310 JJ238 JJ259 GV01 33373 40710

Lupilichi Metangula Ponta Messuli Malema Malema Ribaue Mecula Marrupa Cuamba Cuamba Cuamba Cuamba Marrupa Marrupa Marrupa Majune Mecula Milange Lichinga Cuamba Gurue Lichinga Cuamba Gurue Milange Macaloge Milange Mandimba Metangula Majune Metangula Lichinga Majune Lichinga Lichinga Cuamba Macaloge Macaloge Majune Lichinga Cuamba Mecufi Ribaue Mecufi Mecufi Gurue Gurue Ribaue Montepuez Mecufi

Lithology

Xenocryst/detrital Zrn Ma 3

1135 1234 1134 1437 1437 1438 1237 1337 1436 1436 1436 1436 1337 1337 1337 1336 1237 1635 1335 1436 1536 1335 1436 1536 1635 1235 1635 1435 1234 1336 1234 1335 1336 1335 1335 1436 1235 1235 1336 1335 1436 1340 1438 1340 1340 1536 1536 1438 1339 1340

Txitonga Ponta Messuli Ponta Messuli Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Marrupa Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Unango Ocua Ocua Ocua Ocua Nampula Nampula Nampula Nampula Nampula

Rhyolite Granite Metapelitic gneiss Granodioritic gneiss Syenitic gneiss Granitic gneiss Leucogranite Leucosome Mangeritic gneiss Granitic gneiss Granitic gneiss Leucogneiss Granitic gneiss Granitic gneiss Granitic gneiss Granitic gneiss Granitic gneiss Augen gneiss Granite Quartz mangerite Quartz mangerite Granite Granite Leucogneiss Amphibolite Syenitic gneiss Banded granulite, felsic Charnockite Granitic gneiss Charnockite Leucogneiss Granitic gneiss Quartz mangeritic gneiss Quartz monzonitic gneiss Banded granulite Quartz mangerite gneiss Granodioritic gneiss Leucogneiss Quartz monzonitic gneiss Metapelitic gneiss Banded granulite Boudin-neck pegmatite Mylonitic leucogneiss Mafic granulite Banded gneiss, felsic Granitic gneiss Phenocryst granite Phenocryst granite Phenocryst granite Granitic gneiss

Zrn Ma

Intrusion

Metamorphism

Zrn Ma

Zrn Ma

714 1056 2199

735

2488

1092

2596

±21

to

2074

±12

±14

±8

to

1062

486 504 521 547

±27 ±11 ±15 ±14

753

±13

946 968 1005 1011 1016 1025 1026 493 501 512 512 514 517 519 753 799 827 949 975 991 1008 1013 1032 1034 1037 1039 1040 1047 1062

±11 ±10 ±19 ±16 ±10 ±12 ±9 ±35 ±29 ±36 ±4 ±35 ±12 ±6 ±12 ±8 ±20 ±13 ±31 ±16 ±9 ±10 ±10 ±14 ±10 ±11 ±8 ±8 ±13

±54 532 538

1950

±15

549 549

±13 ±22

569

±16

942

1946

±11

558

±11

547

±16

±21

548

±5

569

±9

535

±11

952

±16

962 945 957 551 951 952 972

±18 ±33 ±27 ±6 ±35 ±38 ±47

536

±6

557

±16

530

±26

946 527

±100 ±8

±13 ±10 540

936

±53

507 508 508 511 511

±3 ±4 ±2 ±12 ±12

Method

Ref.

Fig.

4

5

6

ICPMS ICPMS CAMECA ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS SHRIMP ICPMS ICPMS SHRIMP SHRIMP CAMECA SHRIMP ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS ICPMS CAMECA ICPMS ICPMS ICPMS ICPMS CAMECA ICPMS ICPMS ICPMS ICPMS SHRIMP CAMECA CAMECA ICPMS ICPMS

f a a a a a a a a a a a a a a a a a a a d a a a a a a a a a a a a a a a a a a a a e a c e d d d a a

Mnz Ma

±17 ±11

±11

±20

Mnz Ma

±7

529

±10

Fig. 3d Fig. 3b and c Fig. 8e Fig. 8f Fig. 8a Fig. 7h Fig. 7a Fig. 7c Fig. 6c Fig. 6b Fig. 4c Fig. 4b Fig. 5f–h Fig. 4a Fig. 5c Fig. 3h Fig. 8d Fig. 7g Fig. 8g Fig. 7f Fig. 8b Fig. 8c Fig. 7e Fig. 7b Fig. 7d Fig. 3g Fig. 5d and e Fig. 3f Fig. 5b Fig. 3e Fig. 4d Fig. 4h Fig. 4e Fig. 6a Fig. 4g Fig. 5a Fig. 4f Fig. 6d and e Fig. 6f and g

B. Bingen et al. / Precambrian Research 170 (2009) 231–255

1

Complex/group

Fig. 8h

Fig. 10a Fig. 10b 235

±19 517 543 1019 493 520

±11 ±23 ±18 ±8 ±8

513

Mnz Ma

Granitic–syenitic gneiss Granitic gneiss Augen gneiss Granitic gneiss Granitic gneiss Quartz diorite gneiss Leucogneiss

Zrn Ma

1042 1048 1057 1060 1072 1087 1123

Zrn Ma Zrn Ma 3

±9 ±1 ±9 ±17 ±8 ±16 ±9

Zrn Ma

Metamorphism Intrusion Xenocryst/detrital Lithology

2

Mecufi Mocuba Montepuez Ribaue Mocuba Milange Milange

1

40781 672GJ3 33296 26879 33566 33568 33564

1340 1636 1339 1438 1636 1635 1635

Nampula Nampula Nampula Nampula Nampula Nampula Nampula

Complex/group Map sheet Sample

Table 1(Continued)

(1) Samples: samples ordered with ascending intrusion age for each complex; complexes listed from NW to SE; (2) one square degree map sheet numbered with upper left corner, for example map 1338 has upper left corner at 13◦ S–38◦ E; (3) Zrn: zircon, Mnz: monazite; (4) analytical methods: ICPMS: LA-ICPMS; CAMECA: SIMS CAMECA 1270; SHRIMP: SIMS SHRIMPII; Evapor: Kober method TIMS; (5) reference: (a) this work; (b) Kröner et al. (1997); (c) Engvik et al. (2007); (d) Jacobs et al. (2008a,b); (e) Viola et al. (2008); (f) Bjerkgård et al. (2009); (6) figure with concordia diagram.

Fig. 9d and e Fig. 9c Fig. 9b Fig. 9g Fig. 9a

a b a a a a a

6 5

ICPMS Evapor ICPMS CAMECA ICPMS SHRIMP SHRIMP

Mnz Ma

4

Ref. Method

Fig. 9f

B. Bingen et al. / Precambrian Research 170 (2009) 231–255

Fig.

236

zoic basement. Sample 31779 from this pluton yields an intrusion age of 1056 ± 11 Ma (Fig. 3d). The Ponta Messuli Complex is overlain by the low-grade Txitonga Group (the Cobué Group of Pinna et al., 1993), along a N–S-trending, E-dipping tectonized contact. The Txitonga Group consists mainly of metagreywacke, metasandstone and micaschist, metamorphosed under greenschist to lower-amphibolite-facies conditions. High-Ti metagabbro and greenschist bodies associated with minor rhyolite flows attest to bimodal magmatism coeval with development of the basin. A conformable rhyolite flow gives a zircon crystallization age of 714 ± 17 Ma (Bjerkgård et al., 2009), an age which constrains both deposition of the Txitonga Group to be Cryogenian (850–635 Ma), and deformation/low-grade metamorphism to be Pan-African. 5. Unango and Marrupa Complexes The Unango Complex is exposed to the south of the Maniamba Graben and extends into Malawi to the southwest (Fig. 2; Bloomfield, 1968; Kröner et al., 2001). The Marrupa Complex lies to the east of the Unango Complex, and extends into Tanzania to the NE (Kröner et al., 2003). In the north of the mapped area (Milepa area, Fig. 2), the Marrupa Complex is seen to overlie the Unango Complex along a well-defined, NW–SEtrending, NE-dipping, tectonic contact. In the south (Cuamba area), the contact between the two complexes is poorly defined, and it becomes increasingly diffuse southwards, towards the Lurio Belt. The Unango Complex mainly consists of amphibolite- to granulite-facies, variably foliated orthogneisses, generally of felsic to intermediate composition, including augen gneisses. Extensive units of quartz-mangeritic to charnockitic orthogneiss are described (mainly around and north of Lichinga; Fig. 2) and are regarded as orthopyroxene-bearing magmatic rocks with charnockitic affinity. The granitoids are generally alkali-calcic to shoshonitic, magnesian to ferroan in composition, and have variable I-, A- and, uncommonly, S-type geochemical signatures (Norconsult Consortium, 2007b). There is a conspicuous lack of low-K lithologies, except for minor enderbites and mafic rocks. The Unango Complex contains conformable amphibolite- to granulitefacies paragneiss units between orthogneiss bodies. The relative age relationship between paragneisses and orthogneisses cannot be established from field relations, though paragneiss enclaves in metaplutonic rocks suggest that the paragneisses are at least locally older. Ediacaran low-grade carbonate-rich metasedimentary rocks, known as the Geci Group (Fig. 2), are reported in three tectonic lenses hosted in the Unango Complex (Pinna et al., 1993; Melezhik et al., 2006). Depositional primary structures are locally well preserved, but elsewhere the carbonates are highly deformed and contacts with the enclosing orthogneisses are tectonic. Strontium and carbon isotopic data in carbonates constrain deposition of the sequence from sea-water within two possible time intervals at 630–625 or 595–585 Ma (Melezhik et al., 2006). The Marrupa Complex consists generally of amphibolite-facies felsic orthogneisses, associated with minor volume of migmatitic paragneisses and metasediments. Weakly peraluminous granite to granodiorite gneisses dominate, though low-K tonalitic gneisses are common. Orthogneisses range from I- to A-type granite signatures. Charnockitic orthogneisses are recorded only locally. Cryogenian granite to syenite plutons are present in the Unango and Marrupa complexes (Lulin, 1984). Cambrian plutons are divided in two geographically distinct suites, the Niassa and Malema suites (Norconsult Consortium, 2007b).

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Fig. 3. U–Pb geochronological data. The figures include, for each sample, the sample number, the lithology, a concordia diagram (1 error ellipses on data points and selected age provided at 2 level), the analytical method + the number of analysed crystals, a representative CL image, the location of the sample including the map sheet number (one degree sheet 1338 has the upper left corner at 13◦ S–38◦ E) and UTMWGS84 coordinates, some key whole-rock geochemical parameters (SiO2 , K2 O and Zr), the list of minerals present in the rock (in decreasing order of abundance) and some comments on the texture of the rock.

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Fig. 4. U–Pb geochronological data.

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Fig. 5. U–Pb geochronological data.

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Fig. 6. U–Pb geochronological data.

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Fig. 7. U–Pb geochronological data.

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Fig. 8. U–Pb geochronological data.

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Fig. 9. U–Pb geochronological data.

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Fig. 10. U–Pb geochronological data.

5.1. Stenian–Tonian orthogneisses 5.1.1. Orthogneisses without metamorphic zircon/monazite Zircon and/or monazite were analysed from thirteen samples of orthogneiss from the Unango Complex, and nine samples from the Marrupa Complex. Seven of these samples show no evidence for metamorphic zircon rims or monazite, though five of them have an amphibolite-facies foliation. In the Unango Complex, granitic gneiss 31229 gives a magmatic intrusion age of 1013 ± 10 Ma (Fig. 3e). Two charnockites, without apparent foliation, have intrusion ages of 991 ± 16 Ma (sample 22782, Fig. 3f) and 949 ± 13 Ma (sample 33510, Fig. 3g). In the Marrupa Complex, three granitic gneiss samples yield intrusion ages at 1026 ± 9 Ma (samples 33206, Fig. 3h), 1016 ± 10 Ma (sample 33253, Fig. 4a) and 1005 ± 19 Ma (sample 33405, Fig. 4b), and one leucogneiss at 968 ± 10 Ma (sample 33498, Fig. 4c).

Fig. 11. Relative probability curve summarizing U–Pb geochronological data for high-grade metamorphism and magmatism in the Marrupa, Unango and Nampula Complexes. For the Nampula Complex, two curves are proposed, one for the study area only, and one for the whole exposure of the complex. Data from this work, Kröner et al. (1997), Costa et al. (1994), Engvik et al. (2007), Grantham et al. (2008) and Jacobs et al. (2008a). The time bracket for deposition of the marine Geci Group follows Melezhik et al. (2006).

5.1.2. Orthogneisses with Tonian metamorphic zircon Seven samples of the Unango Complex and one sample from the Marrupa Complex, all of them situated north of 14◦ S latitude, show Tonian metamorphic zircon. Sample 34273 from the Unango Complex is a quartz-mangeritic orthogneiss showing foliation-parallel aggregates of orthopyroxene + clinopyroxene + hornblende + biotite, fringed by fine-grained garnet. This rock is interpreted as a magmatic charnockite, subsequently deformed under high-grade conditions, leading to formation of garnet. Zircon crystals show a core–rim structure with a thick metamorphic rim commonly forming replacement textures. Oscillatory-zoned cores yield an intrusion age of 1032 ± 10 Ma while the rims yield an age of 962 ± 18 Ma, interpreted as the timing of metamorphism (Fig. 4d). Sample 31261 is a granulite showing cm-scale banding. Layers of intermediate composition contain an equilibrium medium-pressure granulite-facies assemblage of plagioclase + quartz + Kfeldspar + biotite + hornblende + orthopyroxene + garnet, while the felsic layers lack orthopyroxene. Zircons show a core–rim structure. Magmatic oscillatory-zoned cores have an age of 1037 ± 10 Ma, while metamorphic rims have a distinctly younger age of 957 ± 27 Ma (Fig. 4e). The metamorphic rims probably relates to formation the granulite-facies assemblage. Six other granitic to charnockitic samples have a similar geochronological signature to samples 34273 and 31261, though the age separation between core and rim is statistically less significant. Sample 40403, is a weakly deformed quartz monzonite, with K-feldspar megacrysts partly mantled by plagioclase (rapakivi texture), and containing minor metamorphic garnet. Zircon cores define an intrusion age of 1062 ± 13 Ma and rims an imprecise age

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of 972 ± 47 Ma (Fig. 4f). Sample 31870 is a K-feldspar megacrystic granodioritic gneiss collected in a weakly deformed unit. Zircon cores are dated at 1040 ± 8 Ma and rims at 951 ± 35 Ma (Fig. 4g). The rock displays well-preserved magmatic textures, implying that metamorphism was not associated with pervasive deformation. Sample 34288 is a weakly deformed quartz monzonite. Zircons show a thin metamorphic rim. Zircon cores provide an age of 1034 ± 14 Ma, and the rim an age of 945 ± 33 Ma (Fig. 4h). Sample 31971 is a megacrystic leucogneiss, from the NW margin of the Unango Complex. Zircon cores have an age of 1047 ± 8 Ma and the rims an age of 952 ± 38 Ma (Fig. 5a). Sample 31863 is a leucogneiss layer in a tightly banded amphibolite-facies sequence, also from the NW margin of the Unango Complex. The zircon cores are dated at 1008 ± 9 Ma and the rims at 952 ± 16 Ma (Fig. 5b). Sample 34284 is an amphibolite-facies biotite-bearing granitic gneiss collected in the central part of the Marrupa Complex. The sample contains zircons with a core–rim texture. The rim generally grew at the expense of the core, as evident from convex replacement textures. Magmatic oscillatory-zoned cores yield an age of 1025 ± 12 Ma, while the rims are distinctly younger at 942 ± 21 Ma (Fig. 5c). 5.1.3. Orthogneisses with Ediacaran–Cambrian metamorphic zircon or monazite Five samples show Ediacaran–Cambrian metamorphic zircon or monazite. In the NW of the Unango Complex, sample 31965 is a slightly peraluminous, muscovite-bearing, amphibolite-facies granitic gneiss, showing a partial cataclasis. The sample contains a poorly abundant zircon population with oscillatory- to sectorzoned low-U magmatic cores surrounded by U-rich metamorphic rims. It also contains few entirely metamorphic neoblasts of U-rich zircon and monazite. Zircon cores yield an imprecise intrusion age of 975 ± 31 Ma (Fig. 5d). Analyses of the metamorphic rim scatter around 500–550 Ma. One analysis of a zircon neoblast is concordant at 535 ± 11 Ma, and provides the best age estimate for the metamorphic overprint (Fig. 5d). The monazite age of 530 ± 26 Ma (Fig. 5e) supports this interpretation. Sample 33252 is a slightly peraluminous garnet–muscovitebearing granitic gneiss collected in the center of the Marrupa Complex (Marrupa area, Fig. 2). Magmatic oscillatory-zoned zircon yields an intrusion age of 1010 ± 16 Ma (Fig. 5f). Zircon crystals in this sample lack metamorphic rims. Metamorphic monazite, however, yields an age of 547 ± 16 Ma (Fig. 5g), implying Pan-African amphibolite-facies metamorphism. Four analyses in one xenocrystic zircon crystal give an age of 2488 ± 12 Ma (Fig. 5h), suggesting the occurrence of earliest Palaeoproterozoic material in the source of this orthogneiss. In the southern part of the Unango Complex (Cuamba area, Fig. 2), sample 33507 is a weakly foliated quartz mangeritic gneiss. Though it lacks orthopyroxene, it contains mesoperthitic feldspar megacrysts, suggesting that the magmatic protolith of this rock was a hypersolvus granite. SIMS data on zircon cores define an intrusion age of 1039 ± 11 Ma (Th/U = 1.1), while analyses of the wide rim yields an age of 551 ± 6 Ma (Th/U = 0.55) for the Pan-African amphibolite-facies or possibly granulite-facies metamorphic overprint (Fig. 6a). In the southern part of the Marrupa Complex, sample 33502 is a biotite–amphibole granitic gneiss. LA-ICPMS analyses of zircon with a core–rim structure spread in the concordia diagram, with a bimodal distribution (Fig. 6b). The oldest mode defines the age of the magmatic core at 946 ± 11 Ma and the youngest mode yields the age of the metamorphic rim at 569 ± 16 Ma. Analyses spreading between the two modes are interpreted as mixtures of core and rim material, or alternatively may record partial re-crystallization of the core. Another sample from the same area, sample 33503, is a granitic gneiss similar to other granitic gneisses formed during the 1026–946 Ma interval. Zircon in this sample is fractured

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and/or metamict and was not analysed. Monazite is present. Monazite crystals commonly show weak concentric to patchy zoning on BSE images (Fig. 6c). Nevertheless, analyses of monazite are well clustered and give a crystallization age of 558 ± 11 Ma (Fig. 6c), recording amphibolite-facies metamorphism. 5.2. Paragneisses The most extensive paragneiss unit in the Unango Complex crops out along the border with Malawi. This is the Metangula Group of Pinna et al. (1993), renamed Chala gneiss (Fig. 2; Norconsult Consortium, 2007b). It consists of commonly mafic migmatitic banded gneiss, with minor quartzite and metapelite layers. Sample 22780 is a metapelite from the Chala gneiss, containing a coarse-grained, granulite-facies assemblage of sillimanite + Kfeldspar + garnet. The sample contains small zircons with a detrital core (not analysed) and metamorphic monazite. Monazite crystals are apparently homogeneous on BSE images (Fig. 6d). Analyses define a (scattered) discordia line, which indicates that some monazite crystals includes at least two age domains. An upper intercept age of 946 ± 100 Ma can be calculated from the data, providing an approximate age for the oldest age domains (Fig. 6d). Analyses around the lower intercept form a coherent group (MSWD = 1.3) with an average 207 Pb/206 Pb age of 529 ± 10 Ma and define the dominant monazite crystallization event (Fig. 6d and e). Consequently, available data provide evidence for two high-grade metamorphic events, at ca. 946 and 529 Ma respectively, but do not allow a decision as to which event is responsible for the granulite-facies assemblage.Sample 33512 is a felsic layer in a thin unit of banded granulite in the southern part of the Unango Complex (Cuamba area). The outcrop is characterized by widespread garnet blastesis. The sample contains zircon crystals with a variably zoned detrital core, surrounded by a commonly thick metamorphic rim. Sector-zoned metamorphic zircon neoblasts are abundant, as well as monazite. Seven SIMS analyses in detrital zircon cores range from 2596 ± 8 to 1062 ± 54 Ma (Fig. 6f), pointing to a varied Neoarchaean to Mesoproterozoic sourcing for the sedimentary protolith. Deposition of this clastic sediment took place after 1062 ± 54 Ma, the age of the youngest analysed detrital grain. Zircon rims and neoblasts give a Pan-African age of 536 ± 6 Ma (Fig. 6g), probably dating the formation of the granulite-facies assemblage. Monazite yields an equivalent crystallization age of 527 ± 8 Ma (Fig. 6h). Sample 33406 is an approximately 10-cm wide, coarse-grained, foliation-parallel leucosome, collected from an amphibolite-facies, biotite-rich, banded-gneiss outcrop in the central part of the Marrupa Complex. It was originally collected to test for Irumidian partial melting. Prismatic, high-U zircon in this sample, however, yields a crystallization age of 548 ± 13 Ma (Fig. 7a). This age is interpreted as recording the crystallization of the leucosome and the timing of migmatitization. 5.3. Cryogenian magmatic rocks The ca. 5 km long Monte Chissindo pluton is situated close to the NW margin of the Unango Complex in the shoulder of the Maniamba Graben. It is composed of foliated quartz syenite to silica-undersaturated nepheline syenite, and hosts subeconomic Nb–Ta mineralization. Sample 31997 is a quartz-free, scapolite-bearing syenite gneiss collected from the pluton. It is heterogranular and partly cataclastic. It contains abundant U-poor zircon. (Uncorrected) SIMS analyses of oscillatory- to sector-zoned zircon yield an intercept age of 799 ± 8 Ma, interpreted as the intrusion age of the pluton (Fig. 7b). Sample 33508, collected in the southern part of the Marrupa Complex, is a mesoperthitic amphibole-bearing felsic orthogneiss with an A-type chemical signature. Prismatic oscillatory-zoned zir-

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con is surrounded by a thin metamorphic rim. LA-ICPMS data show a bimodal age distribution with a first mode at 753 ± 13 Ma corresponding to analyses of the Magmatic core, and a second mode at 549 ± 22 Ma corresponding to analyses of the metamorphic rim (Fig. 7c). Analyses sampling the rim–core interface spread between these two modes. Sample 33508 represents a Cryogenian pluton with probable charnockitic affinity. The sample is similar in many respects to the Stenian–Tonian (1026–943 Ma) host gneisses, and consequently field data are inconclusive as to the extent and limits of this pluton and at larger scale to this Cryogenian magmatism. Two samples collected close to the Lurio Belt provide data that are difficult to interpret in any unambiguous way. Sample 33572 is a low-K, fine-grained, felsic layer in a granulite-facies banded gneiss, possibly representing an orthogneiss of volcanic origin or a paragneiss. It contains zircon with a core–rim structure. SIMS analyses of U-rich zircon cores define a discordant cluster of analyses with an average 238 U/206 Pb age of 827 ± 20 Ma (Th/U = 0.39, Fig. 7d). One analysis of an inherited oscillatory-zoned core is close to concordia at 1092 ± 14 Ma (207 Pb/206 Pb age, Th/U = 0.64). If it has any geological significance, the age of 827 ± 20 Ma represents a minimum age for crystallization of the magmatic protolith, or a minimum age for deposition of the sediment protolith. The thick U-poor zircon rim provides a robust concordia age at 569 ± 9 Ma (Th/U = 0.02) recording the granulite-facies overprint in this sample (Fig. 7d). Sample 33598 is an equigranular amphibolite collected close to sample 33572. It shows cm-scale pods and layers of leucosome. The sample contains rounded, sector-zoned zircons, probably related to migmatitization. SIMS analyses of these zircons provide a concordia age at 548 ± 5 Ma, recording the timing of metamorphism and migmatitization (Fig. 7e). A high-U, probably metamict, discordant core yields a 238 U/206 Pb age of 753 ± 12 Ma. No geological event is attributed to this date, as this core may be exotic to the amphibolite protolith (e.g. Möller et al., 2003). 5.4. Cambrian Niassa Suite Unfoliated (post-kinematic) subcircular plutons and ring complexes of alkaline granite and syenite, referred to as the Niassa Suite, are exposed in the central part of the Unango Complex south of Lichinga (Fig. 2). Samples 31260 and 31225 are representative of two typical granite plutons a few km in diameter. LA-ICPMS data on low-U zircon yield intrusion ages of 514 ± 35 and 501 ± 29 Ma respectively (Fig. 7f and g). Though imprecise, the data demonstrate a Cambrian, i.e. late-Pan-African, age for this magmatism. In the central and northern part of the Marrupa Complex, Pan-African plutons are uncommon, although sample 33205, collected in a minor body of unfoliated biotite–hornblende granite, yields an intrusion age of 547 ± 14 Ma (Fig. 7h).

The Malema Suite also includes large sub-circular concentrically zoned plutons, commonly monzonitic to charnockitic. They are well delineated on maps of airborne magnetic and radioelement data. Though they can be regarded as post-kinematic and their core is generally unfoliated, these plutons are often elongated parallel to the WSW–ENE regional fabric in the Lurio Belt. Sample 33509 was taken from the >50 km long Serra Messolua pluton. This large zoned charnockitic pluton straddles the boundary between the Unango and Marrupa Complexes. The sample is a quartz mangerite collected from close to its margin. It contains abundant low-U zircon, from which LA-ICPMS analyses provide an intrusion age of 512 ± 36 Ma (Fig. 8g). This age is imprecise, but equivalent to the age of 513 ± 4 Ma obtained by SIMS on a smaller and similarly charnockitic pluton located some 50 km along strike (Fig. 2, sample WB295; Jacobs et al., 2008a). 6. Nampula Complex The extensive Nampula Complex lies south of the Lurio Belt and underlies the Unango and Marrupa complexes below this structure (Fig. 2). The boundary with the Lurio Belt corresponds to a WSW-ENE trending, conspicuous negative magnetic anomaly, which closely matches the outcrop of a variably sheared leucogneiss (Viola et al., 2008). The mylonitic facies of this leucogneiss is assigned to the Ocua Complex. Sample 26878 was collected from a thin mylonitic sheet of the leucogneiss (thus assigned to the Ocua Complex), and characterized by a strong NW dipping planar fabric largely defined by flattened quartz ribbons. The sample contains prismatic, U-rich, magmatic zircon without metamorphic rim and thus no apparent record of the mylonitization event. The zircon dates intrusion of the leucogranite protolith at 538 ± 10 Ma (Fig. 8h), which also provides a maximum age for mylonitization. The Nampula Complex itself consists of Stenian amphibolitefacies (locally granulite-facies) gneisses, which are generally migmatitic and felsic, intruded by Cambro-Ordovician plutons. It displays a broad regional WSW–ENE-trending structural grain. In synforms, the gneisses are locally overlain by Ediacaran–Cambrian clastic metasediments, referred to as the Mecuburi and Alto-Benfica Groups (Fig. 2). These groups include amphibolite-facies metasandstone, metaconglomerate and a diagnostic sillimanite-rich nodular gneiss (Norconsult Consortium, 2007b). In this study, only the northern margin of the Nampula Complex was mapped and sampled. The geology south of the study area is described in Sacchi et al. (1984), Costa et al. (1992) and Grantham et al. (2008). The regional nomenclature, updated by Grantham et al. (2008), is adopted here. 6.1. Stenian gneisses

5.5. Cambrian Malema Suite Abundant Pan-African syn-kinematic to post-kinematic felsic plutons, referred to as the Malema Suite, are exposed along the Lurio Belt. The syn-kinematic plutons may be difficult to distinguish from the hosting Mesoproterozoic gneisses. As a result, they are difficult to delineate with accuracy. Six samples of variably foliated plutonic rocks were dated. Samples 26802, 33499, 33586, and 33591 are foliated granites with intrusion ages of 521 ± 15, 517 ± 12, 519 ± 6, and 493 ± 35 Ma respectively (Fig. 8a–d). Sample 26859 is a highly deformed granodioritic gneiss, in which zircons show a core–rim structure. Analyses of cores and rims are statistically indistinguishable, and provide an imprecise pooled age of 486 ± 27 Ma (Fig. 8e). This age approximates the magmatic crystallization of the rock. Sample 26892 represents a small (<2 km) deformed body of quartz-free syenite, which gives an intrusion age of 504 ± 11 Ma (Fig. 8f).

The oldest recorded rocks in the Nampula Complex are variably banded and migmatitic orthogneisses of granodioritic, tonalitic to leucogranitic composition, referred to as the Mocuba Suite. Two samples of this suite have been dated with the U–Pb method south of the study area at 1148 ± 1 and 1129 ± 9 Ma (Kröner et al., 1997; Grantham et al., 2008). The Mocuba Suite is structurally overlain by a supracrustal sequence known as the Molócuè Group. This Group is mainly made up of banded/migmatitic biotite gneiss, interpreted as paragneiss, interlayered with fine-grained leucogneiss, possibly representing metarhyolite, and locally quartzite, metapelite, marble and calc-silicate gneiss. South of the study area, two leucocratic orthogneiss samples, interlayered with the Molócuè Group, gave ages of 1115 ± 1 and 1090 ± 22 Ma (Kröner et al., 1997; Grantham et al., 2008). This suggests that the Mocuba Suite can be viewed as a slightly older basement upon which the Molócuè Group was deposited.

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Sample 33564 is a leucocratic, migmatitic granitic gneiss of the Mocuba Suite, in the study area (Fig. 2). Zircon in the sample shows a core–rim texture. SIMS analyses define a scattered discordia line. The six oldest cores yield a mean 207 Pb/206 Pb age of 1123 ± 9 Ma, regarded as the timing of intrusion of the granite protolith (Fig. 9a). One single analysis of a zircon rim is concordant at 520 ± 8 Ma, and provides an estimate for the timing of amphibolite-facies migmatitization. The Nampula Complex contains voluminous felsic orthogneisses, forming sub-concordant, sheet-like plutons, which are less deformed and less pervasively migmatitic than the Mocuba Suite and Molócuè Group. Two suites of such gneisses are described by Grantham et al. (2008). The tonalitic Rapale Gneiss is restricted in extent. The extensive Culicui Suite shows, generally, a highK, ferroan, A-type geochemical signature, though low-K rocks are also reported (Norconsult Consortium, 2007b). Intrusion of the Rapale gneiss is dated by two samples at 1095 ± 8 and 1091 ± 14 Ma (Grantham et al., 2008). Ten orthogneiss samples of the Culicui Suite, collected south of the study area, gave a range of dates from 1092 ± 42 to 1028 ± 7 Ma (Costa et al., 1994; Kröner et al., 1997; Grantham et al., 2008). Five high-K orthogneiss samples typical of the Culicui Suite, distributed all along the northern margin of the Nampula Complex, were analysed in this study. Sample 33566 is an amphibole-bearing granitic orthogneiss. Zircons in this sample are characterized by a core surrounded by a, generally thin, high-U metamorphic rim. Cores define an intrusion age of 1072 ± 8 Ma (Fig. 9b). Four analyses in two crystals yield a distinctly younger age at 1019 ± 18 Ma. This age probably records formation of the rim and arguably provides evidence for a Stenian amphibolite-facies event. No Ediacaran–Cambrian age signature is recorded in this sample. Sample 26879 is a K-feldspar megacrystic granitic gneiss, containing zircon with a thin U-rich metamorphic rim. SIMS data define a discordia line (Fig. 9c). The upper intercept age at 1060 ± 17 Ma is constrained by zircon cores and interpreted as the magmatic intrusion of the gneiss. The lower intercept age at 543 ± 23 Ma is mainly constrained by one analysis of a thin rim, and yields an imprecise estimate for Pan-African metamorphism. Sample 33296 is a slightly peraluminous garnet-bearing Kfeldspar megacrystic granite gneiss. LA-ICPMS analyses of zircon with core–rim structure show a spread in the concordia diagram. The age of the magmatic core is estimated at 1057 ± 9 Ma, by pooling the oldest analyses. It dates intrusion of the megacrystic granite (Fig. 9d). The youngest available analysis at 517 ± 11 Ma dates the metamorphic rim. Monazite gives an equivalent age of 513 ± 19 Ma and dates amphibolite-facies metamorphism (Fig. 9e). Sample 40781 is a granitic to syenitic gneiss, collected from an inselberg in the coastal plain in the NE corner of the Nampula Complex (Fig. 2). The sample is heterogranular, and was presumably cataclased during Mesozoic faulting. Magmatic prismatic zircon yields an intrusion age of 1042 ± 9 Ma for the syenite (Fig. 9f). No zircon rims are detected. Sample 33568 is a migmatitic, plagioclase megacrystic, low-K, quartz diorite gneiss, tentatively attributed to the Culicui Suite (Fig. 2). SIMS analyses define a discordia line with intercepts at 1087 ± 16 and 483 ± 39 Ma (Fig. 9g). The upper intercept age is controlled by zircon cores of magmatic origin, and reflects intrusion of the protolith of the orthogneiss. Four concordant analyses performed on zircon rims define a concordia age at 493 ± 8 Ma (Fig. 9h), interpreted as the best estimates for the timing of amphibolitefacies metamorphism and partial melting. 6.2. Cambrian Murrupula Suite The Nampula Complex exposes abundant Pan-African synto post-kinematic high-K granitoids forming plutons, dykes and

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sheets of the Murrupula Suite. Deformed (syn-kinematic) granitoids are difficult to distinguish from pre-Pan-African gneisses. Zircon dates from ten plutons, ranging from weakly deformed to unfoliated, and located south of the study area, are reported by Grantham et al. (2008). They range in age between 533 ± 5 and 495 ± 2 Ma. In the present study area, zircon data on three samples are reported by Jacobs et al. (2008a), namely a sheared orthogneiss (sample JJ238), a small phenocrystic granite sheet (sample JJ259) and a large (>60 km long) weakly deformed phenocrystic granite pluton roughly aligned parallel to the regional fabric (Ribáuè pluton, sample GV01; Fig. 2). The three samples have ages of 507 ± 3, 508 ± 4 and 508 ± 2 Ma respectively. Two samples of small plutons are reported here. Sample 33373 is a porphyritic granite with a weak, possibly magmatic, fabric, and sample 40710 a weakly deformed leucocratic granitic gneiss. Both samples have equivalent intrusion ages at 511 ± 12 Ma (Fig. 10a and b). 7. Discussion 7.1. Palaeoproterozoic orogenic events in NE Mozambique To the NW of the Maniamba Graben, the Ponta Messuli Complex represents a Palaeoproterozoic metasedimentary basement, affected by a granulite-facies metamorphism at 1950 ± 15 (Fig. 3b and c). To the SE of the Maniamba Graben, no Palaeoproterozoic rocks have been identified in the Unango and Marrupa Complexes. The occurrence of one earliest Palaeoproterozoic inherited zircon in one peraluminous granitoid sample (sample 33252, Fig. 5h) and the Hf isotopic signature of zircon in several samples (unpublished data) nevertheless attest to the presence of an older Palaeoproterozoic to Archaean crustal component also in parts of the Unango and Marrupa crust. The Ponta Messuli Complex resembles typical crust of the Usagaran and Ubendian Belts forming the foreland of the PanAfrican orogens along the Congo-Tanzania Craton (Lenoir et al., 1994; Ring et al., 1997; Möller et al., 1998; Johnson et al., 2005; De Waele et al., 2006a,b). In northern Malawi and Zambia, the Usagaran–Ubendian Belt is characterized by magmatic rocks intruded between 2093 ± 1 and 1930 ± 30 Ma and associated with cordierite-garnet granulites (Ring et al., 1997; Vrána et al., 2004). Syn-kinematic granites in these granulites bracket metamorphism and deformation between 1995 ± 1 and 1969 ± 1 Ma. In Tanzania, the Usagaran Belt locally hosts eclogite-facies rocks, metamorphosed between 1999 ± 2 and 1977 ± 14 Ma (Möller et al., 1995; Reddy et al., 2003). A younger metamorphic age of 1925 ± 24 Ma is also recorded in Tanzania (Sommer et al., 2005). 7.2. Stenian–Tonian orogenic events in NE Mozambique NE Mozambique shows widespread evidence for Stenian– Tonian events. The Ponta Messuli Complex is intruded by granite plutons, one of them dated at 1056 ± 11 Ma (Fig. 3d). The Unango Complex is mainly made up of a felsic orthogneiss basement, commonly charnockitic, and crystallized between 1062 ± 13 and 949 ± 13 Ma (13 samples, Fig. 2), with a peak of magmatic activity at ca. 1040 Ma (Figs. 11 and 12a). The Marrupa Complex comprises a slightly younger felsic orthogneiss basement formed between 1026 ± 9 and 946 ± 11 Ma, with peak activity at ca. 1025 Ma (7 samples, Figs. 11 and 12a). The few available detrital zircon data indicate that paragneisses are younger than 1062 ± 54 Ma (sample 33512, Fig. 6f), but do not provide a relative chronology between the ortho- and paragneisses. The scattered, generally high-K, Ito A-type geochemical signatures of the granitoids in both complexes (Pinna et al., 1993; Norconsult Consortium, 2007b) do not permit a precise interpretation of their geotectonic setting. No significant geochemical-petrological-temporal trend is apparent.

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The Stenian–Tonian magmatism in the Unango and Marrupa Complexes is, however, extremely voluminous and represents the main crustal growth event. As a result, the most probable geotectonic environment for such felsic magmatic activity is in a continental volcanic arc setting. Available detrital zircon data demonstrate that Stenian–Tonian metasediments in the Unango Complex (sample 33512, Fig. 6f) were sourced in Archaean to Paleoproterozoic crust, suggesting deposition of the basins at the margin of an evolved continent. U–Pb data on zircon rims and monazite crystals demonstrate Tonian high-grade metamorphism north of the 14◦ S in the Unango and Marrupa Complexes (Fig. 2). Nine dates between 972 ± 47 and 942 ± 21 Ma yield a robust average value of 953 ± 9 Ma for this metamorphism (Figs. 11 and 12b). At least one sample, characterized by metamorphic banding, medium-pressure granulite-facies assemblage and zircon rims at 957 ± 27 Ma (sample 31261, Lichinga area, Fig. 4e), strongly suggests that this metamorphism was associated with ductile deformation, and reached granulite-facies conditions, though no regional scale characterization of structures is possible. This metamorphism is coeval with the intrusion of the final pulse of Tonian felsic magmatism in both complexes. It may not have been the product of a collisional orogeny, but may represent the thermal perturbation following extensive magmatic activity. The Nampula Complex consists of plutonic and volcanic rocks formed between 1148 ± 1 and 1028 ± 7 Ma (Costa et al., 1994; Kröner et al., 1997; Grantham et al., 2008). Data from the present study, collected along the northern margin of the Nampula Complex, range from 1123 ± 9 to 1042 ± 9 Ma (Figs. 2, 11 and 12a). The several magmatic suites recognized within this time interval reflect a change from low-K calc-alkaline signature (ca. 1150–1090 Ma Mocuba and Rapale Suites) to more varied high-K signatures (ca. 1090–1030 Ma Culicui Suite; Grantham et al., 2008; this study). The low-K geochemical signature of the early suites suggests a comparatively primitive volcanic arc geotectonic setting at 1150–1090 Ma, evolving to a more mature arc setting after 1090 Ma. The existence of Stenian high-grade metamorphism is not easy to confirm in the Nampula Complex, due to the pervasive Cambrian overprint (Fig. 12b). Field observations suggest that a phase of deformation and migmatitization occurred between deposition of the Molócuè Group and intrusion of the Culicui Suite. Notwithstanding the difference in composition and rheology between these rock types, and the overlap in the geochronological data, these observations may suggest an age of ca. 1090 Ma for deformation and migmatitization (Grantham et al., 2008). Formation of zircon rims at 1090 ± 34 Ma in one tonalite sample and crystallization of a leucosome at 1063 ± 47 Ma, reported by Grantham et al. (2008), support this interpretation. Data from the present study, nevertheless do not provide additional evidence in favor of this interpretation. Instead, a few zircon analyses provide some evidence for a metamorphic event at 1019 ± 18 Ma (sample 33566, Fig. 9b), i.e. after intrusion of the Culicui Suite. Further analytical work is necessary to clarify this point. 7.3. Stenian–Tonian correlations in their Gondwana context 7.3.1. The Congo-Tanzania Craton perspective The Congo-Tanzania Craton was affected by the Irumidian orogeny in Zambia, forming the Irumide Belt (Fig. 12a and b). This orogeny is characterized by NW-directed tectonic transport under amphibolite-facies condition between 1020 ± 7 and 1004 ± 20 Ma, by voluminous syn- to post-kinematic high-K granite plutonism, mainly between 1055 ± 13 and 1005 ± 7 Ma, and by a final magmatic pulse at 943 ± 5 Ma (De Waele et al., 2006a,b). Evidence for the Irumidian orogeny fades away northeastwards along the CongoTanzania margin. In Malawi, minor A-type granite plutons have

Fig. 12. Summary sketch maps showing the distribution of (a) Stenian–Tonian magmatism, (b) Stenian–Tonian metamorphism and (c) Ediacaran–Cambrian metamorphism and felsic plutonism in NE Mozambique and adjacent areas. References quoted in text.

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been dated between 1119 ± 20 and 1087 ± 11 Ma (Ring et al., 1999), while in Tanzania, Irumidian overprint is lacking. The Southern Irumide Belt lies to the south of the Irumide Belt in Zambia and Malawi (Fig. 12a and b; Johnson et al., 2005, 2006). It displays a Pan-African overprint and thus is regarded as part of the Damara-Lufilian-Zambezi Orogen. The Southern Irumide Belt is characterized by widespread Mesoproterozoic magmatism ranging from 1094 ± 2 to 1023 ± 12 Ma, and possibly continuing to 929 ± 1 Ma (Goscombe et al., 2000; Kröner et al., 2001; Johnson et al., 2006, 2007b; Mänttäri, 2008). Locally, it shows evidence for a Palaeoproterozoic history, including whole-rock Sm–Nd isotopic signatures, xenoliths, xenocrysts, detrital zircons (2100–1700 Ma), and a possible early metamorphic event at 1942 ± 5 Ma (Johnson et al., 2006; Kröner et al., 2001). In Zambia, the Southern Irumide Belt has been divided into several terranes, separated by approximately N–S trending Pan-African shear zones. The easternmost Chipata terrane has granite plutonism dated between 1076 ± 6 and 1038 ± 6 Ma, and underwent granulite-facies metamorphism at 1047 ± 20 Ma (Johnson et al., 2006). East of the Chipata terrane, granulites are widespread in southern Malawi (Andreoli, 1984), but their age (Irumidian vs. Pan-African) remains controversial (Kröner et al., 2001). In terms of their Palaeo- to Mesoproterozoic crustal evolution, therefore, the Unango and Marrupa Complexes can possibly be regarded as the eastern continuation of the Southern Irumide Belt. All show diffuse evidence for Paleaoproterozoic to Archaean crustal components. The widespread Mesoproterozoic felsic plutonism in the Unango and Marrupa Complexes (ca. 1062–946 Ma; Figs. 11 and 12a, b) is broadly similar and coeval with magmatism in the Southern Irumide Belt (1094–929 Ma). The Unango and Marrupa Complexes show robust evidence for an Irumidian amphibolite- to granulite-facies metamorphism dated at 953 ± 9 Ma. This metamorphism is, however, significantly younger than reported in the Southern Irumide Belt (ca. 1050 Ma) and in the Irumide Belt (ca. 1020–1005 Ma). At a larger scale, it is part of the late Grenville-aged orogenic evolution (Boger et al., 2000; Küster et al., 2008). The Irumidian event in the Irumide Belt and Southern Irumide Belt is tentatively interpreted by Johnson et al. (2005, 2006) and De Waele et al. (2006b, 2008) as representing accretion of continental arcs (as exposed in the Southern Irumide Belt) to the margin of the Congo-Tanzania Craton at around 1020 Ma. These continental arcs were attached to an unknown continent or microcontinents before accretion. The data on the Unango and Marrupa Complexes reported in this study are not incompatible with such models, but the discovery of a granulite-facies metamorphic event at 953 ± 9 Ma in the Unango Complex suggests a more complex history involving polyphase Irumidian metamorphism. Alternative models picturing the Stenian–Tonian SE margin of the Congo-Tanzania Craton as a long-lived variably compressional active continental margin are equally possible. The craton margin was intruded by voluminous felsic magmatism between ca. 1095 and 930 Ma, with at least one event of foreland propagation at 1020–1005 Ma (the Irumidian orogeny). Irumidian structures were largely overprinted during the Pan-African orogeny and consequently all of these models are highly speculative. Nevertheless, importantly, all models place the Southern Irumide Belt, and the Unango–Marrupa Complexes at the margin of the Congo-Tanzania Craton before the Pan-African orogeny. The Stenian magmatism in the Nampula Complex (ca. 1148–1028 Ma) is distinctly older than in the Unango and Marrupa Complexes (ca. 1062–946 Ma) and in the Southern Irumide Belt (ca. 1094 and 929 Ma; Fig. 11 and 12a). This may suggest that the Nampula Complex formed independently, as a continental arc at the margin of another continent, or as an exotic volcanic arc. Nevertheless, the timing of docking remains unconstrained.

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From a Congo-Tanzania perspective, the Nampula Complex can be interpreted as an exotic terrane assembled to the margin of the Congo-Tanzania Craton during the Irumidian orogeny (Johnson et al., 2005; De Waele et al., 2008). These protracted Mesoproterozoic histories are typical for the Grenvillian–Sveconorwegian orogenic belts in Laurentia and Baltica (Rivers and Corrigan, 2000; Bingen et al., 2008b). 7.3.2. The Kalahari Craton perspective The Kalahari Craton (Jacobs et al., 2008b) is covered, in Antarctica, by the Mesoproterozoic, almost unmetamorphosed rocks of the Ritscherflya Supergroup (Fig. 1a), dated by volcanic horizons at 1130 ± 7 Ma (Frimmel, 2004). The craton also includes a Large Igneous Province manifested as intraplate mafic magmatism at 1112–1106 Ma (Umkondo Suite; Hanson et al., 2006). The Kalahari Craton is bordered in the south by the Namaqua-Natal Belt (Fig. 1a; Jacobs et al., 1997; Cornell and Thomas, 2006; Eglington, 2006; McCourt et al., 2006). The eastern part of this belt, the Natal Sector (which lies closest to Mozambique), consists almost entirely of Stenian juvenile magmatic rocks formed between 1235 ± 9 and 1025 ± 8 Ma, with the main collisional event taking place at ca. 1135 Ma (Namaquan orogeny). It represents one of the very few areas of Mesoproterozoic crust in this part of Gondwana lacking significant Pan-African overprint. The Maud Belt in Antarctica is regarded as an extension of the Namaqua-Natal Belt and shows intense Pan-African overprint east of the Heimefront Shear Zone (Fig. 1a; Jacobs et al., 2003b). Stenian magmatic activity in the Maud Belt ranges from 1171 ± 25 to 1045 ± 9 Ma (Arndt et al., 1991; Jacobs et al., 1998, 2003b; Paulsson and Austrheim, 2003; Bisnath et al., 2006; Board et al., 2005). Stenian high-grade metamorphism is detected, but age estimates scatter between ca. 1095 and 1025 Ma (same references; Fig. 12a and b). The similarity in the geology and geochronology between the Nampula Complex and the Maud Belt, especially as seen in Central Dronning Maud Land, is well established in the literature (Jacobs et al., 1998, 2003a, 2008a; Jacobs and Thomas, 2004; Manhica et al., 2001; Grantham et al., 2008). Stenian magmatic activity in the Nampula Complex (ca. 1148–1028 Ma; Figs. 11 and 12a, b) largely overlaps with that in the Maud Belt. From a Kalahari Craton perspective, therefore, the Nampula Complex and the NamaquaNatal-Maud Belt can be linked into one single Namaquan orogenic belt, characterized by assembly of terranes at the margin of the Kalahari Craton and tectonic transport onto the Archaean to Palaeoproterozoic nucleus of the Kalahari Craton (Groenewald et al., 1991; Jacobs et al., 1998, 2003b, 2008b; Fitzsimons, 2000; Manhica et al., 2001; McCourt et al., 2006). Consequently, the Nampula Complex can be probably restored to the margin of the Kalahari Craton before the Pan-African orogeny. 7.4. Cryogenian extension After the Irumidian orogeny, an extensional stress regime arose in the Congo-Tanzania Craton (De Waele et al., 2008). The SE margin of the craton, largely located inside the Damara-Lufilian-Zambezi Orogen, exposes extensive Neoproterozoic basins, interpreted as rift and passive margin sequences, which locally host magmatic rocks (Porada and Berhorst, 2000; Key et al., 2001; Johnson et al., 2005). A first phase of rifting included deposition of the Zambezi Supracrustal Sequence, floored by a 880 ± 14 Ma rhyolite, in the Zambezi Belt (Johnson et al., 2007a), and the Roan sequence in the Lufilian Arc. A second phase included deposition of the Kundelungu sequence in the Lufilian Arc, dated by volcanic horizons at 735 ± 5 Ma (Key et al., 2001). Small Cryogenian granitoid plutons are reported in the Zambezi and Southern Irumide Belts (Johnson et al., 2006; Ashwal et al., 2007).

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Data from NE Mozambique are consistent with this largescale picture. Two distinct Neoproterozoic cover sequences were deposited, presumably unconformably, on the felsic basement. To the NW of the Maniamba Graben, the Cryogenian clastic Txitonga Group overlies the Ponta Messuli Complex and includes bimodal magmatism at 714 ± 17 Ma (Bjerkgård et al., 2009). It is interpreted as a continental rift sequence, and it is broadly coeval with the Kundelungu sequence. To the southeast of the Maniamba Graben, the Ediacaran carbonate-bearing Geci Group is interpreted as remnants of a marine carbonate platform deposited between ca. 630 and 585 Ma on the Unango Complex (Pinna et al., 1993; Melezhik et al., 2006). The metamorphic contrast between the low-grade Geci Group and its underlying high-grade gneiss basement implies unroofing and erosion of the basement between 953 ± 9 Ma (metamorphism) and ca. 630 Ma (carbonate deposition). Cryogenian plutonic rocks are sporadic in the Unango and Marrupa Complexes, but absent from the Nampula Complex. A charnockitic pluton dated at 753 ± 13 Ma (Sample 33508, Fig. 7c) demonstrates Cryogenian magmatic activity in the Marrupa Complex, although its volume is poorly constrained. The Unango Complex hosts a volumetrically limited suite of nepheline syenitebearing plutons. The suite includes the Monte Chissindo and the Monte Naumale plutons (Fig. 2), dated at 799 ± 8 Ma (Sample 31997, Fig. 7c) and 830 ± 130 Ma (zircon intercept age, Lulin, 1984) respectively. Small plutons of deformed nepheline syenite are reported in the Zambezi and Southern Irumide Belts (Ashwal et al., 2007). They are generally considered to be Neoproterozoic in age. The Tambani pluton in Malawi is dated at 730 ± 4 Ma (Ashwal et al., 2007). Such silica-deficient alkaline plutonism is diagnostic of continental rift settings. Together, the Tambani and Monte Chissindo plutons add to the growing body of evidence for an extensional stress regime at the margin of the Congo-Tanzania Craton, between ca. 800 and 730 Ma. 7.5. Ediacaran–Cambrian orogenic events in NE Mozambique The Pan-African tectonic evolution in NE Mozambique has been divided into four main stages (Viola et al., 2008). (1) Nappes of the Cabo Delgado Nappe Complex were assembled. The nappes carry evidence for early Pan-African metamorphic events dated between ca. 735 and 610 Ma and Neoproterozoic magmatism ranging from ca. 820 to 600 Ma (Norconsult Consortium, 2007b). (2) The nappes were transported towards the NW over the Mesoproterozoic felsic basement (Marrupa Complex). (3) As a continuation of the same shortening process, folding of the felsic basement together with the Cabo Delgado Nappe Complex took place along NE–SW fold axes. Mafic boudins in the Ocua Complex within the Lurio Belt show evidence for high-pressure (HP) granulites-facies conditions, estimated at 1.57 GPa–950 ◦ C. A zircon age of 557 ± 16 Ma records the timing of the HP event (Fig. 2; Engvik et al., 2007). (4) Convergence was followed by NW–SE directed extension, dated by an interboudin leucosome filling in the Ocua Complex (Fig. 2) at 532 ± 13 Ma and younger in the Lurio Belt (Viola et al., 2008). Geochronological data on Pan-African events in the felsic basement are summarized in Figs. 2, 11 and 12c). High-grade metamorphism is dated between 569 ± 16 and 547 ± 16 Ma in the Marrupa Complex (five samples). The data define a monomodal distribution with an average value at 555 ± 11 Ma. In the Unango Complex, data range from 569 ± 9 to 527 ± 8 Ma (six samples) and in the Nampula Complex from 543 ± 23 to 493 ± 8 Ma (four samples). If data reported in Grantham et al. (2008) are included, the age range for the Nampula Complex extends from 555 ± 12 to 490 ± 8 Ma. The three suites of Pan-African felsic plutons, the Niassa, Malema and Murrupula Suites, are dated between 547 ± 14 and 486 ± 27 Ma, including the data by Grantham et al. (2008) and Jacobs et al. (2008a). The geochronological data reported here contribute to

the regional tectonic model with the following five important constraints. (1) The oldest recorded Pan-African high-grade metamorphism in the three Mesoproterozoic basement complexes is ca. 555 Ma (Fig. 11), matching the age of 557 ± 16 Ma for HP granulitefacies metamorphism in the Ocua Complex (Engvik et al., 2007). The ca. 555 Ma event is interpreted as representing the peak of crustal thickening during Pan-African convergence in NE Mozambique, after deposition of the marine Geci Group (ca. 630–585 Ma) (Melezhik et al., 2006). Maximum burial of the felsic gneiss basement structurally post-dates transport of the nappes of the Cabo Delgado Nappe Complex, and consequently represents a minimum age for nappe transport. With the presently available data, transport of the nappes is bracketed between 596 ± 11 Ma, the age of the youngest pluton specific to the nappes (Norconsult Consortium, 2007b), and 555 ± 11 Ma, the timing of amphibolite-facies metamorphism in the underlying Marrupa Complex. The metamorphic overprint at ca. 555 Ma, typical of the felsic basement, is not recorded in the nappes of the Cabo Delgado Nappe Complex, in accordance with the observation that these nappes were situated at a higher tectonic level than the basement at ca. 555 Ma. (2) The Ponta Messuli Complex and its cover, the Txitonga Group, are not affected by high-grade Pan-African metamorphism, as evident from the zircon and monazite U–Pb data. In the Unango and Marrupa Complexes, Pan-African metamorphism is increasingly pervasive, southwards, towards the Lurio Belt, progressively obliterating evidence of the earlier Irumidian metamorphism. High-grade Pan-African metamorphism was detected only in a minority of the samples collected north of 14◦ S, but is present in most samples collected south of 14◦ S. The conformable leucosome layer dated at 549 ± 13 Ma (sample 33406, Fig. 7a) nevertheless demonstrate that PanAfrican migmatitization and deformation took place, at least in biotite-rich banded lithologies, north of 14◦ S in the Marrupa Complex. Metamorphic monazite at 547 ± 16 Ma in sample 33252 (Fig. 5g) supports the contention that Pan-African metamorphism reached amphibolite-facies conditions on a regional scale. The coexistence of two monazite generations in metapelite 22780 (Fig. 6d) also demonstrates superposition of Irumidian and Pan-African high-grade metamorphism in the Unango Complex north of 14◦ S (Lichinga area). Two samples of banded granulite, situated along the southern margin of the Unango Complex contain metamorphic zircon and monazite between 569 ± 9 and 527 ± 8 Ma and demonstrate that PanAfrican metamorphism peaked in granulite-facies conditions in the environs of the Lurio Belt. (3) In the Marrupa, Unango and Nampula Complexes, U–Pb dates on high-grade Pan-African metamorphism show a clear geographic-tectonostratigraphic trend. Metamorphism in the Marrupa Complex is dated at 555 ± 11 Ma. Increasingly younger estimates are recorded in the Unango Complex (535–527 Ma range) and in the Nampula Complex (520–493 Ma range, Figs. 2 and 11). In the latter, zircon rims provide a robust age of 493 ± 8 Ma for amphibolite-facies migmatitization in sample 33568 (Fig. 9h). An equivalent age of 490 ± 8 Ma is reported for a leucosome by Grantham et al. (2008). Field observations place the Marrupa Complex structurally above the Unango Complex (at least in their northern part), and both the Unango and Marrupa Complexes above the Nampula Complex along the NNW-dipping Lurio Belt. The youngest metamorphism is thus recorded in the structurally lowermost unit. This trend suggests that the lowermost unit was the last to be exhumed. (4) Pan-African felsic plutonism shows a frequency peak at around 510 Ma. This magmatism is post-collisional, as it post-dates

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peak metamorphism related to crustal thickening at ca. 555 Ma. The volume of Pan-African felsic magmatism is minor in the northern and central part of the Marrupa Complex. It increases in the Unango Complex (Niassa Suite), and increases suddenly southwards, along the Lurio Belt (Malema Suite) and in the Nampula Complex (Murrupula Suite). This trend implies widespread crustal melting within, and south of, the Lurio Belt. It is coeval with dated evidence for extensional tectonics at 532 ± 13 Ma along the belt (Viola et al., 2008). (5) Along the Lurio Belt, poorly deformed (post-kinematic) granite plutons overlap in time with highly sheared granitic gneisses (syn-kinematic plutons). For example, sample 26802 is a weakly deformed granite sheet dated at 521 ± 15 Ma (Fig. 8a), while sample 26859, collected from a few km across strike, is part of another sheet with a strong NNW dipping planar fabric. It gave an overlapping age of 486 ± 27 Ma (Figs. 2 and 8 Figs. 2 and 8e). Thus the available age data suggest that deformation along the Lurio Belt was strongly partitioned. The data furthermore suggest that deformation was coeval with granite magmatism and was probably lubricated by melts. The leucogneiss marking the boundary between the Nampula Complex and the Lurio Belt yields an intrusion age of 538 ± 10 Ma (sample 26878, Fig. 8h). This gneiss is highly sheared, implying that it accommodated relative shearing between the Nampula Complex and Unango and Marrupa Complexes along the Lurio Belt after 538 ± 10 Ma. The sense of shear of this unit could not be confirmed by field observations. Nevertheless, age relationships (youngest metamorphism at 493 ± 8 Ma is in the Nampula Complex, sample 33568, Fig. 9h) suggest that shearing along the leucogneiss was extensional and partly accommodated by post538 Ma exhumation of the Nampula Complex to the south of the NNW dipping Lurio Belt. 7.6. East African vs. Kuunga orogeny Along the E–W-trending Damara-Lufilian-Zambezi Orogen (Fig. 12c), Pan-African metamorphism ranges from 573 ± 2 to 522 ± 17 Ma in the Lufilian, Zambezi and Southern Irumide Belts, with a frequency maximum at around 530 Ma (Goscombe et al., 2000; Kröner et al., 2001; John et al., 2004; Ashwal et al., 2007; Johnson et al., 2007b). Whiteschists in the Southern Irumide and Lufilian Arc demonstrate convergence-related highpressure amphibolite-facies conditions (1.3 GPa–750 ◦ C) at 532 ± 2 to 525 ± 2 Ma (Johnson and Oliver, 2002; John et al., 2004). The Lufilian Arc shows NE-directed tectonic transport (Porada and Berhorst, 2000). Syn- to post-kinematic felsic plutonism includes the large Hook granite pluton intruded between 566 ± 5 and 533 ± 3 Ma in the Zambezi Belt (Fig. 1; Hanson et al., 1993), along with abundant syenite to granite plutons intruded between 543 ± 6 and 479 ± 9 Ma in the Southern Irumide Belt (Johnson et al., 2006). The Damara-Lufilian-Zambezi Orogen is a typical product of the Kuunga orogeny. Geochronological data on Pan-African events support a link between the Unango Complex and the Southern Irumide Belt. These include prolonged high-grade metamorphism ranging from 569 ± 9 to 527 ± 8 Ma and post-kinematic granite to syenite plutonism (the Niassa Suite) at around 510–500 Ma (Fig. 2). The transition between the Southern Irumide Belt and its Pan-African foreland, the Irumide Belt, is largely concealed under a Karoo-aged graben in Zambia. In the same way, the transition between the Unango Complex and the Pan-African foreland is concealed under the Maniamba Graben. Along the N–S-trending East African Orogen, in Tanzania, allochthonous nappes, known as the Eastern Granulites, were transported towards the NW onto the Archaean to Palaeoproterozoic felsic crust of the Congo-Tanzania Craton (Fig. 12c; Muhongo, 1994; Fritz et al., 2005). Pan-African metamorphism in the nappes

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peaked in granulite-facies conditions between ca. 655 and 610 Ma (Möller et al., 2000; Muhongo et al., 2001; Sommer et al., 2003; Tenczer et al., 2006). In the underlying felsic crust, the timing of metamorphism is contentious. One sample of granulite (in the Western Granulites) provides a metamorphic age of 640 ± 1 Ma (Sommer et al., 2003), while data in high-pressure amphibolitefacies whiteschists and in localized shear zones range from ca. 582 to 533 Ma (Cutten et al., 2006; Hauzenberger et al., 2007). In this crust, there is also evidence for Archaean high-grade metamorphism (Johnson et al., 2003). Final thrusting of the nappes is dated at around 550 Ma using 40 Ar/39 Ar geochronology (Rossetti et al., 2008). Post-collisional magmatism younger than ca. 540 Ma is absent in Tanzania. In Madagascar, a Neoachaean basement is intruded by Cryogenian, ca. 830–720 Ma, plutons interpreted as part of a continental arc, and by Ediacaran, ca. 630 Ma, “stratoid” granite plutons (Paquette and Nédélec, 1998; Tucker et al., 1999; Handke et al., 1999; Kröner et al., 2000). Dates for Pan-African amphibolite-to granulite-facies metamorphism range from ca. 650 to 520 Ma (Kröner et al., 2000; de Wit et al., 2001). In central Madagascar, the dominant Pan-African tectonic transport is towards the east and it took place before ca. 546 Ma (Nédélec et al., 2003; Tucker et al., 2007). In NE Mozambique, Pan-African nappes of the Cabo Delgado Nappe Complex cover the Marrupa Complex (Fig. 12c). They may also have overlain the Unango Complex (Viola et al., 2008). Though the Marrupa Complex (Stenian–Tonian) is younger than the felsic basement in Tanzania (Archaean-Palaeoproterozoic), the Cabo Delgado Nappe Complex shares a number of similarities with the Eastern Granulites of Tanzania (Muhongo, 1994) and the Vohibory Complex in Madagascar (Fig. 1b; de Wit et al., 2001; Jöns and Schenk, 2008). These similarities include preservation of early Pan-African metamorphism older than 610 Ma, occurrence of mafic granulites, volcanic arc lithologies and characteristic marble units in the metasediment packages (Norconsult Consortium, 2007b; Melezhik et al., 2008). The nappes in Tanzania and NE Mozambique and the Vohibory Complex are interpreted as fragments of a collage of volcanic arcs and microcontinents, formed outboard of the Congo-Tanzania Craton. These were accreted to the margin of the craton during closure of the Mozambique Ocean (Möller et al., 1998; Jöns and Schenk, 2008; Viola et al., 2008). The prevailing model for the Tanzanian segment of the Mozambique Belt assumes that accretion of outboard terranes and collision of a minor continent (“Azania”/Proto-Madagascar) took place during the East-African orogeny, between ca. 655 and 610 Ma (Meert, 2003; Collins and Pisarevsky, 2005; Fritz et al., 2005). Local evidence for high-grade metamorphism at ca. 640 Ma in the felsic basement of Tanzania (the Western Granulites) attributed to the CongoTanzania Craton margin lends support to this model (Sommer et al., 2003; 2005). However, a growing body of data emphasizes the importance of the ca. 550 Ma metamorphism and tectonism in Tanzania (Cutten et al., 2006; Hauzenberger et al., 2007; Rossetti et al., 2008), suggesting that crustal thickening at the margin of the Congo-Tanzania Craton may be linked to the Kuunga orogeny and that effects of the East-African orogeny may be restricted to outboard accreted terranes (the Eastern Granulites and S Madagascar). Additional structural and geochronological data are required to solve this controversy. In NE Mozambique, the Marrupa Complex, which underlies the Cabo Delgado Nappe Complex, does not show evidence for the early metamorphic events in the ca. 735–610 Ma interval. At that time, the Unango Complex was overlain by the marine Geci Group (ca. 630–585 Ma; Melezhik et al., 2006). The timing of nappe transport is bracketed between 596 ± 11 and 555 ± 11 Ma. Crustal thickening in NE Mozambique is consequently part of the Kuunga orogeny, and attributed to final stage of Gondwana assembly between 570 and 530 Ma.

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7.7. Cambrian gravitational collapse south of the Lurio Belt In the Maud Belt, Central Dronning Maud Land (Fig. 12c), Pan-African metamorphism reached granulite-facies conditions between ca. 570 and 530 Ma (Jacobs et al., 1998; Asami et al., 2005; Board et al., 2005). In this area, there is abundant evidence for gravitational collapse of this overthickened crust following peak metamorphism. This includes decompression partial melting, postcollisional magmatism, near-isothermal decompression reactions and extensional tectonics (Jacobs et al., 2003c; Engvik and Elvevold, 2004; Board et al., 2005; Jacobs et al., 2008a). Leucosomes generated during this event have been dated between 519 ± 4 and 504 ± 6 Ma (Jacobs et al., 1998; Paulsson and Austrheim, 2003; Board et al., 2005), and NNE-directed extension has been dated from a late shear zone at 507 ± 9 Ma (Jacobs et al., 2008a). Voluminous granitic to charnockitic plutonism and minor mafic plutonism took place between ca. 530 and 485 Ma, with a frequency maximum at 510–500 Ma (Jacobs et al., 2003c; Paulsson and Austrheim, 2003; Bisnath et al., 2006; Jacobs et al., 2008a). The same arguments are developed by Jacobs et al. (2008a) in order to suggest that the Nampula Complex was also affected by gravitational collapse after ca. 530 Ma, possibly as a result of delamination of an orogenic root (Jacobs et al., 2008a), and that the Lurio Belt represents the northern limit of the collapsed domain. This model is supported by the data reported in this study and is based on the observations of (1) voluminous postcollisional, syn- to post-kinematic, granite plutonism in the Lurio Belt (Malema Suite) and Nampula Complex (Murrupula Suite); (2) amphibolite-facies metamorphism and widespread migmatitization (ca. 520–490 Ma) in the Nampula Complex, largely post-dating the peak of convergence-related metamorphism (ca. 560–530 Ma); (3) evidence of late extensional tectonics along the Lurio Belt at, and after ca. 530 Ma, and exhumation of the Nampula Complex as a large antiform south of the Lurio Belt (Viola et al., 2008). 7.8. Is the Lurio Belt a Pan-African suture zone? Along the Damara-Lufilian-Zambezi Orogen, eclogite bodies with mid-oceanic-ridge basalt compositions are exposed in the Zambezi Belt (Fig. 12c). They attest to subduction of oceanic crust between the Congo-Tanzania and Kalahari Cratons at 595 ± 10 Ma and closure of the oceanic basin after this date (John et al., 2003). In the Damara Belt of Namibia, Gray et al. (2008) argue for northwarddipping subduction below the attenuated margin of the Congo Craton at the end of the Neoproterozoic. This polarity is consistent with occurrence, in Malawi and Zambia, of granite plutonism between 667 ± 1 and 577 ± 1 Ma in the Southern Irumide Belt (upper plate; Kröner et al., 2001; Johnson et al., 2006). The WSW–ENE trending Lurio Belt in Mozambique is the subject of ongoing speculation, as it is ideally located to include the PanAfrican suture between the Kalahari and Congo-Tanzania Cratons (Collins and Pisarevsky, 2005; Grantham et al., 2008). This model is not in conflict with the interpretation of the Lurio Belt as representing the northern limit of a collapsed crustal domain (Jacobs et al., 2008a). Although the Lurio Belt becomes progressively diffuse as it is traced westwards across NE Mozambique, it may be identified in the Nsanje area in southern Malawi (Andreoli and Hart, 1990), and tentatively into NW Mozambique, as the basal shear zone underlying the Tete magmatic suite (Fig. 12c; Grantham et al., 2008). Towards the east in Gondwana it has been tentatively linked to the contact between the Vijayan and Highland Complexes in Sri Lanka (Fig. 12c; Collins and Pisarevsky, 2005). A direct link with major shear zone exposed in S Madagascar is geometrically less probable (de Wit et al., 2001). The Lurio Belt lacks undisputable structural and lithological evidence for a suture zone, such as ophiolitic rocks (Viola et al., 2008).

Nevertheless, two main permissive arguments can be put forward in favour of the suture interpretation. (1) As developed above, the Nampula Complex, south of the Lurio, is best restored to the periphery of the Kalahari Craton before the Pan-African orogeny. In contrast, the Marrupa and Unango Complexes, north of the Lurio Belt, were clearly peripheral to the Congo-Tanzania Craton. (2) High pressure granulites in the Ocua Complex within the Lurio Belt indicate burial to a depth of at least ca. 55 km (1.57 GPa–950 ◦ C; Engvik et al., 2007). Reaction textures and garnet zoning imply that burial was followed by rapid exhumation (≤1 Ma) along a near-isothermal decompression path. Such a short-lived burial-exhumation process may indicate a plate boundary (Engi et al., 2001). Undated, ultramafic rocks and high-pressure granulites, transitional to eclogites, are locally reported along strike of the Lurio Belt in the Nsanje area in southern Malawi (Andreoli, 1984; Andreoli and Hart, 1990). The timing of possible suturing along the Lurio Belt, is constrained to be between ca. 595 Ma (Zambian eclogites at 595 ± 10 Ma, youngest pluton specific to the Cabo Delgado Nappe Complex at 596 ± 11 Ma, marine Geci Group between ca. 630 and 585 Ma) and ca. 557 Ma (high-pressure granulite-facies metamorphism). South of the Lurio Belt, the Mugeba and Monapo klippen overlying the Nampula Complex and the Schirmacher klippe overlying the Maud Belt in Antarctica (Fig. 12c), share a number of characteristics in common with nappes of the Cabo Delgado Nappe Complex, north of the Lurio Belt. These include similar mafic rocks and evidence for granulite-facies metamorphism older than ca. 570 Ma. Metamorphism is dated at 615 ± 7 Ma in the Mugeba klippe (Kröner et al., 1997), 579 ± 11 to 596 ± 5 Ma in the Monapo klippe (Grantham et al., 2008) and 632 ± 8 Ma in the Schirmacher klippe (Ravikant et al., 2004). If these southern klippen are correlated with the Cabo Delgado Nappe Complex across the Lurio Belt, then interpretation of the Lurio Belt as a suture zone requires additional explanation. Two alternative ideas can be put forwards. (1) The Nampula and Unango–Marrupa Complexes were sutured before being overthrust by a single allochthonous far-travelled nappe system with a unique transport direction. If the structural data north of the Lurio Belt are taken as a reference, then the transport direction is northwestwards (Muhongo, 1994; Fritz et al., 2005; Rossetti et al., 2008; Viola et al., 2008). (2) The second alternative is to accept a model of a bivergent orogeny, with the Lurio Belt as a root zone. Nappes were possibly transported to the NW north of the Lurio Belt and to the SE, south of the Lurio Belt. In this model, nappe transport could possibly also be diachronous on either side of the suture. Linear structures at the base of the Mugeba and Monapo klippen do not, however, trend NW–SE but E–W (Sacchi et al., 1984; Grantham et al., 2008). Discussion of these two models is beyond the scope of this paper. Nevertheless, available data show that the Lurio Belt can be regarded as a suture zone, though correlation of nappes across the belt requires an elaborate tectonic model. 8. Conclusions This study leads to the following conclusions, relevant for improved understanding of Gondwana assembly. (1) The Ponta Messuli Complex, exposed NW of the Maniamba Graben is part of the Palaeoproterozoic Usagaran crust, situated along the Congo-Tanzania foreland of the Pan-African orogenic system. It is distinctly older than the crust SE of the Maniamba Graben, supporting the idea that this graben conceals a preKaroo crustal discontinuity. (2) The Unango and Marrupa Complexes grew by voluminous magmatic addition between 1062 ± 13 and 946 ± 11 Ma, probably in a continental arc setting (Figs. 11 and 12). They include evidence of a granulite-facies metamorphism dated at 953 ± 8 Ma. The Unango and Marrupa Complexes were probably assembled to

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(3)

(4)

(5)

(6)

(7)

the margin of the Congo-Tanzania Craton during the Irumidian orogeny, together with terranes in the Southern Irumide Belt. South of the Lurio Belt, the Nampula Complex developed between 1148 ± 2 and 1028 ± 7 Ma (1123 ± 9 to 1042 ± 9 Ma in the study area), significantly before the Unango and Marrupa Complexes (Figs. 11 and 12). It also probably evolved within an arc setting, which matured with time. Pre-Pan-African paleogeographic restoration of the Nampula Complex is uncertain. Nevertheless, it is most elegantly restored at the margin of the Kalahari Craton, as an extension of the Maud Belt of Antarctica. The Unango and Marrupa Complexes share minor Neoproterozoic magmatism with the Neoproterozoic margin of the Congo-Tanzania Craton. This includes nepheline syenitebearing plutons, one of which is dated at 799 ± 8 Ma (Fig. 11). Pan-African high-grade metamorphism took place at 555 ± 11 Ma in the Marrupa Complex, between 569 ± 9 and 527 ± 8 Ma in the Unango Complex, and between 555 ± 12 and 490 ± 8 Ma in the Nampula Complex (543 ± 23 to 493 ± 8 Ma in the study area; Figs. 11 and 12c). In the Unango and Marrupa Complexes its intensity increases southwards towards the Lurio Belt. Crustal thickening postdates deposition of the marine Geci Group (630–585 Ma), it postdates transport of nappes of the Cabo Delgado Nappe Complex over the felsic basement, and it is coeval with high-pressure metamorphism in the core of the Lurio Belt at 557 ± 16 Ma. The data therefore imply that maximum crustal thickening in northeastern Mozambique is part of the main Pan-African ca. 570–530 Ma Kuunga orogeny, related to final assembly of the Indian, East Antarctica, Kalahari, and Congo-Tanzania Cratons to form Gondwana. Several arguments suggest that the Nampula Complex was affected by gravitational collapse after ca. 530 Ma, together with the Maud Belt of Antarctica, and that the Lurio Belt represents the northern limit of this process. Post-collisional felsic magmatism, peaking around 510 Ma, shows a dramatic volume increase in the Lurio Belt and to the south, in the Nampula Complex. The Lurio Belt shows evidence for extensional reactivation at and after 530 Ma and the Nampula Complex gives evidence for amphibolite-facies metamorphism as late as 493 ± 8 Ma. The Lurio Belt may include the Pan-African suture zone between the Kalahari and Congo-Tanzania Cratons. The differences in Precambrian geology and geochronology across the Lurio Belt and the occurrence of high-pressure granulites in the Lurio Belt support this interpretation. Probable correlation of nappes of the Cabo Delgado Nappe Complex with the Mugeba and Monapo klippen across the Lurio Belt, nevertheless, rules out a simple tectonic model.

Acknowledgements The Mineral Resources Management Capacity Building Project, Republic of Mozambique, was funded by the Nordic Development Fund and the World Bank. “Lot 1” was implemented by the Norconsult Consortium including the Geological Survey of Norway, British Geological Survey, National Directorate for Geology of Mozambique, Norconsult AS, and Eteng SA. A number of collaborators contributed efficiently to data collection and interpretations. These include T. Bjerkgård, A. Engvik, P. Feitio, E. Gonzalez, L. Ilyinsky, D. Jamal, C.L. Kirkland, K. Lindén, V. Manhica, V.A. Melezhik, A. Moniz, G. Motuza, Ø. Nordgulen, D. Rosse, J.S. Sandstad, R.A. Smith, E. Tveten, M.J. Whitehouse, and M.T.D. Wingate. The World Bank consultants, F. Hartzer and the Late E. Hammerbeck, contributed to the success of the project. A. Möller and S. P. Johnson reviewed the manuscript constructively. R.J. Thomas and R.M. Key publish with permission of the Executive Director of BGS. The John de Laeter Centre of Mass Spectrometry in Perth is operated by a university-government consortium, with the support of the Aus-

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tralian Research Council. The NORDSIM facility in Stockholm is operated under an agreement between the research funding agencies of Denmark, Norway and Sweden, the Geological Survey and the Swedish Museum of Natural History. This is NORDSIM publication 225. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.precamres.2009.01.005. References Andreoli, M.A.G., 1984. Petrochemistry, tectonic evolution and metasomatic mineralisation of Mozambique belt granulites from S. Malawi and Tete Mozambique. Precambrian Research 25, 161–186. Andreoli, M.A.G., Hart, R.J., 1990. Metasomatized granulites and eclogites of the Mozambique Belt: implications for mantle devolatilization. In: Herbert, H.K., Ho, S.E. (Eds.), Stable Isotopes and Fluid Processes in Mineralization. Geol. Dept. and Univ. Extension, Univ. Western Australia, Publication 23, pp. 121–140. Arndt, N.T., Todt, W., Chauvel, C., Tapfer, M., Weber, K., 1991. U–Pb zircon age and Nd isotopic composition of granitoids, charnockites and supracrustal rocks from Heimefrontfjella, Antarctica. Geologische Rundschau 80 (3), 759–777. Asami, M., Suzuki, K., Grew, E.S., 2005. Monazite and zircon dating by the chemical Th–U–total Pb isochron method (CHIME) from Alasheyev Bight to the Sør Rondane Mountains East Antarctica: a reconnaissance study of the Mozambique suture in Eastern Queen Maud Land. The Journal of Geology 113, 59–82. Ashwal, L.D., Armstrong, R.A., Roberts, R.J., Schmitz, M.D., Corfu, F., Hetherington, C.J., Burke, K., Gerber, M., 2007. Geochronology of zircon megacrysts from nephelinebearing gneisses as constraints on tectonic setting: implications for resetting of the U–Pb and Lu–Hf isotopic systems. Contributions to Mineralogy and Petrology 153, 389–403. Bingen, B., Davis, W.J., Hamilton, M.A., Engvik, A., Stein, H.J., Skår, Ø., Nordgulen, Ø., 2008a. Geochronology of high-grade metamorphism in the Sveconorwegian belt, S. Norway: U–Pb, Th–Pb and Re–Os data. Norwegian Journal of Geology 88, 13–42. Bingen, B., Nordgulen, Ø., Viola, G., 2008b. A four-phase model for the Sveconorwegian orogeny SW Scandinavia. Norwegian Journal of Geology 88, 43–72. Bingen, B., Skår, Ø., Marker, M., Sigmond, E.M.O., Nordgulen, Ø., Ragnhildstveit, J., Mansfeld, J., Tucker, R.D., Liégeois, J.P., 2005. Timing of continental building in the Sveconorwegian orogen, SW Scandinavia. Norwegian Journal of Geology 85, 87–116. Bisnath, A., Frimmel, H.E., Armstrong, R.A., Board, W.S., 2006. Tectono-thermal evolution of the Maud Belt: new SHRIMP U–Pb zircon data from Gjelsvikfjella, Dronning Maud Land, East Antarctica. Precambrian Research 150, 95–121. Bjerkgård, T., Stein, H.J., Bingen, B., Henderson, I.H.C., Sandstad, J.S., Moniz, A., 2009. The Niassa Gold Belt, northern Mozambique—a segment of a continental-scale Pan-African gold-bearing structure? Journal of African Earth Sciences 53, 45–58. Black, L.P., Kamo, S.L., Allen, C.M., Aleinikoff, J.N., Davis, D.W., Korsch, R.J., Foudoulis, C., 2003. TEMORA 1, a new zircon standard for Phanerozoic U–Pb geochronology. Chemical Geology 200, 155–170. Bloomfield, K., 1968. The pre-Karoo geology of Malawi 5. Geological Survey of Malawi, Memoir, p. 166. Board, W.S., Frimmel, H.E., Armstrong, R.A., 2005. Pan-African Tectonism in the Western Maud Belt: P–T–t path for high-grade gneisses in the H.U. Sverdrupfjella, East Antarctica. Journal of Petrology 46, 671–699. Boger, S.D., Carson, C.J., Wilson, C.J.L., Fanning, C.M., 2000. Neoproterozoic deformation in the Radok Lake region of the northern Prince Charles Mountains, east Antarctica; evidence for a single protracted orogenic event. Precambrian Research 104, 1–24. Boger, S.D., Miller, J.M., 2004. Terminal suturing of Gondwana and the onset of the Ross-Delamerian Orogeny: the cause and effect of an Early Cambrian reconfiguration of plate motions. Earth and Planetary Science Letters 219, 35–48. Collins, A.S., Pisarevsky, S.A., 2005. Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth Science Reviews 71, 229–270. Cornell, D.H., Thomas, R.J., 2006. Age and tectonic significance of the Banana Beach Gneiss, KwaZulu-Natal South Coast South Africa. South African Journal of Geology 109, 335–340. Costa, M., Cadoppi, P., Sacchi, R., Fanning, C.M., 1994. U–Pb “SHRIMP” dating of zircons from Mozambique gneiss. Bolletino Società Geologica Italiana 113, 173–178. Costa, M., Ferrara, G., Sacchi, R., Tonarini, S., 1992. Rb/Sr dating of the Upper Proterozoic basement of Zambesia, Mozambique. Geologische Rundschau 81, 487–500. Cutten, H., Johnson, S.P., De Waele, B., 2006. Protolith ages and timing of metasomatism related to the formation of whiteschists at Mautia Hill Tanzania: implications for the assembly of Gondwana. The Journal of Geology 114, 683–698. De Waele, B., Johnson, S.P., Pisarevsky, S.A., 2008. Palaeoproterozoic to Neoproterozoic growth and evolution of the eastern Congo Craton: its role in the Rodinia puzzle. Precambrian Research 160, 127–141. De Waele, B., Kampunzu, A.B., Mapani, B.S.E., Tembo, F., 2006a. The Mesoproterozoic Irumide belt of Zambia. Journal of African Earth Sciences 46, 36–70.

254

B. Bingen et al. / Precambrian Research 170 (2009) 231–255

De Waele, B., Liégeois, J.P., Nemchin, A.A., Tembo, F., 2006b. Isotopic and geochemical evidence of Proterozoic episodic crustal reworking within the Irumide Belt of South-Central Africa, the southern metacratonic boundary of an Archaean Bangweulu Craton. Precambrian Research 148, 225–256. de Wit, M.J., Bowring, S.A., Ashwal, L.D., Randrianasolo, L.G., Morel, V.P.I., Rambeloson, R.A., 2001. Age and tectonic evolution of Neoproterozoic ductile shear zones in southwestern Madagascar, with implications for Gondwana studies. Tectonics 20, 1–45. Eglington, B.M., 2006. Evolution of the Namaqua-Natal Belt, southern Africa—a geochronological and isotope geochemical review. Journal of African Earth Sciences 46, 93–111. Engi, M., Berger, A., Roselle, G.T., 2001. Role of the tectonic accretion channel in collisional orogeny. Geology 29, 1143–1146. Engvik, A.K., Elvevold, S., 2004. Pan-African extension and near-isothermal exhumation of a granulite facies terrain, Dronning Maud Land, Antarctica. Geological Magazine 141, 649–660. Engvik, A.K., Tveten, E., Bingen, B., Viola, G., Erambert, M., Feito, P., de Azavedo, S., 2007. P–T–t evolution and decompression textures of Pan-African high-pressure granulites, Lurio Belt, northeastern Mozambique. Journal of Metamorphic Geology 25, 935–952. Fitzsimons, I.C.W., 2000. Grenville-age basement provinces in East Antarctica: evidence for three separate collisional orogens. Geology 28, 879–882. Frimmel, H.E., 2004. Formation of a late Mesoproterozoic supercontinent: the South Africa–East Antarctica connection. In: Eriksson, P.G., Altermann, W., Nelson, D.R., Mueller, W.U., Catuneanu, O. (Eds.), The Precambrian Earth: Tempos and Events. Developments in Precambrian Geology, 12. Elsevier, Amsterdam, pp. 240–255. Fritz, H., Tenczer, V., Hauzenberger, C.A., Wallbrecher, E., Hoinkes, G., 2005. Central Tanzanian tectonic map: a step forward to decipher Proterozoic structural events in the East African Orogen. Tectonics 24, TC6013, doi:10.1029/2005TC001796. Goscombe, B., Armstrong, R., Barton, J.M., 2000. Geology of the Chewore Inliers, Zimbabwe: constraining the Mesoproterozoic to Palaeozoic evolution of the Zambezi Belt. Journal of African Earth Sciences 30, 589–627. Grantham, G.H., Macey, P.H., Ingram, B.A., Roberts, M.P., Armstrong, R.A., Hokada, T., Shiraishi, K., Jackson, C., Bisnath, A., Manhica, V., 2008. Terrane correlation between Antarctica, Mozambique and Sri Lanka; comparison of geochronology, lithology, structure and metamrophism and possible implications for the geology of southern Africa and Antarctica. In: Satish-Kumar, M., Motoyoshi, Y., Osanai, Y., Hiroi, Y., Shiraishi, K. (Eds.), Geodynamic Evolution of East Antarctica: A Key to the East–West Gondwana Connection. Geological Society, Special Publications 308, London, pp. 91–119. Gray, D.R., Foster, D.A., Meert, J.G., Goscombe, B.D., Armstrong, R., Trouw, R.A.J., Passchier, C.W., 2008. A Damara Orogen perspective on the assembly of southwestern Gondwana. In: Pankhurst, R.J., Trouw, R.A.J., de Brito Neves, B.B., de Wit, M.J. (Eds.), West Gondwana: Pre-Cenozoic Correlations Across the South Atlantic Region. Geological Society, Special Publications 294, London, pp. 257–278. Groenewald, P.B., Grantham, G.H., Watkeys, M.K., 1991. Geological evidence for a Proterozoic to Mesozoic link between southeastern Africa and Dronning Maud Land, Antarctica. Journal of the Geological Society, London 148, 1115–1123. Handke, M.J., Tucker, R.D., Ashwal, L.D., 1999. Neoproterozoic continental arc magmatism in west-central Madagascar. Geology 27, 351–354. Hanson, R.E., 2003. Proterozoic geochronology and tectonic evolution of southern Africa. In: Yoshida, M., Windley, B.F., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geological Society, Special Publications 206, London, pp. 427–463. Hanson, R.E., Harmer, R.E., Blenkinsop, T.G., Bullen, D.S., Dalziel, I.W.D., Gose, W.A., Hall, R.P., Kampunzu, A.B., Key, R.M., Mukwakwami, J., Munyanyiwa, H., Pancake, J.A., Seidel, E.K., Ward, S.E., 2006. Mesoproterozoic intraplate magmatism in the Kalahari Craton: a review. Journal of African Earth Sciences 46, 141–167. Hanson, R.E., Wardlaw, M.S., Wilson, T.J., Mwale, G., 1993. U–Pb zircon ages from the Hook granite massif and Mwembeshi dislocation: constraints on Pan-African deformation, plutonism, and transcurrent shearing in central Zambia. Precambrian Research 63, 189–209. Hauzenberger, C.A., Sommer, H., Fritz, H., Bauernhofer, A., Kröner, A., Hoinkes, G., Wallbrecher, E., Thöni, M., 2007. SHRIMP U–Pb and Sm–Nd garnet ages from the granulite-facies basement of SE Kenya: evidence for Neoproterozoic polycyclic assembly of the Mozambique Belt. Journal of the Geological Society, London 164, 189–201. Holmes, A., 1951. The sequence of Pre-Cambrian orogenic belts in South and Central Africa. In: Sandford, K.S., Blondel, F. (Eds.), Proceedings of the 18th International Geological Congress, vol. 14. Association des Services Géologiques Africains, London, pp. 254–269. Jackson, S.E., Pearson, N.J., Griffin, W.L., Belousova, E.A., 2004. The application of laser ablation-inductively coupled plasma-mass spectrometry to in situ U–Pb zircon geochronology. Chemical Geology 211, 47–69. Jacobs, J., Bauer, W., Fanning, C.M., 2003a. Late Neoproterozoic/Early Palaeozoic events in central Dronning Maud Land and significance for the southern extension of East African Orogen into East Antarctica. Precambrian Research 126, 27–53. Jacobs, J., Bingen, B., Thomas, R.J., Bauer, W., Wingate, M., Feitio, P., 2008a. Early Palaeozoic orogenic collapse and voluminous late-tectonic magmatism in Dronning Maud Land and Mozambique: insights into the partially delaminated orogenic root of the East African-Antarctic Orogen? In: Satish-Kumar, M., Motoyoshi, Y., Osanai, Y., Hiroi, Y., Shiraishi, K. (Eds.), Geodynamic Evolution of East Antarctica: A Key to the East–West Gondwana Connection. Geological Society, Special Publications 308, London, pp. 69–90.

Jacobs, J., Falter, M., Thomas, R.J., Kunz, J., 1997. 40 Ar–39 Ar thermochronological constraints on the structural evolution of the Mesoproterozoic Natal metamorphic province SE Africa. Precambrian Research 86, 71–92. Jacobs, J., Fanning, C.M., Bauer, W., 2003b. Timing of Grenville-age vs Pan-African medium- to high grade metamorphism in western Dronning Maud Land (East Antarctica) and significance for correlations in Rodinia and Gondwana. Precambrian Research 125, 1–20. Jacobs, J., Fanning, C.M., Henjes-Kunst, F., Olesch, M., Paech, H.J., 1998. Continuation of the Mozambique Belt into East Antarctica: Grenville-age metamorphism and polyphase Pan-African high-grade events in Central Dronning Maud Land. The Journal of Geology 106, 385–406. Jacobs, J., Klemd, R., Fanning, C.M., Bauer, W., Colombo, F., 2003c. Extensional collapse of the late Neoproterozoic-early Palaeozoic East African-Antarctic Orogen in central Dronning Maud Land, East Antarctica. In: Yoshida, M., Windley, B.F., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geological Society, Special Publications 206, London, pp. 271–287. Jacobs, J., Pisarevsky, S., Thomas, R.J., Becker, T., 2008b. The Kalahari Craton during the assembly and dispersal of Rodinia. Precambrian Research 160, 142–158. Jacobs, J., Thomas, R.J., 2004. Himalayan-type indenter-escape tectonics model for the southern part of the late Neoproterozoic-early Paleozoic East AfricanAntarctic orogen. Geology 32, 721–724. John, T., Schenk, V., Haase, K., Scherer, E., Tembo, F., 2003. Evidence for a Neoproterozoic ocean in south-central Africa from mid-oceanic-ridge-type geochemical signature and pressure–temperature estimates of Zambian eclogites. Geology 31, 243–246. John, T., Schenk, V., Mezger, K., Tembo, F., 2004. Timing and PT evolution of whiteschist metamorphism in the Lufilian Arc–Zambezi Belt orogen (Zambia): implications for the assembly of Gondwana. The Journal of Geology 112, 71–90. Johnson, S.P., Cutten, H.N.C., Muhongo, S., De Waele, B., 2003. Neoarchaean magmatism and metamorphism of the western granulites in the central domain of the Mozambique belt. Tanzania: U–Pb SHRIMP geochronology and PT estimates. Tectonophysics 375, 125–145. Johnson, S.P., De Waele, B., Evans, D., Banda, W., Tembo, F., Milton, J.A., Tani, K., 2007a. Geochronology of the Zambezi supracrustal sequence, southern Zambia: a record of Neoproterozoic divergent processes along the southern margin of the Congo Craton. The Journal of Geology 115, 355–374. Johnson, S.P., De Waele, B., Liyungu, K.A., 2006. U–Pb sensitive high-resolution ion microprobe (SHRIMP) zircon geochronology of granitoid rocks in eastern Zambia: Terrane subdivision of the Mesoproterozoic Southern Irumide Belt. Tectonics 25, TC6004, doi:10.1029/2006TC001977. Johnson, S.P., De Waele, B., Tembo, F., Katongo, C., Tani, K., Chang, Q., Iizuka, T., Dunkley, D., 2007b. Geochemistry, geochronology and isotopic evolution of the Chewore-Rufunsa Terrane, Southern Irumide Belt: a Mesoproterozoic continental margin arc. Journal of Petrology 48, 1411–1441. Johnson, S.P., Oliver, G.J.H., 2002. High fO2 metasomatism during whiteschist metamorphism, Zambesi belt, Northern Zimbabwe. Journal of Petrology 43, 271–290. Johnson, S.P., Rivers, T., De Waele, B., 2005. A review of the Mesoproterozoic to early Palaeozoic magmatic and tectonothermal history of south-central Africa: implications for Rodinia and Gondwana. Journal of the Geological Society, London 161, 433–450. Jöns, N., Schenk, V., 2008. Relics of the Mozambique Ocean in the central East African Orogen: evidence from the Vohibory Block of southern Madagascar. Journal of Metamorphic Geology 26, 17–28. Key, R.M., Bingen, B., Barton, E., Daudi, E.X.F., Manuel, S., Moniz, A., 2007. Kimberlites in a Karoo graben of northern Mozambique: tectonic setting, mineralogy and Rb–Sr geochronology. South African Journal of Geology 110, 111–124. Key, R.M., Liyungu, A.K., Njamu, F.M., Somwe, V., Banda, J., Mosley, P.N., Armstrong, R.A., 2001. The western arm of the Lufilian Arc in NW Zambia and its potential for copper mineralization. Journal of African Earth Sciences 33, 503–528. Key, R.M., Smith, R.A., Smelror, M., Saether, O.M., Thorsnes, T., Powell, J.H., Njange, F., Zandamela, E.B., 2008. Revised lithostratigraphy of the Mesozoic–Cenozoic succession of the onshore Rovuma Basin, northern coastal Mozambique. South African Journal of Geology 111, 89–108. Kröner, A., Hegner, E., Collins, A.S., Windley, B.F., Brewer, T.S., Razakamanana, T., Pidgeon, R.T., 2000. Age and magmatic history of the Antananarivo block, Central Madagascar, as derived from zircon geochronology and Nd isotopic systematics. American Journal of Science 300, 251–288. Kröner, A., Muhongo, S., Hegner, E., Wingate, M.T.D., 2003. Single-zircon geochronology and Nd isotopic systematics of Proterozoic high-grade rocks from the Mozambique belt of southern Tanzania (Masasi area): implications for Gondwana assembly. Journal of the Geological Society, London 160, 745–757. Kröner, A., Sacchi, R., Jaeckel, P., Costa, M., 1997. Kibaran magmatism and PanAfrican granulite metamorphism in northern Mozambique: single zircon ages and regional implications. Journal of African Earth Sciences 25, 467–484. Kröner, A., Willner, A.P., Hegner, E., Jaeckel, P., Nemchin, A.A., 2001. Single zircon ages PT evolution and Nd isotopic systematics of high-grade gneisses in southern Malawi and their bearing on the evolution of the Mozambique belt in southeastern Africa. Precambrian Research 109, 257–291. Küster, D., Liegeois, J.P., Matukovb, D., Sergeev, S., Lucassen, F., 2008. Zircon geochronology and Sr, Nd, Pb isotope geochemistry of granitoids from Bayuda Desert and Sabaloka (Sudan): Evidence for a Bayudian event (920–900 Ma) preceding the Pan-African orogenic cycle (860–590 Ma) at the eastern boundary of the Saharan Metacraton. Precambrian Research 164, 16–39. Lenoir, J.L., Liégeois, J.P., Theunissen, K., Klerkx, J., 1994. The Palaeoproterozoic Ubendian shear belt in Tanzania: geochronology and structure. Journal of African Earth Sciences 19, 169–184.

B. Bingen et al. / Precambrian Research 170 (2009) 231–255 Ludwig, K.R., 1998. On the treatment of concordant uranium–lead ages. Geochimica et Cosmochimica Acta 62, 665–676. Ludwig, K.R., 2001. Users Manual for Isoplot/Ex Version 2.49 a Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center, Special Pubication No. 1a, Berkley. Lulin, J.M., 1984. Un nouveau gîte à Nb, Ta, U, REE d’origine magmatique en Afrique orientale: le complexe alcalin tectonisé de Meponda, Précambrian, province de Niassa, Mocambique. Ph.D. Thesis. Université d’Orléans, Doc. Bureau de Recherche Géologique et Minère, BRGM 87, Orléans. Manhica, A.D.S.T., Grantham, G.H., Armstrong, R.A., Guise, P.G., Kruger, F.J., 2001. Polyphase deformation and metamorphism at the Kalahari Craton–Mozambique Belt boundary. In: Miller, J.A., Holdsworth, R.E., Buick, I.S., Hand, M. (Eds.), Continental Reactivation and Reworking. Geological Society, Special Publications 184, London, pp. 303–322. McCourt, S., Armstrong, R.A., Grantham, G.H., Thomas, R.J., 2006. Geology and evolution of the Natal belt, South Africa. Journal of African Earth Sciences 46, 71– 92. Meert, J.G., 2003. A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362, 1–40. Melezhik, V.A., Bingen, B., Fallick, A.E., Gorokhov, I.M., Kuznetsov, A.B., Sandstad, J.S., Solli, A., Bjerkgård, T., Henderson, I.H.C., Boyd, R., Jamal, D., Moniz, A., 2008. Isotope chemostratigraphy of marbles in northeastern Mozambique: apparent depositional ages and tectonostratigraphic implications. Precambrian Research 162, 540–558. Melezhik, V.A., Fallick, A.E., Smith, R.A., Rosse, D.M., 2007. Spherical and columnar, septarian, O-18-depleted, calcite concretions from Middle–Upper Permian lacustrine siltstones in northern Mozambique: evidence for very early diagenesis and multiple fluids. Sedimentology 54, 1389–1416. Melezhik, V.A., Kuznetsov, A.B., Fallick, A.F., Smith, R.A., Gorokhov, I.M., Jamal, D., Catuane, F., 2006. Depositional environments and an apparent age for the Geci meta-limestone: constraints on the geologic history of northern Mozambique. Precambrian Research 148, 19–31. Muhongo, S., 1994. Neoproterozoic collision tectonics in the Mozambique Belt of East Africa: evidence from the Uluguru mountains, Tanzania. Journal of African Earth Sciences 19, 153–168. Muhongo, S., Kröner, A., Nemchin, A.A., 2001. Single zircon evaporation and SHRIMP ages for granulite-facies rocks in the Mozambique Belt of Tanzania. The Journal of Geology 109, 171–189. Mänttäri, I., 2008. Mesoarchaean to Lower Jurassic U–Pb and Sm–Nd ages from NW Mozambique. In: Pekkala, Y., Lehto, T., Mäkitie, T. (Eds.), GTK Consortium Geological Surveys in Mozambique 2002–2007. Geological Survey of Finland, Special Paper, 48, pp. 81–119. Möller, A., Appel, P., Mezger, K., Schenk, V., 1995. Evidence for a 2 Ga subduction zone: eclogites in the Usagaran belt of Tanzania. Geology 23, 1067–1070. Möller, A., Hensen, B.J., Armstrong, R.A., Mezger, K., Ballevre, M., 2003. U–Pb zircon and monazite age constraints on granulite-facies metamorphism and deformation in the Strangways Metamorphic Complex (central Australia). Contributions to Mineralogy and Petrology 145, 406–423. Möller, A., Mezger, K., Schenk, V., 1998. Crustal age domains and the evolution of the continental crust in the Mozambique belt of Tanzania: combined Sm–Nd, Rb–Sr, and Pb–Pb isotopic evidence. Journal of Petrology 39, 749–783. Möller, A., Mezger, K., Schenk, V., 2000. U–Pb dating of metamorphic minerals: PanAfrican metamorphism and prolonged slow cooling of high pressure granulites in Tanzania, East Africa. Precambrian Research 104, 123–146. Nédélec, A., Bouchez, J.L., Gregoire, V., 2003. Quartz fabric evidence for early PanAfrican penetrative east-directed shearing in the Itremo Supracrustal Group of central Madagascar. Terra Nova 15, 20–28. Nemchin, A.A., Pidgeon, R.T., 1997. Evolution of the Darling Range Batholith, Yilgarn Craton, Western Australia: a SHRIMP zircon study. Journal of Petrology 38, 625–649. Norconsult Consortium, 2007a. Mineral Resources Management Capacity Building Project, Republic of Mozambique; Component 2: Geological Infrastructure Development Project, Geological Mapping Lot 1; Geophysical interpretation, Report No. B4, National Directorate of Geology, Republic of Mozambique. Norconsult Consortium, 2007b. Mineral Resources Management Capacity Building Project, Republic of Mozambique; Component 2: Geological Infrastructure Development Project, Geological Mapping Lot 1; Sheet Explanation: 32 Sheets; Scale: 1/250000, Report No. B6.f, National Directorate of Geology, Republic of Mozambique. Paquette, J.L., Nédélec, A., 1998. A new insight into Pan-African tectonics in the East–West Gondwana collision zone by U–Pb zircon dating of granites from central Madagascar. Earth and Planetary Science Letters 155, 45–56. Paulsson, O., Austrheim, H., 2003. A geochronological and geochemical study of rocks from Gjelsvikfjella, Dronning Maud Land Antarctica—implications for Mesoproterozoic correlations and assembly of Gondwana. Precambrian Research 125, 113–138. Pedersen, R.B., Dunning, G.R., Robins, B., 1989. U–Pb ages of nepheline syenite pegmatites from the Seiland Magmatic Province N Norway. In: Gayer, R.A. (Ed.), The Caledonide geology of Scandinavia. Graham & Trotman, London, pp. 3–8.

255

Pinna, P., Jourde, G., Calvez, J.Y., Mroz, J.P., Marques, J.M., 1993. The Mozambique Belt in northern Mozambique; Neoproterozoic (1100–850 Ma) crustal growth and tectogenesis, and superimposed Pan-African (800–550 Ma) tectonism. Precambrian Research 62, 1–59. Porada, H., Berhorst, V., 2000. Towards a new understanding of the Neoproterozoic–Early Palaeozoic Lufilian and northern Zambezi Belts in Zambia and Democratic Republic of Congo. Journal of African Earth Sciences 30, 727–771. Ravikant, V., Bhaskar Rao, Y.J., Gopalan, K., 2004. Schirmacher Oasis as an extension of the Neoproterozoic East African Orogen into Antarctica: new Sm–Nd isochron age constraints. The Journal of Geology 112, 607–616. Reddy, S.M., Collins, A.S., Mruma, A., 2003. Complex high-strain deformation in the Usagaran Orogen. Tanzania: structural setting of Palaeoproterozoic eclogites. Tectonophysics 375, 101–123. Ring, U., Kröner, A., Layer, P., Buchwaldt, R., Toulkeridis, T., 1999. Deformed A-type granites in northern Malawi, east-central Africa: pre- or syntectonic? Journal of the Geological Society, London 156, 695–714. Ring, U., Kröner, A., Toulkeridis, T., 1997. Palaeoproterozoic granulite-facies metamorphism and granitoid intrusions in the Ubendian–Usagaran Orogen of northern Malawi, east-central Africa. Precambrian Research 85, 27–51. Rivers, T., Corrigan, D., 2000. Convergent margin on southeastern Laurentia during the Mesoproterozoic: tectonic implications. Canadian Journal of Earth Sciences 37, 359–383. Rossetti, F., Cozzupoli, D., Phillips, D., 2008. Compressional reworking of the East African Orogen in the Uluguru Mountains of eastern Tanzania at c 550 Ma: implications for the final assembly of Gondwana. Terra Nova 20, 59–67. Sacchi, R., Cadoppi, P., Costa, M., 2000. Pan-African reactivation of the Lurio segment of the Kibaran Belt system: a reappraisal from recent age determinations in northern Mozambique. Journal of African Earth Sciences 30, 629–639. Sacchi, R., Marques, J.M., Costa, M., Casati, C., 1984. Kibaran events in the southernmost Mozambique belt. Precambrian Research 25, 141–159. Skår, Ø., 2002. U–Pb geochronology and geochemistry of early-Proterozoic rocks of the tectonic basement windows in central Nordland. Caledonides of northcentral Norway. Precambrian Research 116, 265–283. Sommer, H., Kröner, A., Hauzenberger, C., Muhongo, S., 2005. Reworking of Archaean and Palaeoproterozoic crust in the Mozambique belt of central Tanzania as documented by SHRIMP zircon geochronology. Journal of African Earth Sciences 43, 447–463. Sommer, H., Kröner, A., Hauzenberger, C., Muhongo, S., Wingate, M.T.D., 2003. Metamorphic petrology and zircon geochronology of high-grade rocks from the central Mozambique Belt of Tanzania: crustal recycling of Archean and Palaeoproterozoic material during Pan-African orogeny. Journal of Metamorphic Geology 21, 915–934. Stern, R.J., 1994. Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annual Reviews Earth and Planetary Sciences 22, 319–351. Tenczer, V., Hauzenberger, C.A., Fritz, H., Whitehouse, M.J., Mogessie, A., Wallbrecher, E., Muhongo, S., Hoinkes, G., 2006. Anorthosites in the Eastern Granulites of Tanzania—new SIMS zircon U–Pb age data, petrography and geochemistry. Precambrian Research 148, 85–114. Torsvik, T.H., Müller, R.D., Van der Voo, R., Steinberger, B., Gaina, C., 2008. Global plate motion frames: toward a unified model. Reviews of Geophysics 46, RG3004, doi:10.1029/2007RG000227. Tucker, R.D., Ashwal, L.D., Handke, M.J., Hamilton, M.A., Le Grange, M., Rambeloson, R.A., 1999. U–Pb geochronology and isotope geochemistry of the Archean and Proterozoic rocks of north-central Madagascar. The Journal of Geology 107, 135–153. Tucker, R.D., Kusky, T.M., Buchwaldt, R., Handke, M.J., 2007. Neoproterozoic nappes and superposed folding of the Itremo Group, west-central Madagascar. Gondwana Research 12, 356–379. Viola, G., Henderson, I.H.C., Bingen, B., Thomas, R.J., Smethurst, M.A., de Azavedo, S., 2008. Growth and collapse of a deeply eroded orogen: insights from structural and geochronological constraints on the Pan-African evolution of NE Mozambique. Tectonics 27, TC5009, doi:10.1029/2008TC002284. Vrána, S., Kachlík, V., Kröner, A., Marheine, D., Seifert, A.V., Zácek, V., Baburek, J., 2004. Ubendian basement and its late Mesoproterozoic and early Neoproterozoic structural and metamorphism overprint in northeastern Zambia. Journal of African Earth Sciences 38, 1–21. Whitehouse, M.J., Kamber, B.S., 2005. Assigning dates to thin gneissic veins in high-grade metamorphic terranes: a cautionary tale from Akilia, Southwest Greenland. Journal of Petrology 46, 291–318. Whitehouse, M.J., Kamber, B.S., Moorbath, S., 1999. Age significance of U–Th–Pb zircon data from early Archaean rocks of west Greenland—a reassessment based on combined ion-microprobe and imaging studies. Chemical Geology 160, 201–224. Wiedenbeck, M., Allé, P., Corfu, F., Griffin, W.L., Meier, M., Oberli, F., Von Quadt, A., Roddick, J.C., Spiegel, W., 1995. Three natural zircon standards for U–Th–Pb, Lu–Hf, trace element and REE analyses. Geostandards Newsletter 19, 1–23. Williams, I.S., 1998. U–Th–Pb geochronology by ion microprobe. In: McKibben, M.A., Shanks, W.C., Ridley, W.I. (Eds.), Applications of Microanalytical Techniques to Understanding Mineralisation Processes. Reviews of Economic Geology 7, 1–35.