Nd and Sr isotopic investigation of the
Archean – Proterozoic boundary in north eastern Tanzania:
constraints on the nature of Neoproterozoic tectonism in
the Mozambique Belt
M.A.H. Maboko
Department of Geology,Uni6ersity of Dar Es Salaam,P.O.Box35052,Dar Es Salaam,Tanzania
Received 3 March 1999; accepted 22 December 1999
Abstract
A Nd and Sr isotopic transect across the Mozambique Belt in northern Tanzania shows that, apart from the granulite terranes in the east, the belt is composed of reworked Archean crust which show TDMages ] 2.5 Ga,
similar to those obtained from the Tanzania Craton. Evidence for post-Archean crust is limited to the Eastern granulites which yield TDMages of 1.0 – 1.1 Ga. Typical Pan African biotite Rb – Sr cooling ages of 650 – 490 Ma
are found across the entire belt right up to the craton margin. The strong age gradient recorded by the biotite dates is interpreted as indicating diachronous cooling across the belt with the western parts cooling below 300°C up to 150 Ma earlier than the eastern parts. The oldest biotite ages (648914 Ma), from near the western margin, provide the best minimum estimate for the age of the pervasive amphibolite facies metamorphism that characterises the Mozambique Belt. If the metamorphism and deformation in the belt are a result of a continent – continent collision between West and East Gondwana, then this collision must have occurred prior to 650 Ma. © 2000 Elsevier Science B.V. All rights reserved.
Keywords:Mozambique Belt; Tanzania Craton; Palaeoproterozoic; Nd and Sr isotopes
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1. Introduction
A combination of isotope geochemistry and careful structural mapping provides important constraints on the tectonic evolution and palaeo-geography of metamorphic terranes. This ap-proach is particulary useful for mapping ancient
crustal blocks which have different mantle extrac-tion ages and/or different tectonothermal histories where the effects of multiple deformation, poly-metamorphism and deep erosion complicate or even obliterate most geological indicators which are normally used for terrane boundary identifica-tion. Such an approach has recently been used by Maboko (1995) and Maboko and Nakamura (1996) to isotopically characterise the boundary
E-mail address:[email protected] (M.A.H. Maboko)
between the Tanzania Craton and the Palaeoproterozoic Usagaran Belt in south-eastern Tanzania (see Fig. 1 for an outline of the Precam-brian geology of Tanzania). One of the most important results of these works was to demon-strate that in south eastern Tanzania, the Usagaran Belt represents new crust that was ex-tracted from the mantle at about 2.0 Ga but which had assimilated as much as 40% of pre-ex-isting Archean crust during its formation. Implicit in this argument is that, at least in southern Tanzania, the Usagaran Belt forms the western foreland of the Neoproterozoic Mozambique Belt, a metamorphic belt that occurs along most of the
eastern seaboard of the African continent
(Holmes, 1951). Such an interpretation is also consistent with the work of Wendt et al. (1972) who showed that, in south eastern Tanzania, the Mozambique Belt is composed of reworked Palaeoproterozoic Usagaran rocks whose inten-sity of reworking increases progressively towards the east. Although lithological associations
tradi-tionally assigned to the Usagaran have been de-scribed along the entire length of the Tanzanian sector of the Mozambique Belt (Quennell et al., 1956), it is not clear whether the relationship between the two belts established in south eastern Tanzania applies further north.
The aim of this paper is to isotopically identify the boundary between the Mozambique Belt and the Archean Craton in northern and north eastern Tanzania with a view of establishing the prove-nance of the rocks which now constitute Mozam-bique Belt crust and to better constraining the western limit of Pan African tectonothermal influ-ence. In particular, the study aims at establishing whether the crust underlying the Mozambique Belt is composed of: (i) reworked Palaeoprotero-zoic crust similar to the Usagaran of southeastern Tanzania; (ii) juvenile Pan African crust which is
separated from the Archean Craton by
Palaeoproterozoic crust analogous to the
Usagaran of southern Tanzania; (iii) juvenile Pan African crust directly abutting on to the Archean
Craton; or (iv) reworked Archean crust without an intervening Usagaran analogue. The results of the study have significant implications on the temporal and spatial distribution of continental crust east of the Tanzania Craton and the timing of the Neoproterozoic assembly of Gondwana arising from the continent – continent collision which resulted in the formation of the Mozam-bique Belt (Burke et al., 1977; Burke and Sengor, 1986; Stern, 1994). For this purpose, samples were collected along a traverse running from Kwa Mtoro in the craton, through Kondoa eastwards to the Usambara Mountains in the Mozambique Belt (Fig. 2). This traverse was chosen to take advantage of the previous photogeologically-based structural investigation of the Mozambique orogenic front by Hepworth and Kennerley (1970) and Hepworth (1972).
2. Geological background
2.1. Regional geology
The Archean Tanzania Craton constitute the central nucleus of continental crust in East Africa. Clifford (1970) subdivided the craton into two contrasting domains: a central region underlain by granite, granodiorite, felsic gneisses and
migmatites associated with metamorphic
supracrustal rocks and a northern part composed of the granite-greenstone association of basic to acid volcanics, turbidites, pelites and banded iron formations intruded by granites and overlain by conglomerates, arenaceous and argillaceous sedi-ments and minor volcanics. Post-orogenic granites intrude both domains of the craton and yield Rb – Sr isochron ages of about 2.6 Ga (Bell and Dodson, 1981), while lepidolites from pegmatites near Dodoma yield Rb – Sr ages of around 2.5 Ga, indicating that the cratonic rocks had cooled to below about 300°C by that time.
The Tanzania Craton is flanked by high grade metamorphic rocks of the Palaeoproterozoic Ubendian Belt (2.0 Ga, Lenoir et al., 1994) in western and south western Tanzania, whereas its eastern margin is defined as the zone where the east – west cratonic structural trends are truncated
by the meridional trends of the Mozambique Belt (Holmes, 1951, Fig. 1). In the area around Iringa, which also include the type Usagaran area, de-tailed mapping and geochronological work by Wendt et al. (1972) established that the Usagaran Belt is Palaeoproterozoic (2.0 Ga). The same work revealed that the early Proterozoic rocks were reworked during the Neoproterozoic result-ing in a progressive youngresult-ing of mica K – Ar ages from 1.9 Ga near the craton margin in the west and north west to 0.5 Ga in the north east and south east. The latter mineral cooling ages are typical of the Mozambique Belt which Kennedy (1964) recognised as being part of an Africa and Gondwana-wide network of Neoproterozoic oro-genic belts.
2.2. Geology of the cratonic margin in north eastern Tanzania
The geological and structural relationships along the Archean – Proterozoic boundary near Kondoa have been investigated by Hepworth and Kennerley (1970) and Hepworth (1972). In this area, Hepworth and Kennerley (1970) recognised five tectonometamorphic groups or tectonic do-mains (Fig. 2). The principal lithological and structural characteristics of each domain, starting from west to east, are described below and sum-marised in Table 1.
2.2.1. The Main Batholith
Mozam-Fig. 2. Tectonic domain map of eastern Tanzania (modified from Hepworth, 1972). The boundaries (light solid lines) mark the approximate eastern most limits of the various overprinting events. The thick solid lines are motorable tracks. The solid square in the Usambara Mountains correspond to sites where samples of juvenile Proterozoic crust were collected by Moller et al. (1998) and Maboko and Nakamura (submitted).
biquan orogenic front (Hepworth and Kennerley, 1970).
2.2.2. The Parangan Domain
The granite of the Main Batholith give way eastwards to well layered amphibolite-facies quartzo-feldspathic gneisses, psammites and semipelites which are typically mapped in Survey maps as Usagaran but to which Hepworth and Kennerley (1970) referred to as the Parangan Domain. Fozzard (1960) describes a gradational contact between the Main Batholith and rocks
included in the Parangan Domain by Hepworth and Kennerley (1970). Fozzard (1960) ascribed this relationship to the ‘autochthonous’ formation of the granite of the Main Batholith by anatexis of metasediments now constituting the Parangan Domain. A similar gradational relationship has been described by Selby and Mudd (1965) in the Kondoa map sheet.
isoclinal folds into small upright folds with a strong axial planar crenulation (Hepworth and Kennerley, 1970). According to Hepworth and Kennerley (1970), the structures suggest a differ-ence in deformation style between parallel shear-ing in the Main Batholith and flexure foldshear-ing in the Parangan.
2.2.3. The Kondoan Domain
The Parangan give way eastwards to the Kon-doan Domain which consists of a group of flaggy granitic gneisses, amphibolites, kyanite schists and muscovite schists derived from the Parangan by tectonic reworking (Hepworth and Kennerley, 1970). The tectonic working occurred through intense folding upon south-easterly plunging axes and easterly dipping axial planar foliation sur-faces. Like the Parangan, the Kondoan rocks
have previously been considered to be of
Usagaran age and described as ‘Mozambiquan geosynclinal sediments’ in survey maps. Hepworth and Kennerley (1970), however, interpret their flaggy nature, which is suggestive of a primary sedimentary origin, as being entirely of tectonic origin.
2.2.4. Usagaran Metasediments
Away from the orogenic front, Hepworth (1972) recognised a separate group of Usagaran Metasediments which he considers a true Mozam-biquan cover sequence as distinguished from
tec-tonically re-worked cratonic rocks of the
Kondoan tectonic domain to the west. According to Hepworth (1972), the Usagaran Metasediments consist mainly of biotite and hornblende gneisses with some quartzites and marbles metamorphosed in the amphibolite facies.
2.2.5. The Eastern granulites
The Eastern granulites, which define a semi-continuous, 900 km long north – south trending belt of high pressure granulites, constitute the eastern most tectonic domain within the Mozam-bique Belt (Hepworth, 1972). On the basis of a metamorphic and structural discordancy with rocks of the Kondoan Domain, Hepworth (1972) concluded that the Eastern granulites constitute a separate tectonic domain which was not involved in Kondoan tectonism. He proposed that either the granulite facies metamorphism was post-doan or that the Eastern granulites escaped
Kon-Table 1
The main tectonic domains in north eastern Tanzania (After Hepworth and Kennerley, 1970; Hepworth, 1972)
Tectonic Do- Lithology Remarks
main
Little or no foliation. Deformed by the Bubu event Granite
Main Batholith
resulting in easterly dipping planar surfaces towards the eastern margin. Part of the Tanzania Craton Parangan Granitic gneiss, psammite, semipelite Gradational contact with the Main Batholith.
Charac-terized by isoclinal folding overprinted by small up-right folds with axial planar crenulation. Westernmost domain of the Mozambique Belt
Granitic gneiss, amphibolite, kyanite schist, muscov- Formed by tectonic reworking of Parangan rocks. Re-Kondoan
ite schist working occurred through intense folding upon SE plunging axes and easterly dipping foliation surfaces to form flaggy gneisses
Biotite and hornblende gneiss, quartzite, marble
Usagaran True Mozambiquan cover sequence
metasedi-ments
Table 2 Sm–Nd data
Sm (ppm) 143Nd/144Nd 147Sm/144Nd
Nd (ppm) TDM
Sample o(Nd)
MZ2 28.18 4.018 0.511001910 0.0967 2557 −23.0
3.346 0.510691911 0.0914
22.03 2811
MZ5 −28.6
3.00
MZ6 0.4039 0.510783912 0.0810 2508 −25.9
MZ11 10.57 1.940 0.511190915 0.1105 2615 −20.6
7.151 0.510809913 0.0856
50.29 2562
MZ14 −25.8
25.24
MZ17 3.602 0.150809956 0.0859 2568 −25.8
1.929 0.511306915 0.1286
MZ20 9.028 2932 −19.9
2.739 0.511035911 0.1024
16.09 2636
MZ22 −22.9
42.70
MZ24 7.332 0.511019910 0.1034 2676 −23.3
9.240
MZ25 47.71 0.511324911 0.1166 2575 −18.5
3.794 0.511043910 0.0971
23.51 2515
MZ29 −22.2
8.505 0.510865911 0.0914
MZ30 56.03 2610 −25.2
8.249 0.511019916 0.1004
49.47 2611
MZ32 −23.0
MZ34 63.65 11.616 0.511217915 0.1098 2565 −20.0
doan tectonism perhaps as a slice of deeper basement above which Kondoan deformation passed.
3. Experimental methods
The samples were analysed for Sr and Nd iso-topic compositions as well as Sr, Nd, Rb and Sm concentrations using a Finnigan MAT262 mass spectrometer at the Pheasant Memorial Labora-tory (PML) of the Institute for the Study of the Earth’s Interior at Misasa, Japan. Biotite was obtained from ten of the samples by conventional heavy liquid and magnetic separation techniques and similarly analysed for Sr isotopic composition and Sr and Rb concentrations.
The analytical procedure for chemical separa-tion and mass spectrometry are described in Yoshikawa and Nakamura (1993) for the Rb – Sr technique and Shibata et al. (1989) for the Sm – Nd technique and are essentially similar to those described in Maboko and Nakamura (1996). The isotopic ratios were normalised to 87Sr
/86Sr
=
0.1194 and146Nd
/144Nd
=0.7219. Replicate
analy-ses of the La Jolla Nd standard gave
143Nd/144Nd=0.511920913 (2s, n=9) whereas
the NBS 987 standard gave 87Sr/86Sr=
0.71022498 (2s,n=8). The maximum 2s
uncer-tainty in the Sm/Nd and Rb/Sr ratios derived from long term reproducibilities of standard sam-ples is 2%. Typical blank values are 5 and 10 pg for Sm and Nd respectively and negligible. Corre-sponding values for Rb and Sr are 25 and 35 pg respectively and equally negligible.
4. Results
Results of the Nd isotopic analyses are shown in Table 2. The biotite and whole rock Rb – Sr data are shown in Table 3. Also indicated in Table 2 are Nd depleted mantle ages (TDM) calcu-lated assuming a linear evolution model for the mantle together with present day mantle 143Nd/ 144Nd and 147Sm/144Nd values of 0.513114 and
0.222, respectively (Michard et al., 1985). These parameters result in TDMages which are 0.2 Ga younger than those obtained using the parameters of Goldstein et al. (1984) but which are similar to those calculated according to the method of De-Paolo (1981). Table 2 also lists o(Nd) values
4.1. Crustal formation ages
4.1.1. Craton samples
Samples MZ30, MZ32 and MZ34 are granites
from the Tanzania Craton corresponding to the Main Batholith of Hepworth and Kennerley (1970). The structural and petrographic character-istics of the samples are summarised in Table 4.
Table 3 Rb–Sr data
87Sr/86Sr 87Rb/86Sr Age (Ma)
Sample no. Sr (ppm) Rb (ppm)
1.3834 0.70850497
152.0 318.4
US 9B WR
1.120648914 54.93
US 9B BIO 21.44 406.2 540911
499.1
MZ 2 WR 334.2 0.78091898 1.9411
MZ 2 BIO 24.64 676.7 1.2747291 79.616 491910
0.72334598 0.42628
MZ 5 WR 364.2 53.55
1.3041591 87.174
MZ 5 BIO 13.01 391.2 47099
MZ 11 WR 387.0 57.0 0.71903699 0.426899
165.59 531910
1.9686992 480.4
MZ 11 BIO 8.410
0.75696299 1.034901
MZ 14 WR 207.4 74.0
672.4
5.065 4.30392 384.85 648913
MZ 14 BIO
0.74463198 1.05124
MZ 20 WR 330.2 119.8
3.403896 288.97
MZ 22 BIO 6.674 665.3 643913
MZ 25 WR 257.1 142.6 0.756436922 1.60831
96.896 527910
1.4722491 653.9
MZ 25 BIO 19.56
1.03857699 8.7318
MZ 34 WR 93.21 280.8
Domain Sample no. Description MZ30
Main Batholith Underformed granite, qtz shows undulatory extinction MZ32 Underformed granite, qtz shows undulatory extinction Main Batholith
Main Batholith MZ34 Weakly deformed granite, qtz shows sutured margins and is slightly ribboned
Fine grained quartz-feldspathic gneiss moderately deformed by the Bubu event. Qtz recrys-MZ29
Main Batholith
tallized at margins and weakly ribboned
Coarse granite highly deformed by the Bubu event. Qtz occurs as a mosaic of recrystallized MZ25
Main Batholith
subgrains with sutured margins. Fine grained biotite and coarse epidote in areas of maxi-mum strain is syn-deformational
Main Batholith MZ24 Fine grained, highly deformed quartzofeldspathic gneiss. Host to granite sample MZ 25 MZ22
Parangan Highly deformed biotite gneiss. Qtz strongly ribboned and recrystallized into sub-grains Highly deformed biotite gneiss. Qtz strongly ribboned and recrystallized into sub-grains MZ20
Parangan
MZ17
Parangan Highly deformed biotite gneiss. Qtz strongly ribboned and recrystallized into sub-grains Parangan MZ14 Granitic gneiss. Qtz, plagioclase, microcline and biotite form granoblastic texture.
Metased-imentary
Parangan MZ11 Granitic gneiss similar to MZ 14. Metasedimentary nature not established Kondoan MZ6 Granitic gneiss similar to MZ14. Metasedimentary nature not established
MZ5 Granitic gneiss similar to MZ 14. Metasedimentary nature not established Usagaran
metasediments
Granitic gneiss similar to MZ 14. Metasedimentary MZ2
The granites yield TDM ages of 2.6 Ga and
o(Nd) values of between −20.0 and −25.2,
in-distinguishable from those obtained by Maboko and Nakamura (1996) from the southern part of the Tanzania Craton. Sample MZ29, a fine grained quartzo-feldpathic gneiss, yields a TDM age of 2.5 Ga and an o(Nd) value of −22.2,
similar to the values obtained from the granites. Sample MZ25 is a coarse granite which has been strongly deformed by the Bubu event whereas MZ24 is a fine grained, strongly deformed quartzo-feldspathic gneiss into which the granite MZ25 was intruded. The samples yield TDMages and o(Nd) values (2.6 Ga, −18.5 and 2.7 Ga, −23.3, respectively) which are indistinguishable
from those obtained from the undeformed
granites.
4.1.2. The Parangan Domain
Samples MZ11, MZ14, MZ17, MZ20 and MZ22 were collected from within the Parangan Domain. Their structural and petrographic char-acteristics are summarised in Table 3. All the samples except MZ20 yield, respectively, TDM ages of 2.6 Ga and o(Nd) values of between −20.6 and −25.8, similar to those obtained from samples in the craton while MZ20 yields a slightly older TDM age of 2.9 Ga (o(Nd)= −19.9). Sample MZ 20 has a147Sm
/144Nd ratio of 0.1286 which is close to the upper limit of normal crustal values and may have therefore suffered inter-crustal fractionation of its Sm/Nd ratio thereby increasing its TDM age. Thus, apart from the anomalously old age of MZ20, all the Parangan samples yield TDM ages which are indistinguish-able from those obtained in the craton.
4.1.3. The Kondoan and Usagaran Metasediments Domains
Samples MZ2, MZ5 and MZ14 were collected well within the Mozambique Belt in areas desig-nated by Hepworth (1972) as belonging to either the Kondoan Domain (sample MZ6) or the Usagaran Metasediments (samples MZ2 and MZ5). They are generally granitic gneisses con-sisting of quartz, plagioclase, microcline and bi-otite with a granoblastic texture (Table 4). The samples yield TDMages of between 2.5 and 2.8 Ga
and o(Nd) values of between −23.0 and −28.6,
which again fall within the range of values ob-tained from the craton.
4.1.4. The Eastern granulites
Isotopic data from the Eastern granulites end of the transect in the Usambara mountains has been reported by Maboko and Nakamura (sub-mitted). The mafic to felsic granulites from this area yield TDMages which lie in a very tight range between 1.0 and 1.1 Ga and o(Nd) values of
between 2.8 and 3.6, very different from the corre-sponding values in the amphibolite rocks to the west.
4.1.5. Rb–Sr biotite ages
In an attempt to constrain the westernmost boundary of the Mozambique Belt as defined by Neoproterozoic mineral cooling ages, biotite Rb – Sr ages were obtained on some of the samples from the traverse. Choice of biotite was dictated by the fact that it is a widespread mineral with a relatively low closure temperature (300°C, Jager, 1979) which makes it a sensitive indicator of thermal reactivation. At the same time, the closure temperature is sufficiently high to record true ‘orogenic’ uplift and cooling rather than sim-ple epirogenic adjustment as recorded by apatite fission track dates (Noble et al., 1997). In all cases the ages are calculated from mineral-whole rock pairs and all errors are quoted at the 95% confi-dence interval.
Biotite from two cratonic samples, MZ34 and MZ25, were dated. Biotite from the weakly de-formed granite MZ34 yields an age of 1473929 whereas that from the strongly deformed granite MZ25 yield an age of 527910 Ma.
Biotite from the two Parangan samples MZ22 and MZ20 yield ages of 643913 and 592912 Ma, respectively. The Kondoan samples MZ2, MZ5, MZ11 and MZ14 yield ages which show a general westward increasing trend from 491910 Ma for the most easterly sample (MZ2) to 6489
5. Discussion
5.1. The tectonothermal margin of the Mozam
-bique Belt
The westernmost dated sample MZ34 is a typical late to post-orogenic granite from the Tanzania Craton. Similar granites have been dated at about 2.6 Ga by Bell and Dodson (1981) whereas lepidolite ages of 2.5 Ga have been reported from elsewhere in the craton (Ca-hen et al., 1984). The biotite age of 1.47 Ga must therefore represent more than 50% partial resetting of the biotite’s Rb – Sr systematics. Al-ternatively, for some unknown reasons, the anomalously young age may be spurious. Biotite from sample MZ25, on the other hand, yields a typical Pan-African mineral cooling age of
530 Ma. Textural evidence suggest that the
bi-otite grew during the Bubu deformation.
Although this deformation has been described as being largely brittle, quartz from the sample had obviously experienced ductile deformation fol-lowed by recovery as evidenced by its occur-rence in a fine mosaic of sub-grains with sutured margins. Feldspar shows typical brittle deforma-tion behaviour and, away from the zones of maximum strain, the rock had not lost its granitic appearance. This may suggest that the deformation, and therefore biotite growth, oc-curred at temperatures not much higher than the 300°C closure temperature of biotite to Sr diffusion. If this is the case then the 530 Ma age may closely date the Bubu deformation. On the other hand, if not spurious, the strongly re-set age of sample MZ34, collected within the craton, shows that the thermal effects of the Neoproterozoic activity must have extended even further west into the craton itself. Thus, the data show that, unlike the situation in south eastern Tanzania where the craton is flanked by rocks showing Palaeoproterozoic mineral cooling ages, typical Mozambiquan ages in the Kondoa area directly abut on the mapped cratonic
boundary whereas the thermal effects of
Mozambiquan tectonism may have penetrated well into the craton itself.
5.2. The age of Mozambiquan metamorphism
The Parangan samples MZ20 and MZ22 im-mediately to the east of the craton show ages which are significantly older than the 530 Ma age of sample MZ25 (643913 and 592912) but which are still within range of Mozambi-quan cooling ages (B700 Ma). A similarly older age of 648913 Ma is yielded by the most
westerly of the Kondoan samples (MZ14)
whereas further east the biotite ages gradually decrease from 531910 to 47099 Ma before increasing again to 540911 in the granulites of the Usambara mountains. If sample MZ25 which we interpret to date the Bubu deforma-tion, is excluded, the botite ages in the Mozam-bique Belt proper show a general eastward decrease from 650 Ma just east of the margin before gradually decreasing further eastwards to
470 Ma. Two alternative explanations can be advanced for this apparent decrease: (i) Like the
1470 Ma age from the craton, the ages older than the 530 Ma Bubu deformation are par-tially reset relicts of a pre-Mozambiquan meta-morphic event. According to this view, only
ages younger than the 530 Ma Bubu
defor-mation, which are encountered from the vicinity of Kibaya, some 100 km east of the craton mar-gin, represent true Mozambiquan cooling. (ii) All ages east of the craton margin record true post-Mozambiquan cooling. The eastward de-crease reflects a true cooling gradient. This gra-dient indicates that the eastern part of the Mozambique Belt represents a deeper crustal level which cooled more slowly than the shal-lower western part. According to this view, the
530 Ma Bubu age is younger than the biotite ages to the immediate east of the craton margin because the Bubu deformation is a younger dis-crete event which probably dates tectonic escape towards the end of Mozambiquan tectonic activ-ity. This interpretation is consistent with the structural evidence which indicates that the Bubu deformation cross-cuts all pre-existing fab-rics.
up-per amphibolite facies metamorphism with a typi-cal high grade metamorphic fabric. Coarse-grained biotite lie in the foliation plane which Hepworth (1972) attributed to Mozambiquan de-formation with no evidence of later retrogression. The last metamorphism and deformation in these rocks, therefore, is of Mozambiquan age and must have occurred well above the biotite closure temperature. Thus, it is unlikely that the biotite would preserve pre-Mozambiquan relict ages. For this reason, the first alternative is considered to be unlikely and interpret the age gradient as reflect-ing true Mozambiquan coolreflect-ing.
An interpretation of the older biotite ages as reflecting true Pan African cooling has important implications for the age of Mozambiquan meta-morphism and the final assembly of Gondwana arising from the collision of east and west Gond-wana along the Mozambique Belt (Stern, 1994). Most of the isotopic evidence for the age of Mozambiquan metamorphism has so far been obtained from the granulite facies terranes in the eastern part of the belt (Coolen et al., 1982; Maboko et al., 1985; Muhongo and Lenoir, 1994). In the granulites, zircon U – Pb data suggest
that metamorphism occurred at 700 Ma and
certainly no later than the 630 Ma Sm – Nd garnet ages which are interpreted by Maboko and Nakamura (1995) as dating post-metamorphic cooling. Biotite ages, on the other hand, are much younger in both the granulites and surrounding gneisses (Maboko et al., 1985). The biotite Rb – Sr data presented here, however, show a strong age gradient which has been interpreted as indicating diachronous cooling across the belt with the west-ern part cooling below 300°C up to 150 Ma earlier than the eastern part. For a belt showing such a protracted cooling history, the peripheral areas, which most likely experienced minimal burial during metamorphism and hence the fastest post-metamorphic unroofing and cooling, are best suited for constraining the peak metamorphic age. Therefore, the 648914 Ma age of the most west-erly Kondoan sample (MZ14) is considered to be the best minimum estimate of the age of amphibo-lite facies metamorphism in the Mozambique Belt. This age is, however, only minimally younger than the preferred age of peak granulite facies
metamorphism in the east suggesting that there is no significant age difference between granulite facies metamorphism and the more pervasive am-phibolite facies metamorphism that characterise most of the Mozambique Belt. If the interpreta-tion of the Mozambique Belt as the locus of a continent – continent collision between East and West Gondwana (Burke et al., 1977; Burke and Sengor, 1986; Stern, 1994) is accepted, the com-bined isotopic data from both the periphery of the belt and the granulite complexes constrain this event to have occurred prior to 650 Ma.
5.3. The age of the Mozambiquan crust
The Nd isotopic data show that the mapped boundary between the Mozambique Belt and the Tanzania Craton do not represent a boundary between crustal blocks with different mantle ex-traction ages. The 2.6 – 2.8 Ga TDMages of the craton continue through the orogen with the only fundamental change occurring in the Usambara granulites, \250 km east of the craton margin, where TDM ages of 1 Ga are found. Similar results have been reported by Maboko (1995) and Moller et al. (1998) from traverses across central Tanzania. The apparent continuity of Archean crust into the Mozambique Belt in northern and central Tanzania is in marked contrast to the situation in south eastern Tanzania were there is a clear transition from Archean crust with TDMages of 2.5 – 2.8 Ga, similar to those observed in this study, to a younger Proterozoic Usagaran crust with TDM ages of 2.2 – 2.5 Ga (Maboko and Nakamura, 1996). Several explanations can be advanced for the continuity of Archean TDMages from the craton into the Mozambique belt in northern and central Tanzania:
1. The rocks east of the cratonic margin in north-ern Tanzania are wholly composed of sedi-ments eroded from the Tanzania Craton probably overlying a basement composed of Usagaran crust.
3. Except for the granulite complexes, the Mozambiquan crust is made up of reworked Archean crust of the Tanzania Craton without an intervening Palaeoproterozoic Usagaran analogue.
The first explanation is discounted as being unlikely that a belt which is more than 200 km wide would be entirely composed of sediments from one source terrane only. At least some evi-dence of component mixing away from the craton would be expected as the influence of sediments from more easterly sources increases. Moreover, even orthogneisses, some of which have SiO2 con-tents (56 – 58%, Maboko unpublished data) which are too low to preclude a purely anatectic origin from preexisting sediments yield TDM ages which are indistinguishable from those of Archean granites.
In a Gondwana reconstruction, there are sev-eral Archean blocks in southern India and Antarctica that are potential candidates for colli-sion with the Tanzania Craton to form the Mozambique Belt. However, all these blocks in-clude components with TDMages much older than the 2.9 Ga upper limit of our samples (see the review in Moller et al., 1998). Thus, neither of the possible candidates supports a derivation of the Mozambiquan crust from collisional fragments of another Archean Craton with a crustal age profile similar to that of the Tanzania Craton. This leaves only the last explanation that the Mozam-biquan crust in central and northern Tanzania is largely composed of reworked Archean crust of the Tanzania Craton as the most plausible explanation.
6. Concluding remarks
The distribution of TDMages in northern Tan-zania is in marked contrast to the situation ob-served in south eastern Tanzania were the
Archean Craton is bound by younger
Palaeoproterozoic crust which was later reworked during the Neoproterozoic to form the Mozam-bique Belt. The absence of an Usagaran equiva-lent in northern Tanzania coupled with the
presence of juvenile Neoproterozoic crust in the granulite complexes further shows that there is considerable spatial variation in the protolith ages of rocks involved in the Neoproterozoic tectonism that formed the Mozambiquan Belt. In particular the data suggest a fundamental difference between the southern part of the belt in which Pan African
tectonism overprinted Palaeoproterozoic
Usagaran crust and the central and northern parts where tectonothermal activity appeared to have reworked only Archean Crust.
Another difference between the southern and northern sectors of the Mozambique Belt is in the intensity and spatial extent of Pan African activ-ity. In the south, mineral cooling ages near the craton margin are Palaeoproterozoic (1850 Ma, Wendt et al., 1972) with Pan African ages occur-ring well away from the margin. In contrast, Pan African ages in the north directly abut on the craton which also seems to have been strongly reset locally. This difference indicates a weaker influence of Pan African tectonism in the south suggesting that the main axis of tectonic activity there lied further east.
Acknowledgements
I wish to acknowledge the friendship and sup-port of E. Nakamura and the technical assistance of Chie Sakaguchi, Nobuko Takeuchi and Akio Makishima during different stages of the experi-mental work. The research was conducted under the Visiting Research Scholar Program of the Institute for the Study of the Earth’s Interior, Okayama University at Misasa, Japan. Fieldwork was funded by a grant from the Tanzania Com-mission for Science and Technology.
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