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Home All issues Volume 194 (2023) BSGF - Earth Sci. Bull., 194 (2023) 11 Full HTML
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BSGF - Earth Sci. Bull.
Volume 194, 2023
Article Number 11
Number of page(s) 33
DOI https://doi.org/10.1051/bsgf/2023007
Published online 04 August 2023

© A.Y. Koffi et al., Published by EDP Sciences 2023

Licence Creative CommonsThis is an Open Access article distributed under the terms of the Creative Commons Attribution License (https://creativecommons.org/licenses/by/4.0), which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.

1 Introduction

Tectonic regimes on Earth have undergone a secular change in style mainly due to the cooling of the planet during the Archean-Paleoproterozoic period (3.0–2.0 Ga). The tectonic processes operating on Early Earth are vividly debated in the literature and it remains an open question when plate tectonic processes such as Earth-wide subduction and orogenesis began (Kranendonk, 2010; Condie, 2018; Palin et al., 2020; Brown et al., 2020). Many authors place the transition from the so-called “stagnant-lid” to some kind of plate tectonics in the Meso-Archean period (Cawood et al., 2013; Palin et al., 2020) but some propose that it occurred early in the Earth’s history, at ca. 4 Ga (Maruyama et al., 2018) or, much later, at ca. 1 Ga (Stern, 2005; Hamilton, 2011). The nature of the transition from stagnant-lid to mobile-lid tectonics is also questioned (Palin et al., 2020), in particular whether the crustal record is representative of lithospheric scale processes or of crustal dynamics (Vanderhaeghe et al., 2019).

The Paleoproterozoic Eburnean orogeny (2.25–1.9 Ga) in the West African Craton (WAC) exemplifies this controversy. Some authors advocate an “archaic”, “hot-lithosphere” style dominated by vertical movements and the emplacement of granitic plutons separated by greenstone belts (Condie, 1994; de Wit, 1998; Gapais et al., 2005, 2008; Bédard, 2006; Vidal et al., 2009). Others argue that Paleoproterozoic tectonics shares many similarities with “modern”, “cold-lithosphere” plate tectonics characterized by tangential movements leading to crustal thickening and magmatism (Feybesse and Milési, 1994; Ledru et al., 1994). A critical point for the understanding of the tectonic regime prevailing during the Paleoproterozoic Eburnean orogeny in the WAC is the lack of data on the metamorphic conditions, their evolution, and relative and absolute timing. Metamorphic petrology combined with geochronology provides a record of the secular cooling of the Earth in Archean and Paleoproterozoic domains exposing dominantly high-temperature metamorphic rocks (Brown, 2014). The coexistence of high-pressure eclogite- to granulite-facies metamorphism recording gradients of 15 to 25 °C/km and (ultra) high-temperature granulite-facies metamorphism with gradients of 30 to 40 °C/km in Paleoproterozoic belts appears to characterize the so-called “modern” geodynamic context (Brown, 2014).

The genesis of high-pressure (HP) metamorphic rocks in Archean and Paleoproterozoic domains and the mechanisms leading to their exhumation are debated in different parts of the WAC (western Côte d’Ivoire: Pitra et al., 2010; eastern Burkina Faso: Ganne et al., 2012; NW Ghana: Block et al., 2015; SW Ghana: McFarlane et al., 2019). In the Kouibli zone in the West of Côte d’Ivoire, maximum metamorphic conditions preserved at the contact zone between the Archean and Paleoproterozoic domains are in the granulite facies ranging from ∼ 14 kbar, 850 °C to ∼ 7 kbar, 700–800 °C (Triboulet and Feybesse, 1998; Pitra et al., 2010) and dated between 2050 and 2030 Ma (Th–U–total Pb monazite microprobe ages: Cocherie et al., 1998; Sm–Nd garnet ages: Kouamelan et al., 1997). Archean ages ranging from ca. 2800 to 2700 Ma were documented in monazite cores from the same samples by both evaporation and microprobe Th–U–Pb methods (Cocherie et al., 1998).

The Sassandra-Cavally (SASCA) domain (named after the two rivers that run across the area) represents a contact zone between the Archean and Paleoproterozoic domains of the West African Craton in the SW Côte d’Ivoire (Kouamelan, 1996; Kouamelan et al., 1997, 2015). Neither P–T estimations nor metamorphic ages exist in the SASCA domain where the last mapping was done by Papon (1973). The only LA-ICP-MS U–Pb zircon ages are from San Pedro tonalitic gneiss (SA-3) and yield ages of 3207 ± 7 Ma for magmatic zircon grains and an upper intercept of 3155 ± 17 Ma on some magmatic zircon rims and rounded grains (Kouamelan et al., 2015). Younger ages were not found in the studied sample, leading the authors to a conclusion that the gneiss was unaffected by the Eburnean orogeny.

In order to decipher the P–T conditions and timing of deformation of the SASCA domain, we apply a multimethod approach combining field geology, (micro)structural analysis, metamorphic petrology and LA-ICP-MS U–Pb geochronology on monazite. These data are used to discuss the thermal and tectonic regime and timing of the Eburnean reworking of the Archean crust.

2 Geological setting of the West African Craton

Precambrian domain in the West African Craton crop out in four regions of unequal sizes: the Reguibat Shield in the North, the Leo-Man Shield in the South (Berger et al., 2013; Jessell et al., 2015; Thiéblemont, 2016) and the significantly smaller Paleoproterozoic Kédougou-Kéniéba and Kayes inliers between the two shields (Ledru et al., 1991; Kone et al., 2020; Labou et al., 2020; Diallo et al., 2021) (Fig. 1). The Leo-Man and Reguibat Shields consist of Paleoproterozoic formations in the East and Archean formations in the West separated by transcurrent shear zones (Bessoles, 1977; Camil, 1984; Auvray et al., 1992; Kouamelan, 1996; Potrel et al., 1996; Kouamelan et al., 1997; Thiéblemont et al., 2001, 2004; Peucat et al., 2005; Pitra et al., 2010). Three major orogenic episodes mark the history of the WAC: the Leonian (3.4–3.0 Ga), the Liberian (2.9–2.7 Ga) and the Eburnean (2.25–1.9 Ga) after which the whole craton was stabilized (Thiéblemont et al., 2004; Liégeois et al., 2005; Begg et al., 2009; De Waele et al., 2015; Wane et al., 2018). In the Leo-Man Shield, the Archean domain is called Kenema-Man while the Paleoproterozoic domain to the east is known as the Baoule-Mossi domain (Bessoles, 1977). The Archean Kenema-Man domain consists of two main complexes: a basal complex consisting of migmatites and granulitic gneisses unconformably overlain by a sequence of supracrustal rocks deformed and metamorphosed at 2800 ± 20 Ma (Camil, 1981, 1984; Rollinson and Cliff, 1982; Martin, 1994; Kouamelan, 1996; Rollinson, 2016). The basal complex is intruded by various plutonic rocks including granite, granodiorite, charnockite, as well as basic and ultrabasic rocks dated around 2.8 Ga (Rollinson and Cliff, 1982; Kouamelan et al., 1997; Thiéblemont et al., 2001; Rollinson, 2016). The supracrustal rocks reached medium- to high-pressure granulite facies during the Liberian (∼ 2700 Ma) orogeny (Williams, 1988).

The Paleoproterozoic Baoule-Mossi domain consists of belts of metavolcanic (metarhyolite, metadacite, metaandesite, metabasalt, etc.) and metasedimentary (metapelite, metasandstone, metaconglomerate, quartzite etc.) rocks juxtaposed to granitoid batholiths (Leube et al., 1990; Doumbia et al., 1998; Gasquet et al., 2003; Pouclet et al., 2006; Lompo, 2009; Baratoux et al., 2011). The emplacement and the subsequent reworking of all these lithologies occurred from ca. 2.25 to 2.06 Ga (Parra-Avila et al., 2017). The structural grain of the Baoule-Mossi domain is characterized by a subvertical penetrative foliation, associated with isoclinal folds and locally affected by more open folds and by regional-scale shear zones. This structural framework is interpreted by many authors as a result of horizontal shortening leading to crustal thickening followed by transcurrent tectonics (Lemoine et al., 1985; Lemoine, 1988; Milési et al., 1989, 1992; Feybesse et al., 1990, 2006; Ledru et al., 1991; Boher et al., 1992; Fabre, 1995; Vidal et al., 1996, 2009; Bourges et al., 1998; Allibone et al., 2002; Lompo, 2010; Baratoux et al., 2011; Jessell et al., 2012; Block et al., 2016). Some works suggest an “archaic” style of deformation for the Baoule-Mossi domain, dominated by gravity-driven vertical motions in a hotter lithosphere compared to post-Paleoproterozoic period (Delor et al., 1995; Pouclet et al., 1996; Vidal et al., 1996, 2009; Caby et al., 2000). The “archaic” style is characterized by the development of “dome and basin” structures with isobaric metamorphism in the contact aureoles of the granitoids attributed to the development of gravitational instabilities caused by the presence of dense supracrustal mafic volcanic rocks over buoyant granitoid domains (Pons et al., 1995; Pouclet et al., 1996; Caby et al., 2000; Vidal et al., 2009; Lompo, 2010; Ganne et al., 2014). In this model, upwelling of migmatitic rocks to the surface is buoyancy-driven in a hot and rheologically weak orogenic crust and is associated with crustal thickening (Caby et al., 2000; Ganne et al., 2014). As an alternative model, Chardon et al. (2020) interpret the craton-scale shear zone pattern as a result of distributed shortening and lateral flow of the hot lithosphere leading to an anastomosed transpressional shear zone network. The Eburnean orogeny is also marked by partial melting and melt segregation, which leads to the petrogenesis of a variety of pegmatites, some of which are enriched in REE (Bonzi et al., 2021).

Based on the cold apparent geothermal gradients in the volcano-sedimentary rocks from Burkina Faso, Ganne et al. (2012) suggested that the Eburnean orogenic belt in the Baoule-Mossi domain formed owing to modern-style tectonics marked by subduction. In the Baoule-Mossi domain, greenschist-facies conditions are predominant in the greenstone belts. However, recent work of Block et al. (2015) and McFarlane (2018) have shown large domains of Barrovian-type high-grade metamorphism attributed to continental collision with the peak of metamorphism spanning between 2138 ± 7 Ma and 2127 ± 7 Ma (Block et al., 2015) and 2073 ± 2 Ma (McFarlane et al., 2019) (U–Pb on monazite). In addition, migmatites were reported by Opare-Addo et al. (1993) from southern Ghana. At other places, amphibolite-facies metamorphism is interpreted as contact metamorphism around magmatic intrusions at 1973 ± 33 Ma (Ndiaye et al., 1989; Debat et al., 2003). Medium- to high-pressure granulite-facies rocks characterize the boundary zone between the Archean and Paleoproterozoic domains in the Côte d’Ivoire (Triboulet and Feybesse, 1998; Caby et al., 2000; Pitra et al., 2010) and suggest a Paleoproterozoic granulite-facies metamorphic overprint at 2053 ± 15 Ma and 2031 ± 13 Ma (Sm–Nd garnet, Kouamelan et al., 1997; Pitra et al., 2010). Also, Paleoproterozoic ages of 2041 ± 19 Ma and 2037 ± 14 Ma were obtained in high-grade metamorphic rocks in southeastern Guinea (U–Pb SHRIMP on zircon; Thiéblemont et al., 2004). The ages obtained in the Damang gold deposit (Ghana) suggest a timing of mineralization around 2005 ± 26 Ma (U–Th–Pb EMPA on monazite). The cooling ages obtained by the 40Ar/39Ar method on biotite range from 1980 ± 9 to 1898 ± 11 Ma (White et al., 2014).

In the Kedougou-Kéniéba Inlier, U–Pb monazite metamorphic ages from the Saraya batholith in Senegal yield 2064 ± 4 Ma (Hirdes and Davis, 2002) and a maximum age for the deposition of the sedimentary protolith of the Dialé-Daléma series is given by U–Pb on detrital zircon grains at 2120–2110 Ma (Kone et al., 2020).

thumbnail Fig. 1

Geological sketch map of the West African Craton (modified after the BRGM SIG Africa map and Ennih and Liégeois, 2008; Berger et al., 2013; Thiéblemont, 2016). The position of the present-day margins of the craton is constrained by geophysics (Jessell et al., 2015).

3 Geology of the SASCA domain

The SASCA domain is mainly built of migmatitic grey gneisses commonly intercalated with amphibolites and metasediments, also locally migmatitic (Fig. 2). Metasedimentary rocks occur as staurolite-bearing micaschists and banded paragneisses. In the southernmost part (region of Tabou; Fig. 2), diopside- and hypersthene-bearing metabasic granulitic gneisses predominate. These lithological units are intruded by granites, granodiorites, syenites, as well as associated aplites and pegmatites (Papon, 1973), and crosscut by large ductile structures such as the Greenville-Ferkessédougou-Bobo-Dioulasso shear zone (Lemoine et al., 1985; Baratoux et al., 2011).

Geochronological data on the migmatitic grey gneisses of San Pedro and Monogaga (samples SP3B, SP11 and MON4 in Fig. 2) revealed the presence of Archean crustal relics, with a majority of zircon U–Pb ages ranging from ca. 2800 to 3330 Ma (monozircon evaporation: Kouamelan et al., 1997; LA-ICP-MS: Kouamelan et al., 2015; LA-(MC)-ICP-MS: Koffi et al., 2022). Sm–Nd model age of the magmatic protoliths of the migmatitic grey gneiss in the SASCA is 3.46 Ga (Sample SA-3, Kouamelan, 1996; Kouamelan et al., 1997) and εHf ranges from −9.4 to +3.3 (Koffi et al., 2022). A potassic granite intrusion at San Pedro (sample CIS62 in Fig. 2) was dated at 2084 ± 6 Ma and yields a hybrid εHf isotopic signature with clusters between −4.9 and −8.5, and between +2.2 and +6.5 for inherited zircon grains dated between ca. 2343 to 2100 Ma (Koffi et al., 2022). The presence of Archean relics, negative εNd values at 2.1 Ga in the migmatitic grey gneiss, and negative εHf in the San Pedro potassic granite indicate some interaction between the Archean crust and the juvenile magmas formed during the Eburnean orogeny (Kouamelan et al., 1997; Koffi et al., 2022). The age of 2076 ± 6 Ma obtained on zircon rims from migmatitic gneiss from Monogaga (sample SP11 in Fig. 2) was interpreted as reflecting a metamorphic event. A garnet–cordierite–sillimanite migmatitic paragneiss from Grand-Béréby, deposited after 2069 ± 4 Ma (Koffi et al., 2022), shows Sm–Nd model ages of 2.48 and 2.47 Ga confirming the Archean age of the source of the sedimentary protoliths (samples GDB8, KOU7) (Kouadio, 2017; Koffi et al., 2022).

thumbnail Fig. 2

Litho-structural map of the SW of Côte d’Ivoire (modified after Tagini, 1971; Papon, 1973; Pitra et al., 2010; WAXI report, AMIRA Global, 2018). Blue stars show the location of the studied samples. Samples are described from West to East.

4 Methodology

4.1 Petrology

The chemical composition of minerals was determined using a CAMECA SXFive electron microprobe (Raimond Castaing Microcaracterization Center, Toulouse). The analytical conditions were 15 keV, 10 nA for biotite, chlorite, cordierite, white micas and staurolite, 20 nA for garnet and feldspar, and beam spot size < 2 μm. The acquisition times were 10 s on the peak and 5 s on either side of the peak for the continuous background.

P–T conditions were estimated using equilibrium assemblage phase diagrams (pseudosections) calculated in the chemical model system (Mn)NCKFMASHTO with the software Theriak/Domino (de Capitani and Brown, 1987; de Capitani and Petrakakis, 2010) and the internally consistent thermodynamic data set 5.5 (Holland and Powell, 2003). The full system including Mn is useful for examining lower-grade metamorphic conditions, as garnet is significantly stabilized by Mn to lower temperatures (Mahar et al., 1997). Mineral abbreviations follow Holland and Powell (1998). The solution phases considered, and the mixing models used are from Holland and Powell (1998) – chloritoid (ctd), cordierite (cd), epidote (ep), staurolite (st); Holland and Powell (2003) – plagioclase (pl), K-feldspar (ksp); White et al. (2007) – garnet (g), biotite (bi), liquid (liq); Mahar et al. (1997) – chlorite (chl); Coggon and Holland (2002) – muscovite (mu), paragonite (pa); White et al. (2000) – ilmenite (ilm), hematite (hem) and White and Powell (2002) – magnetite (mt), orthopyroxene (opx). Albite (ab), andalusite (and), kyanite (ky), quartz (q), rutile (ru), sillimanite (sill) and titanite (sphene, sph) are considered as pure end members. Bulk rock compositions of the samples were measured from rock powders at Bureau Veritas Commodities Canada Ltd by ICP-AES (Tabs. 1 and 2). The amount of Fe3+ (with respect to total Fe) was estimated using calculated P/T-X(Fe3+) pseudosections and the nature of the oxide (ilmenite, magnetite or hematite) observed in the samples. At temperatures below the solidus, all pseudosections were calculated considering H2O in excess. The amount of H2O for calculating liquid (melt) bearing equilibria was fixed so that the mineral assemblages are just H2O-saturated (i.e., contain a minimal amount, of the order of 1 vol%, of free aqueous fluid) below the solidus. Mineral endmembers (expressed in mole %) and compositional variables (molar ratios) are: XMg = Mg/(Mg + Fe), XNa = Na/(Na + Ca + K), almandine, Alm = Fe/(Fe + Mg + Ca + Mn), pyrope, Prp = Mg/(Fe + Mg + Ca + Mn), grossular, Grs = Ca/(Fe + Mg + Ca + Mn), spessartine, Sps = Mn/(Fe + Mg + Ca + Mn), anorthite, An = Ca/(Ca + Na + K); a.p.f.u. = atoms per formula unit; “→” denotes core-to-rim evolution.

Table 1

Petrographic summary and location in degree (WGS 84 coordinate system) of the samples studied. Quartz, plagioclase, biotite and ilmenite are present in all samples.

Table 2

Chemical compositions (oxide wt%) of the samples used for the P–T pseudosection calculations. Fe2O3 corresponds to total Fe.

4.2 Geochronology

The internal structure of the monazite crystals was characterized by a JEOL JSM 6360 LV microscope equipped with an X-ray detector connected to an EDS (Energy Dispersive Spectrometry) microanalyzer platform at the Laboratory of Environmental Geosciences Toulouse (GET) in France. All monazite crystals were scanned in BSE (Backscattered electron). LA-ICP-MS in situ U–Pb analyses of monazite were carried out directly on the thin sections at Geosciences Montpellier (France). The laser system consists of a Lambda Physik Compex 102 excimer laser coupled to a Thermofinnigan Element XR ICP-MS (AETE-ISO regional facility of the OSU OREME, University of Montpellier). The instrument was tuned for maximum sensitivity and low oxide production (ThO/Th < 1%). Analytical conditions are identical to those reported in previous studies (e.g., Bruguier et al., 2009; Bosch et al., 2011) where ablation experiments were performed under helium, which enhances sensitivity and reduces inter-element fractionation (Günther and Heinrich, 1999). The helium stream and particles ablated from the sample were mixed with Ar before entering the plasma. Laser spot size was 5 µm for monazite and the laser was operated at a repetition rate of 2 Hz using a 12 J/cm2 energy density. Total analysis time was 60 s with the first 15 s used for background measurement, which was subtracted from the sample signal. Before each analysis the surface of the targeted zone was cleaned with 10 pulses using a spot size larger than the size used for U–Pb analysis. Pb/U and Pb/Pb ratios were calibrated against the Manangotry crystal (Poitrasson et al., 2000). This reference material was used to correct the measured ratios for mass discrimination and inter-element fractionation. In the course of this study the Manangotry crystal was measured twice each five unknowns. The 57 analyses yielded a 207Pb/206Pb weighted average of 0.05866 ± 0.00015 (2σ) which corresponds to an age of 555 ± 6 Ma, in good agreement with the reference age for this material (Montel et al., 1996; Horstwood et al., 2003). U–Th–Pb isotopic data were reduced using the Glitter software (Van Achterbergh et al., 2001) and ages were calculated using Isoplot (Ludwig, 2003) and are quoted in the text at the 2σ confidence level. Data points are considered concordant when their concordance degree is equal or higher than 95% (based on the comparison between the 206Pb/238U and 207Pb/206Pb apparent ages). For these analyses, a 207Pb/206Pb weighted mean age was favored. On the contrary, when data points are discordant (concordance < 95%), discordia lines were calculated (Fig. 3).

High spatial resolution mapping at 15 keV and 20 nA for 5–10 s per pixel on the CAMECA SXFive electron microprobe of Archean grey gneisses monazites (SP3B) larger than 100 µm was performed at a resolution of ca. 200 × 120 pixels per grain to visualize their complex chemical zonations. Chemical maps of Ce, Pb, Th, U, and Y were used to guide subsequent quantitative U–Pb analyses by LA-ICP-MS. The monazite data are summarized in Supplementary Materials (Tables S1, S2 and S3).

thumbnail Fig. 3

Concordia plots of U–Pb age dating for reference material Manangotry and Moacyr monazites (a and b respectively).

5 Results

5.1 Structural evolution

Based on the study of the coastal section between Grand-Béréby and Sassandra, three main phases of deformation (D1–D3) are documented in the SASCA region. The first phase D1 is characterized by a NW-SE oriented S1 subvertical penetrative metamorphic foliations in an axial planar position of tight isoclinal folds. The second phase of deformation D2 is marked throughout the studied area by tight to open F2 folds with subvertical NNE-SSW-trending axial planes parallel to the S2 foliation and sinistral ductile shear zones. The L2 lineation is parallel to the fold axes and plunges between 80° and 30° to the NNE. D2 is interpreted as recording continuous NNW-SSE shortening under a transpressive regime. In the San Pedro area, D1 and D2 structures are overprinted by the deformation phase D3 that is expressed as F3 folds and a penetrative subvertical E-W axial planar schistosity S3 associated with E-W-striking dextral shear zones suggesting a WNW-ESE shortening (Fig. 4). Geological cross-section and equal-area, lower hemisphere stereoplots document the structures related to the D2 deformation phase (Figs. 4b and 4c, respectively). The relationship between structural observations and mineral assemblage will be discussed in the petrography and mineral chemistry section.

thumbnail Fig. 4

Detailed geological map with structural relationships. (a) Structural map showing the orientation of S1, S2 and S3 foliations. L2 lineations are also plotted. (b) Geological cross-section showing the D2 deformation phase. (c) Equal-area, lower hemisphere stereoplots show D2 planar (foliation S2) and linear (lineation L2) deformation features.

5.2 Petrography and mineral chemistry

Five samples representative of the different lithological units of the SASCA domain were studied in detail (location and description in Figs. 2 and 5 and Tab. 1). Representative mineral compositions are reported in Tables 3a and 3b.

thumbnail Fig. 5

Field photographs illustrating typical structural relationships (S1, S2 and S3) in this study. (a) S1 and S2 in garnet–cordierite–sillimanite migmatitic paragneiss Grand-Béréby (GDB8). (b) S1 and S2 in migmatitic grey gneiss San Pedro (SP3B). (c) S1 and S2 in staurolite-bearing micaschist San Pedro (SAN). (d) S2 in garnet–sillimanite–cordierite micaschist Madie (MAD1A). (e) S3 in garnet–staurolite-bearing micaschist Kounoukou (KOU7).
Table 3a

Representative microprobe analyses of selected minerals.

Table 3b

Representative microprobe analyses of selected minerals.

5.2.1 Garnet–cordierite–sillimanite migmatitic paragneiss (GDB8)

Macroscopically, the rock displays a prominent migmatitic layering S1 folded by tight to open folds F2 with ca. 30° NE-plunging fold axes and subvertical axial planes, parallel to a less penetrative foliation S2. Coarse-grained quartzo-feldspathic leucosome layers (1–10 cm thick), rimmed by thin layers of biotite-rich melanosome, and alternating with layers of gneissic mesosome define the synmigmatitic foliation S1 (Fig. 5a). The mesosome contains quartz, plagioclase (An23–25), biotite, euhedral prismatic sillimanite (long up to 0.5 mm), cordierite and ilmenite (Fig. 6a). Biotite and sillimanite define the foliation S2 that wraps around anhedral porphyroblasts of garnet (up to 2.5 mm) and locally biotite that preserve S1 inclusion trails. In garnet, S1 is marked by abundant inclusions of fine needles of sillimanite (Fig. 6a). Garnet is commonly surrounded by cordierite (XMg = 0.79–0.85, Na up to 0.02 a.p.f.u.) and locally by symplectitic intergrowths of biotite and quartz. Garnet is chemically zoned with a rimward decrease of pyrope and XMg, and an increase of almandine and spessartine (Alm58→70 Prp34→23 Grs2-3 Sps4→5, XMg = 0.37→0.26; Fig. 7a). Biotite in the matrix and inclusions in cordierite have a constant XMg of ca. 0.60. Inclusions in garnet show a variable XMg (0.62–0.78). All biotite crystals contain 0.23–0.28 a.p.f.u. of Ti and up to 0.22 a.p.f.u. of F.

The inferred peak mineral assemblage comprises garnet, biotite, sillimanite, plagioclase, quartz, ilmenite, and melt. The significantly smaller size of sillimanite inclusions in garnet with respect to the matrix crystals testifies to a prograde growth of garnet, contemporaneous with S1. Garnet zoning is attributed to diffusional reequilibration during partial resorption. Biotite-plagioclase symplectites around garnet suggest that the resorption may reflect reaction with crystallizing melt (e.g., Holness et al., 2011) during cooling. The growth of cordierite at the expense of garnet points to a pressure decrease during D2.

thumbnail Fig. 6

Photomicrographs (plane polarized light). (a) Sample GDB8. Main penetrative foliation S2 (fine-grained biotite + prismatic sillimanite) wraps around (a) a microlithon with preserved foliation S1 marked by coarse-grained biotite and prismatic sillimanite, and anhedral porphyroblasts of garnet with abundant inclusions of fine sillimanite needles. The dashed line A-B indicates the location of the chemical profile of garnet (Fig. 7a). Sample SP3B: (b) String of monazite crystals aligned parallel to the foliation S2 marked by biotite, which also crystallizes in the fractures and at the rims of garnet porphyroblasts containing inclusions of monazite. A-B indicates the location of the chemical profile of garnet (Fig. 7b). Sample SAN: (c) Biotite and rare fibrous sillimanite mark the main foliation (S2) and (d) andalusite replaced partly by chlorite. Sample MAD1A: (e) Foliation S2 marked by biotite and clusters of fibrous sillimanite. (f) Garnet porphyroblast in contact with rare staurolite crystals. The dashed line A-B indicates the location of the chemical profile (Fig. 7c). Sample KOU7: (g) Staurolite contains quartz inclusion trails that mark the foliation S1, oblique to the matrix foliation S2. (h) A-B – Location of the chemical profile (Fig. 7d) across a garnet porphyroblast with inclusions of ilmenite and quartz.
thumbnail Fig. 7

Chemical profiles of garnet (mole %). Location is shown in Figures 6a (GDB8), 6b (SP3B), 6e (MAD1A) and 6g (KOU7).

5.2.2 Migmatitic grey gneiss (SP3B)

Macroscopically, the rock is a grey banded gneiss with a synmigmatitic foliation S1 marked by the alternation of leucosome and melanosome layers (Fig. 5b). This foliation is transposed into a subvertical, N040-striking foliation S2, marked by alternating quartzo-feldspathic and biotite-rich layers, several mm up to 1 cm thick, which correspond to the leucosome and melanosome, respectively. Locally, significantly coarser-grained quartzo-feldspathic leucosomes, up to 10 cm thick, are present with a spacing generally of 20–50 cm. They are in general parallel to the S2 foliation, but are localized also in necks of foliation boudins and shear bands. The sample comprises garnet, biotite, plagioclase (An29–38), quartz and ilmenite. Biotite (XMg = 0.48–0.54, Ti = 0.14 a.p.f.u.; Fig. 6b) is aligned parallel to the layering of the sample, but also commonly crystallizes in the fractures and at the rims of garnet (Fig. 6b). Garnet (up to 2 mm) is anhedral, almandine-rich and displays an unzoned core and markedly zoned rim, characterized by rimward decrease of pyrope and XMg and increase of spessartine (Alm65→68 Prp16→8 Grs6 Sps 11→21, XMg = 0.19→0.11; Fig. 7b). Monazite commonly forms crystal aggregates elongated parallel to the foliation S2 (Fig. 6b), and locally occurs as inclusions in garnet. Numerous grains of zircon are included in biotite, plagioclase and garnet.

The equilibrium mineral assemblage comprises garnet, biotite, plagioclase, quartz and ilmenite, interpreted as contemporaneous with S1. Garnet zoning suggests diffusional reequilibration due to partial resorption of previously chemically homogeneous crystals during D2.

5.2.3 Staurolite-bearing micaschist (SAN)

Macroscopically, the micaschist displays a composite foliation S2 weakly folded by open folds with S3 axial planes (Fig. 5c). Microscopically, S2 is defined by the preferred orientation of biotite (up to 3 mm; XMg = 0.58–0.62; Al = 1.58–1.74 a.p.f.u.; Ti = 0.06–0.09 a.p.f.u.), chlorite (XMg = 0.62–0.65, Al = 2.71–2.86 a.p.f.u.), plagioclase (An21–23) and quartz (Fig. 6c). The foliation contains magnetite and ilmenite and slightly wraps around euhedral to subhedral synkinematic porphyroblasts of staurolite (up to 2.5 mm; XMg = 0.19–0.22, Ti = 0.09–0.12 a.p.f.u., Mn = 0.17–0.24 a.p.f.u.) and andalusite (up to 10 mm), which commonly contain sigmoidal inclusion trails of quartz, continuous with the S2 foliation (Fig. 6c). Andalusite includes ilmenite, magnetite, quartz, biotite and locally anhedral crystals of staurolite. Rare fibrous sillimanite locally occurs parallel to the main foliation.

5.2.4 Garnet–sillimanite–cordierite micaschist (MAD1A)

The micaschist (Fig. 5d) is rich in quartz and plagioclase (An32–34) and displays a foliation S2 marked by the preferred orientation of biotite (XMg = 0.52–0.60; Al = 1.67–1.72 a.p.f.u.; Ti = 0.05–0.10 a.p.f.u.) and abundant fibrous sillimanite (Fig. 6d). Anhedral garnet (up to 2 mm) has a relatively homogeneous core (Alm72–74 Prp15–17 Grs4–5 Sps7, XMg = 0.17–0.19) and a slightly zoned rim characterized by a rimward decrease of pyrope and XMg, and an increase of almandine and spessartine (Alm75 Prp12 Grs5 Sps8, XMg = 0.14; Figs. 6e and 7c). Rare staurolite (XMg = 0.19–0.20) is either included in plagioclase or at contact with garnet (Fig. 6e). Cordierite (XMg = 0.68–0.70; Na = 0.07 a.p.f.u.) is in contact with fibrous sillimanite and biotite. Accessory minerals include tourmaline, ilmenite, zircon and monazite.

The main syntectonic equilibrium mineral assemblage is inferred to comprise cordierite, sillimanite, biotite, plagioclase, quartz and ilmenite, and typifies low-pressure high-temperature metamorphic conditions. Garnet rims may or may not belong to this assemblage. Garnet zoning suggests diffusional reequilibration due to partial resorption of previously chemically homogeneous crystals. Rare staurolite suggests an earlier metamorphic stage of the Barrovian, medium-pressure type.

5.2.5 Garnet–staurolite-bearing micaschist (KOU7)

A macroscopically dark-coloured micaschist from Kounoukou (Fig. 5e) consists of garnet, staurolite, biotite, plagioclase, quartz, ilmenite and tourmaline (Figs. 6f and 6g). The main foliation S2 is marked by the preferred orientation of biotite (XMg = 0.49–0.53), quartz and plagioclase (An29–32), and is weakly folded by open folds with S3 axial planes. The foliation weakly wraps around subhedral porphyroblasts of garnet and staurolite (up to 3 mm). Poikiloblastic staurolite (XMg = 0.16–0.20) contains abundant quartz inclusion trails that mark the foliation S1, oblique to the matrix foliation S2. The chemical zoning of garnet is characterized by a rimward decrease of pyrope, XMg and spessartine, and an increase of almandine, whereas the proportion of grossular remains constant (Alm74→81, Prp13→11, Grs4, Sps8→4, XMg = 0.16→0.12; Fig. 7d). Garnet chemical zoning is interpreted to reflect prograde growth.

5.3 Metamorphic conditions and P–T paths

5.3.1 Garnet–cordierite–sillimanite migmatitic paragneiss (GDB8)

In the P–T pseudosection (Fig. 8a), the inferred peak mineral assemblage garnet, biotite, sillimanite, plagioclase, quartz, ilmenite, and melt is stable at pressures higher than ca. 7 kbar (limited at low pressure by the appearance of cordierite) and temperatures in excess of 690 °C (limited by the solidus at the low-temperature side). The upper P limit is given by the sillimanite–kyanite transition. The increase of the sillimanite grain size from the garnet core to the matrix testifies to crystallisation during an increase in temperature. The constant amount of grossular indicates that the prograde P–T evolution was parallel to the position of the grossular isopleths. The composition of the garnet core (XMg(g) = 0.34; Grs2–3), most susceptible to represent the peak metamorphic composition, compared to the position of the calculated compositional isopleths of garnet, suggest peak metamorphic pressure and temperature at ca. 10 kbar, 820 °C. The subsequent crystallisation of cordierite at the expense of garnet, compatible with the observed rimward decrease of the garnet XMg (0.37→0.26), suggests progressive reequilibration during a pressure decrease down to ca. 6 kbar, probably followed by cooling below the solidus. The homogeneous composition of matrix biotite (XMg = 0.60) suggests subisobaric cooling, parallel to the corresponding isopleths in the field g–bi–cd–ilm–liq–sill (+q + pl).

thumbnail Fig. 8

P–T pseudosections in the model system NCKFMASHTO (whole-rock compositions expressed in mole %) and compositional isopleths for (a) the migmatitic paragneiss GDB8 and (b) the migmatitic grey gneiss SP3B. Red arrow and red star show the inferred P–T path (GDB8) and the equilibration P–T conditions (SP3B), respectively. Field limits are marked by the phase that appears/disappears. See text for details.

5.3.2 Migmatitic grey gneiss (SP3B)

In the P–T pseudosection (Fig. 8b), the mineral assemblage of the migmatitic grey gneiss SP3B, g-ilm-q-pl-bi with or without melt, is stable at T > 620 °C and P > 5 kbar, limited by the appearance of staurolite towards low temperatures and cordierite towards low pressures. Comparing the calculated compositional isopleths with the observed composition of the homogeneous garnet core (XMg(g) = 0.19, Grs6), inferred to represent the composition equilibrated at the metamorphic peak, the equilibration conditions can be constrained at ca. 8–9 kbar, 650–700 °C. The prograde P–T path could not have been established due to the lack of mineral inclusions or any other relics of prograde assemblages. The chemical zoning of the garnet rim is compatible with decompression and/or cooling, but does not allow precise estimation of the retrograde P–T path.

5.3.3 Staurolite-bearing micaschist (SAN)

The syn-S2 assemblage of the micaschist is stable in the field st–and–pl–bi–chl–ilm–mt (+q + H2O at 560–580 °C, 3.5–4 kbar, limited by the disappearance of chlorite towards high temperatures, the disappearance of staurolite towards low pressures and the appearance of paragonite towards low temperatures and high pressures (Fig. 9a). The observation of rare sillimanite needles in the sample suggests that the rocks would have passed through the stability field of sillimanite before, possibly due to decompression from a higher-pressure syn-S1 peak.

thumbnail Fig. 9

P–T pseudosections and compositional isopleths calculated for the samples (a) SAN in the model system NCKFMASHTO, and (b) MAD1A in the model system MnNCKFMASHTO. Red arrows show the inferred P–T paths. See text for details.

5.3.4 Garnet–sillimanite–cordierite micaschist (MAD1A)

In the pseudosection, calculated in the model system MnNCKFMASHTO (Fig. 9b), the dominant syn-S2 assemblage pl-g-ilm-cd-sill (+q + bi + H2O) corresponds to a field at 5–6 kbar, 620–690 °C. If garnet rims were not in equilibrium with the other minerals, the syn-S2 assemblage cd-sill-pl-ilm would indicate lower pressures and temperatures (dashed arrow in Fig. 9b). The former stability of staurolite with garnet suggests higher pressures (5–9 kbar). The isopleth for grossular (Grs5) points to a pressure range of 7–8 kbar, for temperatures of ∼ 620–650 °C. Other compositional isopleths are not very helpful in constraining better the equilibration P–T conditions. Although the values globally fit with the observations in the inferred stability fields, the range of the observed mineral compositions is larger than those fields. Garnet rim zoning (Alm72→75 Prp16→12 Sps7→8, XMg = 0.18→0.14) is compatible with a P–T evolution dominated by a pressure decrease.

5.3.5 Garnet–staurolite-bearing micaschist (KOU7)

In the pseudosection calculated in the model system MnNCKFMASHTO (Fig. 10), the observed syn-S2 assemblage pl-g-ilm-st (+q + bi + H2O) is stable at 5–8 kbar, 600–650 °C. The intersection of the calculated compositional isopleths suggests that the garnet core (Alm74 Grs4 Sps8) started the crystallisation in the neighbouring lower-temperature chlorite-bearing field at ca. 5.8 kbar and 585 °C. The progressive rimward decrease of the proportion of spessartine, while the proportion of grossular remains constant, suggests a prograde up-pressure up-temperature evolution, sub-parallel to the isopleth of the grossular content. It peaks in the g–st–pl–ilm field where the isopleths corresponding to the composition of the garnet rim (Grs4, Sps4) intersect at ca. 6.6 kbar, 620 °C.

thumbnail Fig. 10

P–T pseudosection (MnNCKFMASHTO) and compositional isopleths calculated from the sample KOU7 (whole-rock composition expressed in mole %). Red arrow shows the inferred P–T paths. See text for details.

5.4 U–Pb in situ dating of monazite

5.4.1 Garnet–cordierite–sillimanite migmatitic paragneiss (GDB8)

In GDB8, monazite crystals (ca. 100 μm) are in contact with or included in crystals of biotite and plagioclase (Fig. 11a). Fifteen analyses were performed on eleven monazite crystals, among which eleven are close to concordant (concordance > 95%) and cluster into two groups with distinct ages. The main group of concordant analyses provides a 207Pb/206Pb weighted mean age of 2040 ± 6 Ma (MSWD = 0.1, N = 8). Three other analyses are significantly younger and yield a 207Pb/206Pb weighted mean age of 2003 ± 12 Ma (MSWD = 0.2, N = 3) (Fig. 12c).

thumbnail Fig. 11

(a) Position of the monazite crystals (bright spots) in the thin section GDB8. (b) Alignment of monazite clusters parallel to S2 in migmatitic grey gneiss SP3B. (c–g) Migmatitic grey gneiss SP3B. (c) Back-scattered electron (BSE) images showing the position of the monazite crystals (bright spots) with respect to the garnet porphyroblasts. (d, e) Microprobe X-ray maps and (f, g) BSE images with analytical spots and associated 207Pb/206Pb dates of monazite grain marked in (a). LC: low concentration; HC: high concentration.
thumbnail Fig. 11

(Continued)
thumbnail Fig. 12

U–Pb dating of monazites by LA-ICP-MS. (a) Concordia diagram showing all analytical points for monazites from the migmatitic grey gneiss SP3B. (b) Zoom of the concordia diagram showing the two age groups at ca. 2000 and 2037 Ma in the sample SP3B. (c–f) U–Pb concordia diagrams for monazites from garnet–cordierite–sillimanite migmatitic paragneiss GDB8 (c) and metapelites SAN, MAD1A and KOU7 (d, e and f respectively). SAN: staurolite-bearing micaschist; MAD1A: garnet–sillimanite–cordierite micaschist; KOU7: garnet–staurolite-bearing micaschist.

5.4.2 Migmatitic grey gneiss (SP3B)

Monazite in SP3B has a variable size, ranging from 10 to 100 µm. It commonly forms clusters aligned parallel to S2 (Fig. 11b). Under BSE imaging, some grains reveal contrasting areas showing complex zonations (Figs. 11f and 11g). X-ray maps were performed on three grains (≥ 50 μm) in order to assess the internal distribution of the chemical elements and target U–Pb analyses. Mapped crystals have distinct zones with a heterogeneous distribution of lead (Pb), thorium (Th), uranium (U) and yttrium (Y) (Figs. 11d and 11e). These irregular intra-grain zonations are interpreted as chemical fronts and suggest a redistribution of the elements at the subgrain scale. Eighty-one analyses were carried out on twenty-eight grains located in different textural positions in the thin section, included in either biotite, plagioclase or garnet (Table S2, see Supplementary Materials). Overall, analyses are distributed in two groups with sub-concordant to concordant data points (concordance ≥ 95%) at around 2040 and 2000 Ma, except for three analyses from the central part of one large grain in contact with garnet (Figs. 11f, 11g, 12a and 12b), which yield either discordant Archean dates (minimum dates at 2626 ± 19 Ma and 2527 ± 20 Ma) or a concordant date at 2439 ± 16 Ma (Fig. 12a). The main group of 46 analyses provides a 207Pb/206Pb weighted mean age of 2037 ± 2 Ma (MSWD = 0.79; N = 46) whereas the second group yields a distinctly younger age of 2003 ± 3 Ma (MSWD = 0.45; N = 28). It is noteworthy that the 2037 ± 2 Ma-old group corresponds to monazites that are aligned parallel to the S2 foliation (Fig. 11b).

5.4.3 Staurolite-bearing micaschist (SAN)

Monazite is small (ca. 10 μm) and either included in andalusite and biotite, or located along grain boundaries. Eighteen analyses have been performed on twelve monazite grains. All analyses cluster close to Concordia (≥ 95% concordance) and define three groups of ages yielding 207Pb/206Pb weighted mean ages of 1978 ± 7 Ma (MSWD = 0.3, N = 5), 1949 ± 5 Ma (MSWD = 0.2, N = 10) and 1913 ± 10 Ma (MSWD = 0.1, N = 3) (Fig. 12d).

5.4.4 Garnet–sillimanite–cordierite micaschist (MAD1A)

Monazite forms small crystals (ca. 10 μm) with no particular zonation. They are found both included within grains and at grain boundaries of the main rock-forming minerals (quartz, biotite, plagioclase, garnet). Nine analyses were performed on six monazite grains. Among the dataset, 6 analyses define a discordia line (MSWD = 0.05; N = 6) with an upper intercept at 2005 ± 31 Ma and a lower intercept close to zero (530 ± 500 Ma). Other three analyses plot slightly to the right of this discordia and provide dates in the range 2032–2034 Ma. The concordance degree of ca. 97% allows to calculate a 207Pb/206Pb weighted mean age of 2034 ± 18 Ma (MSWD = 0.02) (Fig. 12e).

5.4.5 Garnet–staurolite-bearing micaschist (KOU7)

Monazite occurs as small crystals (ca. 10 μm) with no particular zoning. Similarly to sample MAD1A, they are found both included within grains and at grain boundaries of the main rock-forming minerals (quartz, biotite, plagioclase, garnet, staurolite). Eleven analyses were performed on ten grains. Ten analyses are aligned along a discordia that displays an upper intercept age of 2004 ± 16 Ma (MSWD = 0.17), which is the same within errors as the 207Pb/206Pb weighted mean age of the nine concordant to slightly discordant analyses (2002 ± 10 Ma, MSWD = 0.04; N = 9). The latter is adopted as our best estimate for the age of monazite in the micaschist KOU7. One analysis is significantly younger and plots concordantly at 1948 ± 33 Ma (Table S2, see Supplementary Materials, Fig. 12f).

6 Discussion

6.1 Tectono-metamorphic synthesis

The thermobarometric results from the SASCA domain suggest that peak metamorphic conditions correspond to granulite facies at the lowest structural level exposed (GDB8 and SP3B) and to amphibolite facies at the higher structural level (SAN, MAD1A and KOU7) (Figs. 13 and 14). The partially melted garnet–cordierite–sillimanite migmatitic paragneiss of Grand-Béréby (GDB8) records a prograde metamorphism that peaks in the granulite facies (∼ 10 kbar, 820 °C), followed by decompression to ∼ 6 kbar, and cooling. The migmatitic grey gneiss of San Pedro (SP3B) recorded a slightly lower-temperature granulite-facies metamorphic peak at ∼ 8–9 kbar, 650–700 °C. The evolution of the samples SAN, MAD1A and KOU7 is in the amphibolite facies. The garnet–staurolite-bearing micaschist of Kounoukou shows a prograde evolution with a metamorphic peak at ∼ 6.6 kbar, 620 °C. The P–T evolution of the garnet–sillimanite–cordierite micaschist of Madie (MAD1A) is dominated by decompression from ∼ 620–650 °C, 7–8 kbar to 620–690 °C, 5–6 kbar. The staurolite-bearing micaschist of San Pedro (SAN) equilibrated at ∼ 570 °C, ∼ 4 kbar possibly following decompression.

The prograde evolution identified in the garnet–cordierite–sillimanite migmatitic paragneiss of Grand-Béréby (GDB8) corresponds to a burial down to ca. 38 km depth along an apparent geothermal gradient of 25–30 °C/km. The prograde evolution of the garnet–staurolite-bearing micaschist of Kounoukou points to a burial down to ca. 20 km along a similar apparent geothermal gradient of 25 °C/km. Following decompression, most P–T paths align along an apparent geothermal gradient of 40 °C/km. Such metamorphic gradients and clockwise P–T paths are typical of Barrovian-type metamorphism characterizing collisional tectonics (Miyashiro, 1961) and thickening of the Earth’s crust as already proposed for the Eburnean orogenic belt (Milési et al., 1992; Allibone et al., 2002; Feybesse et al., 2006). However, the samples presented in this paper indicate that the SASCA domain is marked by a juxtaposition of rocks with contrasting peak P–T conditions that do not depict a single regular metamorphic gradient (Fig. 13). The contrasting P–T conditions and the structural record of the SASCA domain are best explained by syncompressional crustal-scale folding, and vertical extrusion of lower-crustal material, such a proposed for example for the European Variscan belt (e.g., Schulmann et al., 2008; Štípská et al., 2012; Krýza et al., 2022). On the other hand, the moderately plunging L2 lineation, and axes of F2 folds, with steep axial planar surfaces parallel to regional shear zones, are compatible with a transpressional regime marked by lateral flow (e.g., Lin et al., 1998) typical for hot orogens (e.g., Vanderhaeghe and Teyssier, 2001; Cagnard et al., 2006, 2007; Vanderhaeghe et al., 2020). Alternatively, the structural and metamorphic pattern of the SASCA domain might also result from the development of convective and/or diapiric gravitational instabilities as proposed for the Aegean domain (e.g., Vanderhaeghe et al., 2018; Louis-Napoléon et al., 2020, 2021).

thumbnail Fig. 13

Map of the SASCA domain, SW of Côte d’Ivoire, modified after Papon (1973), showing the position and the monazite age of the different samples studied (blue stars). Granulite and amphibolite facies represent the peak conditions of the corresponding areas. Migmatitic gneisses of Monogaga (MON4 and SP11) and potassic granite of San Pedro (CIS62) are from Koffi et al. (2022).
thumbnail Fig. 14

Synthesis of the P–T conditions of the different rocks studied as well as their P–T path according to metamorphic facies and geothermal gradients (modified after Ernst and Liou, 2008). Abbreviations of metamorphic facies: AM: amphibolite; Amp-EC: amphibolite-eclogite; BS: blueschist; EA: epidote amphibolite; GR: sillimanite-granulite; GS: greenschist; HGR: kyanite-granulite.

6.2 Chronology of tectono-metamorphic events

The ages obtained on the monazites of different samples reveal four groups: (1) oldest ages of ca. 2626, 2527 and 2439 Ma only obtained on one monazite grain from the migmatitic grey gneiss of San Pedro (SP3B); (2) ages of ca. 2037 Ma (MAD1A, SP3B, GDB8); (3) ages of ca. 2000 Ma (SP3B, GDB8, MAD1A, KOU7) and (4) very young ages of ca. 1978, 1949 and 1913 Ma (SAN, KOU7) (Figs. 12 and 13).

The ages of 2626, 2527 and 2439 Ma are attributed to the preservation of an Archean/Early Paleoproterozoic component in the protolith of the migmatitic grey gneiss of San Pedro (SP3B). These results alone do not allow to identify whether this component corresponds to an inherited grain or if it was part of an Archean relictual magmatic-metamorphic mineral assemblage. We note, however, that zircons from the same sample (SP3B) yield only Archean ages, and no Paleoproterozoic component was detected in the zircon population (Koffi et al., 2022). The maximum temperature reached by the sample SP3B (∼ 700 °C) is below the nominal closure temperature for Pb in monazite (Copeland et al., 1988; Cherniak et al., 2004) and preservation of relict monazite with ages older than the Paleoproterozoic is thus not surprising. This is also in agreement with numerous studies that have demonstrated the robustness and complexity of this mineral, which can survive high grade metamorphic conditions in excess of 800 °C (e.g., Rubatto et al., 2001; Bosch et al., 2002; Kelsey et al., 2008; Kelsey and Powell, 2011; Yakymchuk and Brown, 2014; Guergouz et al., 2018). The ages of ca. 2037 Ma are interpreted as dating the Paleoproterozoic metamorphism across the region, in agreement with the position of the monazite grains in the penetrative foliation S2 (Fig. 11b). The association of this foliation with decompression implies that this age dates the exhumation period of the rocks in the SASCA domain.

The ages of ca. 2000 Ma and younger (ca. 1978 to 1913 Ma) are generally found in monazite inclusions in andalusite, which suggests that they could be associated with decompression and cooling (Fig. 9a). These ages could result from two types of processes: (1) (partial) resetting of older crystals by either long-term cooling with periodic opening of the U–Pb system (McFarlane et al., 2006; Kirkland et al., 2016) or isotopic disturbance by fluids (Teufel and Heinrich, 1997; Schmidt et al., 2007; Bosse et al., 2009; Williams et al., 2011; Seydoux-Guillaume et al., 2012; Kirkland et al., 2016); (2) episodic crystallisation of monazite during the cooling period (Kirkland et al., 2016; Guergouz et al., 2018).

6.3 Regional correlations

Similar clockwise P–T paths with peak-pressure conditions corresponding to gradients of ∼ 20–30 °C/km were identified elsewhere in the West African Craton (Pitra et al., 2010; Ganne et al., 2012; Block et al., 2015; McFarlane, 2018; Kone, 2020). In the Kouibli zone of the Kenema-Man domain (Fig. 2), this P–T path is associated with similar ages of 2030–2050 Ma (Sm–Nd garnet, U–Pb monazite; Kouamelan et al., 1997, 2018; Cocherie et al., 1998). Similar ages (2041 ± 19 Ma and 2037 ± 14 Ma, U–Pb zircon) were also reported from southeastern Guinea and interpreted as a late high-temperature metamorphic event (Thiéblemont et al., 2004). In the Kedougou Kenieba Inlier of eastern Senegal, two metamorphic phases were dated at ca. 2085 and 2050 Ma, respectively (Kone, 2020; U–Pb monazite and Sm–Nd garnet).

Elsewhere, however, the associated U–Pb monazite ages are different. They are of ca. 2135 Ma and ca. 2073 Ma in northern and southern Ghana, respectively (Block et al., 2015; McFarlane et al., 2019). In the Damang mine in Ghana, the minimum age for the onset of the exhumation is 2005 ± 26 Ma according to U–Th–total Pb EPMA analyses of metamorphic monazite grains (White et al., 2014). In the same study area, ages ranging from 1980 ± 9 Ma to 1898 ± 11 Ma were obtained by the 40Ar/39Ar biotite method and attributed to a protracted period of slow cooling (White et al., 2014).

The youngest U–Pb monazite ages obtained in this study (ca. 2000, 1978–1913 Ma) were not previously reported in the literature of the West African Craton. In the study area, they are localized in some samples only (GDB8, SAN, KOU7 in Fig. 15). At the same time, they are the only ages found in one of the samples (SAN), where older ages (ca. 2037 Ma) are completely absent. This suggests a dissolution-precipitation of monazite in connection with fluid circulations associated with late reactivation of some of the transcurrent shear zones, consistent with the localized distribution of this overprint.

In South America, metamorphic monazite ages obtained by several methods (U–Pb TIMS, LA-SF-ICPMS, EPMA, SHRIMP, TIMS and Pb–Pb TIMS) show that these Paleoproterozoic terrains were affected by a long metamorphic period in the range from ca. 2100 to 1920 Ma (Machado et al., 1992; Schrank and Machado, 1996; Silva et al., 1996; Noce et al., 1998; Barbosa et al., 2000, 2004; Delor et al., 2003; Vlach et al., 2003; da Rosa-Costa et al., 2008; Fraga et al., 2008, 2009, 2017; Leite et al., 2009; De Roever et al., 2015; Kroonenberg et al., 2016; Medeiros Júnior et al., 2016; Aguilar et al., 2017; Cutts et al., 2018). Similar to the SASCA region, in the northern part of Guiana Shield the dominant deformation D2 is associated with major sinistral strike-slip faults (Vanderhaeghe et al., 1998; Delor et al., 2003). The very young metamorphic ages obtained on monazites (< 2000 Ma) in the São Francisco Craton characterize a long period of slow cooling (1 °C/Ma) down to ca. 1940 Ma (Aguilar et al., 2017) (Fig. 15f). Similarly, high-grade metamorphism is dated at 2046 ± 6 Ma in the Borborema Province, north of the São Francisco Craton (Calado et al., 2019). The south of the São Francisco Craton as well as the Guiana Shield are characterized by a long-lived metamorphism at ca. 2100–1940 Ma including a syn-collisional period of ca. 2100–2070 Ma (Peucat et al., 2011; Aguilar et al., 2017; Fraga and Cordani, 2019). da Rosa-Costa et al. (2008) found very similar results for the late orogenic stage of the collisional orogeny along the Amapá Block. The similarity in the metamorphic age spans suggests a potential link between the southern WAC and South America at the time of the Eburnean Orogeny as proposed by Onstott and Hargraves (1981), Ledru et al. (1994), Nomade et al. (2003) and Traoré et al. (2022).

thumbnail Fig. 15

Relative probability plots of U–Pb monazite ages obtained in this study (a–e) and U–Pb monazite data obtained in South America (f) (Machado et al., 1992; Schrank and Machado, 1996; Silva et al., 1996; Noce et al., 1998; Barbosa et al., 2000, 2004; Vlach et al., 2003; da Rosa-Costa et al., 2008; Fraga et al., 2008, 2009; Leite et al., 2009; Medeiros Júnior et al., 2016; Aguilar et al., 2017; Cutts et al., 2018).

6.4 Geodynamic implications for Archean and Paleoproterozoic orogenic processes

We have documented a polyphased tectono-metamorphic evolution of the metasediments and migmatitic gneisses in the SASCA domain. The reconstructed P–T–t–d paths imply thickening of the crust under moderate apparent geothermal gradients of 20–30 °C/km and maximum burial depth of 20–30 km recorded in the granulite facies samples SP3B and GDB8 and 15–25 km in the samples KOU7 and MAD1A (Fig. 14). This is consistent with burial of supracrustal rocks represented mostly by sediments and volcanics in a collisional setting. The initial burial corresponding to the deformation phase D1 operated prior to ca. 2037 Ma (which is the age of D2). Additionally, a sample of Archean migmatitic gneiss, from the vicinity of Monogaga (sample SP11 in Fig. 13), reveals ∼ 2070 Ma metamorphic overprint documented by younger zircon rims (Koffi et al., 2022).

The metamorphic pressure peak is first followed by a nearly isothermal decompression, interpreted as rapid exhumation, and then by simultaneous decompression and cooling. We note that the ca. 2037 Ma old monazites are predominantly found in samples GDB8, SP3B and MAD1A, which record a retrograde path under amphibolite-facies conditions (MAD1A and GDB8) and that are pervasively affected by the NE-SW trending S2 foliation. Therefore, we suggest that the ca. 2037 Ma ages date the retrograde phase D2.

The monazite metamorphic ages corroborate the conclusion of Kouamelan et al. (1997) and Pitra et al. (2010) from the Kouibli area that the major tectono-metamorphic polyphased event is related to the Eburnean orogeny. The presence of relictual Archean zircon and monazite grains is consistent with the interpretation of the SASCA domain as a so-called “metacraton” (Abdelsalam et al., 2002; Liégeois et al., 2005), which suggests tectonic accretion of Paleoproterozoic and Archean domains affected by metamorphism and deformation during the Eburnean orogeny. Nevertheless, this does not preclude the hypothesis of a sedimentary protolith containing detrital zircon grains.

In the Kouibli area to the N of the study zone, the mechanism of homogeneous shortening accommodated by a combination of shear zones and lateral crustal spreading (Cagnard et al., 2006, 2007; Cruden et al., 2006; Rey and Houseman, 2006) was suggested by Pitra et al. (2010) based on the absence of large-scale thrust systems. Nevertheless, metamorphic and structural mapping of the SASCA area (Fig. 13) shows the juxtaposition of rocks with contrasted metamorphic facies from amphibolite to granulite facies between the zones of Grand-Béréby, Kounoukou and San Pedro. The domains with different metamorphic conditions seem to be delimited by transcurrent shear zones except between SP3B and MAD1A/KOU7. The model (Fig. 16) shows tectonic juxtaposition of units with different peak P–T conditions implying metamorphism at different crustal levels. The character of the P–T–t paths suggests that the tectonic juxtaposition of amphibolite and granulite facies rocks would occur during their exhumation after reaching the peak P–T conditions. Flat-lying thrust zones were not found in the field. Either they were never present or they were transposed into steep foliations and folds during later stages of the Eburnean orogeny. To explain the total vertical displacement between the units, which is in the order of 6–10 km, we advocate large-scale crustal folding (Schulmann et al., 2008; Štípská et al., 2012) and/or gravitational convective-diapiric instabilities (Vanderhaeghe et al., 2018) associated with lateral crustal flow and accompanied by transcurrent subvertical NE-SW trending shear zones during the D2 phase, now defining the structural grain of the whole area.

Collisional setting during the assembly of the Archean and Paleoproterozoic terrains of the WAC was evoked in the work of Camil (1981, 1984), Camil et al. (1983), Triboulet and Feybesse (1998), Thiéblemont et al. (2004), Pitra et al. (2010), in western Côte d’Ivoire, Liberia and Sierra Leone and in the work of Eglinger et al. (2017) in Central Guinea. Many works suggest that Paleoproterozoic formations were deformed in a transpressive context in the late stages of the Eburnean orogeny (Ledru et al., 1991; Feybesse et al., 2006; Baratoux et al., 2011; Jessell et al., 2012; Chardon et al., 2020). However, the proposed exhumation mechanisms of the rocks differ from one area to another. In NW Ghana, the exhumation of the lower crust is attributed to a N-S extension during the gravitational collapse generating anatectic domes (Block et al., 2015). In SW Ghana, the work of McFarlane et al. (2019) describes a transtensional ENE-WSW regime for the exhumation of the middle and lower crust. Neither ductile normal faults similar to the ones described in NW Ghana, nor folds with shallow-plunging fold axes oblique to the orientation of the major shear zones similar to those in SW Ghana, were found in our study area. Therefore, we conclude that the extensional component of deformation during exhumation was either negligible or its record was overprinted by the later intense transpressive deformation. Our study confirms the diversity of metamorphic gradients in the West African Craton documented by several recent studies (Ganne et al., 2012; Block et al., 2015; McFarlane et al., 2019). In some areas, such as Burkina Faso and Ghana, the studied metamorphic rocks recorded an apparent cold geothermal gradient of ∼ 15 °C/km (Ganne et al., 2012; Block et al., 2015); however, in SW Côte d’Ivoire such cold gradients were not found.

In the framework of the ongoing discussion about the character of the orogenic processes at the Archean-Paleoproterozoic transition, P–T conditions combined with (micro)structural observations and geochronology presented in this paper suggest a predominance of diffuse crustal thickening owing to tectonic forces followed by and partly concomitant with gravity-driven lateral flow of the thickened crust. The final stages of the Eburnean orogeny in the SASCA domain might be compared to the orogenic plateau stage of modern orogenic belts as described in Bajolet et al. (2015) or in Vanderhaeghe (2009, 2012).

thumbnail Fig. 16

An evolutionary sketch showing the exhumation of rocks in SW Côte d’Ivoire. CIS62: potassic granite of San Pedro; MON4 and SP11: migmatitic gneisses of Monogaga (Koffi et al., 2022).

7 Conclusion

The SASCA domain in SW Côte d’Ivoire is characterized by the tectonic juxtaposition of granulite-facies migmatitic paragneiss with amphibolite-facies migmatitic grey gneiss and micaschists. The metamorphic conditions recorded in these rocks show a Barrovian-type metamorphism. Granulite-facies rocks reached a syn-D1 metamorphic peak of 10 kbar at 820 °C followed by a retrograde evolution marked by decompression and by cooling. This is consistent with a burial down to 38 km depth before exhumation. For the amphibolite-facies rocks, the syn-D1 prograde evolution (garnet–staurolite-bearing micaschist at Kounoukou with a metamorphic peak at ca. 6.6 kbar, 620 °C) reveals a burial down to 25 km along an apparent geothermal gradient of ∼ 25 °C/km. During subsequent exhumation these high- and medium-grade rocks reached an apparent geothermal gradient of ∼ 40 °C/km. This metamorphic phase was associated with the development of the dominant foliation S2 at ca. 2037 Ma. The exhumation results from a combination of crustal-scale folding and/or convective-diapiric instabilities associated with lateral flow owing to a regional transcurrent regime, which would explain the juxtaposition of rocks that reached significantly different P–T conditions.

Very young monazites (ca. 2000 Ma and 1978 to 1913 Ma) found in all dated samples crystallized during the retrograde path and could result from two types of processes: (1) (partial) resetting of older crystals by either long-term cooling with periodic opening of the U–Pb system or isotopic disturbance by fluids, or (2) episodic crystallization of monazite during the cooling period. This is the very first study in the WAC, which determines such young metamorphic ages allowing a correlation with the Guiana Shield where these ages are more abundant.

Acknowledgements

This project is part of a PhD study of A.Y. Koffi, partially funded by the T2GEM project (Geophysical and Geochemical Technologies for Mineral Exploration) and by the project “West African Exploration Initiative” (WAXI). We wish to gratefully acknowledge AMIRA International and the industry sponsors for their support of the WAXI project (P934B). We are also appreciative of the contribution of the various geological surveys and departments of mines in West Africa as sponsors in kind of WAXI. We would like to thank the scholarship directorate of the Ministry of Higher Education and Scientific Research of Côte d’Ivoire for awarding A.Y.K. a scholarship for a doctoral internship in France to carry out the analyses. The African Network of Institutions in Earth Sciences (ANESI) for the financing of the stay of A.Y.K. in Senegal to improve the quality of this work. The authors acknowledge the French National Research Institute for Sustainable Development (IRD, LMI MINERWA Laboratory for Responsible Mining in West Africa) for funding a part of this research project. We are grateful to P. de Parseval. S. Gouy and T. Agouy for their help in acquiring microanalysis and F. Barrère de Parseval for the thin sections. The authors gratefully acknowledge Lêda Maria Fraga and one anonymous reviewer for their constructive reviews.

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Cite this article as: Koffi AY, Baratoux L, Pitra P, Kouamelan AN, Vanderhaeghe O, Thébaud N, Bruguier O, Block S, Fossou Kouadio HJ-L, Kone J. 2023. A tectonic model for the juxtaposition of granulite- and amphibolite-facies rocks in the Eburnean orogenic belt (Sassandra-Cavally domain, Côte d’Ivoire), BSGF - Earth Sciences Bulletin 194: 11.

Supplementary Material

Table S1: LA-ICP-MS U-Pb analyses for standard material Managotry and Moacyr.

Table S2: LA-ICP-MS U-Pb data for monazites from the studied samples.

Table S3: Summary of monazite ages from South America.

Access here

All Tables

Table 1

Petrographic summary and location in degree (WGS 84 coordinate system) of the samples studied. Quartz, plagioclase, biotite and ilmenite are present in all samples.

In the text
Table 2

Chemical compositions (oxide wt%) of the samples used for the P–T pseudosection calculations. Fe2O3 corresponds to total Fe.

In the text
Table 3a

Representative microprobe analyses of selected minerals.

In the text
Table 3b

Representative microprobe analyses of selected minerals.

In the text

All Figures

thumbnail Fig. 1

Geological sketch map of the West African Craton (modified after the BRGM SIG Africa map and Ennih and Liégeois, 2008; Berger et al., 2013; Thiéblemont, 2016). The position of the present-day margins of the craton is constrained by geophysics (Jessell et al., 2015).
In the text
thumbnail Fig. 2

Litho-structural map of the SW of Côte d’Ivoire (modified after Tagini, 1971; Papon, 1973; Pitra et al., 2010; WAXI report, AMIRA Global, 2018). Blue stars show the location of the studied samples. Samples are described from West to East.
In the text
thumbnail Fig. 3

Concordia plots of U–Pb age dating for reference material Manangotry and Moacyr monazites (a and b respectively).
In the text
thumbnail Fig. 4

Detailed geological map with structural relationships. (a) Structural map showing the orientation of S1, S2 and S3 foliations. L2 lineations are also plotted. (b) Geological cross-section showing the D2 deformation phase. (c) Equal-area, lower hemisphere stereoplots show D2 planar (foliation S2) and linear (lineation L2) deformation features.
In the text
thumbnail Fig. 5

Field photographs illustrating typical structural relationships (S1, S2 and S3) in this study. (a) S1 and S2 in garnet–cordierite–sillimanite migmatitic paragneiss Grand-Béréby (GDB8). (b) S1 and S2 in migmatitic grey gneiss San Pedro (SP3B). (c) S1 and S2 in staurolite-bearing micaschist San Pedro (SAN). (d) S2 in garnet–sillimanite–cordierite micaschist Madie (MAD1A). (e) S3 in garnet–staurolite-bearing micaschist Kounoukou (KOU7).
In the text
thumbnail Fig. 6

Photomicrographs (plane polarized light). (a) Sample GDB8. Main penetrative foliation S2 (fine-grained biotite + prismatic sillimanite) wraps around (a) a microlithon with preserved foliation S1 marked by coarse-grained biotite and prismatic sillimanite, and anhedral porphyroblasts of garnet with abundant inclusions of fine sillimanite needles. The dashed line A-B indicates the location of the chemical profile of garnet (Fig. 7a). Sample SP3B: (b) String of monazite crystals aligned parallel to the foliation S2 marked by biotite, which also crystallizes in the fractures and at the rims of garnet porphyroblasts containing inclusions of monazite. A-B indicates the location of the chemical profile of garnet (Fig. 7b). Sample SAN: (c) Biotite and rare fibrous sillimanite mark the main foliation (S2) and (d) andalusite replaced partly by chlorite. Sample MAD1A: (e) Foliation S2 marked by biotite and clusters of fibrous sillimanite. (f) Garnet porphyroblast in contact with rare staurolite crystals. The dashed line A-B indicates the location of the chemical profile (Fig. 7c). Sample KOU7: (g) Staurolite contains quartz inclusion trails that mark the foliation S1, oblique to the matrix foliation S2. (h) A-B – Location of the chemical profile (Fig. 7d) across a garnet porphyroblast with inclusions of ilmenite and quartz.
In the text
thumbnail Fig. 7

Chemical profiles of garnet (mole %). Location is shown in Figures 6a (GDB8), 6b (SP3B), 6e (MAD1A) and 6g (KOU7).
In the text
thumbnail Fig. 8

P–T pseudosections in the model system NCKFMASHTO (whole-rock compositions expressed in mole %) and compositional isopleths for (a) the migmatitic paragneiss GDB8 and (b) the migmatitic grey gneiss SP3B. Red arrow and red star show the inferred P–T path (GDB8) and the equilibration P–T conditions (SP3B), respectively. Field limits are marked by the phase that appears/disappears. See text for details.
In the text
thumbnail Fig. 9

P–T pseudosections and compositional isopleths calculated for the samples (a) SAN in the model system NCKFMASHTO, and (b) MAD1A in the model system MnNCKFMASHTO. Red arrows show the inferred P–T paths. See text for details.
In the text
thumbnail Fig. 10

P–T pseudosection (MnNCKFMASHTO) and compositional isopleths calculated from the sample KOU7 (whole-rock composition expressed in mole %). Red arrow shows the inferred P–T paths. See text for details.
In the text
thumbnail Fig. 11

(a) Position of the monazite crystals (bright spots) in the thin section GDB8. (b) Alignment of monazite clusters parallel to S2 in migmatitic grey gneiss SP3B. (c–g) Migmatitic grey gneiss SP3B. (c) Back-scattered electron (BSE) images showing the position of the monazite crystals (bright spots) with respect to the garnet porphyroblasts. (d, e) Microprobe X-ray maps and (f, g) BSE images with analytical spots and associated 207Pb/206Pb dates of monazite grain marked in (a). LC: low concentration; HC: high concentration.
In the text
thumbnail Fig. 11

(Continued)
In the text
thumbnail Fig. 12

U–Pb dating of monazites by LA-ICP-MS. (a) Concordia diagram showing all analytical points for monazites from the migmatitic grey gneiss SP3B. (b) Zoom of the concordia diagram showing the two age groups at ca. 2000 and 2037 Ma in the sample SP3B. (c–f) U–Pb concordia diagrams for monazites from garnet–cordierite–sillimanite migmatitic paragneiss GDB8 (c) and metapelites SAN, MAD1A and KOU7 (d, e and f respectively). SAN: staurolite-bearing micaschist; MAD1A: garnet–sillimanite–cordierite micaschist; KOU7: garnet–staurolite-bearing micaschist.
In the text
thumbnail Fig. 13

Map of the SASCA domain, SW of Côte d’Ivoire, modified after Papon (1973), showing the position and the monazite age of the different samples studied (blue stars). Granulite and amphibolite facies represent the peak conditions of the corresponding areas. Migmatitic gneisses of Monogaga (MON4 and SP11) and potassic granite of San Pedro (CIS62) are from Koffi et al. (2022).
In the text
thumbnail Fig. 14

Synthesis of the P–T conditions of the different rocks studied as well as their P–T path according to metamorphic facies and geothermal gradients (modified after Ernst and Liou, 2008). Abbreviations of metamorphic facies: AM: amphibolite; Amp-EC: amphibolite-eclogite; BS: blueschist; EA: epidote amphibolite; GR: sillimanite-granulite; GS: greenschist; HGR: kyanite-granulite.
In the text
thumbnail Fig. 15

Relative probability plots of U–Pb monazite ages obtained in this study (a–e) and U–Pb monazite data obtained in South America (f) (Machado et al., 1992; Schrank and Machado, 1996; Silva et al., 1996; Noce et al., 1998; Barbosa et al., 2000, 2004; Vlach et al., 2003; da Rosa-Costa et al., 2008; Fraga et al., 2008, 2009; Leite et al., 2009; Medeiros Júnior et al., 2016; Aguilar et al., 2017; Cutts et al., 2018).
In the text
thumbnail Fig. 16

An evolutionary sketch showing the exhumation of rocks in SW Côte d’Ivoire. CIS62: potassic granite of San Pedro; MON4 and SP11: migmatitic gneisses of Monogaga (Koffi et al., 2022).
In the text

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