# A window on the amalgamation of Western Gondwana: Geological history of the Ouaddaï massif (E. Chad)

Alexis Plunder, Olivier Blein, Moussa Isseini, Issaka Ousman Al-Gadam, Matthieu Chevillard, Emanuel Djedouboum, Philippe Lach, Abdeltif Lahfid, Jérémie Melleton, Olivier Rouzeau, et al.

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## To cite this version:

Alexis Plunder, Olivier Blein, Moussa Isseini, Issaka Ousman Al-Gadam, Matthieu Chevillard, et al. A window on the amalgamation of Western Gondwana: Geological history of the Ouaddaï massif (E. Chad). Precambrian Research, 2026, 434, pp.108000. 10.1016/j.precamres.2025.108000. hal-05450603

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# A window on the amalgamation of Western Gondwana: Geological history of the Ouaddaï massif (E. Chad)

Alexis Plunder<sup>a,b,*</sup>, Olivier Blein<sup>a</sup>, Moussa Isseini<sup>c</sup>, Issaka Ousman Al-Gadam<sup>d</sup>, Matthieu Chevillard<sup>a</sup>, Emanuel Djedouboum<sup>d</sup>, Philippe Lach<sup>a</sup>, Abdeltif Lahfid<sup>a</sup>, Jérémie Melleton<sup>a,b</sup>, Olivier Rouzeau<sup>a</sup>, Guillaume Vic<sup>a</sup>

## ARTICLE INFO

### Keywords:

- Saharan metacraton
- U-Pb Geochronology
- Gondwana
- Chad
- Cratonic destruction
- Panafrican orogeny

## ABSTRACT

The Saharan metacraton was assembled during pre-Neoproterozoic to Neoproterozoic times and was strongly remobilised during the Pan African orogeny. The Ouaddaï massif in eastern Chad represents the core of the Saharan metacraton, yet its geological history remains poorly documented. In this study, we combine of field observations, petrological analysis of metamorphic rocks, and geochemical and geochronological constraints to reconstruct the tectono-metamorphic evolution of the Ouaddaï massif. The study area is dominated by collision-related plutonic rocks and migmatitic gneisses, within which inliers of older basement occur. Geochemical data indicate a transition from active-margin to collisional magmatism, with a compositional evolution from diorite to granite. This suite is cross-cut by syenogranites and alkaline granites with shoshonitic affinities, interpreted as post-collisional granitoids derived from tonalitic rocks to sedimentary protoliths. Geochronological data (U-Pb on zircon and monazite) point to a ca. 1000 Ma age for orthogneisses inliers. Granitoids of the Ouaddaï massif record emplacement ages from 620 to 590 Ma, coeval with high-temperature metamorphism characterized by peak pressure-temperature conditions of 1.2 GPa and 850 °C (mafic granulite) and around 0.8 GPa and 700 °C (sillimanite-garnet bearing migmatites). Integrating our new results with regional data, we discuss the existence and geodynamic evolution of the Saharan Metacraton. Our findings emphasize the significance of the Saharan Metacraton as a key region for understanding the extensive reworking of cratonic lithosphere during both a Tonian magmatic phase prior and the assembly of Gondwana.

## 1. Introduction

The area located between the Tuareg and the Arabian Nubian shields, referred to as the Saharan metacraton or Central Sahara Shield corresponds to a vast area (5,000,000 km²) and is one of the least geologically studied regions in the world (Abdelsalam et al., 2002; de Wit et al., 2021; Liégeois et al., 2013). This region contains occurrences of Archean and Paleoproterozoic protoliths (Bea et al., 2011; Fezaa et al., 2010; Zhang et al., 2019). Together with isotopic and tomographic data this area was interpreted as a Mesoproterozoic craton that was later extensively remobilised in the Neoproterozoic amalgamation of Gondwana (Abdelsalam et al., 2011; Black and Liégeois, 1993; Liégeois et al., 2013).

The relative position of the former Saharan craton (situated between

the Tuareg Shield / West African Craton, the Congo Craton, the Arabian Nubian Shield and an unknown craton) explains why the ‘metacratonization’ process was so intense in this region (Abdelsalam et al., 2002; Fritz et al., 2013; Johnson et al., 2011; Liégeois et al., 2013; Meert and Lieberman, 2008; Shellnutt et al., 2017). A metacraton as defined by Abdelsalam et al. (2002) and Liégeois et al. (2013) corresponds to an area where a craton has been remobilized in response to an orogenic event but remains recognizable through its geochronological, isotopic and geophysical properties. The region therefore provides a unique opportunity to study and challenge the processes leading to the modification or destruction of cratonic lithosphere due to collision and/or accretion of other blocks.

Unfortunately the inferred cratonic parts are covered by Phanerozoic basins (i.e. the Murzuq, Al Kufrah and Chad basin and can be only

<footnote><sup>a</sup> BRGM, F-45060 Orléans, France</footnote>

<footnote><sup>b</sup> Institut des Sciences de la Terre d'Orléans, UMR 7327, Université d'Orléans, CNRS, BRGM, OSUC, Orléans, France</footnote>

<footnote><sup>c</sup> Université de N'Djamena, Chad</footnote>

<footnote><sup>d</sup> Ministère des Mines et de la Géologie, Chad</footnote>

<footnote>* Corresponding author.</footnote>
inferred with geophysical methods (Abdelsalam et al., 2011; Fishwick and Bastow, 2011; Liégeois et al., 2013; Sobh et al., 2020). The metacratonic area bordering the three basins, however, are exposed at the surface and represent the only remnants of the former craton. In Chad, four metacratonic areas have been identified, namely the Ouaddai, Tibesti, Guéra and Mayo Kebbi massif (Fig. 1). Geological investigation “from scratch” of these metacratonic regions are therefore required to advance our understanding of the metacratonization process.

While the timing and history of the collision between the Saharan craton and the Tuareg and Arabian Nubian shield is well established (e. g. Fezaa et al., 2010; Fritz et al., 2013; Johnson et al., 2011; Liégeois et al., 2013; Robinson et al., 2017), the collision between the Saharan craton and the Congo craton is far less studied (Shellnutt et al., 2020; Toteu et al., 2022). It is constrained to ca. 630–580 Ma, mainly based on post-collisional plutonic and metamorphic rocks either from the Chad or the Central African orogenic belt (Bouyo Houketchang et al., 2009; Djerossem et al., 2020; Kwékam et al., 2013, 2010; Ngako and Njonfang, 2011; Pin and Poidevin, 1987; Shellnutt et al., 2019, 2018, 2017).

The existence of the Sahara metacraton has, however, been questioned in the literature, based on synthesis of existing knowledge and new data from previously unexplored areas (de Wit et al., 2021; Djerossem et al., 2020; Şengör et al., 2022; Shellnutt, 2021; Shellnutt et al., 2025). The Sahara region has alternatively been interpreted as an accretionary orogenic belt or a shield terrane pinched between the Tuareg and Arabian Nubian Shields (Shellnutt et al., 2026, 2025). It has even been proposed that it represent a segment of a large-scale double orocline, the Saharides, that extends from the east of the West African craton to the Arabian region. This belt that formed in the Neoproterozoic by subduction and accretion of arc fragments shaved off the Congo craton margin and accreted together (Şengör et al., 2022).

Regardless of the formation mechanism, the pre-collision evolution stages of the region (i.e. > 590 Ma) remain poorly known and is mainly documented in the Guéra massif and the Mayo Kebbi massif (Isseini et al., 2012; Shellnutt et al., 2017). Amongst the Precambrian massifs of Chad, the Waddai or Ouaddai massif represent the southern edge of the Saharan metacraton (Fig. 1). It is the largest Precambrian remnant in Chad and, together with its equivalent in Darfur region of Sudan, constitutes the largest Precambrian massif of the Saharan metacraton (Figs. 1,2). The Ouaddai massif is therefore a key locality for understanding the amalgamation of Gondwana and the processes leading to the destruction of cratonic lithosphere.

The Ouaddai massif is mainly made of calc-alkaline granites, high-temperature metamorphic rocks (including amphibolites, migmatites, marbles and graphitic schists) as well as gneisses and recent basaltic lava

flows (Djerossem, 2018; Gindy, 1984; Kusnir and Moutaye, 1997; Wolff, 1964).

In this study we present a geochemical, geochronological and petrological dataset from the northern Ouaddai massif, in the context of the updated 1:200 000 Am Zoer geological map (14–15° North and 21–22° East; Chevillard et al., 2023). The data were acquired as part of the “Chad Mining Inventory” project, which included a geochemical stream-sediments survey conducted between 2016 and 2018 on the Am Zoer sheet. We document Stenian-Tonian and Ediacaran magmatism and metamorphism using zircon and monazite U-Pb geochronology. We also constrain the condition of formation of the metamorphic assemblages related to the magmatic intrusions using various empirical thermometer and thermobarometric methods (multi-equilibrium, pseudosection modelling and Raman spectroscopy). Finally, we place this dataset in the context of the collision of the Congo and Saharan cratons, during the broader Neoproterozoic Pan-African orogeny and discuss (i) the distinct geodynamic events recorded in the Ouaddai massif and (ii) the potential for cratonic lithosphere destruction during collision.

## 1.1. The Precambrian massifs of Chad

The Republic of Chad covers ~ 30 % of the Saharan metacraton that is known to be the “poorly known tract of continental crust in the World” (Liégeois et al., 2013; Fig. 2). In Chad, four main Precambrian domains are known: the Tibesti massif to the North, the Guéra massif in the center, the Ouaddai massif to the East and the Mayo Kebbi Massif to the South West (Fig. 2; Abdelsalam et al., 2002; Shellnutt et al., 2017). The first three ones belong to the Saharan metacraton whereas the Mayo Kebbi massif is assigned to the Central African orogenic belt (Archelaus Emmanuel et al., 2024; Toteu et al., 2004).

The Tibesti massif is mainly made of metasediments interbedded with metavolcanics intruded by a granitoid series interpreted as emplaced in an arc system (Kusnir and Moutaye, 1997; Pegram et al., 1976). The pioneering work of Pegram et al. (1976) documents the calc-alkaline Ben Ghnema batholith of the Tibesti massif as a subduction related batholith emplaced ~ 560 Ma using the Rb-Sr method on whole rock. Later, the age of the magmatic event was found to spread over ca. 40 Ma, from to 537 to 580 Ma (El-Makhrouf and Fullagar, 2000; El Makhrouf, 1988; Suayah and Rogers, 1986) and confirmed by U-Pb ages on zircon around 550 Ma (Suayah et al., 2006). Interestingly, the Tibesti massif is the that record metamorphism around ca. 939 ± 52 Ma (whole rock Rb-Sr age; El Makhrouf, 1988).

The Guéra massif mainly consists of volcanic rocks and their

<figure><img src="image_4.png" /><figcaption>**Fig. 1.** Schematic map of Africa showing the cratonic zones and the extent of the Saharan metacraton based on the reworked area (modified after Liégeois et al., 2013; Fritz et al., 2013).</figcaption></figure>
<figure><img src="image_5.png" /><figcaption>**Fig. 2.** Geological map of the Saharan Metacraton and surrounding area. Based on Thiéblemont et al. (2016) and on Milesi et al. (2010).</figcaption></figure>

intrusive equivalent (mainly rhyolite and granite) as well as meta-
morphic rocks (schists, marbles, amphibolites, para and orthogneiss).
Granite and rhyolite from the Guéra massif yielded LA ICP-MS U-Pb ages
on zircon ranging from 570 to 590 Ma (Shellnutt et al., 2017, 2019). A
diorite sample from a drill core in the Doba Basin (ca. 300 km southwest
of the Guéra Massif) yielded a CA-ID-TIMS U-Pb zircon age of 573 Ma
(Shellnutt et al., 2017). Granite from drill core in the same region
yielded U-Pb ages on zircon ranging from 580 to 595 Ma (Shellnutt et al.
2025). All these ages, interpreted as crystallisation ages, indicate a 25
Myr magmatic event representing the accretion of the Congo and
Saharan Craton. Arc related rocks emplaced prior to ca. 590 Ma indicate
subduction, and collision is marked by anataxis and granite emplaced
between ca. 590 to 580 Ma. Later, post-collisional magmatism possibly

due to lithospheric delamination is registered at ca. 580 Ma. In the Guéra
massif, a sample of mildly foliated granite yielded inherited zircon with
ages of 1000 Ma and 1900 Ma (Shellnutt et al., 2020, 2017) that matches
the Nd model ages of other rocks form the Saharan Metacraton (in
Sudan, Niger and Egypt; Abdelsalam et al., 2002). In the Doba Basin,
inherited zircon yielded Paleoproterozoic ages of 1700 and 2100 Ma
(Shellnutt et al., 2025).

The Ouaddaï massif is the largest Precambrian massif in Chad, and together with its counterpart in Sudan (the Darfur massif), and is the most extensive massif in the Saharan Metacraton (Fig. 1) but remains ill studied (de Wit et al., 2021). The Ouaddaï massif is divided in: (i) a southern part, mainly made of a low-to high-grade metasedimentary sequence intruded by granodioritic to leucogranitic plutons and (ii) a
northern part essentially made of migmatite and granodiorite.

The southern part, dominated by metasediments and granitic rocks also includes amphibolites, marble, quartzite (Djerossem et al., 2020, 2021). The sedimentary sequence have maximum deposition ages of ca. 1045 Ma to ca. 635 Ma and Sm-Nd model ages of the amphibolite give a ca. 620 Ma age in a passive margin or volcanic arc context (Djerossem, 2018; Djerossem et al., 2021). Archean and Paleoproterozoic inheritance are observed in the zircon of these metasediments. Orthogneisses give ca. ~1000 Ma U-Pb ages on zircon for their emplacement and were sourced from a ca. 1000 to 2000 Ma crust (de Wit et al., 2021). Amongst the magmatic rocks S-Type granites were recognized, with zircon U-Pb crystallization ages of 635 ± 3 Ma and 612 ± 8 Ma (Djerossem et al., 2020). The metamorphic event associated to anatexis is dated from 627 to 600 Ma using U-Th-Pb on monazite (Djerossem, 2018). The last event is attributed to the emplacement of I-Type potassic calc-alkaline granitoids around 538 ± 8 Ma (U-Pb ages on zircon) associated to partial melting of the sub continental mantle.

The northern region of the Ouaddaï massif, is undoubtedly the least studied area. A few U-Pb dates give emplacement ages between ca. 596 to 613 Ma for granite, gabbro and dacite (de Wit et al., 2021). A granite and a high Si granite (orthogneiss?) give ca. 1 Ga age of crystallisation. However the ages are only based on the measurement of a few, often discordant, zircon grains (de Wit et al., 2021). Isotopic Nd data give T<sub>DM1</sub> between 1.45 and 3.08 Ga.

In the province of Darfur, the sequence is similar but metamorphosed with migmatitic metapelite, quartzite and graphitic schist interbedded with metabasite in the amphibolite to granulite facies (Gindy, 1984). It is intruded by post-tectonic monzogranite (Gindy, 1984). Xenoliths from the Jebel Marra volcano, in Darfur, give one of the few regional age constrain, with Sm-Nd model ages of 1930, 2210 and 2820 Ma attesting the existence of an old crust (Davidson and Wilson, 1989).

To the west of the Chad lineament, the **Mayo-Kebbi massif** is part of the Central African orogenic belt which forms the boundary between the Congo Craton, the West African Craton and the Saharan metacraton (Toteu et al., 2004). It is a mixture of metamorphic rocks (metavolcanics, metasediments and metamafic plutonic) intruded by Neo-proterozoic granitoids, interpreted to represent the transition from an arc setting (723–638 Ma; Penaye et al., 2006; Pouclet et al., 2006) to an anorogenic setting at 572 Ma (Isseini et al., 2012). One sample yielded inherited pre-Neoproterozoic zircon ages indicating interaction with an old basement (Isseini et al., 2012). Also, to the west of the Chad lineament, post-collisional granites in the lake Fitri region yielded U-Pb emplacement ages of ca. 550 Ma (Shellnutt et al., 2017, 2018).

## 2. Field observations

Basis on our field observations, the rocks types exposed in the northern part of the Ouaddaï massif can be divided into four units: (i) a basement composed of magmatic and metamorphic rocks of variable metamorphic grade (greenschist- to granulite-facies); (ii) a magmatic unit mainly consisting of diorites to porphyric K-feldspar granitoids, (iii) a silicic magmatic unit, and (iv) a presumably young volcanic complex (Fig. 3).

The **basement** comprises several tectono-metamorphic units. The most abundant rock types correspond to a high-temperature (HT) metamorphic sub-unit dominated by gneisses, mafic to ultramafic metamorphic rocks, and migmatites. This HT unit is commonly anatectic and where possible, we have distinguished a migmatitic and mafic sub-units (Fig. 3). In map view, the geometry of the anatectic unit and the trajectory of the foliation locally suggest the presence of a metamorphic dome. The migmatites consist predominantly of felsic material, often interlayered with mafic layers (Figs. 4b,d). In some locations, we observed garnet-sillimanite bearing migmatite (Fig. 4c) along with associated felsic melts. The relative proportion of felsic and mafic components vary considerably and, locally the sequences evolve to kilometre-scale massifs composed entirely metamafic rocks (Fig. 4e).

The mafic layers are mostly amphibolite-facies rocks, although rare granulite-facies assemblages are also present, as indicated by the garnet and clinopyroxene association.

Other metamorphic sub-units, also representing the regional basement includes (i) an amphibolite-facies volcano-sedimentary sequence composed of graphitic schist, locally containing sillimanite, and interbedded with amphibolite-facies mafic rocks (Kourek area; Fig. 4g) (ii) variably deformed paragneisses and orthogneisses, occurring south of Am Zoer and around Niéré, for which the metamorphic grade is difficult to estimates from field observation alone (Fig. 3). The occurrence of metre-thick mylonitic shear zones in the Niéré orthogneiss suggests deformation under relatively high temperature conditions. In places, the migmatites are covered by sandstones, from which we collected a sample to investigate the distribution of detrital zircon.

The most regionally widespread unit consists of diorites, granodiorites to porphyritic K-feldspar granites, hereafter referred to as the "the arc-collisional series". In the field, these rocks are generally coarse-grained melanocratic to mesocratic and largely undeformed. We also identified some andesitic volcanic rocks that may represent the volcanic equivalent of this series. This unit locally contains small orthogneisses bodies, which are too small to be represented on the geological map (Fig. 3).

The second magmatic series is made of highly differentiated rocks, including monzogranite and alkali granite. These rocks are coarse-grained, mesocratic to leucocratic, and undeformed. They intrude the arc-collisional and the anatectic unit. Numerous cross-cutting dykes of pinkish microgranite were also observed.

The last unit comprises undated basaltic flows that may correspond to the known Cenozoic- Quaternary volcanism documented in the Tibesti massif, in the Tuareg Shield or in Darfur (Bermingham et al., 1983; Deniel et al., 2015; Liégeois et al., 2005).

## 3. Methods

### 3.1. Geochemistry

A total of 76 rock samples were submitted for whole rock geochemistry to ALS (Ireland; https://www.alsglobal.com). All samples were analysed for major and trace elements by ICP-AES and ICP-MS respectively on the same digested powder. The full dataset is presented in Appendix 1. The loss on ignition (LOI) values are generally < 1.4 wt% and the average of 0.96 wt% for the 76 samples.

### 3.2. Electron probe micro analyser

We performed analyses using a Cameca SXFive, electron probe microanalyzer at BRGM. We used standard analytical conditions as follow: 15 kV acceleration voltage, 12nA beam current, and 3 µm beam size; wavelength-dispersive spectroscopy mode. For calibration, we used andradite (Ca), MgO (Mg) albite (Na and Si), orthoclase (K), Al₂O₃ (Al), Fe₂O₃ (Fe), albite (Na), Cr₂O₃ (Cr) and MnTiO₃ (Mn, and Ti).

### 3.3. Methods for PT estimates

Considering the variability of the investigated samples we used different thermobarometric methods to retrieve the PT conditions of formation. We used when suitable the following empirical thermometer and barometer: Ti in biotite thermometer (Henry et al., 2005), garnet-biotite and garnet-clinopyroxene exchange thermometer (Ferry and Spear, 1978; Powell, 1985), hornblende-plagioclase thermometer (Holland and Blundy, 1994), garnet aluminosilicate plagioclase barometer (GASP; Hodges and Spear, 1982). For two samples, we used the average PT mode of THERMOCALC v3.33. We calculated PT pseudosection (i.e. phase diagram for a fixed bulk composition) using Perple_X 6.6.8 combined with the updated Holland and Powell (1998) dataset. We used the following solution model: garnet, ilmenite, liquid (melt)
<figure><img src="image_6.png" /><figcaption>Fig. 3. Simplified geological map of the Am Zoer region and sample location.</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>Fig. 4. Representative pictures of rocks observed in the field.</figcaption></figure>

after (White et al., 2007), feldspar (Fuhrman and Lindsley, 1988), cordierite, staurolite, chloritoid (Holland and Powell, 1998) chlorite (Holland et al., 1998), biotite (Tajčmanová et al., 2009) orthopyroxene and spinel (White et al., 2002).

## 3.4. Geochronology

We took cathodoluminescence and back scattered images for all zircon and monazite respectively using a TESCAN Mira 3 XMU SEM equipped with a detector for cathodoluminescence at BRGM.

We performed 11 analyses using a quadrupole ICP-MS X series II coupled to a Cetac Excite 193 nm laser at BRGM. We used zircon 91,500 (Wiedenbeck et al., 1995) as a primary standard and the Plešovice zircon as secondary standard (338 ± 1 Ma; Sláma et al., 2008). Two more samples (RMC0245, RCJ0115-1) were analysed using an Attom sector Filed mass spectrometer coupled to NewWave 193 nm laser ablation system at GeoRessources laboratory (Nancy, France). Similarly to BRGM

protocol, zircon 91,500 was used as standard, and Plešovice as quality control sample. The ablation diameter was set to 25 µm and the data reduction was performed with Iolite V3 (Hellstrom et al., 2008) software. We calculated and generated all concordia dates and diagrams using the IsoploR software package (Vermeesch, 2018).

## 4. Geochemistry

Based on the major and trace element composition we divided the dataset in 6 groups (Table 1, Supplementary Material 1).

The first group comprises intermediate to felsic rocks, such as monzodiorite, diorite, granodiorite and monzogranite. Intermediate rocks are characterized by moderate SiO₂ contents, with intermediate to high Fe₂O₃, low to high MgO, CaO, and K₂O contents, and low to intermediate Na₂O contents. These monzodiorites and diorites are metaluminous and mainly magnesian. They are calc-alkalic to alkali-calcic, with an intermediate to high Mg#, and K₂O/Na₂O ratio generally higher
than 0.9 (Fig. 5). Felsic rocks are characterized by high SiO₂ and K₂O contents, with intermediate Na₂O contents, and low Fe₂O₃, MgO, and CaO contents. These granodiorites and monzogranites are metaluminous to peraluminous, mainly magnesian and calc-alkalic to alkali-calcic, with intermediate to high Mg#, and K₂O/Na₂O ratios greater than 0.9 (Fig. 5). These metaluminous and magnesian dioritoids and granitoids represent highly potassic calc-alkaline series. High potassic calc-alkaline dioritoids and granitoids are characterized by fractionated REE patterns ([La/Yb]n = 11–75), with slightly fractionated HREE ([Gd/Yb]n = 1.8–3.9), HREE between 10 and 30 times that of chondrites, without positive or negative Eu anomalies. These REE patterns are homogeneous. The calc-alkaline granitoids have high Sr and Ba contents (Sr > 300 ppm; Ba > 1000 ppm), with low Th and U contents (Th < 20 ppm; U < 4 ppm).

The second group is represented by felsic rocks, such as syenogranite and alkaline granite. These rocks are enriched in SiO₂ and K₂O, with intermediate levels of Na₂O, and low levels of Fe₂O₃, MgO, and CaO. These granitoids are peraluminous and ferroan, ranging from calcic-alkali to alkali, with low to intermediate Mg# and K₂O/Na₂O ratio greater than 1 (Fig. 5). These peraluminous and ferroan granitoids represent the shoshonitic series. Shoshonitic granitoids have been

divided into three sub-group based on trace and rare earth element contents. The first sub-group is characterized by fractionated REE patterns ([La/Yb]n > 15), with slightly fractionated HREE ([Gd/Yb]n = 1.9–2.6), HREE between 3 and 20 times chondrites, with or without slightly negative Eu anomalies. The second sub-group is characterized by fractionated REE patterns ([La/Yb]n > 4), with flat HREE patterns ([Gd/Yb]n = 0.6–2.5), HREE lower than 30 times chondrites, without pronounced negative Eu anomalies. These first two groups present La contents lower than 300 times chondrites. The third sub-group is characterized by a higher enrichment in LREE ([La]n > 300), fractionated REE patterns ([La/Yb]n > 8), with fractionated HREE patterns ([Gd/Yb]n = 1.7–7.0), HREE between 10 and 30 times chondrites, without large negative Eu anomalies. Shoshonitic granitoids have low Sr and Ba contents (Sr < 250 ppm; Ba < 1000 ppm), with high Th and U contents (Th > 20 ppm; U > 4 ppm).

The third group is represented by alkaline granites. These rocks are highly enriched in SiO₂, enriched in K₂O, and have intermediate levels of Na₂O, and low levels of Al₂O₃, Fe₂O₃, MgO, and CaO. They are classified as peraluminous, ferroan, calcic-alkali, with low Mg#, and K₂O/Na₂O ratios between 0.7 and 1.8 (Fig. 5). Alkaline granites are characterized by an HREE enrichment ([HREE]n > 40), more or less flat

<figure><img src="image_8.png" /><figcaption>**Fig. 5.** Major element plot of the magmatic rocks of the North Ouddai massif. (a) modified alkali-lime index (Na₂O + K₂O–CaO) vs. SiO₂ (wt%); (b) Alkali index (molar Al/[Na + K]) versus aluminum saturation index (ASI = molar Al/Ca + Na + K); (c) Feroan versus magnesian granite type. (d) Ternary diagram Al₂O₃/(FeOt + MgO); 3 * CaO; 5 * (K₂O/Na₂O) with the composition of the magmatic rock of the Am Zoer area.</figcaption></figure>
REE patterns with slightly fractionated LREE ([La/Sm]n = 1.5–3.2), with flat HREE patterns ([Gd/Yb]n = 0.8–1.6), without large negative Eu anomalies. Alkaline granites have low Sr and Ba contents (Sr = 7–16 ppm; Ba = 26–119 ppm) and high Nb and Y contents (Nb = 25–52 ppm; Y = 74–203 ppm).

*The fourth group* is represented by biotite-poor leucogranites. These

rocks are highly enriched in SiO₂, enriched in K₂O and Al₂O₃, and have intermediate levels of Na₂O, and low levels of Fe₂O₃, MgO, and CaO. They are classified as peraluminous, ferroan, calcic-alkali, with low Mg# and K₂O/Na₂O ratios between 0.8 and 1.5 (Fig. 5). Leucogranites are characterized by low REE contents (ΣREE < 70 ppm), with more or less fractionated REE patterns ([La/Yb]n = 1–28), with slightly fractionated

<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Selected microphotographs of the samples collected in the Am Zoer area. (a)Entire thin section view of a mafic granulite with close-up on the garnet and clinopyroxene texture showing the development of symplectite (a1-3); (b) amphibolite; (c) Silimanite-garnet bearing migmatite and close up on the textures (c1,2); (d,e) close-up on a migmatite showing the migmatite and leucosome texture; (f) othogneiss sample and close-up on a ca. 250 µm monazite and on a small garnet developed during metamorphism; (g,h) representative views of orthogneisses showing the abundance of zircon and monazite; (i) Thin section view and close-up of a graphitic silimanite-bearing schist. Mineral abbreviations: Bt, biotite; Qz, quartz; Ilm, ilmenite; Pl, plagioclase; Cpx, clinopyroxene; Rt, rutile; Sill, sillimanite; Grt, garnet; Kfs, alkali feldspar; Gr, Graphite; Ms, muscovite; Amp, amphibole; Mnz, monazite; Zrn, zircon.</figcaption></figure>
HREE patterns ([Gd/Yb]n = 0.8–1.6), and HREE less than 20 times that of chondrites. Leucogranites have low Zr contents (Zr < 65 ppm).

The fifth group comprises ultramafic, mafic and intermediate rocks, which we identified in the field as amphibolites or pyroxenite layers within the gneisses or migmatites. Amphibolites have low SiO₂ contents and high Al₂O₃, Fe₂O₃, MgO, and CaO contents, as well as low to intermediate Na₂O and K₂O contents. FeO/t/MgO ratios are higher than 1. Pyroxenite layers have low to intermediate SiO₂ contents, high Fe₂O₃ and MgO contents, low to high CaO contents, and low Al₂O₃, Na₂O and K₂O contents. FeO/t/MgO ratios are less than 0.7. Amphibolites and pyroxene layers are characterized by slightly fractionated REE patterns ([La/Yb]n = 1–15), with slightly fractionated HREE ([Gd/Yb]n = 1.2–3.4), HREE less than 30 times chondrites, without positive or negative Eu anomalies.

The sixth group is represented by felsic gneisses. These gneisses have high SiO₂ contents and intermediate Na₂O and K₂O contents, as well as low to intermediate Fe₂O₃, MgO, and CaO contents. They are classified as peraluminous, magnesian, calcic-alkali, with generally intermediate to high Mg#, and K₂O/Na₂O ratios between 0.6 and 1.3 (Fig. 5). Felsic

gneisses are characterized by fractionated REE patterns ([La/Yb]n = 2–20), with slightly fractionated HREE, less than 20 times chondrites ([Gd/Yb]n = 1.4–3.3), or with HREE higher than 50 times chondrites with flat HREE patterns ([Gd/Yb]n = 1.01–1.02).

## 5. Metamorphic petrology

During the field campaigns, 375 rock samples were collected on the Am Zoer sheet, amongst which > 40 are metamorphic. These were cut in the XZ plane of the finite deformation and investigated at the microscope for petrographic determination. Amongst these samples we selected 11 for chemical analyses, Raman spectroscopy, geochronology and thermobarometric calculations. Below, is a summary of the representative samples investigated for PT estimates. The complete list of minerals from the metamorphic samples and is given in Table 2 and EPMA analyses can be found as Supplementary Material Table S2.

Sample RGV0042 is a retrogressed garnet clinopyroxene granulite-facies mafic lens enclosed within migmatite. The rock is dominated by amphibole, some of which replace the original clinopyroxene (Figs. 6a,

<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Chemical composition of mineral from migmatitic gneisses, orthogneisses, amhibolite and schists. (a) Ternary diagram for pyroxene; (b) ternary diagram for feldspar; (c) ternary diagram for garnet; the small inset show the zoning of a garnet in a migmatitic gneiss; (d) XMg versus Si a.p.f.u. diagram for amphibole; (e) Ti a.p. f.u. versus Si a.p.f.u. in amphibole.</figcaption></figure>
al). Garnet coexists with clinopyroxene, plagioclase, and brown amphibole, as indicated by the sharp grain boundaries and inclusions within garnet crystals. Rare quartz inclusions are present, and pyrite is the main opaque phase. Garnet and clinopyroxene are commonly destabilized to symplectites composed of green amphibole and plagioclase (Figs. 6a,a1,a2,a3). The composition of the clinopyroxene falls within the diopside field, with an average XMg = 0.64 (with XMg = Mg/[Fe + Mg]). Garnet shows a relatively constant composition around Alm<sub>51</sub>Prp<sub>26</sub>Grs<sub>22</sub> (Fig. 6c). We measured two groups of plagioclase: Pl<sub>1</sub>, texturally equilibrated with garnet or clinopyroxene, with XAb = 0.4–0.55, corresponding to andesine (Fig. 7b); a more anorthite-rich Pl<sub>2</sub>, in the symplectites, ranging from labradorite to anorthite. Most amphibole have a tschermakite composition (Amp<sub>1</sub>) whereas symplectic amphibole correspond to magnesio-hornblende (Amp<sub>2</sub>; Fig. 7d).

Sample RMC0116-1 is a biotite-amphibole-bearing gneiss composed of plagioclase, quartz, biotite and amphibole with minor ilmenite. Plagioclase has an albite content of XAb = 0.6–0.65. The amphibole range in composition from magnesio-hornblende to tschermakite (Fig. 7d). The Ti content a.p.f.u. (atom per formula unit) reach 0.25 suggesting crystallization at relatively high temperature, in the presence of a Ti-bearing buffer phase (Ernst and Liu, 1998). Biotite has a constant XMg of ~ 0.55, with a Ti content of ~ 0.1–0.25 a.p.f.u. (normalized to 11 oxygen; Fig. 7a).

Sample RMC0102 is a fine-grained amphibolite composed mainly of plagioclase and amphibole, with minor ilmenite (Fig. 6b). Plagioclase as an albite content that varies from 0.4 to 0.65. The few analysed amphibole crystals show a wide compositional range, from ferro-tschermakite to magnesio-hornblende (Fig. 7d).

Sample RMC0121-1B is a garnet-bearing migmatite of the anatectic unit. The foliation is well developed at both the outcrop and the thin section scale. Part of the sample is dominated by a leucogranitic assemblage (Kfs + Pl + Qz) whereas the remainder consists predominantly of garnet, sillimanite, biotite, quartz, K-feldspar and plagioclase (Figs. 6c,c1,c2). Minor phases include rutile, ilmenite, monazite and zircon (Figs. 6c1,c2). Detailed observation under reflective light shows that ilmenite overgrowths rutile in the matrix and that both phases coexist as inclusion in garnet. Garnet composition vary from Alm<sub>60</sub>Prp<sub>33</sub>Grs<sub>6</sub>Sps<sub>1</sub> in the core to Alm<sub>75</sub>Prp<sub>18</sub>Grs<sub>6</sub>Sps<sub>1</sub> at the rim. Biotite compositions vary according to microstructural context: inclusion in garnet versus matrix grains: Ti ranges from 0.17 to 0.25 a.p.f.u. and a XMg varies from 0.30 to 0.70 (Fig. 7a). Alkali feldspar is sanidine, close to the end-member and plagioclase composition range between oligoclase and andesine (Fig. 7b). Ilmenite or rutile have compositionally close to their ideal end-member, FeTiO<sub>3</sub> and TiO<sub>2</sub>, respectively.

Sample RMC0121-2 is a garnet biotite bearing migmatite found near sample RMC0121-1B. It exhibits a weak foliation defined by biotite. The mineral assemblage is made of garnet, biotite, plagioclase, quartz and K-feldspar in textural equilibrium (Figs. 6d,e). Ilmenite, rutile, and monazite occur as minor phases. Garnet composition range from Alm<sub>69</sub>Prp<sub>18</sub>Grs<sub>7</sub>Sps<sub>7</sub> in the core to Alm<sub>71</sub>Prp<sub>10</sub>Grs<sub>7</sub>Sps<sub>12</sub> at the rim (Fig. 7c). As in the previous sample the composition of biotite depends on its position relative to garnet. The Ti content ranges from 0.17 to 0.22 a.p.f.u. and the XMg varies from 0.47 to 0.60 (Fig. 7a). Plagioclase has an andesine composition and the alkali feldspar has a sanidine composition.

Sample RMC0229, RIA0106 and RAO0051 are orthogneisses. They display a coarse-grained texture (Fig. 6f), with a poorly developed foliation defined by the alignment of biotite and muscovite crystals. The main mineral assemblage consists of plagioclase, alkali feldspar, quartz, muscovite and biotite (Figs. 6f,g,h). Monazite (up to 400 µm in samples RMC0229), zircon and apatite occur as minor phases and are concentrated within the biotite-rich layer (Fig. 6f1). A single garnet crystal was identified in sample RMC0229, coexisting with biotite and quartz (Fig. 6f2). Across the three samples, the mineral compositions appear to be relatively homogeneous (Fig. 7).

Sample RMC0107 is a sillimanite-bearing graphitic schist. The foliation is well developed and defined by the alignment of prismatic sillimanite, which often destabilize to muscovite (Fig. 6i). Graphite occurs as thin flakes, either included in quartz or in primary muscovite (Fig. 6i1,2). Mineral compositions were not studied in detail for this sample.

# 6. Pressure–temperature estimates

## 6.1. Graphitic schists

We calculated the Tmax value for three graphitic schists using the Raman spectra following the procedure of Beyssac et al., (2002) (Table 3). The samples yielded homogeneous Tmax values of 607 °C for sample RMC0107, 578 °C for sample RMC0103 and 587 °C for sample RMC0206 (Figs. 8b,e). The Ti-in-biotite thermometer provide a comparable temperature of 600 ± 24 °C on average for sample RMC0206 (Figs. 8a,b).

Samples RMC0103 and RMC0107 belong to the same formation located to the north of the investigated area. Interestingly, sample RMC0206, situated more than 70 km to the south, records a very similar temperature. Combined with the presence of sillimanite in sample RMC0107, the RSCM temperature allows to give a PT estimates of 0.4 ± 0.2 Gpa and 607 ± 30 °C for the climax conditions (Fig. 8e).

## 6.2. Mafic samples: RMC0116-1 and RMC0102 RGV0042

For two amphibole-bearing samples (RMC0116-1 and RMC0102) the hornblende-plagioclase thermometry yielded average temperature of 770 °C (RMC0116-1) and 647 °C (RMC0102), respectively, using the quartz-present and quartz-absent thermometer of Holland and Blundy (1994) (Fig. 8b).

Sample GV0042 is a retrogressed garnet-clinopyroxene-bearing amphibolite. Estimating the peak metamorphic conditions using a forward approach is problematic due to the uncertainties in reconstructing the peak bulk composition. We therefore applied garnet-clinopyroxene thermometry, amphibole-plagioclase thermometry, and multi-equilibrium, 'average PT' calculations. The garnet-clinopyroxene thermometer of Powell (1985) yields a temperature of 760 ± 50 °C (Fig. 8b). The quartz-absent hornblende-plagioclase thermometer of Holland and Blundy (1994) gives a temperature of 514 ± 50 °C for the formation of the Amp<sub>2</sub> + Pl<sub>2</sub> symplectite developing between garnet and clinopyroxene (Fig. 8b). In addition, a brute force average PT approach was performed calculating all combination of the EPMA analyses. The 523 combinations satisfying the statistical criteria of THERMOCALC, provide a robust average PT estimate of 1.2 ± 0.05 GPa and 850 ± 50 °C for the assemblage Grt + Cpx + Pl<sub>1</sub> + Amp<sub>1</sub> + Qz (Fig. 8c).

## 6.3. Anatexis-related samples: RMC0121-1B and RMC0121-2

We investigated sample RMC0121-1B in details. The garnet-aluminosilicate-plagioclase (GASP) barometry indicates a pressure of 0.85 to 0.98 GPa at 850 °C (Fig. 8e). The Ti-in-biotite thermometer yields temperature of 700 to 775 °C, consistent with the high pyrope content of garnet. Combining the available EPMA analysis, resulted in 300 combinations consistent with the statistical criteria of THERMOCALC, giving an average PT estimate of 0.9 GPa and 850 °C for the assemblage Grt + Pl + Kfs + Bt + Qz + Sill + Rt + Ilm.

A PT pseudosection was calculated for this sample (Fig. 9). We considered that the loss on ignition of the bulk analysis to represent the water content. This is supported by the extremely limited retrogression observed in the sample. The diagram was computed at an oxidation state of zero (i.e. XFe<sup>3+</sup> = 0 with XFe<sup>3+</sup> = Fe<sup>3+</sup> / [Fe<sup>3+</sup>+Fe<sup>2+</sup>]) in agreement with the Fe<sup>3+</sup> estimated in minerals from EPMA analyses. The observed assemblage (including melt) is stable at conditions ranging from 0.6 to 1.2 GPa and 825 to 900 °C within both the rutile- and ilmenite-present
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** Summary of PT estiamtes realized of different samples; (a) Ti a.p.f.u. versus XMg diagram for biotite with temperature estimates (after Henry et al., 2005); (b) RSCM and empirical thermometry temperature (different methods); (c,d) 3D plots of the average PT estimates obtained for samples GV0042 and RMC0121-1B; (e) Summary of all estimates.</figcaption></figure>

fields. Using the composition of garnet and the close association of rutile and ilmenite, we refined the PT conditions to a narrower space with PT condition of 1.0 ± 0.1 GPa and 860 ± 15 °C. An additional diagram calculated with XFe³⁺ = 0.06 reproduces a similar assemblage (Grt + Bt + Sill + Kfs + Pl + Qz + Ilm + melt), but rutile is stabilized only in the kyanite-present field. This hypothesis is therefore discarded.

Using the same approach, we calculated a PT pseudosection for sample RMC0121-2 for which we can refine the PT conditions using the stable assemblage (Grt + Qz + Bt + Pl + Rt/Ilm + melt) and the garnet composition at 0.9 ± 0.1GPa and 725 ± 75 °C (Fig. 9).

## 6.4. Orthogneisses: RMC0229, RIA0106 and RAO0051

Due to their high variance assemblages, precise PT estimates are difficult to obtain for orthogneisses. The Ti-in-biotite thermometer yield

temperature of ~ 650 °C for samples RMC0229 and RIA0106 while sample RAO0051 record a slightly higher temperature of 675 °C (Fig. 8b). The temperature calculated for sample RMC0229 using the garnet-biotite thermometer is of 750 °C on average. It is known that the Ferry and Spear (1978) calibration of the garnet-biotite Fe-Mg exchange may overestimate temperature especially above 600 °C. Applying the Kaneko and Miyano (2004) recalibration results in a temperature of ~ 670 °C, which is more consistent with the Ti-in-biotite temperature calculated for sample RMC0229.

## 7. Geochronology: LA-ICP MS U/Pb ages on zircon monazite and xenotime

To decipher the magmatic and metamorphic events recorded in the Am Zoer Area we chose 3 orthogneisses, 6 granitoids, 2 migmatites and a
<figure><img src="image_12.png" /><figcaption>**Fig. 9.** (a) pseudosection for sample RMC121-1B; (bulk mol. % ox). SiO₂ (57.49), TiO₂ (1.43), Al₂O₃ (15.21), FeO (11.498), MgO (5.58), CaO (1.38), Na₂O (0.92), K₂O (2.42), H₂O (3.85). (b) same pseudosection with garnet isopleths; (c) pseudosection for sample RMC121-2; (bulk mol. % ox). SiO₂ (65.60), TiO₂ (0.85), Al₂O₃ (9.74), FeO (6.39), MgO (5.04), CaO (2.83), Na₂O (2.78), K₂O (1.85), H₂O (2.78) (d) same pseudosection with garnet isopleths; Mineral abbreviations: Bt, biotite; Qz, quartz; Ilm, ilmenite; Pl, plagioclase; Crd, cordierite; Opx, orthopyroxene; Rt, rutile; Zo, Zoisite; Sill, sillimanite; Phg, phengite; Pg, paragonite; Grt, garnet; Kfs, alkali feldspar; Ky, kyanite.</figcaption></figure>

<figure><img src="image_13.png" /></figure>

sandstone for dating. The magmatic and one orthogneiss samples were crushed and sieved. The retrieved zircons have been mounted in epoxy for LA-ICP MS analyses. For the remaining, we performed the analysis on thin sections. All data are available as **Supplementary Material Table S3 and S4**.

The zircon crystals sometimes show oscillatory magmatic zoning, sometimes overgrown by a metamorphic rim. Some crystals show inherited cores. Zoning in monazite is generally more complex. In some sample we identified clear core to rim zoning and used these textural criteria, when possible, to calculate the ages. The result of all date

calculation are presented on figure 10 and 11 with all uncertainties given as 2σ.

## 7.1. Orthogneisses

RCJ0115-1: this orthogneiss from the Niéré massif yielded 60 zircons crystals. The zircon crystals are euhedral, prismatic and typically display magmatic zoning with a Th/U ratio of 0.38–0.91. We made 30 laser spot that allows to calculate a discordia line with a natural lower anchor at 0 Ma and an upper intercept giving a date of 1023 ± 4 Ma (n = 24) that we
take as the emplacement age of this rock prior to metamorphism (Fig. 10a).

For both orthogneisses samples RMC0229 and RIA016 all zircon crystals show similar morphologies with a magmatic core and some thin metamorphic rims. In all case the rims are below 25 µm and we could not analyse them. In both cases, zircon have Th/U values of 0.04–1.22 (RMC0229) and 0.01–1.22 (RIA0106).

For sample RMC0229, out of 26 analyses, 20 allow to calculate a discordia line with a natural anchor at 0 Ma and an upper intercept at a date of 997 ± 13 Ma (Fig. 10b). Using the 7 analyses it is possible to calculate a concordia date at 974 ± 7 Ma that is within error of the upper intercept. We interpret this age to represent the emplacement of the orthogneiss.

For sample RIA016, the dataset is of lower quality, with many analyses that plot far from concordia. Out of 26 analyses, 12 allow to draw a discordia line with an upper intercept at 1016 ± 53 Ma and a poorly meaningless lower intercept at 310 ± 168 Ma (Fig. 10c). Despite its poor definition, we interpret the upper intercept to represent the emplacement age of the orthogneiss.

## 7.2. Magmatic rocks

### 7.2.1. High potassic calc-alkaline dioritoids and granitoids

RMC0075: This monzogranite yielded only a few zircons. Out of the 9 crystal, 25 analyses were performed. Nine analyses allow to calculate a concordia date of 618 ± 5 Ma that we interpret as the age of emplacement of the monzogranite (Fig. 10d). Paleoproterozoic and Archean inheritance is found in xenocryst cores at ca. 1953 ± 31 Ma and 2817 ± 33 Ma.

RMC0219: this quartz monzodiorite of the Goz massif yielded 65 euhedral zircon crystals. We performed 30 spot analysis that plot close to the concordia. We used 27 analyses to calculate a concordia date at 588 ± 3 Ma that we interpret as the age of emplacement of the granodiorite (Fig. 10e).

RLB0037: This granodiorite samples yielded 58 zircon crystals. Most crystals are euhedral, long prismatic with typical magmatic cathodoluminescence patterns. Some crystals show xenocrysts cores. Out of the 30 analyses, 28 cluster around 550–620 Ma. Using only the six most concordant analysis, we calculated a concordia date at 594 ± 5 Ma that we interpret as the emplacement of this granodiorite (Fig. 10f). One xenocryst core gives a ca. 975 ± 20 Ma date that could represent the basement.

RMC0245: we found 45 euhedral zircon crystals in this quartz monzodiorite sample, and we analysed 27 of them that allows us to calculate a concordia date at 609 ± 2 Ma using 22 analysis that we interpret as the emplacement of the rock (Fig. 10g).

RLB0041: Out of the 70 fragmented zircon crystal retrieved for this quartz monzonite sample we analysed 30 spots that show a spread over 620 to 560 Ma. Using 17 analyses we calculated a concordia date of 587 ± 3 Ma (Fig. 10h) that we interpret to represent the age of emplacement of the quartz monzonite.

### 7.2.2. Granitoids within migmatites

RGV0018: we retrieved only 13 zircon crystals for this granitoid sample found in migmatite. Their analysis produced two set of ages. The first one obtained from 4 analyses of the zircons rims is a concordia date at 596 ± 8 Ma (Fig. 10i) and the second one, obtained from the core of three zircon at 676 ± 10 Ma (Fig. 10i). The other analyses were discarded due to fractures or inclusions. The two dates presented are weak, but the difference between core and rims is notable and remains interesting. The young date could correspond to the age of emplacement of this granitoid, and the old date certainly represent inheritance.

## 7.3. Metamorphic rocks

Monazite and xenotime ages.

Monazite of both samples show textural zoning (see Appendix), but not convincing differences were found in terms of dates. For sample RIA0106 we analysed both xenotime and monazite. Out of 32 monazite analyses, two show evidence for old (ca. 850 Ma) date. Within the remaining 30 analyses, 20 allow to calculate a discordia line with a lower intercept at 618 ± 3 Ma (Fig. 10j). Out of the 7 xenotime analyses, four allow to calculate a concordia date at 612 ± 8 Ma (Fig. 10k) that we take as the age of high-temperature metamorphism.

For sample RMC0229, out of 54 analyses, 31 allow to calculate a discordia line with a lower intercept at 591 ± 2 Ma and an upper intercept that goes towards common lead (Fig. 10l). Using 11 analyses, it is possible to calculate a concordia date at 589 ± 4 Ma that is equal to the lower intercept. We take this date to represent the age of high-temperature metamorphism. The dataset shows lead loss that certainly represent the continuation of the HT event or following magmatic event.

## 7.4. Sandstone

A sandstone sample, RMC0055, yielded 105 analyses from 60 zircon crystals. Of these, 77 analyses are free of common lead and within 10 % concordance. Five major zircon age populations are identified at ~2400 Ma, 1894 Ma, 928 Ma, 656 Ma and 534 Ma (Fig. 11). The youngest population yields a depositional age of ~534 Ma, indicating that the sandstone was deposited in the early Cambrian. The data also demonstrate that sedimentary input was derived from sources that included Paleoproterozoic and Archean crust.

## 8. Discussion

We discuss the information gathered in the Am Zoer region. First, we present the interpretation of our geochemical, geochronological data and discuss the implication of the PTt estimates provided in this contribution. We then arrange all information in a cogent geodynamic model and discuss the accretion of cratonic block and the process of metacratonization.

### 8.1. A window on the geological history of the former Saharan craton

#### 8.1.1. Old history (> 900 Ma)

In the field, we identified the alkaline granites as part of an orthogneiss basement (essentially the Niéré massif; RCJ0115). In all discrimination diagrams (Fig. 12), these granites plot within the anorogenic, or within plate granite fields, indicating their different origin compared to the calc-alkaline and shoshonitic granitoids described in section 4. Some undifferentiated gneisses also exhibit similar geochemical characteristics. They occur either as small sized migmatitic gneisses remnants within the arc-to-collisional series (RIA0106) or within migmatites (RMC0229). The orthogneisses are magnesian and peraluminous (Fig. 4). When we plot the analyses in the ternary CaO – K₂O/Na₂O – Al₂O₃/(FeOt + MgO) diagram (Laurent et al., 2014) the orthogneisses separate into two distinct compositional groups, implying a dual origin. The first group is derived from the melting of sedimentary materials, while the second originates from K-rich, mantle-derived magmas.

The zircon crystal core retrieved from three sample yield dates of ca. 1.0 Ga (Fig. 10 a,b,c), which we interpret as the emplacement age of the protolith. This age corresponds to the basement of the Ouddaï massif, which is possibly equivalent to the Bayuda and Zalingei massif in Sudan where similar ca. 920 Ma basement (granites and gneisses) are preserved despite the widespread Pan-African magmatism (Küster et al., 2008; Küster and Liégeois, 2001). These ages are also consistent with those reported for the southern and northern Ouaddaï Massif, obtained from granite (de Wit et al., 2021). Together with Archean to Paleoproterozoic ages obtained from xenocryst, these data attest the existence of an older basement upon which the Panafrican magmatism was superimposed. Evidence of this ancient history is also preserved in sandstone RMC0055,
<figure><img src="image_14.png" /><figcaption>**Fig. 10.** U-Pb concordia diagram of zircon and monazite. (a-c) Orthogneisses emplacement; (d-i) granitoids emplacement; (j-l) high-temperature metamorphism. All errors are 2 σ.</figcaption></figure>
<figure><img src="image_15.png" /><figcaption>**Fig. 11.** Detrital zircon dates and peak distribution on a radial plot for sandstone sample RMC0055.</figcaption></figure>

and in pre-Devonian and present-day sediments attesting that the Saharan area was a source for ca. 2.5, 2.0 Ga and 1.0 Ga (e.g. Be'Eri-Shlevin et al., 2014; Stephan et al., 2019).

### 8.1.2. Active margin history: from arc to collision (700–590 Ma)

Metaluminous and magnesian granitoids form part of a calc-alkaline series. They are highly potassic and correspond to the composition of experimental fusion derived from high-K mafic rocks (Fig. 12). These rocks are classified as I-type granites. In the Th/Nb discrimination diagram of Pearce (2008), dioritoids display an enriched Th/Nb ratio compared to the ocean island basalt / mid ocean ridge basalt array (OIB/ MORB). This signature is indicative of external input, either from a subduction zone or through contamination by continental crust. In the Nb/Th vs. Zr/Nb diagram of Condie (2005), the dioritoids plot within the arc field. Similarly, in the Y + Nb vs. Rb diagram of Pearce et al., (1984), both dioritoids and granitoids plot in the arc field (Fig. 12). Taken together, these results support the interpretation of the dioritoids and granitoids as evidence of a magmatic arc along the southern margin of the Saharan craton previously proposed.

The ages obtained for the dioritoids and granitoids range from 618 to 587 Ma (Fig. 10). We interpret these ages as recording the transition from arc to collisional magmatism, with the collision age being equivalent to that of the Guéra massif (590–595 Ma; Shellnutt et al., 2017) and the Doba Basin (ca. 595 Ma; Shellnutt et al., 2025). This interpretation is consistent with monazite crystallization ages in migmatitic rocks (618–591 Ma) which record collisional PT conditions of 1.0 ± 0.1 GPa and 860 ± 15 °C and 0.9 ± 0.1 Gpa and 725 ± 75 °C.

Our ages are slightly younger than those documented in the Mayo Kebbi massif (Isseini et al., 2012). Together with the literature data, this confirms the idea that the southern edge of the Saharan craton has been an active margin between at least 700 and ~ 590 Ma.

### 8.1.3. Post collisional history (590–540 Ma)

Peraluminous and ferroan syenogranites, as well as alkaline granites, with shoshonitic affinities, are derived either from sediments or tonalitic rocks (Fig. 4). According to the classification proposed by Whalen et al. (1987), they range from weakly fractionated granites to A-type granite. In the Y + Nb vs. Rb diagram of Pearce et al., (1984), they plot in the near the triple point (VAG–Syn-COLG–WPG; Fig. 4), as it is commonly observed for post-collisional granitic rocks (Pearce, 1996).

The emplacement ages of these syenogranites and alkaline granites are constrained only by field relationship. They suggest a thermal pulse following the collision of the Congo and Saharan cratons at ca. 590 Ma. Post-collisional granite are documented in the southern part of the Ouaddï massif at 540 Ma and from 575 Ma to 540 Ma in the Guéra massif (Djerossem et al., 2020; Shellnutt et al., 2019, 2018).

### 8.1.4. Non time constrained geochemical data

Leucogranites, with ferroan and peraluminous affinities (Fig. 4), undoubtedly originate from tonalitic melting and demonstrate high degree of fractionation. They do not clearly relate to other granitoids and the absence of geochronological prevents more definitive conclusions.

The amphibolites and pyroxene-rich layers found in the anatectic unit derivate from the MORB/OIB array in the Th/Nb proxy diagram. They plot in the volcanic arc field with variable continental influence (Fig. 11). Their field occurrence, embedded in gneisses of the anatectic series, supports the hypothesis that this arc-related magmatism is older than the documented Pan-African arc.

### 8.1.5. Relations with surrounding Pan African belts

One key geometric question remains: how the Neoproterozoic orogenic belts of Africa are linked, as discussed by Villeneuve and Rossignol (2024). The scenario we document for the Ouaddaï Massif shares similarities with events described in the Central African Orogenic Belt. The collision between the Congo Craton, the Sahara Craton, and intervening continental slivers occurred at ca. 620–580 Ma, whereas post-collisional granite emplacement and sediment deposition are dated at ca. 540 Ma. Part of this interpretation relates to the possible northern continuation of the Tcholliré–Banyo Shear Zone (TBSZ), which may represent a Neoproterozoic suture zone (Toteu et al., 2022). This structure has been tentatively extended northward toward the Chad Lineament (Shellnutt et al., 2026, 2025), and is thought to represent a possible suture zone (Li et al., 2013). Alternatively, the TBSZ may continue eastward, remaining parallel to the northern margin of the Congo Craton (Fig. 2; Milesi et al., 2010).

Links with other Pan-African belts located on the western and eastern sides of the Saharan Metacraton are more speculative, owing to the lack of direct structural connections and the presence of extensive recent sedimentary cover (Fig. 2).

To the west, the Trans-Saharan Orogen consists of a collage of various terranes assembled during at least two Neoproterozoic phases (Liégeois, 2019). Some of the events documented in the Ouaddaï are coeval with those recorded in the easternmost and closest terranes of the Tuareg Shield (e.g., Eastern Hoggar), where a ca. 575–555 Ma intra-continental collision has been documented in the Djanet Terrane (Fezaa et al., 2010). However, their geodynamic significance differs. While collision following subduction occurred around ca. 600 Ma in the Ouaddaï and the Central African Orogenic Belt, collision in the Eastern Hoggar occurred later and is attributed to reactivation of the Saharan
<figure><img src="image_16.png" /><figcaption>**Fig. 12.** Classification of magmatic rocks of the North Ouaddaï area. Th/Yb vs. Nb/Yb diagram after Pearce, 2008. Granite classification after Whalen et al., (1987). Zr/Nb vs. Nb/Th diagram of Condie (2005). Granitic rock classification using the Rb vs. Y. + Nb diagram after Pearce et al., (1984).</figcaption></figure>
Craton margin. This latter event involved collision between the Saharan Metacraton and the Tuareg Shield, following the ca. 630 Ma collision between the Tuareg Shield and the West African Craton (Fezaa et al., 2010; Liégeois et al., 2013; Liégeois, 2019). A major unresolved issue concerns the lack of geological constraints from the Tibesti Massif.

To the east, the assembly of the Saharan Metacraton with the Arabian-Nubian Shield represents a long-lived tectonic evolution. It began with the closure of an oceanic domain from ca. 750 Ma, followed by a period of sinistral transpressive convergence between ca. 650 and 600 Ma, culminating in final collision at ca. 600–580 Ma (Abdelsalam et al., 1998; Abu et al., 2019; Fritz et al., 2013; Johnson, 2014). Key data required for a more comprehensive, large-scale interpretation are likely preserved in the Sudanese parts of the Saharan Metacraton.

## 8.2. Metamorphism: From arc to collision

The samples, for which we characterized the metamorphic grade and the timing of high-temperature events, provide additional information to that presented in the previously published geological history.

### 8.2.1. PT estimates.

The graphitic metasediments and interbedded amphibolite give consistent temperature estimates of 580–640 °C using RSCM, Ti in biotite, and plagioclase + amphibole thermometry. The presence of sillimanite in the graphitic schist does not permit to constrain the pressure in the absence of any other index mineral. The origin of the metamorphism (i.e., regional vs. contact metamorphism) remains uncertain and requires further investigation. However, the extent of this volcano-sedimentary group, the distance between the two graphite-bearing formations (ca. 70 km) and graphite spectra shape suggest that the metamorphism is regional. This formation may correspond to the metasediments of the southern part of the Ouaddaï massif, which exhibit greenschist- to amphibolite-facies conditions, consistent with the temperatures obtained (Djerossem et al., 2020).

In the anatectic unit, the peak temperature is constrained in the range 650–850 °C, depending on the method and samples used. This variation can be explained by both the geographical distribution of the samples and the structural level of sampling. The three samples that we investigated in details (RMC0121-1B, RMC0121-2 and RGV0042) provide some important insights. Despite originating from the same locality, the two migmatitic samples yield slightly different PT estimates at 1.0 ± 0.1 GPa / 860 ± 15 °C and 0.9 ± 0.1 GPa / 725 ± 75 °C. These two estimates do not overlap, and the small uncertainty of 15 °C computed for sample RMA0121-1B is highly improbable when considering the inherent uncertainties of the thermodynamic database (e.g. Holland and Powell, 1998; Powell et al., 1998). Moreover, we did not account for several well-known geological uncertainties, such as the rock volume used to determine the bulk composition, the fact that the equilibrium volume exceed the size of our sample (Palin et al., 2016), the common behaviour of rocks to be out of thermodynamic equilibrium due to the overstepping of reactions requiring of the nucleation of a new mineral (Pattison and Debuhr, 2015; Waters and Lovegrove, 2002), or the thermal buffering associated with melting processes (Schnor and Diener, 2019).

Taking these factors into account, further detailed investigation of our samples is required to better constrain their PT conditions of formation. At present, our first order estimates are satisfactory for an area that remains largely unexplored. We note that the statistical approach involving the systematic combination of all measured minerals at the probe to calculate the average PT condition with THERMOCALC is in good agreement with the pseudosection for sample RMC0121-1B. This approach may represent an efficient method to investigate the ubiquitous polyphased rocks of orogenic systems, where PT estimates can often be challenging to obtain (François et al., 2018; Goscombe et al., 2020).

For the garnet-bearing mafic sample (RGV0042), we relied only on the average PT approach due to the strong retrogression and the

presence of decompression texture, as attested by amphibole + plagioclase symplectite overgrowing garnet + clinopyroxene. The peak pressure conditions is constrained at 1.2 ± 0.05 GPa and 825 ± 50 °C.

The difference of pressure between the granulite (1.2 GPa) and the garnet-sillimanite migmatite (1.0 GPa) can be explained by: (1) different levels of exhumation of the rock that come from distinct structural level within the same group; (2) the same limitation in the precision of PT estimates as discussed for the anatectic rocks; (3) an older metamorphic history preserved only in the mafic parts of the anatectic complex, as documented during various periods (e.g. Agbosoumonde et al., 2001; Ménot and Seddoh, 1985; Ring et al., 2002; Whitney et al., 2020). Such older history could involve the development of granulite-facies rocks at the base of an active margin, similar to scenarios proposed for your settings such as the fossilized Kohistan magmatic arc (Garrido et al., 2006).

### 8.2.2. Timing of metamorphism and relation to geological history.

We obtained metamorphic ages for two samples. Sample RIA0106, a migmatitic orthogneiss, records temperature conditions of ca. 650 °C, consistent with the presence of melt. The protolith age is estimated at ca. 1016 Ma (see §8.1). Metamorphism is constrained by a discordant U-Pb date monazite at 618 ± 3 Ma and a concordant U-Pb date on xenotime age at 612 ± 8 Ma. Both dates are coherent and consistent within uncertainties. These dates are coeval with the age of arc to collision magmatism.

Sample RMC0229, from the anatectic unit, also records conditions compatible with partial melting. It has inherited zircons with a U-Pb date of 997 Ma, interpreted as the protolith emplacement age. The age of metamorphism is constrained by a discordant U-Pb date on monazite at 591 ± 3 Ma. This high-temperature metamorphism is coeval with the age of arc-to-collision magmatism.

## 8.3. The Saharan metacraton, myth or reality?

The Saharan metacraton was first proposed by Abdelsalam et al. (2002). It existence remains debated with arguments both for and against (Abdelsalam et al., 2002; de Wit et al., 2021; Liégeois et al., 2013; Şengör et al., 2022; Shellnutt, 2021; Shellnutt et al., 2025). Here, we aim to highlight several lines of evidence that are often overlooked in the discussion.

It has been proposed that the Chad lineament, a magnetic anomaly (Braitenberg et al., 2011; Li et al., 2013) represents the northern continuity of the Tcholliré-Banyo shear zone (TBSZ) observed in Cameroon (Penaye et al., 2006) and may correspond to a Neoproterozoic suture zone (Braitenberg et al., 2011; de Wit et al., 2021; Shellnutt et al., 2025). This interpretation is supported by contrasting geochemical signature between rocks found east and west of the Chad lineament, suggesting distinct lithotectonic assignment (Shellnutt et al., 2026, 2025). In Cameroon, the TBSZ separates the juvenile Paleoproterozoic Bamenda-Tignere and Garoua domains to the north from the Adawama-Yadé domain to the south, which contains Archean relics (Delor et al., 2021; Vic et al., 2025). However, the continuity between the two structure remains speculative. East-West-trending shear zones described along the northern margin of the Congo Craton, particularly within the Central African orogenic belt, could alternatively represent the prolongation of the TBSZ (e.g. Toteu et al., 2004).

Additional evidence comes from geophysical measurements. Different modelling approaches yield variable estimates of the lithosphere thickness beneath the Saharan metacraton (Abdelsalam et al., 2011; Begg et al., 2009; Celli et al., 2020; Liégeois et al., 2013; McKenzie, 2020; Pasyanos and Nyblade, 2007; Sobh et al., 2020). At the scale of the Sahara, it has been proposed that the Saharan metacraton consists of three smaller relict craton namely, the Murzuq, Al Kufrah and Chad cratons, buried beneath Palaeozoic to recent sediments, and separated by strongly reworked metacratonic area (Liégeois et al., 2013). Archean crust is documented in the Uweinat region, with
Archean (~3.2 Ga) T<sub>DM</sub> model ages, a Sm-Nd isochron age of 3.2 Ga and U-Pb zircon age ranging from 3.1 to 2.55 Ga (Bea et al., 2011; Zhang et al., 2019). Elsewhere, few geochronological dates are available, and Archean rocks have not been directly identified. Nevertheless, age models frequently indicates sources as old as 2.0 Ga, and in some cases up to ~ 3.0 Ga, for example in the Djanet, Tibesti and Ouaddaï massifs (de Wit et al., 2021; Fezaa et al., 2010; Suayah et al., 2006). Our own data document the presence of ~2.0 Ga and 2.8 Ga inherited zircon grains in arc-related rocks of the Ouaddaï massif. This validates the Archean T<sub>DM</sub> ages that were discarded for the Uddaï massif (de Wit et al., 2021) and show that there is an older basement.

The last point comes from the sedimentary record. Detrital zircon from Cambrian to Ordovician rocks of the Sahara region (southern Libya and Algeria), reveal four dominant age populations: an Archean (~2.5–2.6 Ga), a Paleoproterozoic (with a major peak at ~ 1.8 Ga), a Mesoproterozoic (~1.04 Ga) and a Neoproterozoic population (e.g., Meinhold et al., 2013, 2011; Stephan et al., 2019). In our data, detrital zircon from a Cambrian sandstone deposited on migmatite shows equivalent peaks at ~ 2400 Ma, ~1894 Ma, 928 Ma, and 656 Ma and a deposition age of ~ 534 Ma. This indicates that while deposition of the sandstone, Archean, Paleoproterozoic, Mesoproterozoic and Neoproterozoic-Cambrian source rocks were exposed and contributed to the material supply.

These points constitute for us sufficient suspicion that the Saharan craton existed but was highly reworked as already proposed.

8.4. Destruction of cratonic lithosphere: A tale from the upper plate

Finally, we discuss a possible scenario for the evolution of the southern margin of the Saharan metacraton. The regional known age constraints and our ages are reported on Fig. 13a and discussed below.

At ca. 1000 Ma, the Tonian magmatism corresponds with the initial weakening of the boundary between the Congo Craton and Saharan cratons or just to the weakening of the Saharan lithosphere. Such ages are not well known regionally but preliminary age data document them (de Wit et al., 2021). Remnants are found as inherited grains reworked during Neoproterozoic magmatism of Chad (Djerossem et al., 2020; Shellnutt et al., 2019, 2017), and Lu-Hf isotopic dating provide the ev- idence of a magmatic event around 1.0 Ga in Central Cameroon (Saha- Fouotsa et al., 2021). These ages are also found in the detrital zircon record, either in Palaeozoic sediments of the Saharan metacraton (Stephan et al., 2019) or in the present-day Nile River sediments or coastal sediments (Be'eri-Shlevin et al., 2009; Be'Eri-Shlevin et al., 2014; Iizuka et al., 2013). These ca. 1.0 Ga ages are also recorded in Ediacaran to Palaeozoic sediments of the Variscan orogeny and the Arabian Peninsula. Our data support the suggestion made by several researcher the Saharan metacraton acted as a proximal source of sedi- ments following the formation of Gondwana (e.g. Chelle-Michou et al., 2017; Couzinié et al., 2019).

Tonian magmatism could be the first step to decratonization as proposed in Brazil (Ganade et al., 2021). This could lead the margin of the cratonic block being thinned, even fragmented in smaller pieces and eventually resulting in oceanization (Ganade et al., 2021; Toteu et al., 2022). These marginal blocks would then be involved in subsequent orogenic building as for the Variscan belt (Franke et al., 2017).

Between 700 and 610 Ma the area between the Congo and Saharan cratons was closing (Fig. 13b). Evidence for this is attested by arc magmatism in the Mayo Kebbi massif (Isseini et al., 2012) the Guéra massif (e.g. Shellnutt et al., 2020) and our own data from North Ouad- daï. From 635 Ma a metamorphic event is documented in the Ouaddaï massif. In the southern part this event is associated with the formation of S-type granite (Djerossem, 2018). We interpret this metamorphic event as being linked to the hot thermal structure and deformation found in an active margin, either in an arc or during collision (Currie and Hyndman, 2006; Syracuse and Abers, 2006). Noteworthy, the magmatic arc was formed in crust that was at least ca. 1.0 Ga as demonstrated by the

protolith age of the orthogneisses and the xenocryst found in the Pan- african granitoids. Based on the presence of Archean inherited zircon (up to 2.8 Ga in sample RMC0075) recycled in the arc and collisional rocks, we can even advocate for an older basement. This attests that the Saharan craton recycles older continental crust (e.g. Abdelsalam et al., 2002 and references therein).

Around 610–580 Ma, the margins of the two cratonic domains collided to form the so-called central African orogenic belt (Fig. 13c; Pin and Poidevin, 1987; Toteu et al., 2022, Toteu et al., 2006, 2004). This coincided with the end of arc magmatism on the southern margin of the Saharan craton margin as evidenced in the North Ouaddaï and Guéra massifs (this study; Shellnutt et al., 2017). During this period we document a HT metamorphic event that we relate to the collision of the two continental land masses, which led to the formation of Gondwa- naland (Küster et al., 2008; Meert and Lieberman, 2008).

After 580 Ma (Fig. 13c), the system was probably driven by the thermal relaxation and extension which were induced by a hypothetical yet likely, slab break-off upon collision. This is consistent with the geochemical characteristics of some arc series rock that plot near the triple point in the Nb/Th diagram proxy (Hildebrand and Whalen, 2014; Whalen and Hildebrand, 2019). Large scale thermal relaxation triggered the exhumation of a mid to lower crustal domain that underwent climax metamorphism slightly earlier as well as the beginning of post colli- sional magmatism. This period also marks the regionally recognized transition from a collisional to a post-collisional setting in the Central African Orogenic belt (Ngako et al., 2008) that is also documented in the Guéra massif (Shellnutt et al., 2020). Evidence from the Ouaddaï massif (shoshonitic granites cross-cutting the arc to collisional granitoids in our field area) attests that the region entered a post-collisional setting. This magmatism is well documented in the Guéra massif, with similar rocks ranging that range from ca. 575–540 Ma (Shellnutt et al., 2022, 2017) and in Northwest Cameroon (Archelaus Emmanuel et al., 2024). This episode marks the end of the amalgamation of the Western Gondwa- naland. This magmatism may be related to mantle-derived partial melting due to thermal perturbation induced by slab break-off, which can last for tens of millions years (Freeburn et al., 2017).

Overall, this scenario enables us to examine the processes involved in the destruction of cratonic lithosphere upon collision. Based on our data and the literature previously mentioned, it appears that the southern margin of the Saharan craton was an “Andean type” active margin for at least 100 Ma prior to continental collision with the Congo craton. The thermal regime associated with an active margin (including the back-arc region) is notably hotter than the one of a non-active margin. The destabilisation of the cratonic lithosphere during subduction has been hypothesized for South America, North America, and China (Boyce et al., 2016; He, 2014; Kušky et al., 2014), and this process provide a satisfactory explanation as why the lithosphere beneath the Saharan metacraton was partly removed during the Ediacaran following an initial event that triggered decratonization during the Tonian as attested by the orthogneiss basement. Subduction beneath the southern margin and the subsequent back-arc extension eroded part of the cratonic lith- osphere which was likely already thinned in response to Tonian mag- matism. As this process occurred simultaneously with the accretion of the former Saharan craton with to Tuareg and the Arabian Nubian shields (Fezaa et al., 2010; Fritz et al., 2013; Johnson et al., 2011; Liégeois et al., 2013), we can speculate that the destruction of the cratonic lithosphere, triggered by dehydration melting of the astheno- sphere was driven by a complex combination of thermal magmatic events, and of toroidal and poloidal subduction-induced flow beneath the former Saharan craton.

9. Conclusions

We present some original data constraining the geological knowl-
edge of the Saharan metacraton: (i) there was an active subduction
below the southern margin of the former Saharan craton during the
<figure><img src="image_17.png" /><figcaption>**Fig. 13.** (a) summary of ages obtained in the North Ouaddaï area compared to litterature data in Chad and Cameroun. Compilation of geochronolmogica data after Djerossem et al., 2020, 2021, de Wit et al., 2021, Isseini et al., 2012; Shellnutt et al., 2017, 2018, 2019, 2022, 2025 and Toteu et al., 2022. (b) schematic recon-struction at ca. 750–610 Ma; (c) schematic reconstruction at ca. 610–580 Ma; (d) schematic reconstruction at ca. 580–540 Ma. Abbreviations: Cc, Congo Craton; AYD, Adamawa Yadé domain; Smc, Sahara *meta*-craton.</figcaption></figure>
Neoproterozoic time in the interval 700–610 Ma as attested by magmatic and metamorphic events; this subduction event started to erode the cratonic lithosphere below the Saharan metacraton (ii) The collision of the Saharan craton with the northern margin of the Congo craton is attested by the eastwards termination of arc magmatism in the Guéra and Ouaddaï region and coeval HT-MP metamorphism around 610 Ma. At regional scale this formed the Central African orogenic belt; (iii) we hypothesise that the slab breakoff following collision helped to continue the partial erosion of the cratonic lithosphere. This process was accelerated due to the other lateral collision of the Saharan metacraton with the Western African Craton (i.e. leading to the Trans Saharan orogeny) and the collision with the Arabian Nubian Shield to the East and might explain why the Saharan craton has been more affected by metacratonization than any other areas in the world.

### CRediT authorship contribution statement

**Alexis Plunder:** Writing – original draft, Investigation, Formal analysis, Data curation, Conceptualization. **Olivier Blein:** Writing – original draft, Formal analysis. **Moussa Isseini:** Writing – review & editing. **Issaka Ousman Al-Gadam:** Project administration. **Mathieu Chevillard:** Writing – review & editing, Investigation. **Emanuel Djedouboum:** Writing – review & editing, Project administration. **Philippe Lach:** Data curation. **Abdeltif Lahfid:** Methodology, Formal analysis, Data curation. **Jérémie Melleton:** Writing – review & editing. Validation. **Olivier Rouzeau:** Project administration. **Guillaume Vic:** Formal analysis.

### Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

### Acknowledgements

AP is grateful to C. Duée and G. Wille for SEM and EPMA acquisition. Fruitful discussion during the preparation of this work with T. Fullgraf and L. van Unnik Hoorn were greatly appreciated. We thank Gregory Shellnutt and Soumyajit Mukherjee for inspiring and detailed review. We also thank Michel Villeneuve for his time and effort in reviewing the manuscript.

### Appendix A. Supplementary material

Supplementary data to this article can be found online at https://doi.org/10.1016/j.precamres.2025.108000.

### Data availability

data are available as supplementary material

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