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Home All issues Volume 196 (2025) BSGF - Earth Sci. Bull., 196 (2025) 20 Full HTML
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BSGF - Earth Sci. Bull.
Volume 196, 2025
Article Number 20
Number of page(s) 30
DOI https://doi.org/10.1051/bsgf/2025011
Published online 29 October 2025

© J. Vanardois et al., Published by EDP Sciences 2025

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

Quantitative Pressure–Temperature–time–Deformation (P–T–t–D) paths and geothermal gradients derived from the study of deformed metamorphic rocks are crucial for understanding tectonic and thermal processes (e.g. England and Thompson, 1986; Davy and Gillet, 1986; Gerbi et al., 2006) and for characterizing the thermal structure of orogens, whether cold or hot (e.g. Chardon et al., 2009). Such data serve as a foundation for comparison with numerical modelling predictions, advancing our understanding of orogenic dynamics (e.g., Sizova et al., 2019). However, reconstructing the complete P–T–t–D evolution, from early prograde to late retrograde stages, remains highly challenging, and consequently, such essential data are relatively scarce in the literature. Prograde assemblages are often partially or completely overprinted by diffusion or partially replaced during peak metamorphism or retrogression, obscuring key phases of the metamorphic history.

For robust P–T–t–D reconstructions, samples must meet two critical criteria. First, they must exhibit petrological characteristics that preserve evidence of the metamorphic evolution. Second, they must be anchored within a well-constrained structural framework, with the strain pattern documented at both local and regional scales. This dual requirement enables the correlation of P–T data with deformation stages, including early, syn-, and late-orogenic events. Additionally, temporal calibration must be supported by in-situ dating of accessory minerals, such as zircon, monazite, or titanite, to construct a comprehensive P–T–t–D path.

The currently accepted tectonic and geodynamic model for the Variscan belt evolution posits that crustal thickening was driven by the stacking of three major tectonic units through crustal-scale thrusting (Burg and Matte, 1978; Ledru et al., 1989; Faure et al., 2009; Martínez-Catalán et al., 2021). This model has been continuously refined, with recent petrochronological studies suggesting that high-pressure metamorphism occurred during the Late Devonian to Early Carboniferous (Lotout et al., 2018, 2020; de Hoÿm de Marien et al., 2023). The Variscan orogen is non-cylindrical, with crustal thickening to depths of 55–65 km and associated surface uplift to elevations of 3–5 km thought to have propagated westward between 350 Ma and 325 Ma (Hillenbrand and William, 2022).

In the variscan External Crystalline Massifs (ECMs; Fig. 1A), recent studies have provided better constraints on the P–T–t–D paths of mafic eclogites (Jouffray et al., 2020; Jacob et al., 2021, 2022; Vanardois et al., 2022a; Filippi et al., 2024) and metapelites (Fréville et al., 2018, 2022; Jacob et al., 2021), but with only few constraints on the prograde history. In the ARM (Fig. 1B), a diverse range of prograde P–T paths has been documented, including anticlockwise P–T paths in the southwestern ARM (Fig. 1C) (Dobmeier, 1996, 1998) and clockwise P–T paths in the central ARM (Fig. 1D) (Joye, 1989; Schulz and von Raumer, 1993, 2011). However, these studies rely on empirical thermobarometry and EPMA monazite dating, which have large uncertainties.

To refine our understanding of the petrochronological record of both the upper and lower crust in the ARM, this study reports a detailed reconstruction of the P–T–t–D paths of two metapelite samples (AR736 and AR14) from the southwestern and central ARM. The aim is to provide a more detailed and comprehensive history of their metamorphic and tectonic evolution based on structural and petrological observations, phase equilibrium modelling, trace element geochemistry, and in-situ LA-ICP-MS U–Th–Pb monazite dating.

thumbnail Fig. 1

(A) Paleogeographic reconstitution of the European Variscan Belt at the end of the Variscan orogeny (ca. 290 Ma) modified from Franke et al. (2017). (B) Geological map of the Aiguilles-Rouges and the Mont-Blanc Massifs with location of the two metapelites samples analysed and the LA-ICP-MS U-Th-Pb monazite ages obtained. Various P–T paths proposed for the Aiguilles-Rouges massif: (C) Anticlockwise P–T paths of metapelites in the Prarion and Lac Cornu areas after Dobmeier (1996, 1998). (D) Clockwise P–T paths of metapelites from the Val Bérard and Lac Emosson areas after Schulz and von Raumer (1993, 2011), Joye (1989) and Genier et al. (2008). Black arrows represent sections of P–T paths constrained by thermobarometric results and grey arrows are interpreted sections.

2 Geological setting

The ARM, which belongs to the ECMs (Fig. 1A), is a 45-km-long, NE-SW trending Variscan basement massif, bounded by para-autochthonous and allochthonous Mesozoic and Cenozoic sedimentary rocks (Fig. 1B). The ARM is composed of a gneissic basement consisting of metapelites (micaschists and paragneisses) and Ordovician orthogneisses (von Raumer and Bussy, 2004; Bussy et al., 2011), a Devonian to late Carboniferous basin (Servoz basin; Bellière and Streel, 1980; Vanardois et al., 2022c) and a late Carboniferous basin (Salvan-Dorenaz basin; Capuzzo and Bussy, 2000). Several Carboniferous granites intruded the gneissic basement and the Devonian-Carboniferous Servoz basin (Bussy et al., 2000) (Fig. 1B). In the ARM, overprinting of Alpine greenschist facies metamorphism is scarce and mostly localised in shear zones (e.g. Rolland et al., 2003; von Raumer and Bussy, 2004; Rossi et al., 2005; Boutoux et al., 2016), which is conducive to the preservation of Variscan metamorphism (e.g. Schulz and von Raumer, 1993, 2011; Vanardois et al., 2022c) and deformation (e.g. Simonetti et al., 2020; Vanardois et al., 2022b).

The gneissic basement of the ARM underwent a medium to high-P metamorphism, preserved in kyanite-bearing metapelites (von Raumer, 1983; Schulz and von Raumer, 1993). In the southwestern part of the ARM, the peak metamorphic conditions were estimated to be 1.1 GPa and 650 °C, and 0.8 GPa and 600 °C in paragneisses from the flank of the Servoz syncline, with an anticlockwise P–T path (Fig. 1C) (Dobmeier, 1996, 1998). In the uppermost structural level, Raman spectrometry on carbonaceous material analyses of Devonian − Visean metasedimentary rocks from the Servoz syncline associated with zircon LA-ICP-MS U-Th-Pb dating emphasises a high geothermal gradient at ca. 350 Ma, interpreted as a back-arc setting before the onset of crustal thickening (Vanardois et al., 2022c). Crustal thickening has been attributed to nappe-stacking (Dobmeier, 1998), but structural data to confirm this hypothesis are lacking due to reworking during transcurrent tectonism.

In the metasedimentary rocks from the central part of the ARM (Val Bérard and Emosson areas), a clockwise P–T evolution is recorded with an estimation of maximum pressure conditions at ca. 0.8–1.25 GPa/500–650 °C in metapelites (Fig. 1D) (von Raumer, 1983; Joye, 1989; Schulz and von Raumer, 1993, 2011; Genier et al., 2008). Peak metamorphism in the metasedimentary rocks of Lac d’Emosson occurred at 327 ± 2 Ma, as indicated by monazite ID-TIMS dating (Bussy et al., 2000), and between 330 and 300 Ma, indicated by EMP-CHIME Th–U–Pb monazite dating in the Val Bérard area (Schulz and von Raumer, 2011). Mafic eclogite from the Lac Cornu embedded in migmatites records a pressure peak of 1.75 GPa at 700 °C at ca. 340-330 Ma as determined by zircon and rutile U-Pb LA-ICP-MS dating (Vanardois et al., 2022a).

The gneissic basement locally shows evidence of partial melting (von Raumer and Bussy, 2004; Genier et al., 2008; Vanardois et al., 2022b, 2024a). The crystallisation age of anatectic melts formed during decompression was determined by zircon and monazite ID-TIMS U-Pb dating at 320 ± 1 Ma near Lac d’Emosson (Bussy et al., 2000), 317 ± 2 Ma in the Mont-Blanc massif (Bussy and von Raumer, 1994) and 307 ± 2 Ma in the Fully area in the northeastern part of the ARM (Bussy et al., 2000).

The exhumation of the lower crust was initially accommodated by a horizontal flow (D1) forming shallowly dipping foliations S1 (Vanardois et al., 2022a). This deformation stage was followed by a dextral transpressive regime (D2) forming N-S vertical S2 foliations and C2-C’2 shear zones (e.g. Bellière, 1958; von Raumer and Bussy, 2004; Simonetti et al., 2020; Vanardois et al., 2022b). Transpressional shear zones nucleated at 340–330 Ma (Vanardois et al., 2022b) and were active at least between 320 and 300 Ma (monazite EPMA chemical U-Th-Pb; Simonetti et al., 2020). The tectono-metamorphic evolution of the ARM is also characterised by two well-constrained magmatic pulses. A first Visean magmatic event at 340–330 Ma is responsible for the emplacement of high-K calc-alkaline to shoshonitic plutons such as the Montées-Pélissier and Pormenaz granites (Bussy et al., 2000; Vanardois et al., 2022b). A second Late-Carboniferous event at ca. 305 Ma is associated with the emplacement of peraluminous granites such as the Morcles and Vallorcine granites in the ARM and the Mont-Blanc and Montenvers granites in the Mont-Blanc massif (Bussy and von Raumer, 1994; Bussy et al., 2000; Bussien et al., 2017 (Fig. 1B).

3 Analytical methods

3.1 Major element geochemistry

X-ray maps were acquired using an electron probe micro-analyser (EPMA) instrument (JEOL JXA-8530F HyperProbe) at the University of Lausanne with dwell times of 40 ms, a current of 300 nA and a voltage of 15.0 kV, and an EPMA (JEOL JXA-8200 Superprobe) at the University of Bern with dwell times of 80 ms, a current of 100 nA and a voltage of 15 kV. X-ray maps were standardised to oxide weight percentage maps using an internal standard approach and the XMapTools 4.4 program (Lanari et al., 2014, 2019). Structural formula compositional maps expressed as atoms per formula unit (apfu) were calculated and used to investigate garnet compositional variability at the thin section scale. Measured garnet, muscovite, biotite and staurolite compositions are given in Table 1. Garnet endmembers contents in mol% extracted from XMapTools are given in Table 2. Whole-rock major elements for samples AR14 and AR736 (Tab. 3) were obtained using a PANalytical AxiosmAX X-Ray Fluorescence (XRF) spectrometer at the University of Lausanne (Switzerland), and a micro-XRF (M4 Tornado, Bruker) at the University of Strasbourg (France), respectively. For the XRF analyses, quality control was performed on the reference materials BHVO2 and JA-3.

Table 1

Representative EPMA analytical spots for compositions of garnet (Grt), staurolite (St), biotite (Bt) and muscovite (Ms). Xalm: almandine; Xprp: pyrope; Xsps: spessartine; Xgrs: grossular.

Table 2

Garnet endmembers contents extracted from XMapTools.

Table 3

Bulk compositions of samples AR736 and AR14.

3.2 Trace element analysis

Trace element analysis of garnet, monazite and rutile was conducted using LA-ICP-MS with a Resonetics RESOlutionSE 193 nm excimer laser system, equipped with an S-155 large-volume constant-geometry chamber (Laurin Technic, Australia), at the Institute of Geological Sciences, University of Bern, Switzerland. The laser system was coupled to an Agilent 7900 quadrupole ICP-MS instrument. Ablation was performed in an atmosphere of pure He (0.4 L/min) and N2 (0.003 L/min) mixed with Ar (0.86 L/min) immediately after the ablation cell. Laser parameters, analyzed elements with their relative dwell times, and standards are detailed in Table S1 and S2. For garnet, a spot size of 30 µm was used, with GSD-1 g (primary) and NIST 612 (secondary) as standards. Stoichiometric Al was used as an internal standard for garnet (Al2O3: 21 wt%). The spot size on the monazite crystals was 16 µm and the primary and secondary standards used were NIST-610 and NIST-612, respectively. Stoichiometric P2O5 (30 wt%) was employed as an internal standard. We used Al and Mg contents to ensure that the analyses did not include the host mineral. REE analyses were performed on the same zones of dated monazite crystals to infer the geochemical environment during mineral crystallization. Reproducibility and accuracy were within the 10% for all analysed elements. Data were reduced using Iolite software (Hellstrom et al., 2008; Paton et al., 2011) with the trace-element reduction scheme of Woodhead et al. (2007). REE contents were normalized to the chondritic values of McDonough and Sun (1995) (Tab. S1).

3.3 Zirconium-in-rutile thermometry

The rutile crystallisation temperature of sample AR736 was determined using the Zirconium-in-rutile (Zr-in-Rt) thermometer (see Table S2 for analytical conditions). The instrument used was that described in Section 3.2. The spot size was 24 μm and the primary and secondary standards were NIST-610 and NIST-612. Stoichiometric Ti was employed as the internal standard for rutile (TiO2: 99 wt%). The Kohn (2020) calibration was used because this formulation includes a pressure-dependent term. Rutile partially replaced by ilmenite was avoided as this can affect the Zr content in rutile (e.g. Whitney et al., 2015). Based on the recommendations of Kohn (2020), an absolute uncertainty of ± 30 °C is considered for each data point.

3.4 Titanium-in-Biotite thermometry

Biotite compositions were measured in both samples using an EPMA (JEOL JXA-8530F HyperProbe) at the University of Lausanne. The dwell time was 40 ms, the current was 3 nA and the voltage was 15.0 kV. Due to biotite chloritization, only two analyses of sample AR14 yielded consistent biotite compositions. The biotite crystallization temperature in sample AR14 was determined using the titanium-in-biotite thermometer (Ti-in-Bt) of Henry et al. (2005).

3.5 U-Th-Pb geochronology

Monazite was analysed in-situ within thin sections to preserve the textural relationships between accessory mineral distribution and fabric. Prior to analysis, backscatter electron (BSE) images were acquired for all grains using a scanning electron microscope (SEM) to check spot positions with respect to the internal microstructures, inclusions, fractures or other physical defects.

U-Th-Pb geochronology on monazite was conducted by LA-ICP-MS at the Laboratoire Magmas et Volcans, Clermont-Ferrand (France). Analyses involved ablation of grains with a Resonetics Resolution M-50 powered by an ultra-short pulse ATL Atlex Excimer laser system operating at a wavelength of 193 nm. Detailed analytical procedures are described in Paquette et al. (2014) and detailed in Hurai et al. (2010) and in the Supplementary material (Appendix 1). Data reduction was carried out using the GLITTER® software package from Macquarie Research Ltd (van Achterbergh et al., 2001). Dates and diagrams were generated using Isoplot/Ex v. 2.49 software package (Ludwig, 2001). All uncertainties are given at ± 2σ.

3.6 Phase equilibrium modelling

Pressure–Temperature phase diagrams were computed in the compositional system MnO–Na2O–CaO–K2O–FeO–MgO–Al2O3–SiO2–H2O–TiO2 (MnNCKFMASHT) using Perple_X 7.1.6 software (Connolly, 2005) and the hpver62 thermodynamic database from Holland and Powell (2011). Solution models used in the calculations are biotite, chlorite, chloritoid, cordierite, garnet, melt, orthopyroxene, staurolite and white mica (White et al., 2014a, 2014b) as well as epidote, K-feldspar and plagioclase (Holland and Powell, 2011), ilmenite (White et al., 2000) and spinel (White et al., 2002). Bulk rock compositions of both samples were corrected for the presence of apatite and pyrite using P2O5 and S contents (Tab. 3).

Peak mineral assemblages observed in samples AR736 and AR14 are interpreted to have formed under subsolidus conditions and therefore the phase diagrams were calculated under H2O-saturated conditions. T–X(Fe3+) diagrams were calculated, highlighting the absence of co-stability of plagioclase and rutile in sample AR736 and the disappearance of rutile in sample AR14 as soon as a small fraction of Fe3+ was included in the models (Fig. S1). Therefore, all Fe was considered as to be Fe2+ (i.e. FeO) in the model. This assumption is consistent with the presence of pyrite in the samples (Figs. 2A and 3A).

P–T equilibrium conditions were estimated using the third model quality factor Qcmp described in Duesterhoeft and Lanari (2020). The Qcmp factor evaluates the quality of the fit between measured and modelled mineral compositions, including the relative analytical uncertainty of the measured compositions. This method was applied to garnet and we compared its Xalm, Xsps, Xprp and Xgrs contents (Tab. 2). Given the small amount of garnet in both samples (<5 vol%), the potential effects of garnet fractionation on the reactive bulk composition during growth were not considered (Lanari and Engi, 2017).

4 Results

4.1 Petrography and structures

4.1.1 Bellachat metapelite (AR736)

Sample AR736 is a micaschist from the southwestern part of the ARM gneissic basement (Fig. 1B). It consists of muscovite (40% vol.), biotite (25%), quartz (20%), plagioclase (5%), staurolite (5%) and garnet (5%) with minor amounts of apatite, pyrite, zircon, ilmenite and rutile (Fig. 2A,B). Garnet porphyroblasts, up to 2 mm in diameter, are fractured and sub-euhedral or oblong (Fig. 2A,B) and contains widespread small muscovite inclusions (Fig. 2C). Garnet rims also feature biotite, plagioclase, apatite and rutile inclusions (Fig. 2C-D). Garnet fractures are filled with biotite and muscovite (Fig. 2B). Staurolite porphyroblasts are composed of a dark core (St1) and thin and discontinuous bright zones (St2) (Fig. 2F-G). The staurolite cores are mainly replaced by microscopic muscovite and biotite (Fig. 2G), which is responsible for its dark colour in planed polarized light. The bright St2 zones are free of inclusions and are sometimes observed in the fractures of St1 porphyroblasts. Rutile grains occur as inclusions in staurolite cores (Fig. 2F-G). In the matrix, mica-rich and Qz-Pl-rich layers parallel to the preferential orientation of micas define a foliation corresponding to S2 main regional fabric. In thin sections cut perpendicular to the foliation and parallel to a stretching lineation marked by the alignment and elongation of Qz-Ms-Bt, tails around garnet and staurolite form pressure shadows and Qz-Ms-Bt define S-C structures. These criterions consistently indicate a dextral sense of shear. Inclusions trails in the garnet and staurolite porphyroblasts indicate either dextral rotation or preservation of an anterior folded foliation. Ilmenite is observed in the matrix while rutile is not. Biotite is partially replaced by chlorite (Fig. 2F). Plagioclase is sericitized and contains rutile inclusions (Fig. 2E).

thumbnail Fig. 2

Thin section and back-scattered electron images showing petrological characteristics of the Bellachat micaschist (AR736): (A) Main paragenesis of the sample with garnet and staurolite porphyroblasts in mica-plagioclase-quartz matrix. Sigmoids and tails around garnet highlight dextral kinematics. (B) Main paragenesis of the sample with dextral C-S structures corresponding to the S2 and C2 planar fabrics. (C and D) Inclusions in the outer part of garnet grains. (E) Partial replacement of biotite by chlorite and rutile inclusions in a plagioclase. (F) Optical differences between the dark staurolite core (St1) and the bright staurolite zones (St2). (G) Fractured staurolite core (St1) with bright staurolite zone (St2) in the fracture. Mineral abbreviations are from Warr (2021).

4.1.2 Lac Emosson metapelite (AR14)

Sample AR14 is a micaschist located near Lac Emosson (Fig. 1B). It is composed of muscovite (55 vol%), quartz (15%), biotite (15%), plagioclase (10%), garnet (4%), fibrolitic sillimanite (1%) and accessory minerals such as ilmenite, rutile, zircon, xenotime, pyrite and apatite (Fig. 3). Garnet grains are small (ca. 1 mm wide), rounded to oblong, and exhibit subhedral to anhedral shapes. They contain numerous inclusions of biotite, muscovite, quartz, apatite and plagioclase, as well as small rutile crystals (10–20 μm) that are absent from the matrix. Some garnet grains are fractured and/or replaced by chlorite. In the matrix, mica-rich, and Qz-Pl-rich layers parallel to the preferential orientation of micas and sillimanite define a foliation corresponding to S2 main regional fabric. In thin sections cut perpendicular to the foliation and parallel to a stretching lineation marked by the alignment and elongation of Sil-Qz-Ms-Bt, dextral S-C structures are defined by Qz-Ms-Bt-Sil. Some of garnet grains show asymmetric strain shadows of biotite and quartz, which indicate dextral kinematics. Biotite grains within these strain shadows show less chlorite replacement compared to those in the planar fabrics. Ilmenite is observed in the matrix while rutile is not.

thumbnail Fig. 3

Thin section images showing petrological characteristics of the Emosson micaschist (AR14). (A and B) Main paragenesis of the sample with garnet porphyroblasts in a mica-plagioclase-quartz matrix, with dextral S-C structures. (C and D) Plane light and cross light images of fibrolitic sillimanite in the matrix. Note that the biotite is partially chloritized.

4.2 Minerals chemistry

4.2.1 Major elements

In sample AR736, a rounded sub-euhedral garnet (Fig. 4A) and an oblong-anhedral garnet (Fig. 4B) were mapped. Compositional maps of garnet show a moderate compositional zoning (Fig. 4A-C, Tabs. 1 and 2). In both garnets, composition evolves from core (Grtcore) to mantle (Grtmantle) with an increase of Xalm (0.73 to 0.78) and Xprp (0.15 to 0.16) contents and a decrease of Xsps (0.05 to 0.02) and Xgrs (0.09 to 0.06) contents (Fig. 4). The outer part of the grain displays a narrow rim (Grtrim), about 10–20 μm width, with high Xsps (0.08) and low Xprp (0.09) contents (Fig. 4, Tabs. 1 and 2). Muscovite has low Na contents (0.03 apfu) and Si contents ranging between 3.04 and 3.07 apfu with a mean value of 3.05 (n = 4) (Tab. 1). Staurolite cores (St1) show domains with relatively high-XMg values ranging between 0.14 and 0.17, surrounded domains with lower XMg values of 0.1 (Fig. 5). Bright staurolite zones (St2) show more homogeneous domains with XMg values between 0.12 and 0.15 and narrow edges with lower values (Fig. 5). Biotite and plagioclase crystals show local evidence of chloritization and sericitization.

BMI=weight(kg)height(m2).

In sample AR14, four garnet crystals, two oblong-anhedral (Fig. 6A-C) and two sub-euhedral grains (Fig. 6B), were mapped. All grains show a similar composition evolution a nearly homogeneous core (Grtcore) and mantle (Grtmantle) (Xalm 0.72–0.75; Xsps 0.03–0.05; Xgrs 0.06–0.09; Xprp 0.14–0.18) (Fig. 6 and 7; Tabs. 1 and 2). In the garnet elongation, a rim (Grtrim1) with slightly higher Xgrs and lower Xalm contents (Xgrs 0.07–0.12; Xalm 0.68–0.72) is visible in all grains. The transition from mantle to Grtrim1 is either sudden (Fig. 6A and 7) or progressive (Fig. 5B and 6). In three of the four grains, a thin second rim (Grtrim2) with low Xgrs (0.03–0.05) and high Xalm (0.71–0.77) contents is visible with a sharp transition with Grtrim1 (Fig. 6 and 7, Tabs. 1 and 2). Garnet edges are surrounded by a thin Xsps-rich rim (Grtrim3; Xsps 0.19–0.28). Muscovite has low Na content (0.06 apfu) and Si contents ranging between 3.04 and 3.05 apfu (n = 6) (Tab. 1), one data shows a higher value at 3.08. Biotite shows XMg values ranging from 0.33 to 0.39 with a mean value of 0.36, and Ti contents between 0.10–0.12 (afpu) and a mean value of 0.11 (n = 5) (Tab. 1).

thumbnail Fig. 4

Garnet compositional maps and profiles for the Bellachat micaschist (AR736). End-member fraction compositional maps of (A) round sub-euhedral garnet and (B) oblong anhedral garnet. (C) Chemical profiles of the garnet grains. Xalm : almandine fraction; Xgrs: grossular fraction; Xprp: pyrope fraction; Xsps: spessartine fraction. Representative mineral compositions are provided in Tables 1 and 2.
thumbnail Fig. 5

Compositional map of staurolite in sample AR736. Location of maps are shown in Figures 2F and 2G.
thumbnail Fig. 6

Garnet compositional maps and profiles for sample AR736. End-member fraction compositional maps of four garnet grains. Profiles are visible in Figure 7. Representative mineral compositions are provided in Tables 1 and 2.
thumbnail Fig. 7

Chemical profiles of the garnet grains presented in Figure 6. Fig. 8: WDS maps of Y content and profiles in garnet from samples (A) AR736 and (B) AR14. Locations of LA-ICP-MS analyses in garnet and garnet REE patterns normalized to the chondrite values of McDonough and Sun (1995) from samples (C) AR736 and (D) AR14.

4.2.2 Trace element systematics in garnet and rutile

Qualitative WDS (Wavelength Dispersive Spectrometer) X-ray maps of Y content (Fig. 8A and B) and LA-ICP-MS trace element analyses (Fig. 8C and D) were performed on one garnet grain from each sample. In sample AR736, the Y map shows a relatively high Y content in Grtcore compared to Grtmantle (Fig. 8A). Of the fourteen LA-ICP-MS trace element analyses, ten were located in Grtcore, two in Grtmantle and two in Grtrim. The results confirm the Y distribution with 800–2069 μg/g in Grtcore and 1138–1144 μg/g in Grtmantle. The chondrite-normalized REE patterns of the three garnet zones (Grtcore, Grtmantle and Grtrim) overlap for heavy-REE (HREE), with Grtrim being slightly more enriched in light-REE (LREE). All garnet zones are characterized by steep slopes from the LREE to middle-REE (MREE) (DyN/SmN = 66–183) and moderate slopes from MREE to HREE (LuN/DyN = 2–8) with a moderate negative Eu anomaly (0.61–0.75) (Fig. 8C). The thirty-one analyses of Zr content carried out on thirty-one rutile crystals (Tab. S4) with results from 93 to 283 μg/g with a mean value of 195 μg/g and a standard deviation of 118 μg/g (2σ, n = 31).

In sample AR14, the Y content of garnet is lower than in sample AR736 and slightly decreases from the core to the rim (Fig. 8B). Eleven trace element analyses were also obtained in the garnet grain with five spots in Grtcore, four spots in Grtmantle and two in Grtrim1 (Fig. 8D; Tab. S3). The LA-ICP-MS results confirm the Y distribution in garnet with 234–449 μg/g in Grtcore, 217–255 μg/g in Grtmantle and 86–88 μg/g in Grtrim1. The REE patterns of Grtcore and Grtmantle are similar, with the exception of the higher HREE content in Grtcore (Fig. 8D). They are characterized by a moderate slope from LREE to MREE (DyN/SmN = 25–41) and various slopes from MREE to HREE (LuN/DyN = 1–23) with a strong negative Eu/Eu* anomaly (0.37–0.51). The Grtrim1 REE patterns are characterized by weaker slopes from LREE to MREE (DyN/SmN = 5–6) and negative slopes from MREE to HREE (LuN/DyN = 0.45 (± 0.0004)) with moderate negative Eu/Eu* anomaly (0.86 (± 0.001)) (Fig. 8D; Tab. S3).

4.3 Petrological interpretations

In sample AR736, the garnet core (Grtcore) contains only quartz and muscovite inclusions that are interpreted as a partial preservation of a prograde mineral assemblage (MA: Grtcore + Qz + Ms). The Grtmantle contains numerous inclusions of biotite, muscovite, plagioclase, quartz, rutile and apatite (Fig. 2C-D), which define the peak mineral assemblage (MB). Based on the presence of rutile inclusions observed in staurolite core (Fig. 2F-G), it is proposed that staurolite— corresponding to the preserved core St1—also belongs to the MB assemblage, which consists of Qz + Pl + Ms + Bt + St1 + Rt + Ap + Grtmantle (Tab. 4). During retrograde evolution (MC), staurolite cores are partially replaced by muscovite and biotite (Fig. 2G). Given the scarcity of rutile in the matrix and the presence of ilmenite, it is proposed that rutile was replaced by ilmenite during MC. Late retrograde evolution (MD) is characterized by the crystallization of bright staurolite overgrowths St2 around the staurolite cores St1 or sometimes within fractures affecting the cores (Fig. 2F-G), and by the partial replacement of biotite by chlorite.

In sample AR14, garnet core and its inclusions define the mineral assemblage of the metamorphic peak (Mb: Grtcore-mantle + Ms + Bt + Qz + Pl + Rt + Ap; Tab. 4). Given the absence of rutile in the matrix and the presence of ilmenite, it is very likely that rutile was replaced by ilmenite. Ilmenite and sillimanite are observed only in the matrix, suggesting crystallisation after the Mb mineral assemblage and emphasising posterior metamorphism (Mc). Due to the absence of inclusions within Grtrim1 and Grtrim2, we do not know if they belong to the Mb and/or Mc mineral assemblages. Finally, the replacement of biotite by chlorite highlights replacement during late retrograde evolution (Md). Regarding the anhedral shape of most garnet grains, it is likely they were partially resorbed before or during Md.

In both samples, the Xsps-rich rims (Grtrim and Grtrim3) can be the result of an incorporation of Mn by diffusion rather than a garnet overgrowth (Yardley, 1977; Kohn and Spear, 2000).

Table 4

Petrological interpretations. The prograde mineral assemblage of sample AR736 is based on garnet inclusions and is therefore probably incomplete.

4.4 Estimation of Pressure-Temperature conditions

4.4.1 Bellachat metapelites (AR736)

The incomplete preserved prograde mineral assemblage (MA: Grtcore + Qz + Ms) is modelled in all the computed phase diagram, except at low-P and high-T where the muscovite breakdowns (Fig. 9A). Even if the Grtcore composition (Alm0.74Prp0.13Grs0.09Sps0.04) is not perfectly modelled in our phase diagram, isopleths and Qcmp factor converge towards a mineral assemblage of Grt + Pl + Ilm + Ms + St + Qz + Bt/Chl (Figs. 9B-C  and 10). Considering the uncertainties extracted as standard deviation from XMapTools, Grtcore is best modelled at 550–625 °C and 0.5–0.6 GPa (Fig. 10, Table 2).

The inferred peak metamorphic assemblage (MB: Grtmantle + Bt + Pl + Rt + Qz + Ms + Stcore) is modelled at 600–640 °C and 0.76–0.82 GPa (Fig. 9A, stability field with Grt Pl Bt Ms St Qz Rt). The stability field is delimited by the appearance of kyanite at higher-temperature, the disappearance of plagioclase at lower temperature, the disappearance of biotite at higher-pressure and the disappearance of rutile at lower-pressure (Fig. 9A). Grtmantle isopleths and Qcmp maps indicates P–T conditions of ∼620–640 °C and 0.75 GPa in the stability field of Grt + Pl + St + Ms + Bt + Ilm + Qz ± Rt (Figs. 9B-C  and 10). The Zr-in-rutile temperature values range between 527 and 615 °C at a pressure of 0.8 GPa (Fig. 9D). The sixteen data with the highest Zr contents yield a mean temperature of 602 ± 17 °C (2σ, n = 16). This temperature is consistent with the modelled P–T conditions for the metamorphic peak assemblage (Fig. 9D). Measured Si content of muscovite (i.e. 3.05 afpu, Tab. 1) is consistent with the MB field (Fig. 9E).

In the phase diagram, rutile included in garnet rims and ilmenite from the matrix are modelled above and below 0.8 GPa, respectively (Fig. 9A). The XMg value of St1 ranges from 0.14 to 0.17, indicating pressure conditions between 0.45 to 0.65 GPa, which is outside the predicted rutile stability field (Fig. 9F). The replacement of St1 by Bt + Ms can occur during decompression when the system approaches the stability of sillimanite where the modelled St volume decreases and the Bt volume increases (Fig. 9E-F). The XMg value of St2 ranges from 0.12 to 0.15 and are modelled below 0.6 GPa and above 550 °C (Fig. 9F). Chlorite, which locally replaces biotite, is predicted to be stable at lower temperature conditions of 575–515 °C (Fig. 9A).

thumbnail Fig. 8

WDS maps of Y content and profiles in garnet from samples (A) AR736 and (B) AR14. Locations of LA-ICP-MS analyses in garnet and garnet REE patterns normalized to the chondrite values of McDonough and Sun (1995) from samples (C) AR736 and (D) AR14.
thumbnail Fig. 9

Pressure − Temperature phase diagram constructed for the sample AR736 and contoured with compositional and modal isopleths. Bulk composition used is available in Table 3. (A) P–T mineral assemblage diagram with main reaction curve highlighted. MB, MC and MD mineral assemblage stability fields are in bold. (B and C) Xalm, Xgrs, Xprp and Xsps isopleths of garnet. Dotted arrow represents the rough path inferred from garnet compositions in Tables 1 and 2. (D) Zr-in-Rt thermometry results. Each red line represents a Zr-in-Rt analysis, the red area represents the temperature calculated from the sixteenth data with the highest Zr content. (E) Si apfu isopleths of muscovite and modal proportion of biotite expressed as vol%. (F) XMg compositional and modal isopleths for staurolite.
thumbnail Fig. 10

Results of the Qcmp factor of Grtcore and Grtmantle for the garnets presented in Figure 4. Compositional values used for the calculation are given in Table 2.

4.4.2 Lac Emosson metapelite (AR14)

The Mb mineral assemblage (Grt + Pl + Ms + Rt + Qz + BtQz) defined by garnet and its inclusions is predicted to be stable at pressure conditions above 0.9 GPa and temperatures between 610 and 720 °C (Fig. 11A; stability field with Grt Pl Ms Rt Qz Bt). At higher temperatures, a melt phase is predicted to be stable, whereas plagioclase only appears above 610 °C in this pressure range. At lower pressure conditions, rutile is not predicted to be stable and ilmenite and staurolite are predicted to be stable. Under lower temperature and pressure conditions, biotite and staurolite are no longer predicted to be stable and chlorite appears (Fig. 11A). The Ti-in-Bt thermometer yields a temperature of 578 °C (± 35 °C, 2σ) (Fig. 11A). The observed compositions of the garnet core (Grtcore) and mantle (Grtmantle) are best modelled in the stability field of Grt + Pl + Ms + Rt + Qz + Bt + St between 0.7 and 0.9 GPa and 600 and 675 °C (Fig. 12A-B). The increase in Xgrs and the decrease in Xalm in Grtrim1 indicate a core to rim decompression trend (Fig. 11C-D). Grtrim1 composition is modelled at 0.6 GPa and 585 °C within a mineral assemblage of Grt + Ms + Bt + Ilm + Qz + Pl + St (Fig. 12). The Mc mineral assemblage (Grt + Pl + Ms + Bt + Sil + Ilm) is predicted at pressures between 0.3 and 0.7 GPa and temperatures between 575 and 700 °C (stability field with Grt Sil Ms Pl Ilm Bt; Fig. 11A). This increase of temperature conditions is consistent with the decrease of Xgrs in Grtrim2 (Fig. 11C). The Grtrim2 composition is modelled at 0.55 GPa and 620 °C within the stability field of Mc mineral assemblage Grt + Ms + Bt + Ilm + Qz + Pl + Sil/St (Fig. 12D). Chlorite, which locally replaces biotite, is predicted to be stable at lower temperature conditions of 575–525 °C (Fig. 11A).

thumbnail Fig. 11

Pressure − Temperature phase diagram constructed for the sample AR14 and contoured with compositional and modal isopleths. Bulk composition used is available in Table 3. (A) P–T mineral assemblage diagram with main reaction curve highlighted. Mb, Mc and Md mineral assemblage stability fields are in bold. (B) Si apfu isopleths of muscovite and Ti-in-Bt thermometry results. (C and D) Xalm, Xgrs, Xprp and Xsps isopleths of garnet. Dotted arrow represents the rough path inferred from garnet compositions in Tables 1 and 2.
thumbnail Fig. 12

Results of the Qcmp factor of Grtcore, Grtmantle, Grtrim1 and Grtrim2 for the garnets presented in Figure 6. Compositional values used for the calculation are given in Table 2. The diagram Grtrim2 for Map2-G2 is missing because this compositional zone has not been observed in this grain.

4.5 Monazite U-Th-Pb geochronology and trace element compositions

4.5.1 Bellachat metapelite (AR736)

Monazite is abundant in the matrix, with anhedral grains ranging in size from 50 to 100 μm. It occurs both in the biotite-muscovite layers of the matrix (Mz1, 2, 4, 5, 6, 7, 8 and 17) (Fig. 13A-B), in the fractures within staurolite cores (Mz grains 3, 14, 15 and 16) (Fig. 13C-D) and as inclusions (Mz 11, 12 and 18) in plagioclase crystals (Fig. 13E-F). Notably, monazite was not observed as inclusion in garnet. Most monazite grains lack clear evidence of chemical zoning and are fractured (Fig. 13A, C, E and F ). They contain numerous inclusions of quartz, biotite, plagioclase and muscovite (Fig. 13A-Fig. 13A). A total of twenty-two analyses were performed on fifteen monazite grains. The results show Pb, U and Th contents ranging between ∼1300–2200 μg/g, ∼4000–6300 μg/g and ∼30,000–55,000 μg/g, respectively and with Th/U ratios between 6.4 and 11.7 (Tab. S5). The Tera-Wasserburg diagram displays a complex pattern showing discordant data with a percentage of concordance ranging from 52 to 95 % due to common Pb contamination and with a dispersion of 206Pb/238U dates between ∼350 and 300 Ma (Tab. S5; Fig. 14A). In the 206Pb/238U vs. 208Pb/232Th diagram, most of the data are discordant and do not allow the calculation of single-spot dates (Fig. S2A). A 207Pb-based common Pb correction was applied to the U-Pb dataset using the initial-Pb composition predicted (at 315 ± 10 Ma) by the two-stage crustal evolution model of Stacey and Kramers (1975). The corrected dates range between 339 and 298 Ma (Tab. S5) and reveal a correlation between the common-Pb corrected dates and the textural position of monazite (Fig. 14B). Among the ten monazite crystals from the biotite-muscovite layers, nine yield a weighted average common-Pb corrected date of 307 ± 4 Ma (MSWD = 0.45). However, monazite Mz6 (spot #9), which is also a grain from the Bt-Ms layers and is located between two rutile grains (Fig. 13B) gives an older date of 330 ± 12 Ma. Eight analyses were performed on four monazite grains located in staurolite fractures. Two monazite grains (Mz15 and Mz16 or spots #16, 17, 18 and 19) yield a mean date of 330 ± 6 Ma (MSWD = 0,9; n = 4), while the other two (Mz14 and Mz3; #3, 4, 5 and 15) give a weighted average of corrected dates at 308 ± 6 Ma (MSWD = 0.06; n = 4). Four analyses on three monazite grains (Mz11, 12 and 18), included within plagioclase, yield a weighted average date at 320 ± 12 Ma (MSWD = 1.7; n = 4).

Twenty-nine REE analyses by LA-ICP-MS were carried out on fifteen monazite grains. Due to silicate contamination, only thirteen analyses on seven monazite grains were retained (Tab. S6). Among them, five analyses are obtained on three monazite grains located in Bt-Ms layers (Mz6, Mz7 and Mn17), six analyses come from three monazite grains included in plagioclase (Mz11, Mz12 and Mz18) and two analyses come from a monazite in a staurolite fracture (Mz15). The analyses yield similar normalized spectra characterized by a strong HREE depletion relative to LREE (LuN/LaN = 1.70 × 10−5–8.39 × 10−4), with variations in HREE contents (LuN/DyN = 0.003–0.046), and a limited negative Eu/Eu* anomaly (0.55–0.70) (Tab. S6; Fig. 14C). The Y content shows values comprised between 886 and 10110 μg/g (Tab. S6). The highest Y contents are obtained in the youngest monazite grains, suggesting an increase in Y content over time, while the DyN/LuN ratio shows a rough opposite trend (Fig. 14D).

thumbnail Fig. 13

Back-scattered electron images of dated monazite (Mz) from both samples (AR736 and AR14). Monazites from sample AR736 located in (A and B) Bt-Ms layers, in (C and D) fractures in staurolite and in (E and F) plagioclase. Monazites from sample AR14 with (G) elongated shapes in the foliation, (H) partially replaced by apatite and (I) shape perpendicular to the foliation. Red circles indicate the analytical spots locations. Monazite numbers correspond to Table S5.
thumbnail Fig. 14

U-Pb LA-IC-PMS geochronological results and chondrite-normalized REE for monazite from sample AR736. (A) Tera-Wasserburg plot of monazite U-Pb results for the Bellachat micaschist (AR736). (B) Weighted averages of 207Pb-corrected dates diagram as function of monazite textural location. Error ellipses and uncertainties are ±2σ level. (C) Monazite trace element data from sample AR736. REE patterns normalized to chondrite values from McDonough and Sun (1995). (D) Plots of Y content and DyN/LuN ratios vs. 207Pb-corrected dates diagram. Bt-Ms: monazite in biotite-muscovite layers; St: monazite in fractures affecting staurolite cores; Pl: monazite in plagioclase inclusion.

4.5.2 Lac Emosson metapelite (AR14)

All monazite grains lie in the biotite-muscovite layers and none of them were observed as inclusion in garnet. Monazite crystals are abundant, the majority are ∼50 μm in size with a few up to ∼100 μm. They are commonly subhedral (Fig. 13G), but some of them are anhedral associated with apatite (Fig. 13H). Most monazite grains are elongated parallel to the foliation (Fig. 13G-H). Some of them form clusters perpendicular to the foliation (Fig. 13I). Most monazite grains show no fractures or internal deformation. BSE images show no clear chemical zoning and suggest homogeneous monazite grains (Fig. 13G-I) with some quartz or biotite and muscovite inclusions. Twenty-seven spots on 21 grains were analysed. Pb (∼1400–2000 μg/g), Th (∼38,220–63,000 μg/g) and U (2300–4400 μg/g) content and Th/U ratios (9.5–23.5) are similar to those observed for monazite sample AR736 (Tab. S5). The data have percentage of concordance ranging between 71% and 93% (Tab. S5). In the Tera Wasserburg diagram, these data define a regression line, with a lower intercept at 311 ± 5 Ma (MSWD = 0.9; n = 27) (Fig. 15A). In the 206Pb/238U vs. 208Pb/232Th diagram, the data are concordant and spread from ca. 330 Ma to 310 Ma (Fig. S2B). Together, they give a concordia age of 319 ± 3 Ma (MWSD(C+E) = 2.1; n = 23). The two dates are similar within uncertainty, indicating that there was no significant initial overcorrection of Pb in the Tera Wasserburg diagram that could have led to a rejuvenated date calculation.

Thirty REE analyses by LA-ICP-MS were carried out on nineteen monazite grains. Due to silicate contamination, only eight analyses on four monazite grains (Mz1, Mz2, Mz5 and Mz6) are usable (Tab. S6). All analyses give similar REE patterns (Fig. 15B), with a low Eu/Eu* anomaly (0.63-0.76), low Y content (386–2010 μg/g), HREE depletion relative to LREE (LuN/LaN ∼1 × 10−5–10−4) and variable DyN/LuN ratios (48-309) (Fig. 15B; Tab. S6).

thumbnail Fig. 15

U-Pb LA-ICP-MS geochronological results and chondrite-normalized REE for monazite from sample AR714. (A) Tera-Wasserburg plot of monazite U-Pb results for the Lac Emosson micaschist (AR14). Error ellipses and uncertainties are ± 2σ level. (B) Monazite trace element data of sample AR14. REE patterns normalized to the chondrite values of McDonough and Sun (1995).

5 Discussion

5.1 Reconstruction of P–T paths

In sample AR736, the beginning of the prograde path (MA) can be estimated from the Grtcore composition at 550–625 °C and 0.5–0.6 GPa (Fig. 10), in the predicted stability field Grt Pl Ilm Bt Ms St Qz (Fig. 9A), which is consistent with the partially preserved MA assemblage composed of Grtcore + Qz + Ms. The peak MB mineral assemblage, consisting of Grtmantle + Bt + Ms + Pl + Qz + St + Rt, is predicted to be stable at 0.76–0.82 GPa and at temperatures of 600–640 °C (Fig. 16A). These peak metamorphic conditions are consistent with the lack of evidence of partial melting (Fig. 9A), the muscovite composition (Fig. 9E) and the Zr-in-Rt thermometry (602 ± 17 °C at 0.8 GPa) (Fig. 16A). The Grtmantle compositions are modelled at 620–640 °C and 0.75 GPa, just below the Rt-out curve (Fig. 10). This difference between the P–T conditions defined by the MB assemblage and the Grtmantle composition can be explained by the uncertainties of the thermodynamic models (i.e. ±50 °C and ±0.1 GPa; e.g. Holland and Powell, 2008). Grtcore composition is never perfectly modelled (maximum Qcmp values are 0.8–0.9; Fig. 10A-D), nevertheless, the Grtcore composition indicates P–T conditions just above the garnet-in reaction (Fig. 10), suggesting that intracrystalline diffusion may not have significantly modified the garnet core growth composition in this sample. This hypothesis is consistent with the similar compositional zones defined by the major and trace elements maps (Figs. 4 and 8A).

The staurolite core composition is modelled outside the stability field of rutile (Fig. 9F), whereas rutile inclusions have been observed in the St1 (Fig. 2F-G). Uncertainties from the thermodynamic models could lead to underestimate the pressure conditions of the staurolite crystallization (Forshaw and Pattison, 2023). Therefore, the St1 and St2 compositions are not used to constrain P–T conditions.

The retrograde path (MC) is characterized by the partial replacement of staurolite cores by biotite and muscovite, followed by the crystallization of the staurolite overgrowths. This sequence of crystallization, resorption and subsequent recrystallization is plausible with a clockwise P–T path (Fig. 16A), where St1 crystallized within the rutile stability field, is partially resorbed as P–T conditions approach the alumino-silicate-in reaction, and St2 subsequently crystallizes along the retrograde path.

The late retrograde stage (MD) is defined by the crystallization of retrograde chlorite at temperatures lower than 525–550 °C (Fig. 9A).

In a sample from the same area as AR736, Dobmeier (1996, 1998) described a similar petrographic evolution with: (i) a high-Xsps garnet core containing biotite, muscovite, quartz and plagioclase inclusions, consistent with the MA assemblage; (ii) a low-Xsps garnet mantle with biotite, muscovite, quartz and plagioclase inclusions, consistent with the MB assemblage; (iii) instead of staurolite, Dobmeier (1996, 1998) reported sillimanite in the matrix, which is consistent with the MC assemblage close to the sillimanite-in reaction. (iv) finally, the biotite of the matrix is replaced by chlorite, like in the MD assemblage. Thermobarometric data on this sample yield a prograde P–T path similar to our results (Fig. 1C, Lac Cornu area path). These similarities corroborate the P–T path we propose for sample AR736.

In sample AR14, Mb peak mineral assemblage (Grt + Pl + Ms + Rt + Qz + Bt) conditions are estimated at pressures above 0.9 GPa and a temperature range of 610–720 °C (Fig. 16B). No mineral relics of anterior parageneses (Ma) were found in this sample. Although the Mb mineral assemblage is observed as inclusions in Grtcore and Grtmantle, the compositions of Grtcore and Grtmantle yield P–T conditions within the staurolite stability field below 0.7–0.9 GPa and 590–675 °C (Fig. 12A-B; field Grt + Pl + Ms + Rt + Qz + Bt + St). Grtcore and Grtmantle are nearly homogeneous in composition, and it is very likely that their initial compositions have been partially re-equilibrated by intracrystalline diffusion (e.g. Yardley, 1977; Caddick et al., 2010). Therefore, we propose that the P–T conditions defined by compositions of Grtcore and Grtmantle represent a re-equilibration along the P–T path and not the Mb conditions. Isopleths of Si-content in muscovite at 3.05 apfu pass through the Grt + Pl + Ms + Rt + Qz + Bt + St field and are consistent with P–T conditions defined by the Grtcore-mantle composition (Fig. 11B). One data point yields a higher Si content (3.08 apfu) and might be related to Mb peak assemblage or Ma prograde path (Fig. 11B).

The third portion of the P–T path (Mc, Tab. 4) is characterized by the partial replacement of rutile by ilmenite and by the crystallization of sillimanite. These reactions indicate an increase of temperature of 50–75 °C and a slight decompression of 0.1–0.2 GPa (Figs. 11A and 16C). If we only consider the mineral assemblages, a simple loop P–T path can be proposed (Fig. 16C, path 1). Alternatively, if we consider the Grtrim1 and Grtrim2 compositions, a more complex path can be proposed (Fig. 16, path 2) with a decompression suggested by Grtrim1 composition at 0.6 GPa and 585 °C (Fig. 12C), followed by a temperature increase to 0.55 GPa and 620 °C constrained by Grtrim2 composition (Fig. 12D), within the sillimanite stability field. Based on our garnet maps showing well defined Grtrim1 and Grtrim2 zones, we favour this P–T path defined by the garnet compositions. The late retrograde path (Md) is constrained by the crystallization of chlorite, which is stable at temperatures below 525–575 °C (Fig. 11A).

thumbnail Fig. 16

Summary of the main metamorphic mineral sequence and P–T paths of the studied samples. (A) Reconstructed P–T paths for samples AR736, see text for details. (B) Simplified petrological model for sample AR736 showing the garnet and staurolite evolutions and their relationships with the metamorphic stages, crystallisation of monazite (in orange) and ages: (1) Nucleation and crystallisation of garnet and staurolite along the prograde path and incorporation of inclusions near the peak of pressure; (2) beginning of decompression path and peak of metamorphism associated with staurolite resorption, garnet slight resorption and crystallisation of monazite starting at 330 ± 6 Ma; (3) Garnet breakdowns and monazite crystallisation continues. Dextral shearing partially or totally recrystallized older monazite grains assisted by aqueous fluids until 305 Ma. The staurolite rims crystallize around staurolite cores and in its fractures, biotite is chloritized and garnet incorporates MnO. (C) Reconstructed P–T paths for samples AR14, see text for details. (D) Simplified petrological model for sample AR14 showing the garnet evolution and its relationships with the metamorphic stages, crystallisation of monazite and ages: (1) Nucleation and crystallisation of garnet along an unknown prograde path and incorporation of inclusions; (2) beginning of decompression path associated with intra-crystalline diffusion in Grtcore and Grtmantle followed by Grtrim1 crystallization. Staurolite and monazite may start to crystallize; (3) Temperature increases, leading to the crystallization of sillimanite and Grtrim2 and to the hypothetical breakdown of staurolite. (4) Garnet incorporates MnO and biotite is chloritized. Monazite recrystallization is assisted by dextral shearing and aqueous fluids throughout the entire decompression path.

5.2 Monazite as a time record of deformation

None of the dated monazite grains are included in garnet, which prevents linking directly both minerals crystallization timing. The relative timing of the growth of different domains within garnet and monazite can be investigated based on the partitioning of trace elements between these minerals (e.g. Buick et al., 2006; Regis et al., 2014; Iaccarino et al., 2015). Several studies have shown a systematic decrease in partitioning values for HREE between garnet and monazite with an almost constant negative slope (Fig. 17; Hermann and Rubatto, 2003; Rubatto et al., 2006; Buick et al., 2006). Monazite-garnet trace element partitioning was calculated for all possible combinations of garnet domains and monazite ages for sample AR736 and for combinations of Grtcore, Grtmantle and Grtrim domains and monazite in sample AR14 (Fig. 17). None of the calculated pairs lie within the expected equilibrium partitioning zone, particularly for HREE. These results are interpreted as evidence for disequilibrium and diachronous growth of monazite and garnet in the samples.

In sample AR736, monazite dates range from ca. 340 to 300 Ma depending on the textural position of the dated grains (Fig. 14). A date of 330 ± 6 Ma was obtained on two monazite grains located in fractures affecting staurolite cores. Therefore, we propose that this age represents the resorption of the St1, which happened near the peak of temperature during the decompression path (Fig. 16B). A monazite from the Bt-Ms layers but located between two rutile grains (Fig. 13B) gave a date of 330 ± 12 Ma (Fig. 14), which roughly constrains the stability of rutile, and thus the pressure peak. The geochemistry of monazite from sample AR736 indicates an increase in the amount of HREE over time (Fig. 14C), an increase in coarse Y and a decrease in DyN/LuN (Fig. 14D). These observations are consistent with garnet resorption (Pyle et al., 2001; Foster et al., 2002, 2004; Gibson et al., 2004; Krenn et al., 2009; Lanari and Engi, 2017; Kohn, 2016; Shrestha et al., 2019). Therefore, we propose that monazite started to (re)crystallize at peak temperature conditions, subsequent to garnet growth and continued to do so throughout the retrograde path during the garnet resorption (Fig. 16B). This age is consistent with the monazite U-Pb age of 327 ± 2 Ma (ID-TIMS) obtained on a metapelite from the Emosson area and interpreted as the peak metamorphism by Bussy et al. (2000), and with the monazite LA-ICP-MS U-Th-Pb ages at 340–330 Ma of the Belledonne and Pelvoux micaschists in the ECMs, also interpreted as the age of the peak of metamorphism (Fréville et al., 2018, 2022).

A younger date of 307 ± 4 Ma was obtained on monazite included in the mica-rich layers (Fig. 14). These mica-rich layers mark the dextral S-C structures and are also crystallized in sigmoidal tails (Fig. 2A and B), indicating syn-tectonic crystallization. Monazite grains in these mica-rich layers often contain biotite inclusions, also suggesting syn-tectonic crystallisation or recrystallisation. It also yields a maximum age for the biotite breakdown to chlorite during the MLR stage. Monazite included in plagioclase gives a poorly constrained date of 320 ± 12 Ma. Therefore, we interpret the dates ranging from 330 ± 6 Ma to 307 ± 4 Ma as the ages of the onset and end of decompression metamorphism, respectively (Fig. 16B).

Dating of monazite from the Bt-Ms layers of sample AR14 yields a date of 311 ± 5 Ma (Fig. 15). The Lac Emosson metapelite (sample AR14) is located within the EBSZ, where strain is higher compared to the southwestern part of the ARM (Vanardois et al., 2022b), from which the Bellachat metapelite (AR736) was collected (Fig. 1B). The perpendicular and parallel orientations of monazite in the mylonitic foliation (Fig. 13G-I) indicate that the ca. 310 Ma date represents the age of deformation associated with the ductile shearing of the EBSZ (Fig. 16D). This age is consistent with the emplacement of the syn-tectonic Vallorcine granite at 306 ± 5 Ma (Vanardois et al., 2022b). This age is also similar to the age of 307 ± 4 Ma obtained within monazite grains included in the mica-rich layers of sample AR736, which marks also dextral shearing deformation (Fig. 14). The monazite grains have low Y content (Tab. S6), which could be explained by (i) xenotime crystallisation in the vicinity of garnet and fractioning of the Y content (Spear and Pyle, 2010) or (ii) crystallization in equilibrium with Grtrim2 or Grtrim3.

Monazite may (re)crystallize as a result of fluid–rock interactions occurring relatively late in the metamorphic history (e.g. Didier et al., 2014; Erickson et al., 2015; Roger et al., 2020 and references therein). Fluid-assisted dissolution-precipitation mechanisms are efficient in modifying the chemical and isotopic composition of monazite, even at temperatures as low as ∼300 °C (Hawkins and Bowring, 1997; Townsend et al., 2000). In the Lac Emosson area (sample AR14), vertical transcurrent shear zones channelled external water (Genier et al., 2008; Vanardois et al., 2022b, 2024a). We propose that further deformation, recrystallization, and aqueous fluid circulation within these shear zones likely triggered monazite recrystallization in the mineral matrix (Fig. 16D), whereas monazite grains preserved within staurolite fractures in sample AR736 remained unaffected by recrystallization (Fig. 16B).

5.3 Pre-orogenic thermal conditions and geodynamic setting

The study of sample AR736 allowed us to determine the pressure-temperature conditions recorded by the rock at the very beginning of its prograde evolution. The P–T conditions of 0.5–0.6 GPa and 550–625 °C suggest that the geothermal gradient was high (∼30–35 °C/km) prior to the onset of the crustal thickening, which is consistent with a thin crust and a high mantle flux. A high-geothermal gradient has been described in the adjacent Servoz volcano-sedimentary basin in the southwestern ARM and dated at ca. 350 Ma (Vanardois et al., 2022c). This high thermal state, prior to the crustal thickening, has been interpreted as a back-arc domain in a suprasubduction setting between 370-350 Ma in the ECMs (Guillot et al., 2009; Vanardois et al., 2022c; Jacob et al., 2023; Fréville et al., 2024; Gunia et al., 2025).

We interpret the early high-geothermal gradient recorded by our sample AR736 as the consequence of an already warm crust originated from a late-Devonian back-arc setting as documented in the Moldanubian Bohemian massif (e.g. Schulmann et al., 2009), in the southern Vosges mountains (e.g. Skrzypek et al., 2014) and in the French Massif Central (e.g. Pin and Paquette, 1997, 2002).

5.4 Prograde tectono-metamorphic evolution

The results reported in this study constrain a prograde path and provide peak pressure conditions that are very similar to those documented by Dobmeier (1998, sample A). However, a clockwise P–T path is obtained rather than an anticlockwise path as suggested by Dobmeier (1998) (Fig. 18A). Our quantitative prograde paths show a low dT/dP gradient (about 15 °C/0.1 GPa) from 0.55 GPa to 0.8 GPa with a peak pressure attained at around 340 Ma (Fig. 16A). Considering a time span of 10 Ma between the back-arc setting (350 Ma; Vanardois et al., 2022c) and the pressure peak conditions (340 Ma), and a crustal density of 2.7 g cm−3, the rate of this crustal thickening is around 1 mm year−1. These low dT/dP gradient and crustal thickening rate are consistent with P–T path and rate values described in nappe or thrust stacking in natural samples (e.g. Florence et al., 1993) and in numerical models (e.g. Copley and Weller, 2022). Nappe stacking is a likely process to explain crustal thickening in the ARM. No structural evidence of Variscan thrusts or nappes have been observed in the ARM, but some are described in the Belledonne massif (Fernandez et al., 2002; Fréville et al., 2018). In addition, microstructural features recorded by prograde garnet in sample AR736 emphasizes syn-kinematic crystallization with a dextral shearing (Fig. 2A). Thus, besides nappe stacking, transcurrent deformation was an acting process during late crustal thickening. This result is consistent with recent results indicating the beginning of transcurrent tectonics, with nucleation of the regional scale dextral shear zones in the partially molten orthogneisses, at the late prograde evolution (Vanardois et al., 2024a). Similar, to what it is proposed for the Belledonne-Pelvoux area (Fréville et al., 2022) we suggest that due to strain partitioning in the orogenic crust, bulk contraction is accommodated by nappe stacking in the upper-middle crust and by dextral transpression in the middle-lower crust.

thumbnail Fig. 17

Monazite-garnet pairs showing potential equilibrium partitioning of trace elements for (A) sample AR736 and (B) sample AR14. Grey outline indicates approximate equilibrium partitioning conditions (Buick et al., 2006; Hermann and Rubatto, 2003; Rubatto et al., 2006).
thumbnail Fig. 18

Comparison of the P–T paths with previously published data. (A and B) Comparison of our P–T paths with the ones published in the southwestern and central parts of the ARM. Same references than in Figure 1. (C) Comparison of P–T–t–D paths of this study (thick arrows) and lower crust eclogites (Vanardois et al., 2022a). Grey stars indicate timing constraints on the P–T–D paths. (D) Temporal evolutions of the three P–T paths: 350-340 Ma: crustal thickening; 340-330 Ma: longitudinal flow in partially molten lower crust and beginning of dextral transpression in the middle crust; 330-305 Ma: juxtaposition and exhumation of upper and lower crusts by the dextral shear zones.

5.5 β-shape path

Our geochronological data align well with previously published results on the ECMs, suggesting peak metamorphism followed by a retrograde path between approximately 340 and 305 Ma (Bussy et al., 2000; Schulz and von Raumer, 2011; Rubatto et al., 2010; Fréville et al., 2018, 2022; Jouffray et al., 2020; Vanardois et al., 2022a; Jacob et al., 2021, 2022; Philippi et al., 2024). A significant finding from sample AR14 is the record of two successive prograde evolutions, separated by a period of isothermal decompression associated with slight cooling (Fig. 18B).

This type of path, characterized by two prograde phases (one at high pressure and the other at lower pressure), is commonly observed in orogens and is referred to as a “β-type P–T–t path”. It is attributed to a late heating event driven by factors such as shallow break-off of the subducted slab, radiogenic heat production from continental block accretion, mantle delamination, or a slowdown in exhumation (Sizova et al., 2019 and references therein).

Our petrochronological data indicate that monazite (re)crystallization at 311 ± 5 Ma postdates Grtrim1 formation but could be synchronous with Grtrim2 crystallization. This age coincides, within uncertainty, with the emplacement ages of the Vallorcine and Mont-Blanc granites (Fig. 1B; Bussy et al., 2000; Vanardois et al., 2022b), which may have served as the heat source responsible for the brief temperature increase in the upper-middle crust, as described in the Variscan Velay dome (Barbey et al., 2015).

5.6 Timing of exhumation and strain Partitioning in the East Variscan Shear Zone

In sample AR14, the Grtrim1 and Grtrim2 domains are aligned along the elongated garnet grains, which are oblong and parallel to dextral planar fabrics. This observation suggests a syn-tectonic dissolution-reprecipitation process (e.g. Álvarez-Valero et al., 2005). Additionally, sillimanite is located within dextral shear bands, emphasizing the temperature increase during the decompression path. Finally, Ti-in-Bt thermometry on biotite marking the planar fabrics of the EVSZ yields a temperature of 578 °C ± 35 °C, which is consistent with the quartz c‐axis fabric opening‐angle temperatures between 550 and 630 °C obtained on sheared rocks from the Emosson area (Simonetti et al., 2020) (Fig. 11B). These findings indicate that dextral shearing along the EVSZ was active in the middle crust of the ARM) throughout its exhumation, consistent with the conclusions of Schulz and von Raumer (1993). The study of ARM eclogites reveals a two-stage exhumation history (Vanardois et al., 2022b). The first stage involved horizontal flow, leading to significant unroofing due to a non-horizontal component of the flow (e.g. Trap et al., 2011). The second stage was marked by dextral shearing, characterized by a stretching lineation plunging approximately 35° to the north, with occasional steeper plunges (von Raumer and Bussy, 2004; Vanardois et al., 2022a, 2022b). Along these vertical shear zones, horizontal displacements of 50–60 km could account for vertical exhumation of about 30 km. Vanardois et al. (2022b) suggest that deformation partitioning with dextral shearing and horizontal flow occurred as early as 340–330 Ma (Fig. 18D). Our study shows that less deeply buried middle-crustal rocks record the same peak metamorphism age (340–330 Ma). These synchronous P–T records indicate spatial partitioning of movements within the orogenic crust, with horizontal flow domains coexisting alongside domains dominated by vertical motion. Similar deformation partitioning has been documented in other transcurrent settings, such as the Pelvoux massif (Fréville et al., 2022), the Variscan Pyrenees (Cochelin et al., 2017, 2021), and the French Central Massif (Rabin et al., 2015; Roger et al., 2015; Rey et al., 2017; Trap et al., 2017; Vanardois et al., 2024b). We propose that dextral transcurrent deformation in the ARM initiated around 340–330 Ma and was accompanied by deformation partitioning into vertical and horizontal fabrics (Fig. 18D). Continued transcurrent shearing broadened the dextral shear zones (Simonetti et al., 2020; Vanardois et al., 2022b), which coalesced into an anastomosed system forming the EVSZ. The late retrograde P–T paths of samples AR14 and AR736 are consistent with the final retrograde evolution of the Lac Cornu eclogites (Fig. 18C; Vanardois et al., 2022a), indicating that the middle and lower crusts were juxtaposed between 300 and 305 Ma. These results demonstrate that the EVSZ was active throughout the exhumation of the ARM’s middle crust and remained active until the end of the Variscan orogeny (ca. 305–300 Ma). This timeframe coincides with the flow of partially molten lower crust. This prolonged transcurrent tectonic activity suggests that the numerous transcurrent shear zones in the Variscan belt are not merely late-orogenic structures but played a significant role in the geodynamic evolution, particularly in the exhumation of the orogenic crust, from the end of continental collision to the close of the Variscan orogeny.

6 Conclusion

Petrological observations, Zr-in-Rt thermometry and phase equilibrium models were used to constrain the P–T path of two metapelites from the Aiguilles Rouges massif. Garnet compositional mapping reveals that the samples underwent a series of growth stages, followed by a phase of partial resorption. In the southwestern ARM sample, the garnet compositions preserved the record of prograde metamorphic conditions, which were followed by a clockwise path recorded by staurolite crystallization − resorption − overgrowth sequence. Our results also document a high-geothermal gradient prior to the crustal thickening, that is interpret as the opening of volcano-sedimentary basins during a back-arc setting, as described in several ECM and other Variscan massifs of the Moldanubian zone. On the other hand, garnet grains from the central part of the ARM are affected by intracrystalline diffusion, which only allows for the recording of the retrograde part of a clockwise path, which is subsequently disrupted by the emplacement of the late-orogenic granites. Petrological observations and trace element chemistry indicate that the monazite crystallisation is coeval with garnet breakdown during the decompression stages, spanning from 330 ± 6 Ma to 307 ± 4 Ma. Deformation and associated fluids drained into the large-scale dextral shear zone recrystallized monazite grains within the matrix and reset the ages. A comparison with the previously obtained P–T path for the lower crust suggest that a juxtaposition of the middle and lower crustal domains occurred, followed by a shared final exhumation during dextral transpression between ca. 330 Ma and 305 Ma. These results highlight a long lasting dextral transcurrent deformation that began at least near the peak of pressure of the middle crust until its final exhumation at the end of the orogeny. This deformation had significant impacts on the geodynamic evolution, including strain partitioning in the orogenic crust and exhumation. The findings suggest that the numerous transcurrent shear zones described in the Variscan belt may have an important role on the exhumation of the orogenic crust since the end of the crustal thickening.

Acknowledgments

Sadly, our colleagues and friends Jean-Louis Paquette and Didier Marquer passed away in June 2022 and August 2024, respectively, before the completion of this manuscript. Their research has contributed significantly to our knowledge of orogenic evolution, especially for the Variscan cycle and this paper is dedicated to their memory. We will miss their kindness and knowledge. This work was supported by the BRGM through the Référentiel Géologique de la France program (RGF program) And from the European Research Council (ERC) under the European Union’s Horizon 2020 research and innovation program (grant 850530). We thank Didier Convert-Gaubier from the PEA2t platform (Chrono-environnement, University Bourgogne Franche-Comté, UMR CNRS 6249, France) for performing thin section preparation. We thank Marc Ulrich for performing micro-XRF analyses. This work was supported in part by the French RENATECH network and the MIMENTO Technology Center of FEMTO-ST. We thank L. de Hoÿm de Marien and an anonymous reviewer for their constructive comments, and L. Jolivet and O. Vanderhaeghe for their editorial assistance.

Supplementary Material

Fig. S1: P–XFe2O3 diagrams of (A) sample AR736 and (B) sample AR14, at 600 °C showing the strong impact of Fe2O3 on the stability of rutile and ilmenite. Note the absence of field without ilmenite and the absence of co-stability between plagioclase and rutile, except at very low XFe2O3 content.

Fig. S2: U-Th-Pb diagrams (208Pb/232Th vs 206Pb/238U) of samples AR736 (A) and AR14 (B). Error ellipses and uncertainties are ± 2σ level.

Table S1: Operating conditions and instrument setting for trace elements analyses on garnet and monazite.

Table S2: Operating conditions and instrument setting for trace elements analyses on rutile.

Table S3: Garnet trace element analyses.

Table S4: Zr-in-Rt thermometry.

Table S5: Analytical results of LA-ICP-MS U-Th-Pb dating.

Table S6: Monazite trace element analyses.

Appendix 1: Operating conditions and instrument setting for U-Th-Pb analysis.

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Cite this article as: Vanardois J, Trap P, Goncalves P, Roger F, Lanari P, Piccoli F, Marquer D, Paquette J-L, Melleton J. 2025. Quantitative P–T–t–D Paths: A Key to decipher orogenic dynamics in the Variscan Aiguilles Rouges Massif (External Crystalline Massifs, Western Alps), BSGF - Earth Sciences Bulletin 196, 20. https://doi.org/10.1051/bsgf/2025011

All Tables

Table 1

Representative EPMA analytical spots for compositions of garnet (Grt), staurolite (St), biotite (Bt) and muscovite (Ms). Xalm: almandine; Xprp: pyrope; Xsps: spessartine; Xgrs: grossular.

In the text
Table 2

Garnet endmembers contents extracted from XMapTools.

In the text
Table 3

Bulk compositions of samples AR736 and AR14.

In the text
Table 4

Petrological interpretations. The prograde mineral assemblage of sample AR736 is based on garnet inclusions and is therefore probably incomplete.

In the text

All Figures

thumbnail Fig. 1

(A) Paleogeographic reconstitution of the European Variscan Belt at the end of the Variscan orogeny (ca. 290 Ma) modified from Franke et al. (2017). (B) Geological map of the Aiguilles-Rouges and the Mont-Blanc Massifs with location of the two metapelites samples analysed and the LA-ICP-MS U-Th-Pb monazite ages obtained. Various P–T paths proposed for the Aiguilles-Rouges massif: (C) Anticlockwise P–T paths of metapelites in the Prarion and Lac Cornu areas after Dobmeier (1996, 1998). (D) Clockwise P–T paths of metapelites from the Val Bérard and Lac Emosson areas after Schulz and von Raumer (1993, 2011), Joye (1989) and Genier et al. (2008). Black arrows represent sections of P–T paths constrained by thermobarometric results and grey arrows are interpreted sections.
In the text
thumbnail Fig. 2

Thin section and back-scattered electron images showing petrological characteristics of the Bellachat micaschist (AR736): (A) Main paragenesis of the sample with garnet and staurolite porphyroblasts in mica-plagioclase-quartz matrix. Sigmoids and tails around garnet highlight dextral kinematics. (B) Main paragenesis of the sample with dextral C-S structures corresponding to the S2 and C2 planar fabrics. (C and D) Inclusions in the outer part of garnet grains. (E) Partial replacement of biotite by chlorite and rutile inclusions in a plagioclase. (F) Optical differences between the dark staurolite core (St1) and the bright staurolite zones (St2). (G) Fractured staurolite core (St1) with bright staurolite zone (St2) in the fracture. Mineral abbreviations are from Warr (2021).
In the text
thumbnail Fig. 3

Thin section images showing petrological characteristics of the Emosson micaschist (AR14). (A and B) Main paragenesis of the sample with garnet porphyroblasts in a mica-plagioclase-quartz matrix, with dextral S-C structures. (C and D) Plane light and cross light images of fibrolitic sillimanite in the matrix. Note that the biotite is partially chloritized.
In the text
thumbnail Fig. 4

Garnet compositional maps and profiles for the Bellachat micaschist (AR736). End-member fraction compositional maps of (A) round sub-euhedral garnet and (B) oblong anhedral garnet. (C) Chemical profiles of the garnet grains. Xalm : almandine fraction; Xgrs: grossular fraction; Xprp: pyrope fraction; Xsps: spessartine fraction. Representative mineral compositions are provided in Tables 1 and 2.
In the text
thumbnail Fig. 5

Compositional map of staurolite in sample AR736. Location of maps are shown in Figures 2F and 2G.
In the text
thumbnail Fig. 6

Garnet compositional maps and profiles for sample AR736. End-member fraction compositional maps of four garnet grains. Profiles are visible in Figure 7. Representative mineral compositions are provided in Tables 1 and 2.
In the text
thumbnail Fig. 7

Chemical profiles of the garnet grains presented in Figure 6. Fig. 8: WDS maps of Y content and profiles in garnet from samples (A) AR736 and (B) AR14. Locations of LA-ICP-MS analyses in garnet and garnet REE patterns normalized to the chondrite values of McDonough and Sun (1995) from samples (C) AR736 and (D) AR14.
In the text
thumbnail Fig. 8

WDS maps of Y content and profiles in garnet from samples (A) AR736 and (B) AR14. Locations of LA-ICP-MS analyses in garnet and garnet REE patterns normalized to the chondrite values of McDonough and Sun (1995) from samples (C) AR736 and (D) AR14.
In the text
thumbnail Fig. 9

Pressure − Temperature phase diagram constructed for the sample AR736 and contoured with compositional and modal isopleths. Bulk composition used is available in Table 3. (A) P–T mineral assemblage diagram with main reaction curve highlighted. MB, MC and MD mineral assemblage stability fields are in bold. (B and C) Xalm, Xgrs, Xprp and Xsps isopleths of garnet. Dotted arrow represents the rough path inferred from garnet compositions in Tables 1 and 2. (D) Zr-in-Rt thermometry results. Each red line represents a Zr-in-Rt analysis, the red area represents the temperature calculated from the sixteenth data with the highest Zr content. (E) Si apfu isopleths of muscovite and modal proportion of biotite expressed as vol%. (F) XMg compositional and modal isopleths for staurolite.
In the text
thumbnail Fig. 10

Results of the Qcmp factor of Grtcore and Grtmantle for the garnets presented in Figure 4. Compositional values used for the calculation are given in Table 2.
In the text
thumbnail Fig. 11

Pressure − Temperature phase diagram constructed for the sample AR14 and contoured with compositional and modal isopleths. Bulk composition used is available in Table 3. (A) P–T mineral assemblage diagram with main reaction curve highlighted. Mb, Mc and Md mineral assemblage stability fields are in bold. (B) Si apfu isopleths of muscovite and Ti-in-Bt thermometry results. (C and D) Xalm, Xgrs, Xprp and Xsps isopleths of garnet. Dotted arrow represents the rough path inferred from garnet compositions in Tables 1 and 2.
In the text
thumbnail Fig. 12

Results of the Qcmp factor of Grtcore, Grtmantle, Grtrim1 and Grtrim2 for the garnets presented in Figure 6. Compositional values used for the calculation are given in Table 2. The diagram Grtrim2 for Map2-G2 is missing because this compositional zone has not been observed in this grain.
In the text
thumbnail Fig. 13

Back-scattered electron images of dated monazite (Mz) from both samples (AR736 and AR14). Monazites from sample AR736 located in (A and B) Bt-Ms layers, in (C and D) fractures in staurolite and in (E and F) plagioclase. Monazites from sample AR14 with (G) elongated shapes in the foliation, (H) partially replaced by apatite and (I) shape perpendicular to the foliation. Red circles indicate the analytical spots locations. Monazite numbers correspond to Table S5.
In the text
thumbnail Fig. 14

U-Pb LA-IC-PMS geochronological results and chondrite-normalized REE for monazite from sample AR736. (A) Tera-Wasserburg plot of monazite U-Pb results for the Bellachat micaschist (AR736). (B) Weighted averages of 207Pb-corrected dates diagram as function of monazite textural location. Error ellipses and uncertainties are ±2σ level. (C) Monazite trace element data from sample AR736. REE patterns normalized to chondrite values from McDonough and Sun (1995). (D) Plots of Y content and DyN/LuN ratios vs. 207Pb-corrected dates diagram. Bt-Ms: monazite in biotite-muscovite layers; St: monazite in fractures affecting staurolite cores; Pl: monazite in plagioclase inclusion.
In the text
thumbnail Fig. 15

U-Pb LA-ICP-MS geochronological results and chondrite-normalized REE for monazite from sample AR714. (A) Tera-Wasserburg plot of monazite U-Pb results for the Lac Emosson micaschist (AR14). Error ellipses and uncertainties are ± 2σ level. (B) Monazite trace element data of sample AR14. REE patterns normalized to the chondrite values of McDonough and Sun (1995).
In the text
thumbnail Fig. 16

Summary of the main metamorphic mineral sequence and P–T paths of the studied samples. (A) Reconstructed P–T paths for samples AR736, see text for details. (B) Simplified petrological model for sample AR736 showing the garnet and staurolite evolutions and their relationships with the metamorphic stages, crystallisation of monazite (in orange) and ages: (1) Nucleation and crystallisation of garnet and staurolite along the prograde path and incorporation of inclusions near the peak of pressure; (2) beginning of decompression path and peak of metamorphism associated with staurolite resorption, garnet slight resorption and crystallisation of monazite starting at 330 ± 6 Ma; (3) Garnet breakdowns and monazite crystallisation continues. Dextral shearing partially or totally recrystallized older monazite grains assisted by aqueous fluids until 305 Ma. The staurolite rims crystallize around staurolite cores and in its fractures, biotite is chloritized and garnet incorporates MnO. (C) Reconstructed P–T paths for samples AR14, see text for details. (D) Simplified petrological model for sample AR14 showing the garnet evolution and its relationships with the metamorphic stages, crystallisation of monazite and ages: (1) Nucleation and crystallisation of garnet along an unknown prograde path and incorporation of inclusions; (2) beginning of decompression path associated with intra-crystalline diffusion in Grtcore and Grtmantle followed by Grtrim1 crystallization. Staurolite and monazite may start to crystallize; (3) Temperature increases, leading to the crystallization of sillimanite and Grtrim2 and to the hypothetical breakdown of staurolite. (4) Garnet incorporates MnO and biotite is chloritized. Monazite recrystallization is assisted by dextral shearing and aqueous fluids throughout the entire decompression path.
In the text
thumbnail Fig. 17

Monazite-garnet pairs showing potential equilibrium partitioning of trace elements for (A) sample AR736 and (B) sample AR14. Grey outline indicates approximate equilibrium partitioning conditions (Buick et al., 2006; Hermann and Rubatto, 2003; Rubatto et al., 2006).
In the text
thumbnail Fig. 18

Comparison of the P–T paths with previously published data. (A and B) Comparison of our P–T paths with the ones published in the southwestern and central parts of the ARM. Same references than in Figure 1. (C) Comparison of P–T–t–D paths of this study (thick arrows) and lower crust eclogites (Vanardois et al., 2022a). Grey stars indicate timing constraints on the P–T–D paths. (D) Temporal evolutions of the three P–T paths: 350-340 Ma: crustal thickening; 340-330 Ma: longitudinal flow in partially molten lower crust and beginning of dextral transpression in the middle crust; 330-305 Ma: juxtaposition and exhumation of upper and lower crusts by the dextral shear zones.
In the text

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