U – Pb dating of metamorphic minerals: Pan-African
metamorphism and prolonged slow cooling of high pressure
granulites in Tanzania, East Africa
Andreas Mo¨ller
a,b,*, Klaus Mezger
b,1, Volker Schenk
aaMineralogisch-Petrographisches Institut,Uni
6ersita¨t Kiel,24098Kiel,Germany bMax-Planck-Institut fu¨r Chemie,Postfach3060,55020Mainz,Germany
Received 1 November 1999; accepted 4 May 2000
Abstract
U – Pb monazite and zircon ages reveal that the high pressure granulites from eastern Tanzania were metamor-phosed during a Pan-African tectonothermal episode. These mineral ages range from 610 to 655 Ma and indicate that peak metamorphic conditions were diachronous in the different granulite domains. U – Pb titanite and rutile ages define integrated cooling rates of 2 – 5°C/Ma for all investigated granulite areas, and suggest a common process for the post-metamorphic histories of the different granulite areas. Prolonged slow cooling-rates are consistent with near-isobaric cooling in the deep crust after the metamorphic peak. The process responsible for crustal thickening during heating did not produce isostatic instability and fast erosion-driven or tectonic exhumation. The thermal history determined in this study is not consistent with the collision of East- and West-Gondwana as the cause of granulite facies metamorphism. Palaeomagnetic data have shown that this collision did not occur until 550 Ma, when the Pan-African granulites in Tanzania had already cooled below 500°C. The high pressure granulites of eastern Tanzania are thus interpreted as having attained their metamorphic peak prior to the final amalgamation of Gondwana, probably in an active continental margin setting. © 2000 Elsevier Science B.V. All rights reserved.
Keywords:Mozambique Belt; Pan-African orogeny; U – Pb geochronology; Monazite; Titanite; Rutile
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1. Introduction
Metamorphic pressure – temperature – time (P– T–t) paths provide essential constraints for any models that relate metamorphism to tectonic pro-cesses. The different tectonic settings can be in-dicative of the plate-tectonic scenario that led to metamorphism and the formation of an orogenic belt. In order to unravel the evolution of a
com-* Corresponding author. Present address: Institut fu¨r Geo-wissenschaften, Johannes Gutenberg Universita¨t Mainz, Post-fach 3980, D-55099 Mainz, Germany. Tel.: + 49-6131-3925584; fax: +49-6131-3924769.
E-mail address:[email protected] (A. Mo¨ller).
1Present address: Institut fu¨r Mineralogie, Universita¨t
Mu¨n-ster, Corrensstr. 24, 48149 Mu¨nMu¨n-ster, Germany.
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 124
plex orogenic belt such as the Mozambique Belt (MB) of East Africa, the direct coupling of geochronologic data with petrologic information (Appel et al., 1998) and crustal residence ages (Mo¨ller et al., 1998) is of paramount importance. In this study, both the prograde and retrograde thermal histories are reconstructed using U – Pb ages obtained on metamorphic minerals with dif-ferent closure temperatures including monazite, titanite and rutile. Because most of the minerals sampled in this study were extracted from gran-ulite facies metasediments they can be considered to be most likely of metamorphic origin. This study also compares published U – Pb zircon ages and K – Ar, Ar – Ar, Rb – Sr on hornblende, biotite and muscovite data from different granulite
ter-ranes in Tanzania for their consistency with new U – Pb ages. The scarcity of age data for the Pan-African orogen of East Africa has led some authors to use ages determined on different gran-ulite complexes for an integrated interpretation of the whole orogenic belt (Maboko et al., 1985, 1989; Muhongo and Lenoir, 1994). However, it can be shown that it is important to know the age of metamorphism for each area separately for P–T–tpath construction, because rock units jux-taposed today may have been at different crustal levels and experienced differentP–Thistories, but the same tectonic and metamorphic processes. Samples from 17 locations (metapelitic gneisses, orthogneisses, marbles and calcsilicates) within the MB were chosen to cover the different parts of the respective granulite complexes in eastern Tanzania.
2. Geologic setting: granulite complexes in the Pan-African Belt of Tanzania
One of the most influential contributions that shaped the understanding of the African Precam-brian geology was the definition of the Mozam-bique Belt by Holmes (1951). He recognised the discontinuity of geological structural trends be-tween the Tanzania craton and its eastern hinter-land and showed that these areas had to be younger than the craton. Subsequently Shackleton (1967) proposed that the MB has a complex his-tory and suggested that the belt is composed of Archaean basement and several younger metasedi-mentary sequences. The MB then served as one of the classical examples for rejuvenation (i.e. no new crustal material added during orogenic cycle) of Archaean and Early Proterozoic basement (Watson, 1976). However, it was also proposed that the MB is a product of late Precambrian plate collision following ocean closure (Burke et al., 1977 McWilliams, 1981).
Stern (1994) proposed the term ‘East African orogen’ for the areas covered by the older terms
‘Arabian – Nubian shield’ and ‘Mozambique
Belt’’, because it is appropriate to view the whole area as the product of one Neoproterozoic
Wilson-cycle (see inset of Fig. 1). The Arabian – Nubian shield contains large tracts of Pan-African juvenile crust and abundant ophiolites and is in-terpreted by Stern (1994) as a collage of accreted terranes. In contrast, the MB with its high-grade gneisses resembles the deeply eroded root of an orogen formed by a single collision event between East- and West-Gondwana. The MB experienced further uplift during Phanerozoic rifting, some of it associated with the development of the East African Rift. This interpretation supports the model of Hoffman (1991), i.e. the MB was formed by fan-like closure of a previously existing Mozambique ocean, with the hinge of the fan somewhere in South Africa. Since this fan never fully closed, crustal shortening was most intense in the southern part of the belt. Stern (1994) argues further that the exposure of granulites at the surface in Kenya and Tanzania is evidence that crustal thickness of the orogen was greatest and collision most intense in these areas, because today the granulites are found within the crust of normal thickness of approximately 35km (e.g. KRISP Working Party, 1995).
Within Tanzania, geochronological results
show that the Mozambique Belt of Holmes (1951) has to be subdivided into a Pan-African (late Proterozoic) domain to the east and an Usagaran
(=Ubendian, Early Proterozoic) domain to the
west (Fig. 1). A tentative subdivision in southern Tanzania was based on progressively older Rb – Sr biotite ages towards the west (Wendt et al., 1972; Priem et al., 1979) interpreted as the result of the decreasing Pan-African thermal overprint on the Early Proterozoic rocks (Gabert and Wendt, 1974). U – Pb dating of metamorphic monazite and titanite from eclogite-facies rocks places the main metamorphic event in the Usagaran domain at 2000 Ma (Mo¨ller et al., 1995). Appel et al. (1998) suggest that distinctive decompression tex-tures in the Usagaran Belt and cooling textex-tures in the Pan-African granulites can be used to distin-guish the two belts.
To distinguish the two metamorphic events we endorse the use of the terms ‘Pan-African Belt of East Africa’ or the ‘East-African Orogen’ pro-posed by Stern (1994) for the Pan-African gran-ulite facies gneisses of eastern Tanzania and use
the name ‘Usagaran Belt’ or ‘Ubendian –
Usagaran Belt’ for the region where the main metamorphic event occurred at about 2 Ga (Fig. 1). The term Pan-African is used in this study for the time span from about 650 to 550 Ma, relevant to and encompassing metamorphic events in the circum-Indic region related to the formation of Gondwana.
The Pan-African Belt in Tanzania consists of Archaean to Proterozoic rocks (e.g. Mo¨ller et al., 1998) metamorphosed under granulite facies con-ditions (e.g. Bagnall, 1963; Sampson and Wright, 1964; Coolen, 1980; Appel et al., 1998) during the Pan-African orogeny (e.g. Coolen et al., 1982; Maboko et al., 1985, this study). Some of the granulite complexes (Fig. 1) apparently form fault bounded mountain ranges, interpreted as tectonic klippen (e.g. Shackleton, 1986), namely the Pare and Usambara Mountains (Bagnall, 1963; Bagnall et al., 1963) and the Uluguru Mountains (Samp-son and Wright, 1964).
Previous petrologic and geochronological stud-ies have been carried out mainly on the Furua complex (Coolen, 1980; Coolen et al., 1982), the Wami River complex (Maboko et al., 1985), and
the Uluguru Mountains (Muhongo, 1990;
Maboko et al., 1989). The granulite complexes within the Mozambique Belt exhibit striking simi-larities in lithology, structure and grade of meta-morphism (Coolen, 1980; Appel et al., 1998). Petrologic studies reveal similar peak
metamor-phic conditions of 810940°C and 9.5 to 11 kbar
and a similar P–T path for an extensive area
within the Pan-African Belt including the Pare, Usambara and Uluguru Mountains granulite complexes and some adjacent lowland areas (Ap-pel et al., 1998).
3. Analytical methods
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 126
colour. Titanite was cleaned in pure alcohol in an ultrasonic bath for about 15 min, washed in warm distilled 3 N HCl for about 10 min to remove surface contamination, and twice rinsed in tilled water. Monazite was washed in warm dis-tilled water only prior to dissolution. Rutile was washed in warm 0.5 N HF for about half an hour, zircon in hot 6 N HCL for about 15 min.
Uranium and Pb concentrations were
deter-mined by isotope dissolution with a 233U/205Pb
mixed spike, added before dissolution to allow optimum homogenisation with the sample. Ele-ment concentrations in weighed mineral fractions are known to about 0.2%, calculated from analy-tical errors alone. All zircon-, monazite- and
ru-tile-fractions were digested in 3 ml Savillex®
screw-top beakers in a Krogh-style or Parr®
Teflon® bomb within a screw top steel container
at 210°C. Monazite dissolved in 0.5 ml 7 N HNO3
and 0.5 ml 6.2 N HCl after 1 – 3 days. Rutile dissolved within a few days in a mixture of 0.5 ml
concentrated HF and five drops of 7 N HNO3.
Titanite fractions were digested overnight in the oven in a mixture of 0.5 ml concentrated HF and
ten drops of 7 N HNO3 after boiling for 12 – 24 h
on the hot-plate. The zircon fraction dissolved in
concentrated HF and ten drops of 7 N HNO3 in
the oven within 10 days. Dissolution was checked optically for each sample, under a microscope where necessary.
Uranium and Pb were separated with
ion-ex-change Teflon® columns filled with about 0.5 ml
of DOWEX AG 1X8® anion exchange resin (e.g.
Krogh, 1973; Tilton, 1973). Pb chemistry for mon-azite, rutile, titanite, and feldspar employed the HBr – HCl method, whereas Pb from zircon was separated with HCl. Uranium was separated with
the HCl – HNO3 method. Five total procedural
blanks were determined between 44 and 123 pg with an average of 80 pg. The Pb-isotope ratios
measured for the blank were: 206
Pb/204
Pb: 18.53;
207Pb
/204Pb: 15.69; 208Pb
/204Pb: 35.90.
Isotope ratios were measured on a Finnigan MAT 261 mass-spectrometer in multi-collector static mode on Faraday cups, using single Re filaments. A secondary electron multiplier (SEM)
was used for measuring 204Pb when high ratios
made it necessary, and for some U analyses in
dynamic mode. Pb was loaded with H3PO4 and
silica-gel (Cameron et al., 1969). The measured Pb isotopic ratios were corrected for fractionation
with a mass discrimination factor of 0. 1%/amu,
based on 23 analyses of 50 ng of equal atom SRM-982, measured during this study in compari-son with the values recommended by Todt et al. (1996). Reproducibility of the207Pb/206Pb ratio of
the SRM-982 standard (average: 0.466512) was 0.033%. Within-run reproducibility was much
higher, with an average of 0.0021% at 2s
confi-dence level. The measurements of 206
Pb/204
Pb ra-tios with 204
Pb on the SEM were corrected with a factor of 1.0038, determined from five measure-ments of SRM-982. Most U was measured as
oxide after loading with H3PO4 and silica-gel.
Based on repeated analyses of 100 ng SRM-U500 standard, a mass fractionation correction factor of
0.01%/amu was applied to samples measured in
static mode and a correction factor of 0.3%/amu
to SEM dynamic measurements. Reproducibility for the 235U/238U ratio of the standard (static
mode) was 0.29%, with an average within-run reproducibility of better than 0.04%.
For some samples, U was loaded with graphite
dispersed in a water/alcohol-solution and
mea-sured as U+
at temperatures between 1650 and 1740°C. Reproducibility estimated from seven U500 standards loaded with graphite was 0.28% for Faraday cup in static mode. Fractionation was
corrected with a factor of 0. 1%/amu. Mass
frac-tionation was strongly time-dependent with these graphite loaded samples and care was taken to heat up all samples in the same manner and avoid acquisition times longer than approximately five blocks of 20 measurements each.
4. Closure temperature estimates
For a valid interpretation of mineral ages, infor-mation on the closure temperature (Tc) for parent/
daughter systems is essential. The closure
temperature is defined as the temperature of the system at the time given by its apparent ages
(Dodson, 1979). ThisTcdepends on grain size and
Table 1
Summary of approximate closure temperaturesa
Mineral Tc(°C) Grain size (mm) References U–Pb system
100–300 Schenk (1980, 1990), Bingen and van Breemen (1998), Parrish and 800 (peak
Monazite
metamorphism) Whitehouse (1999), this study 200–30 000
Titanite (630–730) Mezger et al. (1991), Gromet (1991), Cherniak (1993), Scott and St-Onge (1995), Zhang and Scha¨rer (1996)
200–500
650 This study
130–430 Mezger et al. (1989) 380–420
Rutile
K–Ar and Ar–Ar system
160
450–500 e.g. Harrison (1981) Hornblende
–
Muscovite 350–400 e.g. Hanson and Gast (1967) –
300 e.g. Harrison et al. (1985)
Biotite
Microcline 150–200 125–250 Harrison and McDougall (1982)
Rb–Sr system
Muscovite 450–500 – e.g. Harrison and McDougall (1982) –
350 e.g. Harrison and McDougall (1982) Biotite
aT
cvalues are chosen for selected minerals at different grain sizes for slow cooling rates of 1–10°C/Ma.
For some parent/daughter systems in some
minerals experimental data is available (e.g. U – Pb in titanite, Cherniak (1993); K – Ar in horn-blende, Harrison (1981) U – Pb in monazite, Smith and Giletti (1997)). For other systems only empirical values are available and many of them may need further refinement. A correlation of experimental results with well controlled natural geologic settings is still wanting for many
miner-als. Table 1 summarises Tc for minerals relevant
to this study and the choice preferred by the authors, which is pivotal for the interpretation of the geochronological data and the cooling his-tory.
4.1. Monazite
It is generally accepted that theTcof monazite
is at least 700°C for slowly cooled rocks.
How-ever, there is ample evidence that the Tcmay be
significantly higher as indicated by field data from the Hercynian crustal section of Calabria, Italy, (Schenk, 1980, 1990) or the Valhalla com-plex in British Columbia (Spear and Parrish, 1996). A single grain U – Pb study by Bingen and van Breemen (1998) in amphibolite to granulite facies rocks shows that monazite growth ages
can be preserved through 850°C metamorphism under dry conditions. A study by Parrish and
Whitehouse (1999) also suggests higher Tc.
Re-cent experiments on Pb diffusion rates in monaz-ite by Smith and Giletti (1997) suggest that
circular or elongate monazite grains of 100 mm
radius should have closure temperatures of only 630 – 720°C in regions which cool at rates
be-tween 1 and 10°C/Ma. The authors caution that
uncertainties in their closure temperature calcula-tions may be as high as 140°C. Comparison with the examples from geochronological field studies in granulites (see above) leads us to conclude that the calculations of Smith and Giletti (1997) underestimate the closure temperature of monaz-ite and are not accurate enough for application.
We conclude that monazite ages from this study may be interpreted as growth ages and thus date the peak of the granulite facies meta-morphic event in eastern Tanzania which reached
temperatures of 810940°C (Appel et al., 1998)
in most areas. Estimates of a lower Tc may then
be due either to growth of new monazite at tem-peratures below its Tcor, alternatively, to
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 128
4.2. Titanite
Experimental as well as empirical estimates for the Tc of the U – Pb system in titanite are
avail-able. Mezger et al. (1991) estimated a closure temperature of 630°C for titanite crystals of 1 cm
diameter at a cooling rate of 2°C/Ma from field
studies in the Adirondack Mountains. In other parts of the Grenville Orogen, titanite preserved
their U – Pb ages although the surrounding
gneisses were later migmatised, which indicates
thatTcmay be at least as high as 650°C (Mezger
et al., 1992) for larger grains. Experimental stud-ies of Cherniak (1993) yield a closure temperature of approximately 630°C for a diffusion radius of
500 mm at 2°C/Ma cooling rate. Cherniak (1993)
thus concluded that effective diffusion radius may be smaller than grain size.
Evidence for a higher closure temperature of titanite in slowly cooled rocks was presented by Scott and St-Onge (1995). Their combination of thermobarometry and U – Pb dating suggests that the Tc of 100 mm – 1 mm diameter titanite lies in
the range 660 – 700°C, higher than all previous estimates. This conclusion is now supported by other studies (e.g. Corfu (1996), Verts et al. (1996)) and by discordance patterns observed in rocks which experience brief thermal events, where discordant titanite data can be interpreted with the episodic Pb-loss model (e.g. Tucker et al. (1986), Haggart et al. (1992)). Similar evidence was presented by Gromet (1991), where titanite
grains of 500 to 2000mm diameter showed strong
discordance to an upper intercept and even
titan-ite grains of 250mm diameter showed some
inher-itance although this rock experienced only about 650°C during a metamorphic event, possibly re-lated to the brevity of the overprint. From inher-ited magmatic titanite in a syenite intrusion, Zhang and Scha¨rer (1996) deduced a closure tem-perature for volume diffusion of Pb in excess of 710°C. They suggest that titanite is always closed to Pb at its crystallisation temperature and that the closure temperature concept may be mislead-ing for metamorphic titanite, a contention not supported by the data of this study. Important factors in all these studies are the time of titanite growth relative to the onset of cooling and the
duration of the metamorphic event in case the titanite had formed previously. Both may limit the ability to determine a closure temperature from these mineral ages for slowly cooled terranes.
Most titanite fractions analysed in this study consist of whole grains with a diameter of 200 –
500 mm. Based on the studies of Cherniak (1993),
Gromet (1991), and Scott and St-Onge (1995) Tc
of 650°C for titanite from granulites-facies rocks with slow cooling rates has been used in this study as a conservative estimate.
4.3. Rutile
An estimate for the closure temperature for Pb in rutile was given by Mezger et al. (1989), based
on comparison with K – Ar and 40Ar
/39Ar ages of
hornblende and biotite. It was suggested that the
closure temperature is ca. 430930°C for slow
cooling rates of 2 – 10°C/Ma. This value is consis-tent with U – Pb ages on rutile obtained by Scha¨rer et al. (1986). Most published U – Pb ages of rutile, are concordant or only slightly discor-dant. Inheritance of older age information is only likely when a metamorphic overprint does not exceed greenschist-facies conditions (Mo¨ller et al., 1995).
4.4. Mica and amphibole
Table 1 shows that Ar – Ar ages from horn-blende are expected to be close to, but younger than U – Pb titanite ages from the same area and older than U – Pb ages of rutile. Similarly, Ar – Ar ages of biotite are expected to be close to, but younger than, U – Pb ages obtained on rutile. This is important in the discussion of inherited or excess Ar, which appears to be very common in biotite and hornblende of granulites from the Mozambique Belt (data of Priem et al. (1979) and Maboko et al. (1989), discussed below).
5. Results
Sample locations are illustrated and geochrono-logical results summarised in Fig. 2. Major and trace element analyses for most of the samples are reported by Appel (1996). Descriptions of sample
locations and mineral assemblages can be found in Table 4.
Several reasons — geological and analytical — may explain discordant analyses (Pb-loss, over-growth core relations, incomplete dissolution). This study uses Pb-loss as the most likely
interpre-tation and unless specifically stated the 207
Pb/
206
Pb of discordant analyses is the most likely minimum age for this mineral fraction. Reversely discordant results are quite common with
monaz-ite and often indicate excess 206
Pb (from short-lived 230
Th) due to preferential incorporation of Th over U during monazite growth (Scha¨rer, 1984), with monazite remaining below its closure temperature. For such reversely discordant re-sults, this study uses the207
Pb/235
U age as the best estimate for the true age since this ratio is not affected by excess 206
Pb. Because concordant re-sults could also be the result of a combination of
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 130
Fig. 3. Concordia diagram for monazite from the Pare and Usambara Mountains and the Umba Steppe. The results indicate an age difference of about 15 Ma. Two pairs of discordant and near concordant monazite from the Pare and Usambara Mountains are shown (T115, A108). For both locations, monazite with the higher U content are more discor-dant than monazite with lower U content.
5.1. Monazite and zircon
Most of the monazite used in this study was separated from metasedimentary rocks. The mon-azite occur as pale to bright yellow spheres or ellipsoids free of inclusions and range in diameter
from 100 to 600 mm.
Two monazite fractions from the Pare Moun-tains (metapelites A16 and T115b) plot slightly above concordia (Fig. 3). Their207Pb/235U ages of
64192 Ma are interpreted as the true ages of
formation. Another fraction from metapelite T115
is discordant with a similar 207
Pb/206
Pb age of
64092 Ma.
Monazite from metapelite T137 and metagrani-toid T121 from the Usambara Mountains are also reversely discordant and yield 207Pb
/235U ages of
62192 and 62492 Ma, similar to the207Pb/206Pb
age of 62492 Ma obtained from concordant
monazite fraction A108b (Fig. 3). A second, dis-cordant fraction of monazite from metapelite
A108a has a slightly higher 207Pb/206Pb age of
62992 Ma. Monazite in metapelite T137 has
been observed in thin section to occur mostly as large (\300 mm) grains attached to high-Ti
gran-ulite facies biotite and hence as part of the high grade assemblage. Backscatter electron (BSE) imaging indicates strong ‘patchy’ zoning, reflect-ing zonation in Th content (Fig. 4). U – Pb datreflect-ing
by laser-ICP-MS yielded a 206Pb
/238U age of
618915 Ma from four analyses on two grains
(Mo¨ller and Jackson, unpublished data) support-ing the multi-grain-isotope dilution results of this study. Two other analyses indicate some Pb loss, but no evidence of older growth phases within the monazite has been found.
A fraction of monazite grains from semipelite
A 144 in the Umba Steppe yields a 207Pb/206Pb
age of 60992 Ma (Figs. 3 and 5), the discordance
(1.5%) may be attributed to recent Pb-loss, possi-bly due to weathering.
Four fractions of monazite were analysed from three metapelite samples of the northern, north-eastern and north-eastern Uluguru Mountains (Fig. 6). Two monazite fractions from sample P1 show no differences in colour or grain size and differ only slightly in their Th and U contents. Fraction P1a
is 1% discordant and has a 207Pb/206Pb age of
Fig. 4. Backscatter-electron image of monazite in metapelite T137 showing pronounced patchy zoning probably related to growth inhibition at low H2O activity under granulite facies
conditions.
reverse discordance and Pb-loss, their 207Pb
/206Pb
has also to be taken as a minimum age. For age calculations of some mineral fractions (rutile and some titanite) very sensitive to corrections for Pb blank and initial common Pb, the206Pb/238U age
66993 Ma, whereas fraction P1b is 0.6% re-versely discordant with a207Pb/235U age of 6579
2 Ma. Sample P1b could have been affected by some late disturbance, and the reverse discor-dance observed may only be the remainder of an originally much higher discordance. The age of
P1a may either indicate the presence of an inher-ited Pb component (detrital core) or reflects the age of monazite growth during prograde meta-morphism. We consider the latter case more likely and the result of fraction P1a is, therefore, not used to calculate cooling rates for this location. However, this result may have some bearing on the discussion of the age and duration of high grade metamorphism in this part of the Uluguru Mountains.
Monazite from the graphite-rich metapelite P9
has a high a 208
Pb/206
Pb ratio of 31.4 and plots
significantly above concordia (2%) with a 207
Pb/
235
U age of 64692 Ma. The monazite fraction
analysed from the eastern Uluguru Mountains (T28) is also slightly reversely discordant with a
207Pb
/235U age of 65392 Ma (Fig. 6). The results
from three metapelite samples of the northern and eastern Uluguru Mountains thus span an age
range of about 11 m.y. between 64692 and
65792 Ma (Table 2; samples P1b, P9, T28).
Monazite and zircon were separated from a meta-qtz-diorite (T46) originating from the north-western Uluguru Mountains. This meta-qtz-dior-ite shows evidence of all three deformational phases observed in the surrounding granulite-fa-cies rocks (Appel et al., 1998). Its emplacement, therefore, must have been pre-peak-metamor-phism and pre-deformation. Two monazite frac-tions from the sample have different U – Th – Pb contents but a very similar age. The smaller mon-azite grains. with a higher Th/U ratio are slightly reversely discordant at a207Pb/235U age of 62492
Ma. The larger monazite size fraction has a lower
208Pb/206Pb ratio and is slightly discordant at a 207Pb/206Pb age of 62592 Ma. The zircon
frac-tion analysed from this sample consists of 19
long-prismatic, clear and euhedral grains (\l40
mm) without evidence of older cores when
ob-served in transmitted light. The U – Pb result is slightly discordant (1%) with a 207
Pb/206
Pb age of
62693 Ma, overlapping the age of the monazite
fractions (Fig. 6).
5.2. Titanite
The geological map (see Fig. 2) shows very few marbles and calcsilicate rocks in the Pare and
Fig. 5. Concordia diagram for monazite, titanite and rutile from the Umba Steppe. Different fragments of the same titanite grain from sample A141 yield similar207Pb/206Pb ages
close to the 207Pb/206Pb age of 60992 Ma obtained for
monazite from sample A144.
A U–Pb isotope dataa
Sample, mineral Rock type Wt. (mg)b Pb (ppm) U (ppm) Isotopic ratios Ages (in Ma) Discordancef(%)
206Pb/207Pbc 208Pb/206Pbd 207Pb/206Pbd 206Pb/238Ue 207Pb/235Ue 206Pb/238U 207Pb/235U 207Pb/206Pb
Pare Mountains
21.50 0.082368 0.10473 0.88055 64292
0.35 389 64192 63994 0.6
Metapelite
A16, Mnz 185 597
0.8082 0.061064 0.09612
T115a, Mnz Metapelite 0.32 470 3050 58700 0.80872 59292 60292 64092 −7.5 29.87 0.063380 0.10495 0.88088 64392 64192
2110 63594
790 280 1.3
T115b, Mnz Metapelite 0.15
610
Calcsilicate 5.49 19.4 130 0.6612 0.083598 0.09663 0.80160 59592 59892 61093 −2.5 A26, Tit
167.4
Metabasite 4.52 18.2 98 0.9988 0.146452 0.08937 0.74589 552910 56699 623916 −11.4 T114, Tit
0.0181 0.061373 0.08635 0.69205 53492 53492
2140 53594
A16, Rt Metapelite 13.31 0.55 6.86 −0.1
0.2226 0.058517 0.08517 0.68126 52792 52892
T115, Rt Metapelite 10.89 2.9 31.1 9750 53094 −0.6
Usambara Mountains
23.75 0.064764 0.10173 0.84275 62592
0.40 1260 570 62192 60692 3.0
Metapelite
T137, Mnz 2640
4210 28900 3.189 0.061022 0.10041 0.84068 61793 62093 62992 −2.0 0.22
A108a, Mnz Metapelite 1560
3020 22600 3.382 0.060666 0.10182 0.85021 62592 62591 62492 0.2 0.12
A108b, Mnz Metapelite 1180
100.3 0.061912 0.10178 0.84821 62592 62492
4470 61992
T121, Mnz Charnockite 0.22 4260 480 0.9
0.3119 0.178682 0.08191 0.64797 50892 50794
T137, Rt Metapelite 4.00 4.89 39.5 118.3 506917 0.3
0.0446 0.070951 0.08384 0.66902 51992 52092
830 52594
T139, Rt Meta-qtz-di. 8.30 0.67 8.2 −1.1
A108, Rt Metapelite 12.78 1.14 13.7 4050 0.0718 0.059622 0.08526 0.68112 52792 52791 52793 0.0
Umba Steppe
1170 9730 5.949 0.061302 0.09740 0.80778 59992 60192 60992 −1.6 0.50
A144, Mnz Grt-Bt gneiss 690
0.0181 0.061365 0.07358 0.61103 45892 48492
12300 61292
A141a, Tit Marble 2.82 174 2540 −25.2
0.0502 0.061741 0.09333 0.77705 57592 58492
A141b, Tit Marble 2.38 111 1240 7060 61792 −6.8
0.0371 0.071373 0.08306 0.65759 51492 51392
928 50897
A144a, Rt Grt-Bt gneiss 3.70 2.9 36 1.3
A144b, Rt Grt-Bt gneiss 15.52 2.4 32 16800 0.0025 0.057805 0.08318 0.66102 51595 51594 51693 −0.2
Uluguru Mountains
0.0810 0.062140 0.10075 0.84223 61992
T46, Zrng Meta-qtz-di. 0.59 37.1 370 6100 62092 62693 −1.2
18.52 0.064354 0.10180 0.84866 62592 62492
2140 62093
Diorite 0.7
T46a, Mnz sm 0.08 1340 770
8300
Diorite 0.18 2460 2170 11.88 0.061860 0.10106 0.84457 62192 62292 62592 −0.8 T46b, Mnz la
1040 6110 7.995 0.063559 0.10828 0.92330 66392 66492 66993 −0.9 0.25
P1a, Mnz Metapelite 880
8.514 0.063433 0.10740 0.90936 65892 65792
5060 65492
910 1020 0.6
P1b, Mnz Metapelite 0.20
2190
Metapelite 0.23 960 320 31.37 0.065035 0.10612 0.88956 65092 64692 63293 2.9 P9, Mnz
7370
Metapelite 0.19 510 980 4.621 0.062217 0.10683 0.90205 65492 65392 64892 1.0 T28, Mnz
0.9208 0.092502 0.09415 0.78521 58096 58897
453 621922
P8a, Tit Calcsilicate 4.71 43.6 260 −6.6
280 553 0.7989 0.086545 0.10183 0.84817 625911 62498 61894 1.1 P8b, Tit Calcsilicate 2.79 48.1
0.4862 0.081095 0.09893 0.82422 60895 61095
681 61997
Marble −1.7
T25a, Tit 5.78 15.8 115
630
Marble 4.76 10.8 112 0.0504 0.082931 0.09638 0.80358 59394 59993 62194 −4.4 T25b, Tit
831
Marble 4.74 17 121 0.5163 0.077363 0.09989 0.83277 61497 61595 62093 −1.0 T25c, Tit
0.0479 0.105575 0.09956 0.82986 61296 61495
317.5 62096
P88a, Tit Calcsilicate 4.43 12.2 116 −1.3
0.4831 0.100277 0.09992
P88b, Tit Calcsilicate 3.83 19.5 136 359.1 0.83243 61493 61592 62195 −1.1 0.0314 0.063062 0.08853 0.71904 54792 55091
2740 56492
2.1 25.3 −3.0
P1, Rt Metapelite 10.27
1910
Metapelite 12.02 2.5 25.5 0.2263 0.064848 0.08909 0.71589 55094 54893 54093 1.8 P9, Rt
25.3 2120 0.0976 0.062982 0.08065 0.63767 50095 50194 50595 −0.9 T28, Rt Metapelite 12.2 2.0
aAges and errors (2s) are calculated with Pbdat and ISOPLOT for Excel 2.0.6 software, after Ludwig (Ludwig, 1980, 1994); Zrn, zircon; Mnz, monazite; Tit, titanite; Rt, rutile; sm, small; la, large. bMost Mnz weights estimated from size, error about 10–20%, other samples weighed toB1% error.
cMeasured ratio.
dMeasured ratio, corrected for spike, 80 pg Pb blank, 0. 1% mass fractionation per a.m.u.
eCorrected for spike, 80 pg Pb blank, 0. 1% mass fractionation per a.m.u. and common Pb composition determined from leached coexisting K-feldspar or plagioclase (Mo¨ller et al., 1998). fDiscordance of result expressed as deviation of the207Pb/206Pb age from the206Pb/238U age=[207Pb/206Pb age/206Pb/238U age/100]−100.
gNinteen clear, pink, euhedral, prismatic grains,\140
Fig. 7. Concordia diagrams for titanite.
isotope composition of A26 as no analysis of coexisting plagioclase was available. The possible error associated with the correction is small con-sidering the narrow range of common Pb compo-sition of feldspars from the Pare and Usambara Mountains (Mo¨ller et al., 1998). The result sug-gests that cooling through the closure temperature of titanite in the South Pare Mountains occurred in the same age range as in the North Pare Mountains.
In the Umba Steppe interlayered calcsilicate rocks and marbles were found at the Umba river. A coarse-grained sample (A141) yielded an opaque to very dark brown titanite grain of more than 0.5 cm diameter. Two fragments from the core of the grain have variable U and Pb concen-trations and degree of discordance (Table 2, Fig.
5) but similar 207Pb/206Pb ages of 61292 and
61792 Ma. The high U content of the titanite
fragments may have caused structural damage
and Pb-loss. It can be concluded that the 207Pb/
206Pb ages record the age of peak metamorphism
and the effective closure temperature of this titan-ite grain is higher than the ca. 730°C calculated for grains with 0.5 cm diffusion radius by Cher-niak (1993) or alternatively that metamorphic
temperatures did not exceed Tc after titanite
growth.
Suitable titanite-bearing samples were only found in the eastern part of the Uluguru Moun-tains (P8 and T25) close to the locations of the monazite and rutile samples. An additional sam-ple was taken from the southeast Uluguru Moun-tains (P88). To evaluate the reproducibility of U – Pb ages of titanite from the eastern Uluguru Mountains, several fractions of grains were analysed for each sample. Their 207Pb/206Pb ages
overlap within error at 618 – 621 Ma. The variably discordant titanite fractions can be fitted on a single regression line (Fig. 7b) despite being taken form localities about 60 km apart. The combined intercept of titanite in the eastern Uluguru
Moun-tains at 61992 Ma is interpreted as the age at
which at least this part of the Uluguru Mountains (the crystalline limestone group of Sampson and Wright (1964)) cooled through the closure temper-ature of titanite at ca. 650°C.
Usambara Mountains. Only two suitable samples of titanite-bearing rocks could be collected (A26, T114) from the Pare Mountains. Sample A26 from the North Pare Mountains is a calcsilicate
gneiss with the metamorphic assemblage Grt+
Cpx+Hbl+Pl+Qtz+Cc+Scp+Tit9Kfs. The
titanite is reddish-brown and does not exhibit colour zonation. The 207Pb/206Pb age of 61093
Ma (Fig. 7a) is interpreted as the minimum age for closure at ca. 650°C. Titanite in metabasite T114 from the South Pare Mountains is pale
brown and only 150 to 200mm in diameter. It has
a low 206
Pb/204
Pb ratio of 167.4 and yields an
imprecise 207Pb/206Pb age of 623916 Ma.
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 134
5.3. Rutile
The rutile fractions used for U – Pb age determi-nations were obtained from metapelitic samples that also yielded monazite (A16, T115, T137,
A108, A144, P1, P9, T28). Rutile in the
metapelitic samples were mostly elongate grains (aspect ratio higher than 4) or fragments thereof. In all of these samples rutile is part of the high-pressure granulite-facies assemblage together with garnet, sillimanite/kyanite, plagioclase, quartz9
ilmenite. An exception is qtz-dioritic enderbite T139 from the western Usambara Mountains which contains large, dark, short rutile grains
with an average diameter \250 mm. Rutile from
a single sample often spans a range of colours from translucent reddish brown to almost opaque and dark-brown to black. Care was taken to pick grains of similar size and colour for each rutile fraction (4 – 16 mg) to avoid mixing of different chemical compositions and possibly different dif-fusion behaviours.
Concordant rutile fractions yield 207Pb/206Pb
ages of 53594 and 53094 Ma for the North and
South Pare Mountains, respectively (Fig. 8a). Re-sults of two rutile fractions from meta-qtz-diorite T139 and metapelite A108 from the Usambara
Mountains have indistinguishable207
Pb/206
Pb ages
of 52594 and 52793 Ma, respectively but
sig-nificantly different206Pb
/238U ages of 5 1992 and
52792 Ma. Rutile from sample T137 is still
younger with a 206Pb
/238U age of 50892 Ma, but
it is slightly discordant and has a large uncer-tainty in207Pb/206Pb (506917 Ma) due to its low
proportion of radiogenic Pb. Sample T139 was collected just 10 km from metapelite T137 and belongs to a suite of meta-qtz-diorites, some of which cross-cut the foliation of the surrounding granulites. Its rutile age is 11 Ma older than that of rutile from sample T137 and could possibly be explained by the larger diffusion radius of the stubby rutile grains from the meta-qtz-diorite.
Two rutile fractions from the Umba Steppe were picked from the same sample of semipelitic gneiss as the monazite (A144) and yield similar
206Pb
/238U ages of 51595 and 51492 Ma (Fig.
5). The two rutile fractions from metapelite sam-ples P1 and P9 of the northern Uluguru Moun-tains are normally and reversely discordant, respectively. There is no reason to assume that rutile could be affected by excess206Pb since rutile
generally has extremely low Th contents. Unless other geological problems are responsible for the discordance, the best estimate of the ‘true’ age is
probably the 206
Pb/238
U age (see discussion
above). The 206
Pb/238
U ages are indistinguishable
within error at 55094 and 54792 Ma. They are
interpreted to date the time the rocks cooled
below Tc. Rutile from metapelite T28 is
concor-dant and has a206Pb
/238U age of 50095 Ma (Fig.
8b). The slightly discordant ages at ca. 550 Ma from the northern Uluguru Mountains are also about 15 Ma older than the oldest rutile age
obtained in the Pare and Usambara Mountains, a similar age difference as observed with the monazite.
6. Discussion
6.1. Discussion of pre6ious geochronological results
The geochronologic results from this study can be combined with published mineral ages and
P–T estimates to derive quantitative P–T–t
paths for the different granulite segments of the
Mozambique Belt. Previously published
geochronological data for the granulites of eastern Tanzania are summarised in Table 3. They are recalculated using modern decay constants when necessary and re-interpreted (using the closure temperatures summarised in Table 1). All ages are combined to discuss and compare the cooling histories of the different Pan-African granulite complexes in Tanzania.
Zircon ages are available from four of the granulite complexes (Coolen et al., 1982; Maboko et al., 1985; Muhongo and Lenoir, 1994). These ages span a period of 70 Ma between 645 Ma and 715 Ma and were originally interpreted as the time of high grade metamorphism. Upper discor-dia intercepts at around 700 Ma (Maboko et al. (1985), Fig. 9a) are re-interpreted as intrusion ages. However, the U – Pb results on the large magnetic and size fractions yield mostly short discordias intersecting concordia at a low angle
(Maboko et al., 1985) and are, therefore,
imprecise.
Studies which allow direct comparison of Rb – Sr with K – Ar data from the same terrane (An-driessen et al., 1985; Priem et al., 1979) or direct comparison of K – Ar and Ar – Ar data (Maboko et al., 1989) reveal that many of the K – Ar and some Ar – Ar ages of biotite and hornblende are too old and may be influenced by excess Ar (see Fig. 9b and c). Biotite Rb – Sr data and muscovite K – Ar and Ar – Ar data, however, are consistent with other geochronological results.
For the Wami River granulites correlation of the re-interpreted results yields a slow integrated
A
Summary of published geochronological data for Pan-African granulites in E-Tanzaniaa
Area Rock type Method Mineral Age (Ma) Remarks, re-interpretation Reference
Muhongo and Lenoir u.i., metamorphic
645910 Zrn
S-Pare Mountains Enderbite U–Pb
(1994)
503920b Cahen and Snelling
Pegmatite K–Ar Bt
Granulitic gneiss K–Ar Hbl Excess Ar Cahen and Snelling
20 km W of N-Pare
u.i., intrusions Maboko et al. (1985) 882–1600, 701–715
Zrn U–Pb
Ortho-granulites
Wami River
Ortho-granulites U-Pb Zrn 476–538, 620 – 642 u.i., Pb-loss, metamorphic Maboko et al. (1985)
Ortho-granulites Rb–Sr Bt 458–485 Five results Maboko et al. (1985)
Maboko et al. (1985) 7198
69594 u.i., intrusion
Uluguru Mountains Anorthosite U–Pb Zrn
(1994) 63397, 618916
Anorthosite, Sm–Nd Grt-wr Two results Maboko and Nakamura
(1995) Ortho-granulite
53693, 550910, Three results Muhongo (1990)
Rb–Sr Bt
Ortho-granulites
577930
(K–Ar) 702914 Maboko et al. (1989)
Ar–Ar Hbl 62893
Excess Ar (Ar–Ar) Maboko et al. (1989)
K–Ar Bt 978920, 1537933
560911, 529910 –, –,580, –, no plateau
48792, 49592 (K–Ar) 488910, 500910 Maboko et al. (1989)
Ar–Ar Ms
Maboko et al. (1989) (K–Ar) 43499, 45299
Ar–Ar Kfs 42292, 45092
Approx. ages at 115, 80 and
Ortho-granulite Fission track Ap 300, 80, 30 Noble et al. (1997)
20°C
Coolen et al. (1982)
Furua complex Grt-2Px granulite U–Pb Zrn 652910 l.i., metamorphic
Priem et al. (1979) Bt Av. 463 (388–511) Fourteen Results
Rb–Sr
Furua complex Ortho-granulites
Two results Priem et al. (1979)
Rb–Sr Ms 524910, 531912
Surrounds
414–560 Fourteen results Priem et al. (1979)
K–Ar Bt
Priem et al. (1979)
K–Ar Ms 482914, 483914, Three results
487915
Andriessen et al. (1985)
K–Ar Hbl 614–665 Six results
aNote: biotite and hornblende K–Ar data of Andriessen et al. (1985) and Priem et al. (1979), and biotite Rb–Sr (Maboko et al., 1985) are too numerous to list
cooling rate between 2 and 4.5°C/Ma (Fig. 9a) for
granulites which have experienced similar P–T
conditions to the Pare, Usambara and Uluguru Mountains (Appel et al., 1998). Maboko et al. (1989) used K – Ar as well as the40Ar –39Ar
step-heating technique on hornblende, biotite, muscov-ite and K-feldspar from samples collected at the NW edge of the Uluguru Mountains, close to the northern edge of the anorthosite complex and to location T46 of this study (Fig. 2). Interpretation of the data of Maboko et al. (1989) is complicated by the fact that only one of the samples in their study (yielding the hornblende analysis and one of the biotite results) is from the granulite complex itself, whereas the other three samples (biotite, muscovite and K-feldspar) were collected from felsic gneisses and migmatites surrounding the granulite complex. Growth of muscovite in these rocks may be retrograde (related to rehydration reactions) in rocks which re-equilibrated to lower Tthan the granulite complex itself. Therefore, the muscovite ages may have to be regarded as mini-mum ages. Despite their higher closure tempera-ture, K – Ar muscovite ages are without exception significantly younger than K – Ar biotite ages (Maboko et al., 1989). High K – Ar and Ar – Ar ages on hornblende and biotite from the granulite sample and biotite K – Ar and Ar – Ar results of migmatite samples are interpreted to reflect excess Ar or were already interpreted in this way by Maboko et al. (1989). The hornblende Ar – Ar age is identical to the zircon and monazite ages (T46, this study) as well as to two garnet Sm – Nd isochrons ages for granulites from the area (Maboko and Nakamura, 1995). Because instan-taneous cooling by about 350°C down to 450°C (shaded area in Fig. 9b) is considered very un-likely for these deep crustal (ca. 10 kb) granulites which show no petrological evidence for decom-pression, this result is also interpreted to reflect excess Ar. A tentative cooling path of about
3°C/Ma is indicated by a dashed line (Fig. 9b).
The garnet Sm – Nd isochron age of 618916 Ma
on texturally late, undeformed garnet coronas in the anorthosite (Maboko and Nakamura, 1995) is important evidence in support of a single gran-ulite facies event, because it indicates that garnet growth during retrograde isobaric cooling
oc-curred just after or contemporaneously with zir-con and monazite growth in meta-qtz-diorite T46. In the Furua complex, the combination of a
lower U – Pb intercept of zircon at 652910 Ma
(Archaean upper intercept; Coolen et al. 1982) and Rb – Sr biotite and muscovite ages (Priem et al., 1979) yields a prolonged history of slow inte-grated cooling at a rate of about 2.5°C/Ma for the
granulite-facies rocks, and the surrounding
migmatitic gneisses (Fig. 9c). The results of K – Ar on hornblende and biotite (Andriessen et al., 1985; Priem et al., 1979) are again interpreted as unreliable possibly because of the presence of excess Ar. The reconstruction of the cooling his-tory is thus based on Rb – Sr biotite and muscov-ite ages and K – Ar results on muscovmuscov-ite.
The geochronological data available prior to this study and their interpretation did not provide a conclusive picture of the precise age of Pan-African metamorphism in Tanzania, its spatial distribution and the post-metamorphic cooling history. One of the problems appears to be the ubiquitous presence of excess Ar in hornblende and biotite.
6.2. Cooling histories of the granulites
Interpretation of the published U – Pb data on zircon for the Wami River complex, the Furua complex and data from the Pare Mountains (Muhongo and Lenoir, 1994), in combination with the new mineral data presented here, sug-gests that the granulite-facies event reached its peak between 620 Ma and about 640 – 650 Ma. The range of ages from published U – Pb studies on zircon is consistent with the results of this study. The data allow the tentative reconstruction of cooling histories for the granulites of the Wami River complex and the Furua Complex and possi-bly the NW Uluguru Mountains (Fig. 9), indicat-ing similar prolonged slow coolindicat-ing with integrated cooling rates of 2 – 5°C/Ma over time intervals of 140 – 240 Ma.
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 138
1998), consistent with a U – Pb zircon upper
inter-cept age of 645910 Ma on a granulite facies
gneiss from the Pare Mountains (Muhongo and Lenoir, 1994). Monazite from the Usambara Mountains yields ages of 625 – 630 Ma which are about 15 Ma younger than the time of peak metamoiphism in the Pare Mountains. The mon-azite and titanite ages of 617 – 620 Ma from the Umba Steppe semipelites and marbles are inte-grated to date the peak of metamorphism, esti-mated at 750 – 800°C by Mo¨ller (1995) and significantly younger than the time of metamor-phism determined for the Pare and the Usambara Mountains.
Within the northern and eastern Uluguru Mountains monazite ages span an age range of at least 11 m.y., possibly 23 m.y. It is unclear whether the different ages reflect true differences in the time of peak metamorphism or whether they can, in part, also be attributed to inheritance or preservation of prograde growth, the latter being the preferred interpretation. The results may be seen as the age envelope in which to place granulite facies metamorphism in this part of the Uluguru Mountains. More obvious diachronism of Pan-African metamorphism is indicated by the 20 m.y. younger monazite and zircon ages on meta-qtz-diorite sample T46 in the NW Uluguru Mountains than of monazite in the northern and eastern Uluguru Mountains (Fig. 10b). The mini-mum age difference is 20 m.y. Congruence of the monazite and zircon age of T46 (where the zircon fraction is interpreted as the product of crystalli-sation of partial melt during the earliest stages of cooling from peak of metamorphism) is taken as strong evidence for peak metamorphism at this
time (and a high Tc of monazite). The upper
zircon intercept of 69594 Ma from the
anorthosite (Muhongo and Lenoir, 1994) may be interpreted as the age of crystallisation of the anorthosite body (Fig. 9b and Fig. 10b). It can be concluded that there appear to be differences in the timing of high grade metamorphism between different granulite complexes in the Pan-African Mozambique Belt of Tanzania on a relatively small scale of less than 100 km.
Using a Tc of 650°C for titanite results in an
initial cooling rate of about 5°C/Ma for the Pare and the Uluguru Mountains. Diachronism in the thermal history within the granulites of north-eastern Tanzania is supported by rutile (and some biotite) cooling ages which show the same pattern of regional age distribution as the monazite in many areas (Fig. 2 and Fig. 10). Rutile ages are about 100 Ma younger than monazite ages,
indi-cating an integrated cooling rate of about 4°C/
Ma. This supports the interpretation that the age difference observed in the monazite results has a geological cause associated with Pan-African metamorphism (Fig. 10a) and the subsequent un-roofing history. The Umba Steppe and the Usam-bara Mountains underwent a cooling history
Muhongo and Lenoir, 1994) and thus provide a reliable estimate for the time of the Pan-African metamorphism in the Mozambique Belt of Tanza-nia. The observed cooling paths are consistent with the anti-clockwise isobaric cooling
(ACW-IBC) P–Tpath deduced from petrological
obser-vations and thermobarometry (Appel et al., 1998). The cooling histories of the granulite terranes, are parallel but offset (Fig. 11). supporting the notion that there are real age differences between mountain ranges (Pare Mountains and Usambara Mountains and Umba Steppe) and within a single mountain range (Uluguru Mountains). These dif-ferences may be taken as evidence that these age domains are separated by important but hidden faults, or may be explained by variations in the location of an external heat source. The progres-sive slowing of cooling rates is consistent with the hypothesis that a granulite-facies event is at least in part driven by an external heat source, e.g. the intrusion of magmas into the lower crust (Anovitz and Chase, 1990; Oxburgh, 1990) and slow uplift rates.
7.2. Tectonic scenario for granulite metamorphism in eastern Tanzania
Previous geochronological results for the Tan-zanian granulites were interpreted to indicate fast uplift following tectonic crustal thickening after plate-collision during the formation of Gondwana (e.g. Maboko et al., 1985, 1989; Muhongo and Lenoir, 1994; Maboko and Nakamura, 1995). A collision process would cause rapid decompres-sion after or during the thermal peak of metamor-phism, leaving behind decompression textures in the rock record. Such rapid decompression is invariably associated with fast cooling rates
(Eng-land and Thompson, 1984; Bohlen, 1991) of \
20°C/Ma as shown by data from the Alps and
Himalayas (e.g. von Blanckenburg et al., 1989) and contrast with the slow integrated cooling rates in the Tanzanian granulites. It is concluded that a continent collision scenario is not compat-ible with mineral texture (Coolen, 1980; Mo¨ller, 1995; Appel et al., 1998), fluid inclusion (Herms and Schenk, 1998) and geochronological evidence (this study).
Fig. 11. Compilation of the integrated cooling paths of gran-ulite complexes from the Pan-African Belt of NE Tanzania: Pare and Usambara Mountains and Umba Steppe (dark paths) and different parts of the Uluguru Mountains (light paths). The range of cooling paths in the studied granulite areas is indicated by the light shaded area. Paths constructed from previous geochronological data for the Wami River (1: Maboko et al., 1985) and Furua complex granulites (2: Coolen et al., 1982; Andriessen et al., 1985; Priem et al., 1979) shown for comparison (black paths) fall on the same swath of cooling paths.
similar but time-parallel to that of the Pare Mountains, with initial cooling rates of about 5°C/Ma.
7. Conclusions
7.1. The P–T–t e6olution of Pan-African high pressure granulite terranes in Eastern Tanzania
The ages of the high pressure granulite peak-metamorphism in eastern Tanzania can be con-strained to an interval from 655 to 615 Ma and varies significantly in the five different areas (the Pare Mountains, the Usambara Mountains, the Umba Steppe, the NW Uluguru Mountains and the N and E Uluguru Mountains). The ages ob-tained from monazite and zircon in this study are in the lower range of previously published U – Pb
zircon ages for the Wami River Complex
(Maboko et al., 1985) or coincide closely with zircon ages obtained for the Pare Mountains and
A.Mo¨ller et al./Precambrian Research104 (2000) 123 – 146 140
Slow cooling and all other evidence can be best reconciled with a scenario in which Pan-African granulite facies metamorphism in eastern Tanza-nia was caused by underplating and intrusion of magmas into the crust, causing heating and burial (e.g. Oxburgh, 1990). Cessation of magmatic ac-tivity initiated cooling at rates of about 5°C/Ma which slowed progressively. Post-orogenic col-lapse was either slow or evidence for it could not
be detected by integration of the present
geochronologic data. An active continental mar-gin setting is a likely tectonic setting consistent with the chemistry of the meta-igneous granulites (Appel, 1996).
The combined geochronologic evidence from the granulite terranes provides strong evidence that the final collision of East- and West-Gond-wana has not directly caused granulite-facies metamorphism in the Pan-African domains of Tanzania. Crustal shortening and uplift of the granulites due to continent collision is younger than the peak-metamorphic ages determined in this study. The available data suggest that the granulites of eastern Tanzania had already cooled to 450°C or below at about 550 Ma. This inter-pretation requires refinement of recent models for the geodynamic evolution of the Mozambique Belt of East Africa and the circum-Indic Pan-African orogen (e.g. Stern, 1994) on the whole.
7.3. Age distribution of Pan-African high grade terranes in central Gondwana: implications for plate tectonic scenarios?
It is well accepted that the Pan-African Belt in Central Gondwana is the site of a suture from the collision of East- and West-Gondwana (e.g. Burke et al., 1977). However, there are currently two different plate-tectonic interpretations for the for-mation of granulites in the Pan-African Belt of Tanzania (and East Africa on the whole).
The first model proposes the collision of East-and West-Gondwana causing high-grade meta-morphism in East Africa before 600 Ma and is similar to the collision and ocean-closure model of Stern (1994). The age of metamorphism in central Gondwana has been used repeatedly to constrain the age of collision, based on the
as-sumption that metamorphism was caused directly by the collision event. Kro¨ner (1993), for example, concluded that the granulite-facies metamorphic events in eastern and north-eastern Africa could be explained by ocean closure, terrane accretion and continental collision of East- and West-Gondwana during the time-span between 600 and 800 Ma, consistent with the scenario of the SWEAT-hypothesis discussed by Dalziel (1991) and Moores (1991).
The second model, put forward by Meert et al. (1995) based on palaeomagnetic evidence and Ar – Ar dating of Neoproterozoic lavas from Tan-zania, argues for polyphase accretion of terranes to form the Gondwana supercontinent, because they observed no common polar-wander-path for East- and West-Gondwana before 550 Ma. They proposed specifically that the formation of Gond-wana took place in two distinct orogenic events. The Pan-African orogen is considered to represent the earlier collision of India, Madagascar, Sri Lanka, part of Antarctica and the Kalahari cra-ton with the Congo cracra-ton and the Arabian – Nu-bian Shield between 800 and 650 Ma, during closure of the northern part of the Mozambique ocean. The second event occurred further to the south at 550 Ma and resulted in collision and amalgamation of the earlier ‘Pan-African’ orogen with the rest of Antarctica and Australia to form Gondwana at 550 Ma.
Neither of the two variations on the collision model for the Pan-African Belt can be fully sup-ported by the present study as it leads to the conclusion that high-grade metamorphism prior to 600 Ma did not require a concurrent continent collision. The collision of East- and West-Gond-wana did not cause the formation of granulites in eastern Tanzania and, therefore, the age of high-grade metamorphism cannot be used directly to date the collision event, an approach that has been used in all previous interpretations.
Madagascar (Andriamarofahatra et al., 1990; Guerrot et al., 1993; Ne´de´lec et al., 1994; Paquette et al., 1994; Tucker et al., 1999), South India (Buhl et al., 1983; Choudhary et al., 1992; Un-nikrishnan-Warrier et al., 1995) and also in the Lu¨tzow-Holm Complex of Antarctica (Shiraishi et al., 1992; 1994). Granulite-facies metamorphism in Sri Lanka appears to have a complex history with evidence for a metamorphic event at 610 Ma, but magmatic activity lasting until 550 Ma (Baur et al., 1991; Ho¨lzl et al., 1994). Except for the
results from Somalia, these ages are consistent with the second of two collision phases in the
model of Meert et al. (1995). Only complete P–
T–t histories of these terranes will help to deter-mine whether collision was indeed the cause of metamorphism.
Pre-600 Ma granulites occur in eastern Tanza-nia (this study, Coolen et al. (1982), Maboko et al. (1985), Muhongo and Lenoir (1994)), inter-preted by Appel et al. (1998) to follow
ACW-IBC-paths. Other ACW-IBC-granulites occur in
Antarctica (Shiraishi and Kagami, 1992). Pre-600 Ma granulites are also exposed in southern Ethiopia (Teklay et al., 1998) and central Kenya (Key et al., 1989), but more data are needed in these two areas to substantiate the existing results as ages of metamorphism. There are different groups of ages for the Pan-African metamorphic event in different areas of the circum-Indic region, but no simple East – West pattern as proposed by the two-phase collision model of Meert et al. (1995) is evident (Fig. 12). MoreP–T–t informa-tion on other granulite terranes is needed to estab-lish whether the apparent diachronism of pre- and post-600 Ma metamorphic ages can be explained best by diachronism in the collision of East- and West-Gondwana similar to the model of Meert et al. (1995), or whether it is the result of complex continental margin processes as diverse as those found in modern orogens.
Acknowledgements
This paper is a contribution to IGCP 348. A.M. thanks the colleagues and staff of the Max-Planck-Institut fu¨r Chemie in Mainz for their kind assistance and helpful discussions. We like to thank P. Appel for his close collaboration in this project and numerous fruitful discussions. We acknowledge the detailed reviews of F. Corfu, A. Lanzirotti and an anonymous reviewer which helped to improve the manuscript significantly. Research permits from the Tanzanian Commis-sion for Science and Technology and support from the University of Dar-es-Salaam (foremost S. Muhongo) and the Geological Survey of Tan-zania are also gratefully acknowledged. Research
Fig. 12. Palaeogeographic map of part of central Gondwana (modified after Windley et al. (1994), adapted from Kriegsman (1995)) with estimates for the age of Pan-African granulite-fa-cies metamorphism. Geochronological data from (1) Ho¨lzl et al. (1994): U – Pb Zrn+Mnz, (2) Baur et al. (1991): U – Pb Zrn (SHRIMP), (3) Shiraishi et al. (1992): U – Pb Zrn (SHRIMP), (4) Shiraishi et al. (1994): U – Pb Zrn (SHRIMP), (5) Choud-hary et al. (1992): Grt – Sm – Nd isochron, (6) Buhl et al. (1983): U – Pb Mnz+Zrn, (7) Unnikrishnan-Warrier et al. (1995): Grt – Sm – Nd isochron, (8) Lenoir et al. (1994): U – Pb Zrn, (9) Key et al. (1989): Sm – Nd+Rb – Sr isochron, (10) this study: U – Pb Mnz+Zrn, (11) Muhongo and Lenoir (1994): U – Pb Zrn, (12) Maboko et al. (1985): U – Pb Zrn, (13) Coolen et al. (1982): U – Pb Zrn, (14) Paquette et al. (1994): U – Pb+
A
Mineral assemblages and sample location descriptions
Kfs Bt Qtz Ms Ky Sil Rt Gr Spl Ore
Sample Area Location Rock-type Grt P1 Sec. Ky Sec. Sil Incl. in Grt Other
Metapelites
Grt P1 Kfs Bt Qtz
A016 N-Pare Eastern side of N-Pare Mountains, 1st small Metapelite Sil Rt Spl Ore Sil Mouniains bridge south of Butu-River
Metapelite Grt P1 Kfs Bt Qtz Sil Rt Gr Ore Sil S-Pare Foot of western slopes, road outcrop at road
T115
to secondary school in Hedaru Mountains
Usambara Kfs Bt Qtz Ky Sil Rt
T137 Road outcrop, Mombo-Lushoto Road, 1.5 Metapelite Grt P1 Spl Ore Sil km NE of Mombo toll station
Mountains
S-Usambara Metapelite Grt P1 Kfs Bt Qtz Ky Sil Rt Gr Ore Ky
A108 Roadcut+blocks on slope below road, 21.8 Mountains km S of Mahanga Mlai, heading to Korogwe
Grt P1 Kfs Bt Qtz sMs Rt Grt-Bt gneiss
A144 Umba Steppe Outcrop S side Umba river, along Ore
Umba-Mwakijembe track, 11.2 km E of Umba ruby mine
Grt P1 Bt Qtz Sil Rt
P001 N Uluguru River outcrop, 100 m below Morningside, ca. Metapelite sKy 10 km south of Morogoro town center
Mountains
Metapelite Grt P1 Kfs Bt Qtz Ky Sil Creek outcrop, 1 km NNE along track from Rt
P009 NE Uluguru Gr sSil Ky
Gulioni Dispensary at Morogoro-Matombo Mountains
road, Topo sheet 183-4
Grt P1 Kfs Bt Qtz Ms Ky Rt Gr Metapelite
TO28 SE Uluguru Road outcrop, 8.2 km. S of Mtombozi, Scp
Mountains Sheppards pass
Bt Qtz Cpx Opx Hbl
P1 Rt Scp Tit Ore Other
Orthogneisses,calcsilicate gneisses and marbles Grt Kfs
Grt P1 Kfs Qtz Cpx Hbl Calcsilicate
Roadcut on old Kwa Kihindi-Usangi road, Scp
A026 SE N-Pare Tit Ore
Mountains 800 m from Kwa Kihindi village
Metabasite Grt P1 Qtz Cpx Tit T114 S-Pare Western slope, creek outcrop, N of Saweni
Mountains
Outcrop S side Umba river, along Qtz Cpx Hbl
Umba Steppe Scp Tit Ore Cc
A141 Calcsilicate
Umba-Mwakijembe track, S of Gerevi hills
P1 Kfs Bt Qtz Hbl Charnockite
Quarry at crossroads in Lukosi, road to
T121 Usambara Ore
Mountains Shume
Meta-qtz-diorite
‘Anorthosite’ on QDS, old Mombo-Lushoto P1 Kfs Bt Qtz Cpx Opx Hbl
Usambara Rt
T139 Ore
Mountains road, 100 m E of first pass from sisal plantation in Mombo
T025 SE Uluguru Road outcrop Matombo-Mvuha road, 2.3 km Marble Tit Cc Mountains south of Mtornbozi village
Calcsilicate Grt P1 Bt Cpx
P008 NE Uluguru Creek outcrop, 1 km NNE of Hbl Tit
Morogoro-Matombo road, at Madabala Mountains
school, Topo sheet 183/4
Kfs Qtz Cpx
S Uluguru Hbl
P088 Creek outcrop, 100 m E of road to Lubasazi, Calcsilicate Scp Tit ca. 15 km N from Mvuha-Duturni road,
Mountains
Topo sheet 201/4
W Uluguru River outcrop W of road Morogoro-Mgeta, Meta-qtz-dioriteGrt P1 Bt Qtz Hbl
T046 Zoi
was financially supported by the Deutsche Forschungsgerneinschaft (DFG) through grants Sche 265-2/5 and Sche 265-6/1.
Appendix
See Table 4.
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