Download - High-pressure granulite facies metamorphism in the Pan-African belt of eastern Tanzania: P-T-t evidence against granulite formation by continent collision

J. metamorphic Geol., 1998, 16, 491–509

High-pressure granulite facies metamorphism in the Pan-African beltof eastern Tanzania: P–T–t evidence against granulite formation bycontinent collisionP. APPEL (e-mai l : [email protected] -kie l .de) , A . MOLLER* AND V. SCHENKMineralogisch-Petrographisches Institut und Museum der Christian-Albrechts-Universitat zu Kiel, 24098 Kiel, Germany*Current address: Department of Applied Geology, University of New South Wales, Sydney, 2052 NSW, Australia

ABSTRACT To constrain the tectonic history of the Pan-African belt in Tanzania, we have studied the P–T evolutionof granulites from northern and eastern Tanzania representative for a large part of the southern Pan-African belt of East Africa (e.g. Pare, Usambara, Ukaguru and Uluguru Mountains). Thermobarometry(conventional and multireaction equilibria) on enderbites and metapelites gives 9.5–11 kbar and810±40 °C during peak metamorphism at 650–620 Ma. This is consistent with the occurrence of bothsillimanite and kyanite in metapelites and of the high-P granulite facies assemblage garnet–clinopyroxene–quartz in mafic rocks. Peak metamorphic conditions are surprisingly similar over a very large area withN-S and E-W extents of about 700 and 200 km respectively. The prograde metamorphic evolution in theentire area started in the kyanite field but evolved mainly within the sillimanite stability field. Theretrograde P–T evolution is characterized by late-stage kyanite in metapelites and garnet–clinopyroxenecoronas around orthopyroxene in meta-igneous rocks. This is in agreement with thermobarometric resultsand isotopic dating, indicating a period of nearly isobaric and slow cooling prior to tectonic uplift. Theanticlockwise P–T path could have resulted from magmatic underplating and loading of the lower conti-nental crust which caused heating and thickening of the crust. Substantial postmetamorphic crustalthickening of yet unknown age (presumably after 550 Ma) led subsequently to the exhumation of high-P granulites over a large area. The results are consistent with formation of the Pan-African granulites atan active continental margin where tonalitic intrusions caused crustal growth and heating 70–100 Maprior to continental collision. The P–T –t path contradicts recent geodynamic models which proposedtectonic crustal thickening due to continental collision between East and West Gondwana as the causeof granulite formation in the southern part of the Pan-African belt.

Key words: anticlockwise P–T path; mineral analyses; Mozambique Belt; Pan-African high-P granulites;thermobarometry.Mineral abbreviations are from Kretz (1983).

1995; Stern, 1994) involving the accretion of islandINTRODUCTION

arcs in the north (Egypt, Sudan) and the widespreadformation of Pan-African granulites in the southSince the definition of the Mozambique Belt in East

Africa by Holmes (1951), which was based on N-S (Tanzania, Mozambique, Malawi). The collisionmodel for the southern part of the belt is based ontrending structures younger than those of E-W trend

in the Tanzania Craton (Fig. 1), the presence of palaeomagnetic evidence (McWilliams, 1981) and agedeterminations, interpreted as dating a granulite faciesgranulites in this belt have played a major role in its

interpretation. event between 715 and 650 Ma (Coolen, 1980;Maboko et al., 1985; Muhongo & Lenoir, 1994).The network-like patterns of orogenic belts around

the Bangweulu Block and Tanzania Craton were Further support for the collisional model was givenby Shackleton (1976) and Berhe (1990), based oninitially interpreted as ensialic orogenies with rejuven-

ation of Archean or Early Proterozoic crust (Watson, their interpretation of some occurrences of mafic rocksuites as Pan-African ophiolites. However, no isotopic1976; Shackleton, 1973). It was argued that the

exhumation of granulites, thought to occur only in evidence yet confirms their Pan-African age andgeochemical studies from some of these bodies inthe well-exposed mountain ranges of the Pan-African

belt, resulted from strong vertical tectonics, leading to northern Tanzania do not support their ophioliticnature (Prochaska & Pohl, 1983). A further problemthe formation of horst-like structures. In more recent

geodynamic models, the Pan-African mobile belt of of the plate tectonic collision model is the lack ofhigh-P/low-T metamorphic rocks of Pan-African ageEast Africa is attributed to collision between East and

West Gondwana (Maboko et al., 1985; Meert et al., in the southern part of the belt. There has also been

491© Blackwell Science Inc., 0263-4929/98/$14.00Journal of Metamorphic Geology, Volume 16, Number 4, 1998, 491–510

492 P. APPEL ET AL .

lower pressures (Meinhold, 1970; Schenk et al., 1995).Both metamorphic styles found in the Usagaran areclearly distinguishable from those in the adjoiningPan-African belt as discussed below.

This study focuses on the metamorphic evolution ofthe Pan-African belt in Tanzania. Since recent geodyn-amic interpretations of the Pan-African belt are inconflict with existing geochronological data and arein part based on assumptions, a petrological studyhas been undertaken to constrain the geodynamicsetting on the basis of the P–T –t evolution of thegranulites.

STRUCTURAL AND PETROGRAPHIC OVERVIEW

Our investigation concentrates on metapelites andorthogneisses from the well-exposed mountain rangesof north-east and east Tanzania including a 250 kmtraverse oblique to the strike of the Mozambique Beltin the Pare and Usambara Mountains (Figs 1 & 2). Asecond traverse 250 km to the south runs through theUluguru Mountains. More than 1000 thin sectionsfrom about 450 outcrops were evaluated to compilemaps of the regional distribution of mineral assem-blages on a large scale (Fig. 2).

The subparallel lithological banding of most rockFig. 1. Simplified geological map of eastern Tanzania, that types is in part due to repetition by isoclinal foldingshows the main geological units and the locations of the study under granulite facies conditions. An early high-gradeareas. The north-south extent of the Usagaran Belt is deformation is visible in the large anorthosite bodiesunknown. Occurrences of high-pressure granulites within the

in the Uluguru Mountains and the Pare MountainsPan-African belt are shaded.and in the enderbitic intrusive rocks of the belt. Theregional schistosity in the anorthosite is parallel tothe axial plane of folded mafic dykes cutting throughno petrological study of the granulites to establish

whether their P–T –t paths are consistent with the anorthosite (Fig. 3a). The limbs of thin maficdykes show smaller parasitic folds that are partlyformation by tectonic crustal thickening processes.

Palaeomagnetic data (Meert et al., 1995) indicated sheared off the limb and thus form isolated maficmaterial in the anorthosite (Fig. 3a). Since thisthat the closure of the southern Mozambique Ocean

did not occur before 550 Ma. This is 70–100 Ma later isolated mafic material (orthopyroxene) is completelysurrounded by late-stage garnet–clinopyroxene reac-than the granulite facies event recently dated with

monazite and zircon between 650 and 620 Ma (Moller tion rims, a typical feature for the early stages of theretrograde evolution in many orthogneisses, it canet al., 1996).

In eastern Tanzania the Archean (2.6–2.5 Ma) be concluded that these intrusive rocks share thewhole deformation and metamorphic history withTanzania Craton is surrounded by the Early

Proterozoic (2.0–1.8 Ga) Ubendian/Usagaran Belt the country rocks. The same conclusion can bedrawn from mafic minerals of the anorthosites, which(Fig. 1). The extension of the Usagaran Belt to the

north as well as the location of its border against the are stretched parallel to the fold axes and retain thegarnet–clinopyroxene reaction rims described above.adjoining Pan-African belt in the east is not well

known. The fact that, at the south-eastern corner of Thus, the growth of these late-stage reaction rimsoutlasted deformation. Refolding of the granulitethe craton, biotite cooling ages in Usagaran rocks

become progressively younger towards the south-east facies banding into upright, open folds with axialplanes that locally contain migmatitic leucosomes, is(Meinhold, in: Wendt et al., 1972) has been used as an

argument for thermal rejuvenation of pre-existing attributed to the late stage uplift history of the belt.This uplift may also be responsible for the numerousProterozoic crust by the Pan-African orogeny. Recent

petrological work and isotopic dating in the Usagaran shear zones (Fig. 3b), in which garnet–hornblende–scapolite bearing assemblages locally replace theBelt showed a 2 Ga high-P/low-T metamorphic event

that led locally to eclogite formation during subduction prograde granulite facies mineralogy. Late stagecharnockite patches in intrusive granitic rocks whichof an oceanic lithosphere (Moller et al., 1995).

Cordierite–sillimanite gneisses found south-east of the are typical for Pan-African granulites from Sri Lankaand southern India cut and obliterate earlier granulitehigh-P belt, suggest locally higher temperatures and

HIGH-PRESSURE GRANULITES IN TANZANIA 493

Fig. 2. Geological map of the Uluguru Mountains (based on Sampson & Wright, 1964) in east Tanzania and the Pare andUsambara Mountains in northern Tanzania (based on Quarter Degree Sheets published by the Geological Survey of Tanganyika)and regional distribution of mineral assemblages of orthogneisses. Numbers are given for those samples discussed in the text. Amap with all sample numbers and locations is available for electronic retrieval via the www site of the Journal of MetamorphicGeology.

facies banding (Fig. 3c). A further important feature rock types. K-feldspar is very rare. Common accessoryminerals are ilmenite and magnetite; pyrite and copper-of the tectonometamorphic evolution of the Tanzanian

granulites are mafic rocks which cut the granulite sulphide minerals are scarce.The predominant mineral assemblage for orthogne-facies schistosity, but may develop a later foliation

(Fig. 3d). The presence of the high-P, garnet– isses of all granulite areas shown in Fig. 1 is garnet–clinopyroxene–plagioclase–quartz (Fig. 2), demon-clinopyroxene–quartz assemblage in these mafic rocks

indicate that they intruded prior to tectonic uplift strating that high-P granulite facies conditions wereattained over a very large area extending from thewhen the granulites were still in the deep crust.Furua Complex in the south (Coolen, 1980) to theNorth Pare Mountains.

Orthogneisses

Basic to intermediate orthogneisses (with prevailingMeta-anorthosites

quartz-normative tonalitic compositions) comprisemore than 90% of the exposed rock types in the Pan- Meta-anorthosites of the South Pare Mountains and

the western Uluguru Mountains are both characterizedAfrican belt of Tanzania (Fig. 2). Most contain plagio-clase–clinopyroxene–hornblende assemblages, with or by the occurrence of clinopyroxene and garnet as the

main mafic minerals but may contain minor orthopy-without quartz and garnet (Fig. 2). Garnet-bearingtwo-pyroxene granulites are more widespread in the roxene and locally scapolite. Titanomagnetite is

common and may be associated with green spinel.Pare and Usambara Mountains than in the UluguruMountains, where quartz- and/or orthopyroxene-free Closely associated with the anorthosite of the Uluguru

Mountains are small bodies of garnet–clinopyroxeneassemblages are more common. In addition to themineral assemblages shown in Fig. 2, biotite may occur rocks, interpreted as eclogites (Sampson & Wright,

1964; Muhongo & Lenoir, 1994). These rocks havein more felsic rocks and scapolite in calcium-rich basic


Fig. 3. (a) Folded mafic dyke within anorthosite. Regional schistosity parallel to axial plane of the fold. (b) Ductile shear zone withrotated blocks, exhibiting an earlier granulite facies foliation. (c) Intrusive charnockite ( light), disrupting the earlier granulite facieslayering (pyroxene bearing dark bands). (d) Intrusive mafic dyke cutting the earlier schistosity.

served as evidence for subduction-related Pan-African are ubiquitous throughout the whole area, whereasilmenite is only present in the Pare and Usambarametamorphism (Muhongo, 1994), but a high-P/low-T

origin is not supported by the mineral chemistry Mountains. Cordierite has not been found in anymetapelites of the Pan-African belt. Some samples(see below).from the eastern Uluguru Mountains contain muscov-ite (Table 1) within the metamorphic foliation which

Metapelitesseems to indicate a prograde formation of muscovite.Zinc-rich spinel occurs in the feldspar–quartz matrixMetapelites occur in the Pare and Usambara

Mountains as well as in the Uluguru Mountains of several samples from the Pare Mountains.intercalated with orthogneisses (Fig. 2). Surprisingly,many of the metapelites in all studied areas contain

PETROGRAPHY AND MINERAL CHEMISTRYboth sillimanite and kyanite in the peak metamorphicassemblage (Table 1). Most metapelites are garnet-

Analytical proceduresand biotite-rich. Those from the Pare and UsambaraMountains commonly contain K-feldspar whereas it is Mineral compositions were determined with a

CAMEBAX electron microprobe at the University ofabsent in many samples from the Uluguru Mountains.The regional distribution of mineral assemblages Kiel at 15 kV accelerating potential, using the PAP

(Pouchou & Pichoir, 1984) correction program.(Table 1) and inclusion textures in garnets point to aregional difference in the metamorphic conditions. In Plagioclase was analysed with a beam diameter of c. 8

mm. Scanning measurements were carried out over anthe eastern Uluguru Mountains, the typical peakmetamorphic assemblage is quartz–plagioclase–bio- area of c. 15×15 mm on clinopyroxene, showing fine,

but microscopically visible exsolution lamellae.tite–garnet-kyanite±K-feldspar whereas sillimanite-bearing assemblages are more common in the other Representative mineral analyses are given in Table 2–5,

a full dataset is available for electronic retrieval via theinvestigated areas (western Uluguru Mountains, Pareand Usambara Mountains). Graphite, rutile and pyrite WWW site of the Journal of Metamorphic Geology.


Table 1. Mineral assemblages ofmetapelites. All contain plagioclase andquartz in addition. Information aboutinclusions in garnet are only given forsillimanite, kyanite, spinel and staurolite.

Sample no. Kfs Bt Ms Grt Sil Ky St Spl Scp Rt

Uluguru Mountains

T2–1 X X X (X)

T2–2 X (X) X I (X) X

T2–4 X X (X) X XI X

T2–5 X X X XI X

T2–6 X X X I (X) X

T6–1+2 X (X) X

T6–5 X X X

T6–6 X X XI I I X

T6–8 X X (X) X X

T6–9 X X X X X

T6–10 X X X X X X

T10–1 X X X X I X X

T11–1 X X X X X X

T12–1 X X X XI

T12–5 X X X (X)I I X X

T22–2 X X (X)I X

T28–1 X X X X X X X

T34–1 X X

P1 X X X (X) X

P8–1 X X X X I I X

P9–2 X X X XI XI X

P13–2 X X (X) X XI (X) X

P26–1 X X X XI X

P26–2 X X XI XI X

P36–1 X X X X

P40–1 X X X XI X

P52–3 X X X X X

P52–4 X X XI X

P89–2 X X X

Pare- and Usambara Mountains

A16–1 X X X XI X X

A16–2 X X (X) X XI X X X

A16–4 X X X XI X

A17–4 X X X XI X

A23–1 X X XI X

A33–1 X X XI X

A39–3+5 X X X XI X

A40–2 X XI X

A46–1+4 X X XI X

A48–1 X X XI X

A101–5+6 X X (X) X XI

A101–7 X X X XI

A108–1 X X X X XI X

A108–3 X X X I X

A108–6 X X X X X

A113–1+7+8 X X X XI X

A115–2 X X X XI X

A128–1 X (X) X X

A128–4+5 X X XI X

A129–11 X X XI X

T82–1 X (X) X XI X

T82–2 X (X) X X X

T86–2 X X (X) X XI X

T86–3 X X X XI X X

T98–2 X X X X

T98–3 X X X X X

T98–5 X (X) X X X X

T98–6 X X XI X

T112–3 X (X) X XI X X

T114–6 X X X X X

T115–9 X X X XI X

T115–10+12 X X X XI X X

T115–11 X X X

T115–13 X X XI X X

T124–3 X X (X) X

T124–4 X X X X

T137–1 X X X XI X X X

T137–2+3 X X X XI X

X, matrix phase; (X), late stage phase; 1, inclusion in garnet; all samples contain plagioclase and quartz in addition.

gneisses, (2) clinopyroxene bearing gneisses and (3)Enderbitic orthogneisses

orthopyroxene bearing gneisses.Garnet commonly occurs in two generations in theFor the comparison of mineral compositions, enderbitic

orthogneisses are grouped into (1) two-pyroxene two-pyroxene gneisses. Large subhedral garnet–


Table 2. Representative analyses of garnet from two-pyroxene granulites and metapelites.

Rock Type Two-pyroxene Granulite Metapelite

Area

Uluguru Mts Ukaguru Pare Mts Usambara Mts Uluguru Mountains Pare Mts Us. Mts

Analysis no. 18-core 60-core 24-core 2-core 30-core 6-core 26-core 73-core 64-core 75-core 11-core 86-core 47-core 2-core 64-core 29-core 115-core 86-rim 85-core

Sample no. P25–2 P30–6 P103–1 T6–19 T19–1s A150–6 A26–3 T77–1 T94–3 A101–2 T134–1 T138–1 T6–6 T2–6 T11–1 T28–1s T86–3 T86–3 T137–3

SiO2 38.85 38.57 38.57 38.51 38.55 39.00 38.45 38.73 38.53 38.74 38.16 38.79 40.18 40.32 38.45 38.51 38.95 38.82 39.34

TiO2 0.07 0.02 0.06 0.12 0.05 0.03 0.08 0.09 0.09 0.00 0.07 0.13 0.01 0.01 0.02 0.00 0.01 0.01 0.01

Al2O3 21.45 21.02 20.76 20.77 21.83 21.81 20.89 21.01 20.89 21.02 20.64 21.07 22.56 22.25 21.6 21.65 21.27 21.80 21.94

Cr2O3 0.00 0.03 0.04 0.01 0.03 0.00 0.00 0.00 0.03 0.03 ND 0.09 0.00 0.01 0.00 0.03 ND ND ND

FeO* 26.73 26.52 26.67 26.20 26.51 25.31 29.37 25.42 27.96 26.47 25.57 26.10 24.02 22.82 29.73 27.04 25.47 26.39 28.27

MgO 6.31 6.60 6.30 6.27 5.81 7.49 4.57 7.61 5.59 6.62 6.65 6.67 12.32 13.49 7.19 7.68 9.28 10.03 9.58

MnO 0.95 1.16 2.20 1.40 1.10 0.72 1.17 0.71 0.89 1.18 1.42 0.59 0.24 0.25 0.26 0.44 0.49 0.56 0.36

CaO 6.55 5.96 6.14 6.80 6.92 6.07 6.85 6.82 6.84 6.52 6.76 6.78 1.73 1.76 2.55 4.32 4.32 1.44 1.48

Total 100.91 99.88 100.92 100.08 100.80 100.43 101.38 100.39 100.79 100.58 99.27 100.30 101.06 100.91 99.80 99.67 99.79 99.05 100.98

Si 6.00 6.02 6.02 6.02 5.97 6.00 6.01 5.99 6.01 6.01 6.00 6.02 6.00 6.00 6.01 5.99 6.00 6.00 6.00

Al 3.91 3.87 3.80 3.82 3.98 3.95 3.85 3.83 3.83 3.84 3.82 3.85 3.97 3.90 3.98 3.97 3.86 3.97 3.95

Ti 0.01 0.00 0.01 0.01 0.01 0.00 0.01 0.01 0.01 0.00 0.01 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Cr 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 ND 0.01 0.00 0.00 0.00 0.00 ND ND ND

Fe(2+)* 3.45 3.46 3.47 3.42 3.43 3.62 3.84 3.29 3.65 3.44 3.36 3.40 3.00 2.84 3.88 3.52 3.28 3.41 3.61

Mg 1.45 1.54 1.46 1.46 1.34 1.72 1.06 1.75 1.30 1.53 1.56 1.54 2.74 2.99 1.67 1.78 2.13 2.31 2.18

Mn 0.12 0.15 0.29 0.19 0.14 0.09 0.16 0.09 0.12 0.16 0.19 0.08 0.03 0.03 0.03 0.6 0.06 0.07 0.05

Ca 1.08 1.00 1.02 1.14 1.15 1.00 1.15 1.13 1.14 1.08 1.14 1.13 0.28 0.28 0.43 0.72 0.71 0.24 0.24

xFe 0.70 0.69 0.70 0.70 0.72 0.65 0.78 0.65 0.74 0.69 0.68 0.69 0.52 0.49 0.70 0.66 0.61 0.60 0.62

xAlm 0.57 0.56 0.56 0.55 0.57 0.54 0.62 0.53 0.59 0.55 0.54 0.55 0.50 0.46 0.65 0.58 0.53 0.57 0.59

xGrs 0.18 0.16 0.16 0.18 0.19 0.17 0.19 0.18 0.18 0.17 0.18 0.18 0.05 0.05 0.08 0.12 0.12 0.04 0.04

xPrp 0.24 0.25 0.23 0.24 0.22 0.28 0.17 0.28 0.21 0.25 0.25 0.25 0.45 0.49 0.28 0.29 0.34 0.38 0.36

xSps 0.02 0.03 0.05 0.03 0.02 0.02 0.03 0.02 0.02 0.03 0.03 0.01 0.01 0.01 0.00 0.01 0.00 0.01 0.01

Normalization, 24 oxygens; *all iron as FeO; ND, not determined.

Table 3. Representative analyses of plagioclase from two-pyroxene granulites and metapelites.

Rock Type Two-pyroxene Granulite Metapelite

Area

Uluguru Mts Ukaguru Pare Mts Usambara Mts Uluguru Mountains Pure Mts Us. Mts

Analysis no. 34 60 16 25 27 19 18 22 57 92 35 21 22 114 67-inc. 114-inc. 32-m.c. 86

Sample no. P25–2 P30–6 P103–1 T6–19 T19–1 A156–6 A26–6 A26–3 T77–1 T94–3 A101–2 T134–1 T138–1 T2–6 T11–1 T28–1s T86–3 T137–3

SiO2 61.09 61.93 61.20 60.92 60.43 60.90 61.94 56.98 61.02 60.13 59.27 59.71 60.75 60.57 57.60 55.53 61.68 62.11

Al2O3 24.73 23.92 24.79 25.42 25.76 24.37 24.5 25.49 24.37 25.28 26.54 26.17 24.13 25.61 27.26 27.59 23.86 23.61

Fe2O3 0.02 0.07 0.08 0.03 0.04 0.14 0.07 0.07 0.03 0.00 0.03 0.04 0.00 0.12 0.07 ND ND 0.00

CaO 6.85 6.12 6.78 6.29 7.20 6.09 5.97 8.85 6.53 7.29 8.40 7.85 6.33 7.37 9.37 10.75 5.31 5.46

Na2O 7.64 7.83 7.59 7.76 7.63 7.86 7.92 6.05 7.51 7.28 6.65 6.68 7.64 7.29 5.87 5.31 8.38 8.28

K2O 0.30 0.37 0.50 0.27 0.25 0.63 0.46 0.42 0.46 0.26 0.32 0.50 0.26 0.21 0.28 0.27 0.42 0.26

Total 100.63 100.24 100.94 100.69 101.31 99.99 100.86 97.86 99.92 100.24 101.21 100.95 99.11 101.17 100.81 99.45 99.65 99.72

Si 2.70 2.74 2.70 2.69 2.66 2.71 2.73 2.61 2.72 2.67 2.62 2.64 2.72 2.67 2.56 2.51 2.75 2.76

Al 1.29 1.25 1.29 1.32 1.34 1.28 1.27 1.38 1.28 1.32 1.38 1.36 1.27 1.33 1.43 1.47 1.25 1.24

Fe(3+)* 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 ND ND 0.00

Ca 0.33 0.29 0.32 0.30 0.34 0.29 0.28 0.43 0.31 0.35 0.40 0.37 0.30 0.35 0.46 0.52 0.25 0.26

Na 0.66 0.67 0.65 0.66 0.65 0.68 0.68 0.54 0.65 0.63 0.57 0.57 0.66 0.62 0.51 0.47 0.72 0.71

K 0.02 0.02 0.03 0.02 0.01 0.04 0.03 0.03 0.03 0.03 0.02 0.02 0.03 0.02 0.01 0.02 0.02 0.02

XAn 0.33 0.30 0.32 0.31 0.34 0.29 0.29 0.44 0.32 0.35 0.40 0.38 0.31 0.35 0.47 0.52 0.25 0.26

XAb 0.66 0.68 0.65 0.68 0.65 0.68 0.69 0.54 0.66 0.63 0.58 0.59 0.68 0.63 0.51 0.47 0.72 0.72

XKfs 0.02 0.02 0.03 0.02 0.01 0.04 0.03 0.03 0.03 0.02 0.02 0.03 0.02 0.01 0.02 0.02 0.02 0.02

Normalization, oxygens; *all irons as Fe2O3; ND, not determined; inc., inclusion in garnet; m.c., matrix-core.

porphyroblasts (garnet I) with inclusions of plagioclase granulites has a restricted range of compositions withXGrs in the range of 0.16–0.18, XFe between 0.63 andand quartz are interpreted to have grown during the

prograde stages of metamorphism. Late stage garnet 0.76 (Fig. 5a), and XSps is generally <0.02. The maindifference of late stage garnet II compared with garnet(garnet II) is highly poikilitic with abundant inclusions

of quartz and opaque phases. This garnet commonly I is the somewhat higher grossular content (2–3 mol%; Fig. 5b). Garnet in assemblages that contain onlyforms rims around iron oxides (mostly ilmenite and

magnetite) or may overgrow older cores of garnet I clinopyroxene has high XGrs (up to 0.29) and avariation in XFe between 0.48 and 0.82 (Fig. 5a). Theseporphyroblasts (Fig. 4a). Garnet I in two-pyroxene


Table 4. Representative analyses of orthopyroxene and clinopyroxene from two-pyroxene granulites and one garnet-clinopyroxenecumulate (T52-7).

Mineral Clinopyroxene Orthopyroxene

Area

Uluguru Mountains Ukaguru Pare Mts Usambara Mts Uluguru Mountains Ukaguru Pare Mts Us. Mts

Analysis no. 27 9 18 16 26 47 30 12 42 70 93 52 11 20 8 25 27 107

Sample no. P25–2 P30–6 P103–1 T19–1 T52–7 A156–6 A26–3 T77–1 T94–3 T134–1 T138–1 P30–6 P103–1 T6–19 T19–1 A156–6 T77–1 A101–2

SiO2 51.82 51.49 51.42 51.91 51.83 52.35 51.56 50.45 51.08 50.92 52.10 51.40 51.53 51.97 50.72 52.34 50.87 51.66

TiO2 0.39 0.22 0.29 0.28 0.20 0.24 0.34 0.32 0.33 0.36 0.31 0.06 0.07 0.10 0.09 0.08 0.07 0.08

Al2O3 2.76 3.38 3.81 2.56 4.53 2.47 2.83 3.75 3.41 4.14 2.56 1.64 1.76 1.13 0.98 1.12 2.18 1.99

Cr2O3 0.00 0.00 0.00 0.12 0.16 0.07 0.00 0.02 0.00 0.00 0.01 0.01 0.00 0.01 0.01 0.07 0.01 0.02

FeO* 11.94 10.96 11.14 10.90 6.01 7.81 15.14 9.56 12.39 11.77 11.52 26.94 26.78 26.55 30.45 23.11 26.05 26.87

MgO 12.00 11.17 11.22 11.51 13.32 14.01 10.84 12.15 10.67 11.89 13.52 18.92 18.43 19.11 17.03 22.38 19.97 19.22

MnO 0.16 0.16 0.44 0.28 0.13 0.10 0.21 0.09 0.09 0.26 0.12 0.35 0.67 0.45 0.64 0.21 0.41 0.46

CaO 20.44 20.95 21.49 21.98 22.34 22.01 18.92 22.56 21.06 20.46 19.55 0.33 0.47 0.55 0.56 0.53 0.44 0.49

Na2O 0.72 1.36 1.15 0.81 1.06 0.82 0.94 0.88 1.12 0.82 0.51 0.06 0.01 0.02 0.03 0.02 0.00 0.02

Total 100.23 99.69 100.96 100.35 99.58 99.88 100.78 99.78 100.15 100.62 100.20 99.71 99.72 99.89 100.51 99.86 100.00 100.81

Si 1.95 1.94 1.92 1.95 1.92 1.95 1.95 1.90 1.93 1.91 1.95 1.96 1.97 1.98 1.96 1.96 1.93 1.95

Al(IV) 0.05 0.06 0.08 0.05 0.08 0.05 0.05 0.10 0.07 0.09 0.05 0.04 0.03 0.02 0.04 0.04 0.07 0.05

Al(VI) 0.07 0.09 0.09 0.06 0.12 0.05 0.07 0.07 0.08 0.09 0.06 0.04 0.05 0.03 0.00 0.01 0.03 0.04

Ti 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Cr 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Fe(2+)* 0.38 0.35 0.35 0.34 0.19 0.24 0.48 0.30 0.39 0.37 0.36 0.86 0.86 0.84 0.98 0.72 0.83 0.85

Mg 0.67 0.63 0.63 0.64 0.73 0.78 0.61 0.68 0.60 0.66 0.75 1.08 1.05 1.08 0.98 1.25 1.13 1.08

Mn 0.01 0.01 0.01 0.01 0.00 0.00 0.01 0.00 0.00 0.01 0.00 0.01 0.02 0.01 0.02 0.01 0.01 0.02

Ca 0.82 0.85 0.86 0.88 0.89 0.88 0.77 0.91 0.85 0.82 0.78 0.01 0.02 0.02 0.02 0.02 0.02 0.02

Na 0.05 0.10 0.08 0.06 0.08 0.06 0.07 0.06 0.08 0.06 0.04 0.00 0.00 0.00 0.00 0.00 0.00 0.00

XFe 0.36 0.36 0.36 0.35 0.20 0.24 0.44 0.31 0.39 0.36 0.32 0.44 0.45 0.44 0.50 0.37 0.42 0.44

XJd 0.03 0.02 0.00 0.02 0.03 0.00 0.03 0.00 0.01 0.00 0.02

XAcm 0.02 0.08 0.08 0.04 0.04 0.07 0.04 0.12 0.07 0.06 0.02

XCa-Tsch 0.04 0.06 0.08 0.04 0.08 0.05 0.04 0.10 0.06 0.08 0.04

XFs 0.18 0.13 0.13 0.15 0.07 0.09 0.22 0.09 0.30 0.15 0.17

XEn 0.34 0.31 0.31 0.32 0.37 0.39 0.30 0.34 0.39 0.33 0.38

XWo 0.39 0.39 0.39 0.42 0.40 0.41 0.36 0.40 0.00 0.36 0.37

Normalization, 6 oxygen; all iron as FeO; endmembers were calculated for clinopyroxene after recalculation of mineral formula to six oxygens and four cations by a matrix method.

garnets are only weakly zoned with increasing XFetowards the rim.

Plagioclase is in the range An25–43 (Table 3). LargeTable 5. Representative analyses of staurolite from metapelitesof the Uluguru Mountains. All analyses were obtained from matrix-porphyroblasts have high XAn and are weaklystaurolite-inclusions in garnet. zoned with a decrease of 2–4 mol% towards the rim.

Inverse plagioclase zoning was rarely detected. NoAnalysis no. 30 33 112

Sample no. T6–6 T6–6 T12–5 compositional difference between plagioclase in two-pyroxene assemblages and orthopyroxene-free- or

SiO2 27.64 26.49 27.39clinopyroxene-free-assemblages was found. SmallTiO2 0.67 0.85 0.75

Al2O3 55.85 55.36 42.14 plagioclase grains crystallized during the formation ofCr2O3 0.10 0.22 0.11 garnet–clinopyroxene coronas around orthopyroxeneFeO 7.10 8.29 12.82

shows elevated anorthite contents compared withMgO 4.22 3.68 3.11

MnO 0.01 0.02 0.07 early-formed matrix plagioclase (5–8 mol% higher).ZnO 3.45 3.33 1.21

Hornblende exhibits equilibrium textures with otherNa2O 0.06 0.06 0.03

K2O 0.02 0.01 0.02 matrix phases indicating a prograde formation.Total 99.12 98.31 97.65 Titanium content of hornblende coexisting with ilmen-

ite and/or rutile is mostly between 0.2 and 0.3 pfu butSi 7.55 7.37 7.87

Ti 0.14 0.18 0.16 may increase up to 0.4 in a few cases. As observed forAl 17.98 18.16 17.66 garnet, the hornblende composition depends on theCr 0.02 0.05 0.03

mineral assemblage: XFe of hornblende in orthopyrox-Fe 1.62 1.93 3.08

Mg 1.72 1.53 1.33 ene-free rocks extends over a range of 0.43–0.68;Mn 0.00 0.00 0.02

whereas, it is confined to the narrow range of 0.53–0.60Zn 0.70 0.68 0.26

Na 0.03 0.03 0.02 in two-pyroxene assemblages.K 0.01 0.00 0.01 Prograde clinopyroxene is coarse grained (0.5–3 mm),XFe 0.49 0.56 0.70

has a distinct bright to yellowish green pleochroismNormalization, Si+Al=25.53 (Holdaway et al., 1986). and shows well-equilibrated textures with matrix


Fig. 4. (a) Garnet porphyroblast with relic kyanite inclusion (P26–1). Long dimension of the field of view is 4.3 mm. (b) Largegarnet and kyanite, both with prograde inclusions of sillimanite. Sample P9–2. Long dimension of the field of view is 1.7 mm.(c) Formation of late stage kyanite and biotite at the expense of garnet (T11–1). Long dimension of the field of view is 4.3 mm.(d) Late-stage kyanite as in Fig. 5(b), surrounded by later sillimanite (P9–2). Long dimension of the field of view is 1.7 mm.(e) Garnet porphyroblast in a two-pyroxene granulite (P81–3). Late-stage garnet growth is documented by garnet rims (garnet II)around orthopyroxene and poikilitic garnet II-rims at the porphyroblast. Long dimension of the field of view is 4.3 mm.(f ) Formation of late-stage garnet and clinopyroxene, rimming pre-existing orthopyroxene in a charnockitic sample (T8–93). Longdimension of the field of view is 1.7 mm.

phases. Late stage clinopyroxene is developed in some orthopyroxene (Fig. 5a). Jadeite component is negli-gible while acmite component may reach 10 mol%.samples as a narrow rim around orthopyroxene mostly

together with garnet rims (Fig. 4b). XFe of clinopyrox- Ca-Tschermaks component is generally in the rangeof 8–10 mol%. Compositional zoning is only recog-ene varies only between 0.23 and 0.37 in two-pyroxene

rocks but can be lower in assemblages without nized in large porphyroblasts, where diopside- and


diopside with a negligible jadeite component (3 mol%;Table 4, T52–7). Petrographically the diopside ishomogenous and shows no signs of plagioclase orpyroxene exsolutions, thus excluding equilibrationunder eclogite facies conditions. They most probablyrepresent granulite facies cumulates which are commonin many anorthosite-complexes.

Metapelites

Sillimanite and kyanite occur as inclusions in garnetand as matrix phases (Fig. 4c,d). The majority of theinclusions is sillimanite, commonly showing preferredorientation due to synmetamorphic growth. The overabundance of these sillimanite inclusions is evidencethat prograde metamorphism proceeded mainly withinthe stability field of sillimanite. Scarce kyaniteinclusions in garnet (Fig. 4c) are interpreted as theoldest relics of the prograde metamorphism, thatformed prior to sillimanite. In the eastern UluguruMountains the crystallization relationships during theprograde evolution are at variance with the otherareas, since kyanite is here the most common alumosil-icate inclusion in garnet.

Fig. 5. (a) Ca–Fe–Mg plot of coexisting orthopyroxene, A second generation of kyanite in metapelites of theclinopyroxene and garnet from different rock types. whole investigated area was formed by the retrograde(b) Composition of cores of garnet-porphyroblasts (garnet I)

reaction:and late-stage garnet (garnet II). Late-stage garnets alwaysshow higher XGrs than garnet cores. Kfs+Grt+H2O=Qtz+Ky+Bt

In many samples this reaction led to the completeconsumption of K-feldspar and a partial breakdownCa-Tschermaks component increase towards the rim,

while XFe slightly decreases. of garnet (Fig. 4e). A good example for preserved age-relationships between sillimanite and younger kyaniteOrthopyroxene is anhedral and commonly rimmed

by late stage poikilitic garnet and/or clinopyroxene. has been found in sample P9, where large garnet andretrograde kyanite–porphyroblasts contain progradeNo significant compositional zoning was found.

Chemical variation between different samples is small sillimanite I inclusions (Fig. 4d). In the same sample asecond generation of sillimanite (Sil II) has formed at(Fig. 5a). The XFe ranges from 0.36 to 0.51, Al2O3-

content is between 1.0 and 2.5 wt%. Orthopyroxene in the rims of retrograde kyanite blasts (Fig. 4f ).From the described textures the general alumosilicateclinopyroxene-free assemblages tends to have XAl

between 0.04 and 0.05, in two-pyroxene assemblages succession in the investigated area is Ky I(inclusion)�Sil I (inclusion and matrix)�Ky IIthe range is significantly higher.(retrograde)�Sil II (retrograde). With the exception ofthe kyanite I inclusions in garnet, the crystallization

Meta-anorthosites and related rockssequence reconstructed from many samples is alsoobserved in one single sample (P9–2).Mafic minerals of the meta-anorthosites are all

Mg-rich: e.g. garnet (XFe=0.45), orthopyroxene (XFe= Ellipsoidal garnet–porphyroblasts (2–5 mm) areabundant in metapelites and commonly have poikilitic0.26) and clinopyroxene (XFe=0.18). Relics of mag-

matic orthopyroxene are rare and where present, form cores. The most abundant inclusions in garnet aresillimanite, kyanite, quartz, rutile and in minor amountsthe cores of clinopyroxene–garnet augen. Plagioclase

is normally unzoned. Its anorthite-content varies plagioclase and biotite. Core composition of garnet isin the range Alm38–72, Prp20–57 and Grs04–07, butbetween An35–72, but is typically in the range An40–55.

Meionite-rich scapolite (Mei75–80) occurs in many grossular may reach 11 mol% in samples with pre-served growth zonation (Table 2). Spessartine compo-samples from both anorthosite complexes.

Small bodies of mafic rocks associated with the nent is generally very low (<1 mol%). Zoning patternsof most analysed garnets are similar with largeanorthosite of the Uluguru Mountains, which have

been described as eclogites, contain the mineral homogeneous cores and increasing XFe near the rims.Exceptions have been found in a few samples withassemblage garnet–clinopyroxene–titanite–rutile–horn-

blende–zoisite beside small amounts of quartz and large garnets exhibiting strongly decreasing grossularcontents towards the rim.plagioclase. The clinopyroxene in these rocks is


Feldspar in the matrix is mostly plagioclase, normallyunzoned with An16–38 (Table 3). Plagioclase inclusionsin garnet–porphyroblasts are generally less sodic thancore compositions of matrix-plagioclase, with differ-ences in anorthite contents between 7 and 27 mol%.K-feldspar composition lies in the range of Kfs82–92.

Matrix biotite shows strong variation in XFe fromabout 0.50–0.82. TiO2 contents are in the range of3.8–4.5 wt% (0.18–0.30 pfu), but may be as high as5.5 wt%. AlIV is about 0.16–0.36. Muscovite, found inthe foliation of some metapelites in the eastern UluguruMountains, has near stoichiometric Si contents ( lessthan 3.2 pfu).

Staurolite has been found as rounded inclusions ingarnet of two samples from the Uluguru Mountainsonly (T6–6, T12–5). The staurolite in T12 has a normalXFe of 0.70, whereas that of T6–6 is extremely Fe-poor(Table 5). Such Mg–staurolites are only known from afew other granulites (e.g. Schreyer et al., 1984) and high-P rocks (e.g. Simon et al., 1997; for a compilation seeLattard & Bubenik, 1995). Sample T6–6 is quartziticand contains sillimanite, kyanite, garnet and minoramounts of plagioclase and K-feldspar. Staurolite inthis sample occurs together with kyanite in the sameinclusion. This texture is interpreted to have formed bythe divariant reaction St+Qtz=Ky+Grt+H2O. Thestaurolite is ZnO-rich (2.5–3.5 wt%) and TiO2-poor(<1 at %). The distribution coefficient for Fe–Mg Fig. 6. AFM projection from K-feldspar for metapelites from

the Uluguru Mountains (a) and the Pare and UsambaraKGrt–hStD

is near one, but in most cases XFe in stauroliteMountains (b).(XFe=0.49–0.55) is slightly higher than that of the

including garnet (XFe=0.51–0.53 in core). The relativelyhigh variation in XFe may be due to the fact that some phic temperatures and points to different water

fugacities as the main cause for the shift of the three-staurolite grains—but not all—are completely separatedfrom the including garnet by a rim of kyanite. This may phase fields. The most Fe-rich sample (T2–2) was

collected from a late stage metapelitic shear zone.have prevented Fe–Mg equilibration between stauroliteand garnet. XFe of such isolated staurolite is higher (up Since garnet–biotite thermometry of this biotite-rich

sample indicates high equilibration temperatures ofto 0.55) than XFe of staurolite in direct contact withgarnet. Another explanation for the scatter in XFe of about 700 °C, the Fe-rich composition of the coexisting

phases may reflect high water activities in the shearstaurolite could be that manteling of kyanite preventeddecomposition of staurolite and quartz, whereas zone.

Compared with normal Mg-rich metapelites dis-unmanteled staurolite shifted to more Mg-rich composi-tions due to the progress of the continuous reaction. cussed above, Fe-rich rocks with the GRAIL assem-

blage (garnet–sillimanite–quartz–alumosilicate–rutile)are rare and were only found in the Pare MountainsPhase relations(Fig. 6, T82–1). The width of the two phase fieldgarnet–rutile (+sillimanite+quartz+K-feldspar) isMetapelites containing the low-variance assemblage

garnet–K-feldspar–plagioclase–biotite–alumosilicate– similar to that found in the south-east HighlandComplex of Sri Lanka (Raase & Schenk, 1994) andquartz exhibit a strong variation of Fe/Mg-ratios in

coexisting garnet and biotite. This is obvious from the indicates overstepping of the reaction Bt+Ilm+Qtz=Grt+Rt+Kfs+H2O in the Pare and UsambaraAFM-diagram (Fig. 6) where the three-phase fields

garnet–biotite–sillimanite of samples from the Pare Mountains.and Usambara Mountains are further to the Mg-sidecompared with the samples from the Uluguru

THERMOBAROMETRYMountains. The only exception from this generalpattern are samples from a locality in the northern The widespread occurrence of two-pyroxene–garnet

granulites within the Pan-African belt of east TanzaniaUluguru Mountains (T2–5, T2–6).Nevertheless, the Fe–Mg distribution between garnet provides favourable conditions for regional thermobar-

ometry because these rocks contain assemblages suitedand biotite shows no correlation with increasingMg-enrichment. This argues against varying metamor- for a large set of well-calibrated and widely applied geo-


thermobarometers. Pressure and temperature estimates prograde growth zoning is rarely preserved, core–compositions of minerals were used in most cases towere made by data processing with mineral formulae

calculated without any correction for ferric iron. calculate the peak metamorphic P–T conditions.Garnet–orthopyroxene–plagioclase–quartz bar-Pressure estimates for enderbitic gneisses were

calculated with garnet–clinopyroxene–plagioclase– ometry gives pressures between 9 and 12 kbar at areference temperature of 800 °C. Pressures calculatedquartz and garnet–orthopyroxene–plagioclase- quartz

net-transfer reactions, and for metapelites containing from the Mg-reaction (reaction (1) in Table 6) arehigher (mostly in the range 1–2.5 kbar) than thosethe assemblage garnet–sillimanite/kyanite–plagioclase–

quartz with the GASP (3An=Grs+2Als+Qtz) reac- from the Fe-reaction (Fig. 7). This behaviour is inaccordance with results of Perkins & Chipera (1985),tion. Table 6 gives a complete list of the equilibria used

and references for their calibrations. Our preferred who showed that pressures based on the Mg-reactioncan be exceptionally high for Fe-rich compositions.estimates of maximum temperatures are based on the

exchange of Mg and Fe between clinopyroxene (or Garnet–clinopyroxene–plagioclase–quartz bar-ometry yields slightly lower pressures in the sameorthopyroxene) and garnet. The calibrations of Krogh

(1988) and Ellis & Green (1979) were used for garnet– samples when compared to results of orthopyroxene-barometry. Differences in pressure between the twoclinopyroxene thermometry, the calibrations of Harley

(1984), Bhattacharya et al. (1991) and Sen & barometers are in the range of 0.5–1 kbar with theGADS-calibration of Perkins & Chipera (1985), butBhattacharya (1984) for garnet–orthopyroxene ther-

mometry. For metapelitic compositions only the only 0.2–0.6 kbar with the calibration of Eckertet al. (1991).garnet–biotite Fe–Mg exchange is applicable for tem-

perature estimates. Because re-equilibration effects tend Calculated pressures for orthopyroxene-free assem-blages are higher than 9 kbar in all samples andto give erroneous results garnet–biotite temperatures

should be interpreted with caution. generally match the pressure estimates for two-pyroxene granulites. However, some samples giveTWEEQU calculations (Berman, 1991) were per-

formed for a number of two-pyroxene granulites with unrealistic high pressures of 13 kbar or more (Table 7).Garnet–clinopyroxene Fe–Mg exchange ther-the assemblage garnet–clinopyroxene–orthopyroxene–

plagioclase–quartz in the system MgO–CaO–FeO–SiO2–Al2O3. The compositions of these minerals canbest be approximated by the end members diopside–ferrosilite–enstatite–pyrope–grossular–almandine–anorthite, giving a set of 11 possible reactions of whichthree are linearly independent. Calculations wereperformed using the version ‘Jun92’ of the thermo-dynamic dataset of Berman (1988). Activity modelsare those of Berman (1990) for garnet and Fuhrman& Lindsley (1988) for plagioclase. An ideal mixingmodel for clino- and orthopyroxene has been chosen.

Conditions of peak metamorphism

Orthogneisses

Representative P–T estimates of samples from allinvestigated areas are presented in Table 7. Since

Table 6. Equilibria used for geothermobarometric calculations.

Grt–Cpx–Pl–Qtz

(1) An+Di=2/3 Grs+1/3Prp+Qtz (‘GADS’)

Perkins & Newton, 1981 (PN); Eckert et al., 1991 (E);

Moecher et al., 1988 (M–Mg)

(2) An+Hd=2/3Grs+1/3Alm+Qtz

Moecher et al., 1988 (M–Fe)

Grt–Opx–Pl–Qtz

(3) An+En=1/3Grs+2/3Prp+Qtz

Bhattacharya et al., 1991 (B–Mg); Perkins & Newton, 1981 (PN–Mg)

Perkins & Chipera, 1985 (PC–Mg)

(4) An+Fs=1/3Grs+2/3 Alm+Qtz Fig. 7. Pressure estimates using the Fe- and Mg-reactions ofBhattacharya et al., 1991 (B–Fe); Perkins & Chipera, 1985 (PC–Fe) the garnet–orthopyroxene–plagioclase–quartz barometer from

Eckert et al., 1991 (E–Fe) the calibrations of Bhattacharya et al. (1991) and Perkins &Metapelites (Grt–Pl–Als–Qtz)

Chipera (1985). Pressures were calculated with core(5) 3An=Grs+2Als+Qtz (‘GASP’)

compositions at a reference temperature of 800 °C. Only theKoziol & Newton, 1988; Newton & Haselton, 1981results for two-pyroxene granulites are shown.


mometers based on the calibrations of Krogh (1988) indicator of metamorphic grade. The Tanzanian horn-blendes coexist with ilmenite±rutile, ±magnetite andand Ellis & Green (1979) record temperatures between

750 and 910 °C in two-pyroxene assemblages (Table 7). have high Ti-contents of 0.2–0.3 pfu, comparable togranulite facies hornblende from South India and SriTemperatures calculated from the calibration of Ellis

& Green (1979) are close to the TWEEQU-tempera- Lanka for which metamorphic temperatures of about700–750 and 850 °C have been estimated (Raase et al.,tures, while the calibration of Krogh (1988) generally

yields temperatures about 60 °C lower. Orthopyroxene- 1986; Schumacher et al., 1990). Since apart from hightemperatures low fO2 will also increase the Ti-contentfree samples show remarkably lower temperature

differences (10–50 °C) between the Ellis & Green- and of hornblende (Spear, 1981), the temperature estimatesfor these other areas cannot be directly applied to thethe Krogh-calibrations than for two-pyroxene rocks

where differences are in the range of 50–70 °C. The Tanzanian granulites.Bhattacharya et al. (1991) orthopyroxene–garnet tem-peratures are similar to the Ellis & Green and

MetapelitesTWEEQU results, whereas the Harley (1984) garnet–orthopyroxene calibration yields 50–100 °C lower tem- Figure 9 shows pressure estimates for metapelites of

the Pare, Usambara and Uluguru Mountains, calcu-peratures. Only the latter values are in contradictionwith sillimanite stable during peak pressures and lated with the GASP-calibration of Koziol & Newton

(1988). Peak metamorphic conditions were obtainedtemperatures.P–T estimates of metamorphism with the TWEEQU using garnet rim compositions, which may have lower

XGrs than garnet cores, and core compositions of themethod were performed on 10 garnet–pyroxene granu-lites with unzoned minerals and exhibiting textural mostly unzoned plagioclase. Again, the results for all

areas plot close to the kyanite–sillimanite boundaryequilibria (Fig. 8). Weighted averages are given inTable 7. All P–T estimates for the peak metamorphic and pass through the field of peak metamorphic

conditions as obtained by garnet–pyroxene thermobar-conditions are in a narrow range. In accordance withresults from conventional thermobarometry most ometry. However, some results fall into the kyanite

stability field despite the fact that sillimanite was thepressure estimates are between 10.5 and 11 kbar, whilethe range of most temperature estimates is 820–890 °C. stable aluminosilicate during peak metamorphism.

With the exception of one Ca-rich sample (T28) theThese P–T values plot close to the kyanite–sillimaniteboundary and are thus nicely in agreement with pressure overestimate is apparently less than 1 kbar

(Fig. 9).petrographic observations. Regional differences of peakmetamorphic conditions could not be detected by Information about the prograde P–T path was

gained from samples in which the garnet has preservedmeans of thermobarometry, although pressures tendto be slightly lower (0.5–1 kbar) in the Pare and strong grossular zonation (e.g. Figure 10), high in the

core (Grs11 in T86–3) and decreasing towards the rimUsambara Mountains than in the Uluguru Mountains.In addition to quantitative thermobarometry, the (Grs04). The same compositional trend is observed in

the coexisting plagioclase: Inclusions in the garnet-Ti-content of hornblende can be used as a qualitativecore are anorthite-rich (An52 in T86–3) while matrixplagioclase is An25. This compositional behaviourcorresponds to an increase of Keq for the GASP-reaction from about 90 for combinations of garnet-and plagioclase-core to 210 for combinations of garnet-rim with plagioclase-matrix. This is evidence that thekyanite–sillimanite boundary was crossed during pro-grade garnet growth (Figs 9 & 10).

According to uncertainties of garnet–biotite ther-mometry of granulites the Fe–Mg distribution betweenunzoned garnet cores and matrix biotite of the samplesscatter in KGrt–Bt

Dbetween 0.27 and 0.36. Corres-

pondingly, garnet–biotite temperatures scatter in therange of 750–870 °C and 760–920 °C with the cali-bration of Thompson (1976) and Ferry & Spear (1978),respectively.

Fig. 8. Peak-metamorphic conditions calculated with Late stage assemblagesTWEEQU-thermobarometry (Berman, 1991) for two-pyroxenegranulites from the Pare, Usambara and Uluguru Mountains. Conditions of postpeak metamorphism were calculatedTriangular fields give the area between the reaction curves.

with compositions of garnet II, late-stage plagioclaseRetrograde conditions calculated with late-stage corona garnet,and late-stage clinopyroxene in orthogneisseslate-stage clinopyroxene, late-stage plagioclase and

orthopyroxene-rim compositions of sample A156. (Fig. 4a,b; Table 7).

HIG

H-P

RESSU

RE

GR

AN

ULIT

ESIN

TAN

ZA

NIA

503

Table 7. Representative pressure and temperature estimates from selected samples. Pressures were calculated at a temperature of 800 °C, temperatures at a pressure of 10 kbar.Late-stage metamorphic pressures calculated for 700 °C, late-stage temperatures for 9 kbar. Abbreviations of applied equilibria and authors see Table 6.

Grt–Cpx Grt–Opx Grt–Opx–Pl–Qtz Grt–Cpx–Pl–Qtz

Sample Combination Assemblage TEG TK TH TSB TB TTWQ PPC–Mg PPC–Fe PB–Mg PB–Fe PPN–Mg PE–Fe PM–Mg PM–Fe PPN PE PTWQ

Uluguru Mountains

P21–2 core Grt–Cpx–Opx–Pl–Qtz 844 779 744 804 795 — 11.1 10.4 10.0 9.0 10.0 10.0 10.5 10.6 9.6 10.6 —

P21–2 late stage Grt–Cpx–Pl–Qtz* 741 683 9.9 9.2 8.5 8.9

P22–2 core Grt–Cpx–Opx–Pl–Qtz 820 761 762 830 815 817 11.5 10.7 10.6 9.3 10.2 10.3 10.8 11.9 9.6 9.9 11.7




P30–6 rim Grt–Cpx–Pl–Qtz* 773 707 8.9 8.7 9.1 9.6










T2–8 core Grt–Cpx–Opx–Pl–Qtz 784 728 744 838 820 — 10.1 10.3 10.0 8.3 9.7 9.7 9.5 11.5 8.7 9.1 —

T6–19 core Grt–Cpx–Opx–Pl–Qtz 750 687 747 810 804 820 11.5 10.9 10.5 9.4 10.2 10.3 10.8 12.6 9.6 10.0 10.8

T16–1s core Grt–Cpx–Opx–Pl–Qtz 822 770 698 749 759 — 12.0 11.1 10.0 10.0 9.9 9.9 12.0 11.3 10.3 10.6 —

T16–1s late stage Grt–Cpx–Pl–Qtz* 677 616 11.6 11.1 9.4 9.9




MST130 core Grt–Cpx–Pl–Qtz 939 924 12.3 15.3 11.6 12.0

P5 core Grt–Cpx–Pl–Qtz 773 741 14.3 17.1 12.7 13.1

P11–3 core Grt–Cpx–Pl–Qtz 819 770 11.6 11.5 9.6 9.9


P59–G core Grt–Cpx–Pl–Qtz 944 928 14.3 14.2 12.9 13.2



T9–2 core Grt–Cpx–Pl–Qtz 803 777 12.2 14.7 11.1 11.5





P66–G core Grt–Opx–Pl–Qtz 727 779 778 10.7 11.1 10.5 9.4 10.1 10.1

P88–7 core Grt–Opx–Pl–Qtz 816 908 863 12.0 10.3 11.0 9.1 10.7 10.7

T16–17 core Grt–Opx–Pl–Qtz 742 794 788 9.9 10.7 11.6 12.7 9.6 9.7

Ukaguru Area

T8–2–93 late stage Grt–Cpx–Opx–Pl–Qtz* 702 635 691 738 — 10.0 10.1 10.4 8.5 10.0 9.5 9.8 11.2 8.8 9.3 —

A156–6 late stage Grt–Cpx–Opx–Pl–Qtz* 712 639 690 738 747 688 10.0 10.4 10.7 8.8 9.8 9.8 9.9 11.0 8.9 9.3 9.4

Umba Steppe

A141–1 core Grt–Opx–Pl–Qtz 741 794 788 9.6 10.4 9.9 8.4 8.6 8.6

A147–1 core Grt–Cpx–Pl–Qtz 839 799 11.9 12.7 9.9 10.0

Pare Mountains

T77–1 core Grt–Cpx–Opx–Pl–Qtz 834 781 831 925 880 891 9.9 10.8 10.7 8.1 9.9 10.0 10.1 10.3 8.9 9.3



A26–3 core Grt–Cpx–Opx–Pl–Qtz 813 758 705 760 766 — 11.8 10.5 9.7 9.6 9.4 9.5 11.4 11.9 9.6 9.0 —

Usambara Mountains

A101–2 core Grt–Cpx–Opx–Pl–Qtz 864 810 769 841 823 10.9 10.3 10.2 8.7 9.8 9.9 10.5 11.1 9.4 9.7

T122–1 core Grt–Cpx–Opx–Pl–Qtz 888 843 764 822 820 10.8 10.1 9.9 8.5 9.6 9.6 10.8 11.1 9.5 9.9



Kitumbi Quarry, 60 km south of Korogwe (Usambara)

A154–3 core Grt–Cpx–Pl–Qtz 792 732 11.8 12.5 9.6 10.1


Fig. 9. (a) Pressure estimates for metapelites of the Pare and Usambara Mountains and (b) of the Uluguru Mountains that arecalculated with the calibration of Koziol & Newton (1988). Each line type corresponds to a specific sample. Temperature range forprograde pressure estimates is based on the assumption that garnet growth starts at temperatures of about 500 °C.

The results of garnet–clinopyroxene thermometry mineral inclusions, but only rarely by mineral zonation.Because of kyanite inclusions in metapelitic garnet,are between 550 and 750 °C (Fig. 11). Comparing

results of the same garnet–clinopyroxene calibration, which may occur together with staurolite, it isconcluded that the early prograde part of the P–Tlate-stage temperatures are always 100–150 °C lower

than those for peak assemblages in the same rock. The path was within the stability field of kyanite andstaurolite. The composition of garnet cores coexistinglate-stage (GADS) pressures are in the range 8–10 kbar

and thus only slightly lower than peak pressures with these inclusions allow the calculation of progradeGASP-equilibria. These are in the kyanite stability(Fig. 11). Thermobarometry on late stage assemblages

thus point to a near-isobaric cooling path after peak field (T66, Fig. 9a), in accordance with the petrographicobservations and high density fluid inclusions foundmetamorphism. This is consistent with textures

reflecting garnet growth due to the reaction in staurolite (Herms & Schenk, 1998). Most of theprograde evolution must have taken place in the

Opx+Pl=Grt+Cpx+Qtz (Fig. 4b),stability field of sillimanite because of the abundanceof sillimanite inclusions in garnet of most areas (stagewhich has a slope of about 55 °C/kbar (Green &

Ringwood, 1967; Harley, 1989). Good examples of 2, Fig. 12). This interpretation is supported by thegarnet zoning pattern of a sample from the Parelate-stage garnet–clinopyroxene coronas between

orthopyroxene and plagioclase have been found in Mountains (T86–3, Fig. 10). The grossular-rich coreincludes calcic plagioclase (An52), whereas the grossu-charnockitic rocks about 100 km to the west of the

Uluguru Mountains (Fig. 4b). These rocks, which lar-poor rim coexists with An25 of the matrix. ResultingGASP-equilibria indicate that the kyanite–sillimanitecontain no prograde garnet, developed the well-

preserved coronas after any penetrative deformation. boundary was crossed during prograde metamorphism.Peak metamorphic conditions are close to theLocal equilibrium seems to have been attained in the

corona of sample A156 as testified by the TWEEQU- sillimanite–kyanite boundary, but most samples liewithin the sillimanite stability field (Fig. 12, stage 3).method, which shows nearly perfect convergence of

the calculated reaction lines at 690 °C/9.5 kbar (Fig. 8). The P–T calculations indicate very similar conditionsfor the peak of metamorphism and a mean geothermalEven the large orthopyroxene–porphyroblasts of these

samples are unzoned with respect to XFe and XAl. gradient of about 22 °C km−1 throughout the investi-gated area.

Textural observations as well as thermobarometricTHE P–T PATHcalculations performed on late-stage mineral assem-blages are evidence for a period of nearly isobaricDue to the high peak temperatures during granulite

facies metamorphism, relics of the early stages of cooling after the peak of metamorphism (Fig. 12, stage4). Textural support for isobaric cooling in metapelitesmetamorphic evolutions are typically provided by


Fig. 11. Estimates of peak metamorphic P–T conditionscalculated for orthogneisses of the Uluguru Mountains andthose 100 km west of the Uluguru Mountains. P–T estimatesof retrograde metamorphism for the same samples werecalculated with compositions of late-stage minerals. Shadedareas refer to P–T estimates from equilibria that involveclinopyroxene, white area to those that involve orthopyroxene.Arrows connect peak and retrograde P–T estimates of thesame samples.

which underwent isothermal decompression but arecompletely missing in the Pan-African belt of Tanzania.

The last part of the P–T path is only rarelydocumented by the late formation of sillimanite at the

Fig. 10. (a) Garnet zoning of a large porphyroblast in aexpense of retrograde kyanite and can be explained bymetapelite from the North Pare Mountains (T86–3). Thisa rapid uplift of the crust after the isobaric coolinggarnet is one of the few examples where a preserved growth

zonation was found. (b) Pressure estimates for the same garnet, stage (Fig. 12, stage 5). The tectonic process responsibleusing the GASP calibration of Koziol & Newton (1988) and for this exhumation of the granulites occurred after athe composition of garnet-core–plagioclase-inclusion in the

period of slow cooling in the lower crust (3–5 °C km−1,core and garnet-rim–matrix plagioclase.Moller et al., 1994) and was therefore unrelated to theprograde granulite facies metamorphism at about620–650 Ma.is the formation of late-stage kyanite and biotite at the

expense of garnet and K-feldspar and inclusions of With the results discussed above, the metamorphicevolution can be described by an anticlockwise P–Tsillimanite in kyanite. In charnockitic rocks, cooling

textures include the formation of late-stage garnet and path (Fig. 12). This P–T path is supported by a studyof fluid inclusions (Herms & Schenk, 1998) in mineralsclinopyroxene rims at the expense of orthopyroxene

and plagioclase. Such textures explicitly indicate falling which can unambiguously be attributed to specificmetamorphic stages. However, conflicting results weretemperatures at approximately constant pressures.

Despite the difficulties that arise from P–T estimates obtained for the last stage of the P–T path, since late,high density fluid inclusions contradict a strongof late-stage mineral assemblages (e.g. Spear &

Florence, 1992), both textural and thermobarometric decompression as implied by the formation of latesillimanite.results indicate a late-stage near-isobaric period of

cooling for the Pan-African granulites in easternTanzania. There is no evidence for decompression,

CONCLUSIONS AND REGIONALwhich could result in specific reaction textures, for CONSEQUENCESexample: (1) cordierite coronas around garnet in meta-pelitic rocks or (2) orthopyroxene+plagioclase sym- This paper concentrates on the deduction of the

metamorphic conditions and the P–T path for a largeplectites rimming garnet in mafic rocks. Textures ofthese types are typical for many granulite facies areas granulite facies terrane in the Pan-African belt of


record a clockwise P–T path with near-isothermaldecompression at the peak of metamorphism, theycannot result from overthickening of the crust duringthe collision of lithospheric plates (process 1). This isin sharp contradiction with the collisional modelproposed by most authors for the generation of thegranulites of the Mozambique Belt (Burke & Dewey,1972; Muhongo & Lenoir, 1994; Stern, 1994). Furtherevidence against a continent–collision origin of thegranulites can be gained from a correlation of recentgeochronological results and palaeomagnetic data: thegranulite facies event in eastern Tanzania occurredbetween 650 and 620 Ma (Moller et al., 1994), whereasthe final collision in this part of central Gondwana didnot take place until 550 Ma (Meert et al., 1995), i.e.70–100 Ma after the peak of metamorphism.

According to a number of authors (England &Thompson, 1984; Bohlen, 1987; Ellis, 1987; Harley, 1989;Bohlen, 1991) granulites with anticlockwise P–T paths,i.e. with heating and moderate pressure increase duringthe prograde evolution and with a retrograde pathinvolving an initial phase of near-isobaric cooling, arerelated to underplating of magma at the base of thecrust and simultaneous intrusions (magma loading) intothe crust. This process may occur in very different

Fig. 12. Counterclockwise P–T path of the granulite facies tectonic settings, such as extensional regimes (e.g. riftmetamorphism in the Pan-African belt of eastern Tanzania.

environments) or active continental margins. In the caseNumbers refer to stages of the P–T path which can be relatedof the Pan-African belt of Tanzania, granulite faciesto textural and thermobarometric data. 1: Early prograde

stage, deduced from kyanite and staurolite-inclusions in garnet metamorphism could be induced by this process, thatand GASP (Grt-core)-pressures. 2: Prograde garnet growth in preceded by some tens of million years, a collisionalmetapelites mainly within the sillimanite stability field. 3: Peak- event which led to postmetamorphic tectonic crustalmetamorphic conditions obtained by geothermobarometry with

thickening and to the exhumation of lower crustal rocks.compositions of peak-metamorphic assemblages. 4: Near-An active continental margin experiencing magmaticisobaric cooling stage after peak metamorphism deduced from

reaction textures in metapelites and garnet–pyroxene gneisses underplating seems the most likely environment for theand from thermobarometric data of late stage assemblages. 5: high-P granulite facies metamorphism of the belt;Late-stage uplift, documented by late sillimanite formed after

however, field evidence for syn-metamorphic intrusionsretrograde kyanite.is scarce. Nevertheless, late-stage charnockitic intrusionsand mafic dykes with a high-P granulite mineralogy arefound in many parts of the belt (cf. Fig. 3c,d).Tanzania. Surprisingly uniform P–T conditions and

similar P–T evolution exist over a large area, extending Anorthosites and leucogabbros of the Uluguru and PareMountains experienced the same deformation history asfrom the north of Tanzania to the Uluguru Mountains

350 km to the south. This area extends to the Furua the country rocks (cf. Fig. 3a). The Uluguru Mountainsanorthosite yields an U–Pb zircon age of 695±4 MaComplex 200 km further to the south of the Uluguru

Moutains where Coolen (1980) and Coolen et al. (Muhongo & Lenoir, 1994) which, if the interpretationas its intrusion age is correct, preceeds metamorphism(1982) deduced similar P–T conditions and described

reaction textures pointing to the same type of P–T by 40–70 Ma (Moller et al., 1994) and consequently, isan unlikely heat source to instigate metamorphism.path as determined here. This finding has some

important geological and geodynamic implications. Unfortunately, no other young Pan-African intrusionshave been dated so far. Our preferred magmatic under-Granulite facies conditions can be brought about by

a number of different tectonic processes of which the plating-loading model is consistent with the petrologicaland isotopic data, but it is not exclusive. A hybrid model,most important are thought to be (1) crustal-thickening

processes, driven by converging plate movements (e.g. where magmatic thickening accompanies tectonic thick-ening of a previously thinned (extended) crust is alsoEngland & Thompson, 1984) and (2) large intrusions

into the base of the continental crust (Wells, 1979; compatible with the P–T path. However, processesleading to an overthickened crust during progradeBohlen, 1987) or (3) magma loading in higher crustal

levels (e.g. Wells, 1979; Brown, 1996), and can also metamorphism are very unlikely since decompressionshortly after peak metamorphism did not occur.develop (4) in areas with an extensional tectonic

regime (Sandiford & Powell, 1986). Since the Pan- The results of isotopic and petrological studies aresummarized in Fig. 13: The eastern part of theAfrican granulites of the Mozambique Belt do not


Fig. 13. Schematic E-W profile through eastern Tanzania from the Tanzania Craton to the coast. See also the map of easternTanzania (Fig. 1) for orientation.

Mozambique Belt sensu Holmes shows Pan-African age we propose to abandon the use of the termMozambique Belt. Instead, we endorse incorporationhigh-P granulite facies metamorphism and an anticlock-

wise P–T path as described in this paper. This type of of the high-grade gneisses of eastern Tanzania into a‘Pan-African Belt of East Africa’ or the ‘East-Africanmetamorphism is in contrast to high-P metamorphism

in the Early Proterozoic Usagaran Belt, which is Orogen’ proposed by Stern (1994), and to use the name‘Usagaran Belt’ or ‘Ubendian-Usagaran Belt’ for thosecharacterized by a clockwise P–T evolution and abun-

dant decompression textures. Outcrops of the oldest gneisses which experienced their last main metamorphicevent in the Early Proterozoic (about 2 Ga). A boundaryknown MORB-type eclogites are evidence for a 2 Ga

old subduction zone in this belt (Moller et al., 1995). between these terranes that is based on petrologicaland geochronological data has yet to be defined, butLate Archean Sm–Nd model ages and primitive Pb

isotope systematics of feldspars from metapelites in the the presence of either decompression or cooling texturesmay prove to be a useful criterion.Usagaran Belt suggest derivation of these rocks from a

common source with rocks from the Tanzania Craton(Moller et al., 1998). However, Nd model ages of ACKNOWLEDGEMENTSgranites (Maboko & Nakamura, 1996) indicate thatsome juvenile material was probably added to the crust This paper is a contribution to the IGCP project 348

(The Mozambique and related belts). It forms part ofduring the Usagaran-Ubendian orogeny. During thelate Proterozoic (1–1.5 Ga) juvenile crustal material the doctoral dissertations of P. A. and A. M. The

authors thank D. Ackermand and B. Mader in Kielwas added only in the Pare and Usambara Mountainsand in the eastern Uluguru Mountains, whereas in for assistance with microprobe analyses, the Division

of Mines and Geology in Dodoma and Morogoro andother areas of the Pan-African belt Archean crust wasreworked. Mixing of older crust and juvenile crustal S. Muhongo and the University of Dar-es-Salaam for

their assistance and administrative support. We thankmaterial occurred only in the western part of theUluguru Mountains in the eastern Mozambique Belt P. Raase for many discussions and his participation in

fieldwork. The Tanzanian Commission for Science and(Moller et al., 1998). These results show that theUsagaran and Pan-African metamorphic belts are Technology (UTAFITI) is gratefully acknowledged for

research permits. The project was financed by Deutschecharacterized by different metamorphic evolutions butboth contain reworked crustal material. Forschungsgemeinschaft (DFG) grants Sche 265–2/5

and 265–6/1. E. H. Brown, S. L. Harley and J. C.Petrological and isotopic data from Tanzania indicatethat two different orogenic belts exist within the Schumacher are thanked for critical reviews and

D. Robinson for the editorial handling of theMozambique Belt sensu Holmes (1951). For the pur-pose of distinguishing metamorphic events of different manuscript.


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