Precambrian Research 101 (2000) 25 – 47
A Paleoproterozoic ultramafic-mafic assemblage and
associated volcanic rocks of the Boromo greenstone belt:
fractionates originating from island-arc volcanic activity in
the West African craton
Didier Be´ziat
a,*, Franc¸ois Bourges
a, Pierre Debat
a, Martin Lompo
b,
Franc¸ois Martin
a, Francis Tollon
aaLaboratoire de Mine´ralogie et Cristallographie,UMR5563,Uni6ersite´ Paul Sabatier,39Alle´es Jules Guesde, 31000Toulouse Cedex,France
bDe´partement de Ge´ologie,FAST,Uni
6ersite´ de Ouagadougou,Ouagadougou,Burkina Faso
Received 31 March 1999; accepted 5 November 1999
Abstract
The Loraboue´ Birimian ultramafic-mafic assemblage, located in the Boromo greenstone belt (Burkina Faso), is interpreted as the remains of a magma chamber that crystallized at the base of an island arc. The ultramafic rocks exhibit an heteradcumulate texture, being generally made up of wehrlites and more rarely dunites. The crystallization sequence inferred from the cumulates is olivine+chromite followed by clinopyroxene+amphibole9 orthopyrox-ene9biotite. The gabbroic rocks are mainly layered and grade into more differentiated facies with sub-pegmatitic texture, containing up to 70% modal plagioclase including zircon and, more commonly, apatite crystals. Textural relationships and mineral phase chemistry are indicative of crystallization at elevated pressures (\8 kbar), the parental magma being generated by a moderate to high degree of partial melting of a mantle source affected by previous metasomatic events above a subducted oceanic slab. The Loraboue´ volcanic formations exhibit a range of geochemical features. They consist dominantly of calc-alkaline basalts, pyroclastites and rhyolite and, more rarely, of basalt, dolerite and gabbro of tholeiitic affinity. These different types of basalt, as well as the dolerite and the isolated massive gabbro, show the classic features of arc magmatic suites, namely LILE and Pb enrichment, depleted HFSE patterns and high Ce/Nb and Th/Nb ratios. Thus, the calc-alkaline plutonic and volcanic assemblages of Loraboue´ could represent the roots of an island arc and the associated coeval volcanic rocks. The Paleoproterozoic crust of the West African craton was heterogeneous and was not the consequence of a single process of genesis. As some modern igneous province, the Birimian crust was generated by both volcanic arc accretion and oceanic plateau accretion. © 2000 Elsevier Science B.V. All rights reserved.
Keywords:Greenstone belt; Paleoproterozoic; West African craton; Cumulate; Island-arc; Crustal growth
www.elsevier.com/locate/precamres
* Corresponding author..
E-mail address:dbeziat@cict.fr (D. Be´ziat)
1. Introduction
Paleoproterozoic rocks of the Man shield (Fig. 1a), referred to as the Birimian Group (Junner, 1954), form a major part of the West African Craton (Bessoles, 1977), to the east and north of the Archean Liberian cratonic nucleus (Camil, 1984). The Birimian terrains form narrow sedi-mentary basins and linear or arcuate volcanic belts intruded by various generations of granitoids (Leube et al., 1990; Pons et al., 1995; Hirdes et al., 1996). They correspond to a period of accretion around 2.1 Ga (Abouchami et al., 1990; Boher et al., 1992; Taylor et al., 1992) during the 2.1 – 2.0
Ga Eburnean orogeny (Bonhomme, 1962;
Lie´geois et al., 1991). The supracrustal sequence was folded and metamorphosed under green-schist-facies conditions.
The lithostratigraphic succession of the Birim-ian greenstone belt formations, although previ-ously much debated (e.g. Feybesse and Milesi, 1994; Hirdes et al., 1996; Pouclet et al., 1996), now seems to be well established. The Birimian greenstone belt succession consist of a thick se-quence of basalt, locally pillowed, as well as
do-lerites, and gabbros displaying a tholeiitic
composition, interlayered with immature detrital sediments and carbonates, overlain by a detrital sedimentary pile (volcanoclastics, turbidites, mud-stones and carbonates), including interbedded calc-alkaline volcanics (references in Sylvester and Attoh, 1992; Hirdes et al., 1996; Pouclet et al., 1996).
On the other hand, the processes of growth of the Proterozoic continental crust has been the subject of much speculation: island-arc accretion, intracontinental rifting and magmatic
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 27
Fig. 1. (Continued)
ing and overplating or continent-continent colli-sion (e.g. Taylor and McLennan, 1985; Condie, 1992). In the Man shield, these processes have been characterized essentially from geochemical and isotopic data on the volcanic sequence of the Birimian mafic suite which is, in fact, made up of volcanic rocks but also of plutonic complexes (ultrabasites and gabbros) sometimes outcropping over wide areas, e.g. the ultrabasic bodies from the eastern border of the Mako series, Eastern Senegal (Bassot, 1966; Debat et al., 1984; N’Gom, 1995) and from the Ashanti belt in Ghana (Loh
and Hirdes, 1996). Because of the absence of radiometric age dates and the structural disloca-tion, the relations between such intrusive com-plexes and their enclosing volcanic rocks remain ambiguous (Mile´si et al., 1989; N’Gom, 1995; Hirdes et al., 1996).
crystalliza-Fig. 1. (Continued)
tion so that to discuss the possible environments of formation of the series. This volcanic and plutonic assemblage may represent:
1. an association of basalt and layered intrusions, 2. all of the crustal members of an ophiolitic
suite,
3. the plutonic roots of the lower levels and the coeval volcanic rocks of an island arc.
2. Geology
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 29
which trends north-south and is 500 km long and 40 km wide (Feybesse et al., 1990; Lompo et al., 1991). The Boromo greenstone belt is 2.15 – 2.10 Ga old (Feybesse and Milesi, 1994) and belongs to the Paleoproterozoic Eastern subprovince defined by Hirdes et al. (1996). This terrain was chosen because of the presence of numerous bod-ies of ultramafic-mafic rock (Syndicat BRGM-BUVOGMI, 1983; Fig. 1c) together with drill cores obtained by the SOREMIB Company (SOREMIB, 1990) which allow us to observe some unaltered samples and their relationships. The Loraboue´ prospect forms an area over 4 km long and 2 km wide (Fig. 1d). All rocks have been subjected to greenschist-facies metamorphism and hydrothermal alteration, and are affected by a N-S-trending steeply dipping regional foliation. The Boromo belt hosts several types of mineral-ization (Mile´si et al., 1992) including a massive Zn-Ag sulphide deposit at Perkoa (Marcoux et al., 1988; Napon, 1988; Ouedraogo, 1989). Dis-seminated gold is associated with albitite and listvenite at Loraboue´ (Be´ziat et al., 1998), and with albitite at Larafella (Bamba et al., 1997). As with most of the gold of West Africa, gold-bear-ing quartz veins are observed at Poura (Oue-draogo, 1989; Sanogo, 1993).
3. Petrography
The original sequence is composed of volcanic rocks interlayered with metasediments, within which are developed assemblages of mafic and ultramafic rocks and some isolated gabbroic bod-ies. The volcanic sequence is dominated by fine-grained basalt flows, with volumetrically minor differentiates occurring as metafelsic and pyro-clastic rocks, dolerite dykes up to several metres thick and thin rhyolitic dykes. The basalts are generally aphyric and yield typical greenschist-fa-cies alteration assemblages. They can be greatly affected by carbonate alteration. Dolerites exhibit the same mineral assemblage as basalt, but an intersertal to ophitic texture is commonly pre-served. Small cross-cutting rhyolitic dykes are typ-ically composed of plagioclase phenocrysts and corroded quartz (Fig. 2A).
The plutonic suite consists of ultramafic rocks associated with gabbroic rocks, as well as gab-broic rocks cropping out as isolated bodies in restricted areas. The ultramafic rocks exhibit an
heteradcumulate texture, with sub-rounded
olivine crystals (up to 4 mm across) poikilitically enclosed by megacrysts (2 – 3 cm) of pyroxene (Fig. 2B) and brown amphibole. Varying propor-tions of the cumulus and post-cumulus minerals generally give rise to wehrlitic and, more rarely, to dunitic cumulates. The post-cumulus minerals are essentially represented by clinopyroxene and brown amphibole (Fig. 2C), as well as accessory orthopyroxene (Fig. 2D) and biotite. The latter occurs as a selvage around amphibole crystals (Fig. 2E). Chromite is located both in olivine and post-cumulus minerals (Fig. 2B, C and D). The degree of alteration is highly variable, but olivine is generally pseudomorphed into serpentine ac-companied by veinlets of magnetite (Fig. 2D) and
occasionally heazlewoodite (Ni3S2).
Orthopyrox-ene and brown amphibole are altered into tremo-lite, which is in turn replaced by chlorite; chromite is locally transformed to ferritchromite and biotite to chlorite (Fig. 2E).
The gabbroic rocks associated with the ultra-mafic cumulates are usually layered. Composi-tional layering is made up of modal variations, with layers of wehrlitic composition grading into layers of gabbroic composition devoid of olivine and chromite crystals, in which plagioclase is much more abundant (Fig. 2F). In the gabbroic suite, amphibole is markedly more abundant than clinopyroxene. Locally, gabbros show some dif-ferentiated facies with sub-pegmatitic texture
cor-responding to late-stage segregation. These
dioritic facies contain up to 70% modal plagio-clase (up to 1 cm across) including zircon, and more frequently, apatite crystals. Clinopyroxene is partially recrystallized to clinozoisite, and ilmenite replaced at the margins by titanite (Fig. 2G).
The gabbros occurring as isolated bodies con-sist of two petrographic types:
1. layered gabbro, similar to the gabbro associ-ated with ultramafic rocks;
Fig. 2. Sample photographs and photomicrographs showing: A. Rhyolitic dykes cross-cutting tholeiitic basalts, typically with plagioclase (Pl) phenocrysts and corroded quartz (Qz); B. Heteradcumulate texture of a wehrlite; serpentinized olivine (S.Ol), chromite (Chr); C. Post-cumulus clinopyroxene (Cpx) and brown amphibole (Am) in wehrlite; D. Relic grains of fresh olivine (Ol) enclosed in orthopyroxene (Opx); magnetite (Mt); E. Post-cumulus amphibole and biotite (Bi, partially pseudomorphosed into chlorite) in wehrlite; F. Layered gabbro; G. Differentiated gabbro; ilmenite (Ilm); H. CaKa X-ray map of a clinopyroxene containing thin lamellae of chlorite.
with the layered gabbro-magnetite with thin lamellae of exsolved ilmenite.
4. Mineralogy
More than a hundred electron-microprobe analyses were performed using a CAMEBAX SX 50 instrument at the Laboratoire de Mineralogie, Toulouse (beam current of 10 or 20 nA depending on the mineral’s resistance to beam damage, 15 kV, with counting times of 10 s at peak positions and 5 s for backgrounds). The data were cor-rected using PAP (SX50) procedures (a list of the standards used is available from the authors upon
request). Concentrations of major elements are accurate to 1 wt% of the element present, whereas concentrations of minor elements are less accu-rate, being reproductible to 0.1 wt%.
Representative compositions of olivine, pyroxe-nes, spinel, and amphiboles obtained only from mafic and ultramafic rocks are given in Tables 1 – 4. Relics of primary igneous phenocrysts are not observed in the volcanic rocks.
4.1. Oli6ine
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 31
Fig. 2. (Continued)
with NiO reaching a maximum of 0.40 wt% (Table 1). In terms of both Ni and Fo contents, the olivine compositions are consistent with a mantle-derived origin (Sato, 1977), being compat-ible with the compositions of olivine in cumulates crystallizing from the melt fraction.
4.2. Pyroxenes
Selected compositions of different types of py-roxene are listed in Table 1. In wehrlite, the orthopyroxene is enstatite of mean composition
(Wo5 En80 Fs15), while the clinopyroxene varies
from Mg-augite (Wo32 – 44.5 En48 – 58 Fs7.5 – 10) to
diopside (Wo47 En47 Fs6) for clinopyroxene
con-taining thin lamellae of chlorite along the (100) crystallographic planes (Fig. 2H Fig. 3). This
difference of composition could result from the
reaction Mg-augitediopside+orthopyroxene,
the chlorite lamellae representing relictual or-thopyroxene exsolution lamellae. The
clinopyrox-ene in the gabbros is Mg-augite, with lower Cr2O3
and Al2O3 contents (Fig. 4) than in the
clinopy-roxene of wehrlites. The strong positive
correla-tion between Al2O3 and Cr2O3 suggests that
charge balance is maintained by a CaCrAlSiO6
-type substitution. CaO contents are highly vari-able in all clinopyroxenes, which may reflect submicroscopic orthopyroxene exsolution lamel-lae resulting in large variations in Wo and corre-spondingly also in En components. However, the
range in magnesium number (mg*=100*Mg/
(Mg+Fe2+)) is more limited in both the
clinopy-roxenes (0.90 – 0.85) and orthopyroxenes
Table 1
Representative microprobe analyses of olivine and pyroxenesa
1
Analysis 2 3 4 5 6 7 8 9
W W W W LG
Host rock type W EG W W
Cpx Cpx Opx Chlorit.-Cpx Cpx
Cpx Cpx Ol Ol
SiO2 51.97 53.29 52.94 55.61 54.19 52.60 52.96 39.96 39.69
0.08 0.22 0.06 0.03
TIO2 0.19 0.12 0.16 -
-2.48 2.62 1.77 1.20 2.42
2.68 1.71
Al2O3 -
-4.58
FeO* 4.04 5.87 9.51 3.83 5.60 8.14 14.70 14.62
0.89 0.92 0.56 0.47 0.64
Cr2O3 1.03 0.04 0.05 0.05
0.21 0.15 0.17 0.19 0.30
0.16 0.38
MnO 0.23 0.16
16.98
MgO 18.76 18.95 30.41 16.50 16.32 12.17 44.66 45.28
19.11 18.12
CaO 21.88 2.14 23.29 21.67 23.70 0.06 0.02
0.37 0.21 0.03 0.34 0.32
0.24 0.82
Na2O -
-0.00
K2O 0.01 0.00 0.01 0.04 0.03 0.02 -
-99.24 100.00 100.27 100.08 100.02
Total 99.71 100.10 99.99 100.16
39.38 36.84 4.12 47.16
Wo 44.47 44.24 50.12 Fo 0.84 0.85
En 48.01 53.78 53.61 81.36 46.49 46.36 35.81 Fa 0.16 0.15
6.84 9.56
Fs 7.52 14.53 6.36 9.41 14.07
aIn W: wehrlite; LG: layered gabbro; EG: evolved gabbro; Chlorit-Cpx: clinopyroxene containing thin lamellae of chlorite.
In the more evolved gabbros, clinopyroxene is
diopside (Wo50 En36 Fs14). It differs from the
clinopyroxene of wehrlites and gabbros in having
lower mg* (B80), Al2O3 and Cr2O3 (Fig. 4) and
higher CaO, Na2O and FeO contents.
4.3. Spinel
Spinel is mostly encountered in ultramafic cu-mulates, more rarely in the gabbros and basalts.
Selected compositions of different types of
chromite are listed in Table 2. They correspond to chromite, with only one grain having the
compo-sition of magnesio-chromite (Mg/M2+\Fe2+/
M2+, Table 2). Chromite is dispersed throughout
the entire volume of the ultramafic cumulates. It
occurs as 200 – 400mm euhedral to rounded grains
included in olivine, interstitial to olivine, or in-completely enclosed by olivine and in partial con-tact with postcumulus phases (Fig. 2B, C and D). Compositional zoning within individual grains is not observed; only a thin ferritchromite rim is developed along boundaries in individual grains located in secondary serpentine or chlorite phases.
When plotted against Mg ratio (=Mg/(Mg+
Fe2+), chromite exhibits similar Cr contents but a
distinct inverse relationship between Al and Fe3+
,
except those having a Mg ratio B0.2 which
shows more scattered trivalent cation contents
(Fig. 5). TiO2also correlates positively with Fe
3+
(Fig. 6). These compositional variations indicate a spinel-magnetite type-substitution.
No significant differences in chemistry exist be-tween chromites in dunites or wehrlites. However, significant differences exist among interstitial chromites entirely included in either clinopyrox-ene or amphibole. These chromites are distinctly
Al-poor and Fe3+-rich, while most of them are
Mg-rich compared with chromites included or incompletely enclosed by olivine.
Fractional crystallization can be excluded as a major cause of the compositional variation of chromites because of the slight change in Cr content relative to their Mg ratio (Hulbert and Von Gruenewaldt, 1985; Leblanc, 1985). So, the chemical characteristics of chromite are in part a result of hydrothermal alteration processes and metamorphism which affect the host rocks, with
dissolution-recrystallization from magmatic
D
Representative microprobe analyses of spinela
5 6 7 8 9 10
Analysis 1c 1r 2 3 4 11 12 13
W W W W W W LG
W LG
Host rock D D D D B
S.Ol S.Ol-Am Cpx Am S.01-Bi Bi S.Ol
Related silicate S.Ol S.Ol S.Ol-Cpx Cpx Ol Am Trem
0.05 0.01 0.12 0.00 0.01 0.00
0.09 0.00
0.09 0.00 0.45 0.10
SiO2 0.02 0.08
16.24 13.75 20.23 11.78 12.51 6.79
Al2O3 16.68 17.03 16.04 17.32 18.90 19.21 16.87 13.40
34.47 38.12 24.75 48.74 42.11 57.47 30.16
34.12 38.36
FeO* 31.48 33.35 31.97 29.91 35.19
0.41 0.46 0.28 0.29 0.36 0.32 0.24 0.27 0.35 0.74 0.18 1.33 1.68
MnO 0.34
6.33 6.18 11.84 6.27 5.56 2.32 8.93
7.16 1.44
MgO 8.49 7.30 8.75 9.78 0.35
0.66 1.61 0.39 0.14 0.82 3.10 0.50
TiO2 0.72 0.64 1.74 1.64 0.43 0.35 0.20
40.15 38.21 42.93 30.17 37.15 27.26 39.11
37.66 37.75
Cr2O3 39.90 38.75 39.62 38.80 44.49
0.00
0.00 0.00 0.00 0.00 0.00 0.00 0.08 0.20 0.15 0.11 0.00 0.00 0.00
ZnO
0.66 0.20 0.09 0.28 0.22 0.25 0.17
0.43 0.16
0.04 0.07
NiO 0.15 0.17 0.18
98.99
97.85 97.65 98.66 98.06 98.92 98.40 100.67 97.85 98.88 98.04 98.26 96.71 95.48
Total
0.002 0.000 0.004 0.000 0.000 0.000 0.000 0.016
Si 0.001 0.003 0.003 0.003 0.000 0.004
0.633 0.542 0.738 0.470 0.496 0.281 0.731
0.724 0.697
0.659 0.574
Al 0.645 0.664 0.614
0.300 0.408 0.202 0.718 0.494 0.881 0.259
Fe3+
0.248
0.303 0.305 0.325 0.311 0.298 0.144
0.654 0.658 0.439 0.662 0.692 0.807 0.555
0.630 0.877
Fe2+ 0.561 0.618 0.543 0.496 0.925
0.010 0.009 0.006 0.008 0.010 0.022 0.005
Mn 0.011 0.013 0.008 0.009 0.008 0.039 0.052
0.312 0.308 0.547 0.317 0.279 0.121 0.430
0.347 0.075
0.471 0.019
Mg 0.415 0.360 0.424
0.016 0.040 0.009 0.004 0.021 0.082 0.012
Ti 0.018 0.016 0.043 0.040 0.011 0.009 0.005
1.050 1.010 1.051 0.808 0.989 0.757 0.998
0.968 1.046
Cr 1.035 1.014 1.018 0.990 1.277
0.000
Zn 0.000 0.000 0.000 0.000 0.000 0.000 0.002 0.005 0.004 0.003 0.000 0.000 0.000
0.015 0.005 0.002 0.007 0.005 0.006 0.004
0.010 0.004
Ni 0.003 0.001 0.004 0.004 0.002
0.525 0.505 0.525 0.404 0.494 0.378
Cr/M3+ 0.517 0.507 0.509 0.495 0.484 0.499 0.523 0.639
0.317 0.271 0.369 0.235 0.248 0.140 0.365
0.362 0.348
0.670 0.675 0.442 0.668 0.703 0.846 0.561
0.640 0.884
Mg/M2+ 0.420 0.364 0.435 0.482 0.352 0.316 0.550 0.319 0.284 0.127 0.434 0.076 0.019
ac: core; r: rim; D: dunite; W: wehrlite; LG: layered gabbro; B: basalt; spinels included in fresh olivine (Ol), in serpentinized olivine (S.Ol) and in partial contact with clinopyroxene
chromites, leading to the complete late-stage transformation to ferritchromite (Jan et al., 1985; Kimball, 1990; Be´ziat and Monchoux, 1991; Liipo et al., 1995). The low Mg
contents of chromite grains trapped in olivine could result from subsolidus re-equilibration be-tween chromite and olivine (Wilson, 1982; Hatton and Von Gruenewaldt, 1985; Roeder and Camp-bell, 1985; Yang and Seccombe, 1993; Scowen et al., 1991). On the contrary, their generally high Al contents could be inherited from the magmatic stage because Al partitions preferentially into the Al silicates (Eales and Marsh, 1983) causing a relative decrease of Al in the interstitial chromite relative to the cumulus chromite where the only cocumulus phase is olivine.
4.4. Amphiboles
Brown amphibole frequently coexists with clinopyroxene in ultramafic cumulates (Fig. 2C) and layered gabbros, while green amphibole oc-curs in the isolated massive gabbros and the vol-canic sequence. Using IMA nomenclature (1997, Table 3), they range in composition, respectively, from pargasite to edenite and from magnesio-hornblende to actinolite. Tremolite occurs as a selvage in these amphiboles and is also as pseudo-morph of serpentine in contact with chlorite.
Amphibole compositions (Fig. 7) fall into two
distinct groups. Brown amphibole has high AlIV,
Ti, Na and K and low Si. Green and colourless
amphiboles have low AlIV, Ti, Na and K and high
Table 3
Representative microprobe analyses of amphibolesa
4
Analysis 1 2 3 5 6 7
W W W G Di Do B
Host rock
44.18 43.59 43.75 43.04 50.73 41.84
SiO2 44.95
4.06 3.77 3.51 3.37 0.16 0.48
TiO2 2.32
FeO 7.72 7.56 6.57 12.05 9.63 15.55 22.18
0.28
16.37 16.23 16.41 13.66 13.98 6.33
MgO
11.55 11.54 11.50
CaO 11.05 11.79 11.91 11.08
2.47 2.60 3.02
Si 6.409 6.456 6.438
1.584 0.513
1.529 1.562
AlIV 1.591 1.622 1.544
0.769 0.256
AlVl 0.216 0.249 0.259 0.495 0.297
Cr 0.029 0.083 0.119 0.018 0.007 0.001 0.010
0.000 0.000 0.000 0.000 0.059
Fe3+ 0.000 0.000
0.936 0.925 0.811 1.508 1.202 1.919 2.786
Fe2+
0.035
0.001 0.012 0.005 0.020 0.039 0.055
Mn
1.884
1.795 1.809 1.818 1.771 1.886 1.821
Ca
0.568 0.212
0.588 0.711
Na 0.695 0.738 0.864
0.067 0.043 0.068
K 0.109 0.118 0.090 0.128
0.79 0.67 0.72 0.59 0.34
0.79 0.82
mg*
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 35
Fig. 3. Compositional variation of pyroxenes. Orthopyroxene ("), clinopyroxene (2) and clinopyroxene containing thin lamellae of chlorite ( ) from wehrlite; clinopyroxene from layered () and more differentiated () gabbros.
MgO and Cr2O3 (Table 3) are depleted in more
differentiated gabbros relative to the cumulate amphiboles.
4.5. Feldspar
In the different lithological facies, plagioclase is always pseudomorphed into albite.
4.6. Accessory phases
Late magmatic biotite, rimming brown amphi-bole (Fig. 2E) and clinopyroxene, is present within the wehrlites and gabbros. They clearly differ from biotites of metamorphic origin
en-Fig. 5. Compositional variation of chromites. Dunite: included chromite in serpentinized olivine (), partly in contact with clinopyroxene (|stop10|) and entirely within clinopyroxene (). Wehrlite: included chromite in serpentinized olivine (2), fresh olivine ( ), clinopyroxene or amphibole (") and biotite ( ); partly in contact with clinopyroxene or amphibole ( )
and biotite ( ). Gabbro: chromite included in serpentinized olivine () and amphibole (). Basalt ( ).
Fig. 4. Cr2O3vs. Al2O3 in clinopyroxenes.
Si. The reasonably good linear correlation shown
in Fig. 7 between AlIVand Na
+K parallel to the
tremolite-pargasite line indicate that the substitu-tion mechanism operative in both amphiboles is
principally of the edenite type (Na+K)A+AlIV
A+Si. A simple Mg for Fe2+ substitution is
D
Representative whole-rock chemical analysesa
5
1 2 3 4 6 7 8 9 10 11 12 13 14 15
Sample
LG EG EG IG Do CAB And Ca-CAB Rh
D W W LG LG ThB
52.6 59.6 58.0 47.5 48.0 51.5 57.8
44.7 44.0
SiO2 36.1 37.1 40.3 49.5 66.0 46.1
0.08 0.06 0.17 0.46 0.80 1.37 3.06 1.08 1.24 0.60 0.70 0.46 0.32 1.17
TiO2 0.20
13.2 15.2 16.4 14.1 14.2 13.9 14.0
Al2O3 3.0 5.4 5.7 9.9 17.9 15.6 14.8 14.0
0.15 0.15 0.15 0.16 0.11 0.08 0.24 0.20 0.22 0.12 0.11 0.11 0.03 0.17
MnO
0.37 0.21 0.01 0.40 0.09 0.13 0.15 1.30
K2O 0.03 0.01 0.01 0.01 0.39 1.35 0.01
0.13 0.18 0.22 0.19 0.16 0.18 0.23
0.05 0.08
99.01 99.14 99.93 98.07 99.61 98.53 99.09 99.05
Total 100.05 97.57 96.88 97.36 98.11 98.81 98.11
750 43 265 37 62 109 40
0.42 0.05 0.07 0.07 0.28 0.81 20.00 2.20 2.60 2.80 3.50 0.70 2.80 2.70
Ta
5.10 11.90 2.72 0.81 1.01 1.29 1.36
Nb 0.30 0.10 0.50 0.60 2.20 0.69 3.39 1.03
1.41 4.52 1.36 0.12 0.12 0.30 0.34
0.33 0.10
0.05 0.20 0.24 0.85 0.09
Th 0.04
0.11
0.02 0.02 0.07 0.07 0.44 1.26 22.23 7.24 6.76 11.87 9.97 4.49 15.74 5.39
U
11.61 36.87 .06 2.40 2.30 3.32 2.98
La 0.59 0.65 1.73 2.04 4.79 1.35 4.10 2.01
28.61 73.40 25.12 10.14 10.32 13.00 13.01
10.53 5.52
0.00 16.06 9.70
Ce 1.40 3.84 4.45
1.32
0.16 0.18 0.49 0.59 4.08 8.83 6.22 2.39 2.50 2.52 2.91 1.16 2.73 2.66
Pr
Nd 0.70 0.76 2.11 2.50 5.76 18.33 34.33 1.47 0.93 0.97 0.79 0.94 0.39 0.69 1.03
4.72 7.48 6.54 2.48 3.01 2.17 2.83
1.23 1.05
0.16 2.03 2.94
Sm 0.18 0.53 0.69
0.50
0.06 0.09 0.21 0.24 1.07 1.30 1.22 0.44 0.52 0.34 0.46 0.13 0.26 0.56
Eu
5.61 8.73 8.22 2.92 3.49 2.06 2.89
Gd 0.22 0.21 0.60 0.81 1.54 0.63 1.23 3.57
1.03 1.52 1.80 0.68 0.77 0.46 0.63
0.23 0.13
0.04 0.20 0.81
Tb 0.04 0.12 0.16
1.33
0.23 0.31 0.87 1.12 6.72 9.63 5.22 2.00 2.26 1.31 1.82 0.35 0.57 2.26
Dy
1.53
Ho 0.06 0.07 0.20 0.27 0.29 2.19 0.74 0.29 0.33 0.20 0.27 0.05 0.07 0.32
4.50 6.20 4.66 1.97 2.27 1.36 1.79
0.88 0.37
0.63 0.80 0.39 2.07
Er 0.18 0.22
0.12
0.03 0.03 0.10 0.12 0.64 0.83 0.70 0.30 0.35 0.21 0.27 0.07 0.08 0.31
Tm
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 37
Fig. 6. Ti versus Fe3+ in chromites.
SiO2plot (Fig. 8), the Loraboue´ basaltic rocks are
clearly distributed on both sides of the Miyashiro (1974) discriminant boundary. A first group made up of dolerites, isolated massive gabbros and some basalts is tholeiitic, while the second group, made up of basaltic rocks, evolved gabbros and rhyolites is calc-alkaline. The chronology can be established on the field, dykes of rhyolite cross-cutting basalts of the first group (Fig. 2A, anal. 14 in Table 4). Due to their cumulative nature, the
Fig. 7. Compositional variation of amphiboles. AlIV, Na+K and Ti in atoms p.f.u.
countered in the volcanic sequence in having
higher TiO2 (1.5 – 3 wt%) and Cr2O3 (0.5 – 0.7
wt%) contents. Mn-rich ilmenite (\5 wt% MnO)
partially replaced by titanite, apatite and zircon are common accessories in the more differentiated gabbros (Fig. 2G). Magnetite and Ni-Fe sulphide (heazlewoodite and pentlandite) are common by-products of the serpentinization of olivine in
ul-tramafic cumulates and gabbros. However,
Ti-magnetite is present as a primary phase in the isolated massive gabbros, dolerites and the vol-canic sequence.
5. Geochemistry
Some thirty rocks exposed in the Loraboue´ prospect were analysed by XRF spectrometry for major elements as well as Nb, Rb, Y and Zr, and by ICP-AES for Ba, Co, Cr, Ni and Sr at Chemex Labs, Canada. The concentrations of the remain-ing elements were measured by ICP-MS at the Laboratoire de Ge´ochimie, Toulouse. Fourteen representative analyses are reported in Table 4. Whole-rock major and trace element analyses of cumulate rocks reflect their cumulus mineralogy;
for example, this is shown by the high MgO (\36
wt%), Cr (\600 ppm) and Ni (\1000 ppm), as
well as the low Al2O3 (B3 wt%) and CaO (B1
wt%) contents in the dunite. The moderate mg*
(Mg/(Mg+Fe)=0.40 – 0.66) and moderate levels
of Cr, Ni and Co (B240, 110 and 50 ppm,
respectively) of the basaltic rocks indicate
Fig. 8. FeO*/MgO vs. SiO2plot (Miyashiro, 1974).
On the basis of chondrite-normalized rare earth element (REE) patterns (Fig. 9), two types of basaltic rocks can be also distinguished: a
tholei-itic group with chondrite-normalized La/Yb ratio
(La/Yb)N in the range 2.1 – 3.1, similar to some
tholeiitic basalts of Haute-Comoe´, Ivory-Coast (Group 2 from Pouclet et al., 1996), in the central part of Mako (Sabodala region, N’Gom, 1995) and the Sirba region, Niger (Subgroup 1b from Ama-Salah et al., 1996), and a calc-alkaline group
with (La/Yb)Nhigher than 4 similar to the
calc-al-kaline lavas of Haute-Comoe´, Sirba region, Diale´-Dale´ma supergroups (Bassot, 1987; Dioh, 1995) and of the Bouroum area, Burkina Faso (Zonou, 1987). However, the tholeiitic group clearly differs from some of the tholeiitic Birimian basalts of Haute-Comoe´ (Group 1 from Pouclet et al., 1996), Mako (Abouchami et al., 1990; N’Gom, 1995), the Sirba region (Subgroup 1a from Ama Salah and al., 1996) and Bouroum (Abouchami et al., 1990; Zonou, 1987) which all exhibit flat to slightly light-REE-depleted patterns compatible with an oceanic magmatic signature.
A plot of Al2O3 versus their MgO content in
the different lithological units shows a steady
trend of sharply increasing Al2O3from the
plagio-clase-free Al-poor ultramafic cumulates through the gabbroic and calc-alkaline basaltic rocks to
Fig. 9. Concentrations of the rare-earth elements (REE) for different lithological units of the Loraboue´ area (C-1 chondrite values of Sun and Mc Donough, 1989); rock numbers from Table 4.
Fig. 10. Abundances of Al2O3vs. MgO for suites of samples from the Loraboue´ area.
ultramafic rocks (dunite, wehrlite and layered gabbro) associated with the differentiated gabbros plot in the tholeiitic field but appear clearly linked to the calc-alkaline suite. In the same way, the highly carbonatized basalts (e.g. anal. 12 in Table 4) plot in the same field because of the strong
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 39
Fig. 11. Abundances of U vs. Th for suites of samples from the Loraboue´ area. Note change in scale between upper and lower parts of the figure.
cumulates, to light-REE-enriched (10 – 30×
chon-dritic abundance) for basic rocks. With fractiona-tion of basic magma, the REE abundances increase, the light-REE become fractionated, the heavy-REE show a parallel rise, and the Eu anomaly decreases, becoming negative in the most evolved gabbroic rocks. The latter present the
strongest REE contents with 60 – 80×chondritic
abundance.
If normalized to the composition of N-type MORB, the different types of Loraboue´ basalt, as well as the dolerites and isolated massive gabbros, all show the classic features of arc magmatism, namely large-ion lithophile element (LILE) en-richment and high field-strength element (HFSE) depletion, with a large enrichment spike in Pb (Fig. 12). Nevertheless, we note lowest contents, notably for the LILE, in the tholeiitic samples.
6. Discussion
In the Paleoproterozoic of West Africa, there are two generations of volcanism which are up to
80×106
years apart (c. 2185 and 2105 Ma) and both contain tholeiitic as well as calc-alkaline assemblages (Hirdes et al., 1996). Several authors have proposed that ultramafic-mafic assemblages are associated with tholeiitic basalts whereas granitic rocks are associated with calc-alkaline suite (e.g. Abouchami et al., 1990; Boher et al., 1992; Sylvester and Attoh, 1992). Now,
1. in several volcanic belts in Ghana and Coˆte d’Ivoire, synvolcanic granitoids occur within tholeiitic piles,
2. in this study and in Ghana, in the Ashanti belt (Loh and Hirdes, 1996), the largest ultramafic bodies of the country occurs in a calc-alkaline suite.
6.1. Petrogenetic implications
In the ultramafic-mafic assemblage, the crystal-lization sequence inferred from the cumulates is
olivine+chromite followed by clinopyroxene+
amphibole9orthopyroxene9biotite.
Orthopy-roxene is rare as discrete grains in the wehrlites, but is common as narrow exsolution lamellae in Fig. 12. N-Type MORB-normalized multi-element diagram.
Normalization values of Sun and Mc Donough (1989).
the more evolved gabbros (Fig. 10). Plotting the data of incompatible elements (e.g. U versus Th, Fig. 11), we note that most of the points repre-senting mafic-ultramafic samples lie very close to a straight line passing through the origin, consis-tent with fractional crystallization (Joron and Treuil, 1977, 1989). REE patterns (Fig. 9) are flat
clinopyroxene. Plagioclase is never present in the wehrlites, and appears only in the gabbros. The high Mg-number of both clinopyroxene (0.90 – 0.85) and orthopyroxene (0.88 – 0.84) reflects the fact that they followed olivine very closely on the liquidus, before the magma was depleted in Mg (Elthon et al., 1982). Jaques and Green (1980) showed that the first phase crystallizing after olivine in a cumulate sequence is determined by the degree of partial melting of the mantle
source: plagioclase corresponding to low,
clinopyroxene to medium and orthopyroxene to high degrees of partial melting. Ishiwatari (1985)
showed also that high TiO2 (0.6 – 0.8 wt%) is
found in clinopyroxenes in plagioclase-type
cu-mulates, moderate TiO2 (0.4 wt%) in
clinopyrox-ene-type cumulates and low TiO2 (0.1 wt%) in
those in orthopyroxene-type cumulates. The early stage of crystallization of both pyroxenes and the
moderate to low TiO2 content of clinopyroxene
(average value of 0.22, st. dev. of 0.14) is consis-tent with the parental magma being generated by a moderate to high degree of partial melting. The comparatively low heavy REE, Y and Sc con-tents of the basaltic rocks (Table 4) could thus be a result of high degrees of melting of a de-pleted source (Baker et al., 1994).
Evidence for relatively high crystallization pressures is suggested by the early crystallization of aluminous chromite and clinopyroxene. Ac-cording to Dick and Bullen (1984), partition co-efficients for Cr in spinel decrease significantly with increasing pressure, leading to lower Cr
contents (Cr/(Cr+Al)B0.6) in spinels formed at
high pressure. In addition, Irvine (1967) sug-gested that, at pressures above the
olivine-plagio-clase stability field (\8 kbar), a basaltic liquid
will crystallize aluminous spinel and pyroxene
rather than plagioclase. The low Cr/(Cr+Al)
ra-tios (0.50 – 0.60) in the earliest crystallized spinels
from Loraboue´ and the absence of olivine+
pla-gioclase assemblages may therefore be indicative of crystallization at elevated pressures. Moreover, the presence of orthopyroxene exsolution lamel-lae in clinopyroxene of diopsidic composition, the low Mg contents of chromite grains trapped in olivine and the absence of zonation in chromite, olivine and pyroxene crystals are
in-dicative of extensive subsolidus reequilibration during slow cooling of the complex at high pres-sures.
Evidence of primary crystallization of brown
amphibole is demonstrated by the textural
(postcumulus phase enclosing olivine grains) and compositional (rich in Cr, Ti and alkalis) charac-teristics, similar to those reported in the layered sequences of ophiolites in the Oman (Lippard et al., 1986), in Halmahera, eastern Indonesia (Bal-lantyne, 1992), in amphibole gabbros, e.g. Ton-sina, Alaska (Burns, 1985; DeBari and Coleman, 1989) and in picritic rocks, e.g. Scourie dykes, northwest Scotland (Tarney and Weaver, 1987) as well as in lower crustal xenoliths from mature island arcs, e.g. the Aleutians (DeBari et al., 1987). So, the presence of primary magmatic am-phibole and biotite (Green, 1982), and the late crystallization of plagioclase (Burnham, 1979; Al-lan and Carmichael, 1984) indicate relatively high water contents in the parental magma com-patible with hydrous partial melting of a mantle source affected by previous metasomatic events.
Thus, a combination of high Ptotal and PH2O
would cause an extended period of fractionation of Al-poor minerals, such as olivine and pyroxe-nes, yielding plagioclase-free ultramafic
cumu-lates. The compositional continuity between
volcanic and plutonic calc-alkaline rocks (Figs. 8 and 10) is consistent with the fractionation of such Al-poor phases leading to a residual magma represented by basalt. In addition, the
clinopy-roxene and amphibole compositional trends
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 41
6.2. Implications for tectonic setting
It is now generally accepted that the compo-nents of individual greenstone belts probably rep-resent a variety of tectonic settings (e.g. Polat et al., 1998). Moreover, various models have been proposed for the crustal growth of the West African craton at 2.1 Ga: intracontinental rift (Ratomaharo et al., 1988; Alric, 1990; Leube et al., 1990), oceanic plateaus such as in the Nauru basin (Abouchami et al., 1990; Boher et al., 1992; Pouclet et al., 1996), intraoceanic island arc
(Sylvester and Attoh, 1992; Ama-Salah et al., 1996) and back-arc basin (Vidal and Alric, 1994; Ratomaharo et al., 1988). Here, we discuss only the tectonic significance of the volcanic and plu-tonic assemblage of Loraboue´ as part of the Boromo greenstone belt, but not the emplacement environment of the Birimian mafic volcanics as a whole.
The volcanic and plutonic assemblage of Loraboue´ could represent a part of the crustal members of an ophiolitic suite (blue schists, sheeted dykes and the tectonized, depleted mantle rocks having not yet been identified). However, occurrences of ophiolite as old as 2.0 Ga are rare, e.g. Purtuniq Ophiolite Complex in the Cape Smith belt (Scott et al., 1991) and Outokumpu and Jormua Ophiolite Complexes (Kontinen, 1987; Vuollo and Piirainen, 1989). In the same way, these igneous rocks may also represent lay-ered intrusions and their related volcanic rocks.
Several aspects of the mineral chemistry as well as the phase relationships in the cumulate rocks indicate that they were not crystallized at pres-sures and temperatures typical of a shallow ig-neous intrusion. For example, typical ultramafic cumulate rocks of oceanic crust crystallize at shal-low depths, with plagioclase crystallizing early in the sequence, directly after olivine and spinel. In
the same way, the penecontemporaneous
Bushveld complex contains layered ultramafic-mafic rocks in which orthopyroxene is the major phase even when clinopyroxene is the minor phase and the plagioclase is mostly an interstitial com-ponent of the lower zone (e.g. Hatton and Von Gruenewaldt, 1990). So, the Loraboue´ assemblage do represent neither typical layered intrusion nor typical ocean crust ophiolite.
Many authors (e.g. Irvine, 1967; Dick and Bul-len, 1984; Leblanc, 1985; Haggerty, 1991; Stowe, 1994; Zhou and Robinson, 1997) have proposed a crystallochemical model that discriminate spinels occurring in different tectonic environments (Fig.
13). As demonstrated above, post-magmatic
changes affect the Loraboue´ chromite in produc-ing chromite of lower Mg ratio and higher Cr
ratio (=Cr/(Cr+Al). So, the possible original
compositions of the Loraboue´ chromite can be inferred from the compositional trend (Fig. 13) Fig. 13. 100Cr/(Cr+Al) vs. 100Mg/(Mg+Fe2+) plot. See
Fig. 14. Plot of Ce/Nb vs. Th/Nb. Data sources from Saunders and Tarney (1991).
and used in this discriminative diagram. Although the Loraboue´ chromite compositions overlap more than one tectonic setting, their possible orig-inal compositions are located in the field of ophio-lites from arc-related crust (New Caledonia, Leblanc, 1987) rather than from mid-ocean ridges. Moreover, they have Cr ratios (50 – 60) and Ti
contents (B0.05 cation units) lower than those of
layered intrusions such as Bushveld where Cr/
(Cr+Al)=70 – 85 (Irvine, 1967) and Ti=0.1 – 0.3
cation units (Stowe, 1994).
LILE (Ba, Th, U) and HFSE (Nb, Ta, Zr, Hf, Ti, Y) element ratios combined with light-REE abundance have been widely used to identify the original tectonic environment. In this study, we do not include alkalis (K, Rb and Cs) because of their mobility during secondary alteration (e.g. Ludden et al., 1982). On the contrary, Th, Ba, Ce and the HFSE are generally considered as immo-bile in even strongly altered basalt, and therefore can be assumed to represent magmatic values with reasonable confidence. For example, Saunders
and Tarney (1991) proposed a Ce/Nb versus Th/
Nb (Fig. 14) plot. This latter also allows compari-sons of the distribution of trace element data in various basaltic rocks: ocean ridge basalt (N-type
and E-type MORB), island arc basalt (IAB), back-arc basalt (BAB) and intraplate basalt (OIB). The high ratios of calc-alkaline types are similar to those of IAB, consistent with the previ-ous light-REE and LILE enrichments and the
depletion in HFSE, even when the lowest Th/Nb
ratios of the tholeiitic types bring them nearer the BAB.
Recently, Albare`de (1998) focussing especially
on the Birimian terranes of West Africa
(Abouchami et al., 1990; Boher et al., 1992), pointed out the importance of accreting oceanic plateaus in crustal growth processes (e.g. Stein and Hofmann, 1994; White et al., 1999). There-fore, we have thought it necessary to plot our
samples in the Ce/Nb versus Ce diagram (Fig. 15),
previously used by Abouchami et al. (1990) in order to show the similarity between the tholeiitic basalts of West Africa and the Nauru basin. In
this diagram, the Ce/Nb ratios of the Loraboue´
D.Be´ziat et al./Precambrian Research101 (2000) 25 – 47 43
site (Zonou, 1987; Dia, 1988; Sylvester and Attoh, 1992; Ama-Salah et al., 1996; this study) or an oceanic plateau site (Abouchami et al., 1990; Bo-her et al., 1992; Pouclet et al., 1996).
As with most of the Birimian greenstone belts in the West African craton, we recognize both tholeiitic and calc-alkaline lavas in the Loraboue´ prospect, but their chemical signature on the whole are compatible with an island-arc environ-ment. So, we consider that all the volcanics of the Loraboue´ area were generated in a same mag-matic environment. In reference to some modern suites (e.g. the Honshu arc, Gust et al., 1997) which include both tholeiitic and calc-alkaline types, the presence of both tholeiitic and calc-al-kaline lavas could indicate a progressive change of the primitive magma type, reflecting the variable contribution of a slab-derived fluid component to the mantle wedge and the corresponding melting processes. There are some striking similarities in composition and mineralogy between plutonic rocks in island arcs and the rocks of the Loraboue´ area. Thus, the rocks of Loraboue´, part of the
Boromo greenstone belt, contain minerals and mineral compositions that appear to be character-istic of cumulate rocks from the roots of island arcs, in contrast to rift-related layered intrusions, and probably reflect the deeper levels of crystal-lization in arcs generally, e.g. the Alaskan Tal-keetna (Burns, 1985; DeBari and Coleman, 1989; DeBari, 1997) and Kohistan (Bard et al., 1980) island arcs.
The presence of massive Zn-Ag sulphide (with accessory barite) deposit at Perkoa, embedded in the Birimian volcano-sedimentary sequence from the same Boromo belt, suggests an arc-related depositional basin. Although albitization takes place in a wide variety of geological settings, the
occurrence of gold disseminated albitite in
Loraboue´ (Be´ziat et al., 1998) and in the neigh-bouring Larafella area (Bamba et al., 1997) also hints at an arc-related geological setting. These occurrences are similar to the gold-bearing al-bitized rocks of Alleghany, California (Bo¨hlke, 1989), Bardoc-Kalgoorlie area, Western Australia (Witt, 1992) and Sudbury-Wanapitei Lake area, Ontario (Schandl et al., 1994), respectively inter-preted as coeval with regional arc magmatic activ-ity or associated with a continent-arc collision. In the same way, Bassot (1997) pointed out that intense albitization has affected locally all the members of the Dale´ma calc-alkaline suite (Pale-oproterozoic of Eastern Senegal).
7. Conclusions
In the Boromo greenstone belt, the Loraboue´ terrain consists dominantly of calc-alkaline basalt genetically linked to the ultramafic and mafic cumulate rocks and, more rarely, the basalt, do-lerite and gabbro of tholeiitic affinity. These dif-ferent types of basalt, as well as the dolerite and the isolated massive gabbro, exhibit the classic features of arc magmatism, namely LILE and Pb
enrichment, HFSE depletion and high Ce/Nb and
Th/Nb ratios. The compositional and
mineralogi-cal similarities between plutonic rocks in island arcs and the studied rocks, the close spatial asso-ciation of the ultramafic and mafic cumulate rocks with intermediate-composition plutonic rocks, Fig. 15. Plot of Ce/Nb vs. Ce.: Birimian tholeiitic basalts of
and the thick sequence of basaltic, metafelsic and pyroclastic rocks strongly suggest that the ultra-mafic and ultra-mafic cumulate assemblage of the Loraboue´ area could represent a magma chamber that crystallized in the lower crust of an island arc. Isolated doleritic and gabbroic bodies with tholeiitic affinity probably represent shallower lev-els in the same arc which were juxtaposed against the volcanic pile during either accretion or tec-tonic events. These data as a whole imply that the primary magma could have been generated through relatively high degrees of partial melting of a metasomatically enriched lherzolite source.
Therefore, the Loraboue´ mafic-ultramafic as-semblage supports the model that, during the Eburnean orogeny, the continental crust grew by accretion of island arcs. Such a geodynamic envi-ronment has been previously suggested for the emplacement of Proterozoic metavolcanics (e.g. Condie, 1989) and, with regard to the West African craton, for Birimian greenstone belts lo-cated in Burkina Faso (Zonou, 1987), Ghana (Sylvester and Attoh, 1992; Davis et al., 1994), Senegal (Dia, 1988) and Niger (Ama-Salah et al., 1996). However, some fragments of oceanic plateau crust have been identified as accreted members of greenstone belt in the West African craton (Abouchami et al., 1990; Boher et al., 1992; Pouclet et al., 1996). The recognition of two tectonic settings suggests that the Birimian crust was heterogeneous and generated by both vol-canic arc accretion and oceanic plateau accretion. Such a hypothesis is supported by data from some modern igneous province, such as the Solomon arc along the southern margin of the Ontong-Java Plateau (e.g. Tejada et al. 1996) and various arcs on the Caribbean Plateau (e.g. White et al., 1999) in which we note that plateaus and arcs occur juxtaposed.
Acknowledgements
This study was supported by the Campus re-search contract ‘Cartographie ge´ologique ap-plique´e a` la recherche minie`re au Burkina Faso.’ The authors gratefully acknowledge SOREMIB for access to prospect properties and C. Halpin
(BILLITON Company) for ongoing field studies in the Loraboue´ prospect. The English style was corrected by M.S.N. Carpenter. Comments by W. Hirdes and J. Tarney improved the manuscript.
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