# West European Variscan Belt dismantling and early fragmentation of Pangea: The key role of the Paleotethys subduction

Bryan Cochelin, Yoann Denèle, Nicolas Saspiturry

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Bryan Cochelin, Yoann Denèle, Nicolas Saspiturry. West European Variscan Belt dismantling and early fragmentation of Pangea: The key role of the Paleotethys subduction. Earth-Science Reviews, 2025, 271, pp.105304. 10.1016/j.earscirev.2025.105304. hal-05326919

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# West European Variscan Belt dismantling and early fragmentation of Pangea: The key role of the Paleotethys subduction

Bryan Cochelin<sup>a,b,c,*</sup>, Yoann Denèle<sup>d</sup>, Nicolas Saspiturry<sup>e</sup>

## ARTICLE INFO

### Keywords:

* Variscan belt
* Orogenic dismantling
* Strain partitioning
* Permian extension
* Pangea
* Paleotethys
* Back-arc extension

## ABSTRACT

This review explores the evolution of the West European Variscan Belt from the Late Carboniferous to the Permian, shedding light on a prolonged period of lithospheric reorganization that coincided with both the final assembly of Pangea and its early dismantling. To understand the balance of forces in the Variscan lithosphere during late- and post-orogenic stages, we integrate magmatic, tectonic, and metamorphic features spanning the 330-270 Ma interval. By considering modern paleogeodynamic reconstructions across Central-Western Europe, we develop a series of 3D tectonic models. These reconstructions emphasize the key role of the enclosed Paleotethys Ocean, which underwent subduction processes in the core of the Pangea supercontinent and interacted with the Alleghanian intracontinental collision to the west. Our model suggests that Paleotethys subduction triggered back-arc extension in the southern Variscan realm and established a lateral free-edge boundary. It allowed mass transfer between collisional and subduction domains from at least the end of the Carboniferous. This lateral free-edge boundary accommodated part of the material transfers from the overriding continental plate during the progressive dismantling of the Variscan belt. The dismantling of the upper plate, which is facilitated by earlier Late-Carboniferous delamination processes, culminates in plate-scale oroclinal bending, in which the accommodation space is controlled by trench-directed deep crustal (and mantle) flow and strike-slip tectonics. Our findings suggest that the rapid post-orogenic re-equilibration of the Variscan crust was not solely governed by gravitational collapse, but was significantly influenced by lateral mass redistribution, likely driven by slab retreat operating since the end of the Carboniferous.

## Contents

<table><tr><td>1.</td><td>Introduction</td><td>1</td></tr><tr><td>2.</td><td>Methodology and dataset for the review of late-Variscan tectonic, magmatic, and metamorphic features</td><td>2</td></tr><tr><td>3.</td><td>Review of the Paleogeographic and tectonic framework of the Variscan Belt</td><td>5</td></tr><tr><td>3.1.</td><td>Paleogeographic evolution</td><td>5</td></tr><tr><td>3.2.</td><td>The litho-tectonic zones</td><td>8</td></tr><tr><td>3.2.1.</td><td>The Rhenohercynian Zone</td><td>9</td></tr><tr><td>3.2.2.</td><td>The Saxothuringian Zone</td><td>9</td></tr><tr><td>3.2.3.</td><td>The Moldanubian Zone</td><td>9</td></tr><tr><td>3.2.4.</td><td>The Southern Variscan Belt</td><td>9</td></tr><tr><td>3.3.</td><td>The late Variscan tectonic evolution</td><td>10</td></tr><tr><td>4.</td><td>Influence of Paleotethys Subduction on Crustal Deformation in the Pangea Overriding Plate</td><td>11</td></tr><tr><td>4.1.</td><td>The N-S trending linear Variscan belt at early Carboniferous time</td><td>11</td></tr><tr><td>4.2.</td><td>West European Variscan belt reorganization: Oroclinal bending and trench-directed crustal flow (320-300 Ma)</td><td>12</td></tr></table>

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

<footnote><sup>b</sup> ISTO, UMR 7327, Univ Orleans, CNRS, BRGM, OSUC, F-45071 Orléans, France</footnote>

<footnote><sup>c</sup> State Key Laboratory for Mineral Deposits Research, Nanjing University, Nanjing 210093, China</footnote>

<footnote><sup>d</sup> Géosciences Environnement Toulouse, Université de Toulouse, CNRS, IRD, UPS, CNES, Toulouse, France</footnote>

<footnote><sup>e</sup> Géosciences Montpellier, Université de Montpellier, CNRS, Montpellier, France</footnote>

<footnote>* Corresponding author at: BRGM, F-45060, Orléans, France.</footnote>

<footnote>E-mail address: b.cochelin@brgm.fr (B. Cochelin).</footnote>
<table><tr><td>4.3.</td><td>Wide rift in Pangea and back-arc extension in the Paleotethys realm (290-250 MA)</td><td>13</td></tr><tr><td>5.</td><td>Processes of dismantling of the Variscan orogen</td><td>13</td></tr><tr><td>6.</td><td>Conclusion</td><td>14</td></tr><tr><td></td><td>Declaration of competing interest</td><td>14</td></tr><tr><td></td><td>Acknowledgment</td><td>15</td></tr><tr><td></td><td>Data availability</td><td>15</td></tr><tr><td></td><td>References</td><td>19</td></tr></table>

# 1. Introduction

How deep mantle, far-field, and body forces interact to control late- orogenic deformation remains an elusive question in earth sciences (Capitanio et al., 2015; Labrousse et al., 2015; Coltice et al., 2019; Jolivet et al., 2021). To date, this issue has remained unresolved due to the limited number of orogenic systems where *in situ* deep crustal processes and surface tectono-sedimentary records can be addressed in tandem. However, the number of studies carried out over the last decades on the West European Variscan belt (WEVB, Fig. 1a) now provides a unique opportunity to address this question.

The WEVB developed during the Variscan-Alleghanian orogenic cycle as because of the accretion of continental blocks, the progressive westward suturing of the Rheic Ocean, and the collision between Laurussia and Gondwana landmasses from the Devonian to the Late Carboniferous (380-320 Ma, Fig. 1b, Rey et al., 1997; Stampfli and

Borel, 2002; Edel et al., 2018; Torsvik and Cocks, 2019; Martínez Catalán et al., 2021). This orogenic cycle ultimately culminated in the formation of the Pangea supercontinent by the Early Permian (ca. 295 Ma; Edel et al., 2018). At that time, Pangea was characterized by the presence of the enclosed Paleotethys Ocean (Domeier and Torsvik, 2014; Le Pichon et al., 2023; Fig. 1b). The Late-Carboniferous to Early Permian dismantling of the WEVB occurred concurrently with the initial fragmentation of Pangea. This geodynamic phase was governed by a complex three-dimensional interplay between deep mantle processes, far-field forces, and body forces. Deep mantle dynamics has been linked to the delamination of the lithospheric root (Gutiérrez-Alonso et al., 2004; Laurent et al., 2017). The influence of the TUZO mantle plume, believed to have been active beneath northern Gondwana from the Ordovician to at least the Permian, has also been proposed to explain long-term sustained magmatism (Simancas et al., 2006; Torsvik et al., 2014; Franke et al., 2017). Far-field forces are associated with the

<figure><img src="image_3.png" /><figcaption>**Fig. 1.** a- Structural map of the WEVB showing the main Pre-Mesozoic basement outcrops in their present-day positions and the principal lithotectonic zones. b- Paleogeographic synthesis maps at 330 and 270 Ma illustrating the positions of the main continental blocks and oceans involved in the WEVB, after Domeier and Torsvik, 2014.</figcaption></figure>
variable directions of convergence between Gondwana and Laurussia, driven by the diachronous subduction of the Rheic and Paleotethys oceans (Edel et al., 2018). Body forces have been invoked to explain both the rapid re-equilibration of the crust and the loss of topographic relief up to the Permian (Ménard and Molnar, 1988) in response to syn- and post-orogenic gravitational collapse (Burg et al., 1994; Costa and Rey, 1995; Faure et al., 2009a; Martínez Catalán et al., 2009; Díez Fernández et al., 2012; Whitney et al., 2015; Rubio Pascual et al., 2016;

Hemelsdaël et al., 2023; Mazur et al., 2026).

Modern global paleogeographic syntheses (Stampfli et al., 2013; Torsvik and Cocks, 2019), combined with high-resolution paleomagnetic data from the 330-275 Ma interval (Edel et al., 2018; Domeier et al., 2021), provide a refined understanding of far-field forces during the Variscan orogenic cycle. These studies highlight the influence of retreating slabs of oceanic or continental lithosphere slabs and provide a well-reasoned argument for the opening of the Paleotethys Ocean at the

<figure><img src="image_4.png" /><figcaption>**Fig. 2.** Geochronological synthesis of magmatic events in the Variscan belt since the Lower Carboniferous. Correspondence between numbers and references/ages is provided in Tables A, B, and C (Appendices).</figcaption></figure>
end of the Silurian (Stampfli et al., 2013). This oceanic domain developed between Gondwana and the North Gondwana continental blocks, both of which are involved in the WEVB. The coeval expansion of the Paleotethys and the convergence between Laurussia and these continental blocks imply the closure of the North European oceanic domains, triggered by southeast-dipping subductions (Lardeaux et al., 2014; Edel et al., 2018; de Hoym de Marien et al., 2023). This convergence ultimately led to the closure of the Paleotethys to the southeast (Fig. 1b; Edel et al., 2018). Based on a comprehensive review of around 230

studies, this paper proposes a 3D evolutionary geodynamic scenario that integrates high-resolution structural, magmatic, and metamorphic features. Modern paleogeodynamic models allow us to reconstruct and analyze the interactions between the southern WEVB and the Paleotethys Ocean from ca. 330 Ma to 270 Ma.

<figure><img src="image_5.png" /><figcaption>**Fig. 3.** Geochronological synthesis of metamorphic events in the Variscan belt since the Lower Carboniferous. Correspondence between numbers and references/ages are in Tables A, B, and C (Appendices).</figcaption></figure>
## 2. Methodology and dataset for the review of late-Variscan tectonic, magmatic, and metamorphic features

To investigate the link between the Paleotethys Ocean and the late evolution of the WEVB, we integrate tectonic, magmatic, and metamorphic features into paleogeographic models covering the 340-270 Ma interval.

The examination of magmatic and metamorphic evolutions is based on a geochronological synthesis (Tables A, B, and C from the Appendices), which integrates an extensive compilation of approximately 165 references, encompassing all major Variscan massifs in Central and Western Europe and North Africa. In addition, nine supplementary sources provide geochronological syntheses for the Alpine massifs, the Pyrenees, the Iberian Peninsula, the French Massif Central, the Corsica-Sardinia block, the Vosges, and the Bohemian Massif. Our synthesis is presented on geological maps (Figs 2 & 3, respectively), integrating all available geochronological data about the distinct phases of the magmatic and metamorphic evolution of the orogenic belt. First, we classified magmatic rocks by their age in relation to the main orogenic phases (Fig. 2), as defined in the reference massifs of the WEVB, the Bohemian Massif, the French Massif Central, the South Armorican Massif, and the NW Iberian Massif. These phases are distinguished as early-collisional (370–350 Ma), syn-collisional (350–320 Ma), late-Variscan (320–290 Ma), and post-Variscan (i.e., post-290 Ma). The dating of metamorphic events has been correlated with the corresponding metamorphic parageneses (Fig. 3), to determine whether they represent relics of HP conditions related to Variscan subductions, high-temperature events that occurred during the early stages of collision and at its peak, or late high-temperature events recognized as late- or post-Variscan, with a transition at 290 Ma. To provide a rigorous geochronological synthesis of magmatic and metamorphic events across the WEVB, our compilation primarily relies on recent U–Pb ages obtained by ID-TIMS, LA-ICP-MS, and SHRIMP analyses on zircon, monazite, and, more rarely, apatite. We consider these published ages to be the most reliable currently available. Nevertheless, they must be interpreted with

caution: LA-ICP-MS ages, often the most recent, and SHRIMP spot analyses may be affected by Pb loss, leading to biased apparent ages. ID-TIMS ages remain the most robust (e.g., Szymanowski et al., 2025), although mixed ages may still arise if newly formed zircons with inherited cores are not carefully selected. However, we consider these limitations to have only a minor impact on our large-scale synthesis, which is based on a substantial body of published data. For certain geological features (e.g., mafic volcanics; metamorphic units), such high-precision U–Pb datasets are not yet available, and results from alternative methods (e.g., Ar–Ar on micas) are incorporated. These still provide the best constraints until more precise data becomes available. Our synthesis is therefore both up-to-date and open to refinement as new investigations address currently poorly constrained domains. Finally, the specific magmatic-metamorphic evolution of the WEVB is highlighted in Fig. 4, where all the late to post-Variscan features are presented.

Our tectonic synthesis (Fig. 5) focuses on the late tectonic phases, defined after 340 Ma, by considering the large-scale transition from a N–S to a strongly folded E–W striking belt, in agreement with paleogeographic models (Edel et al., 2018). It establishes correlations between the evolution of the tectonic regime in sixteen massifs of the WEVB, which are grouped into four major tectonic domains. The Central/Northern domain, affected by limited Alpine overprint, comprises the Bohemian, Central Armorican, Southern Armorican, Western Massif Central, Eastern Massif Central, Montagne Noire, and the Alpine external massifs. The South-westernmost domain, separated from the Central/Northern massifs by the Bay of Biscay, consists of various massifs that make up the Iberian plate: the Pyrenean Axial Zone, Cantabrian Zone, Central Iberia and Ossa-Morena, South Portuguese massifs. The North-Africa Meseta corresponds to a Paleozoic outcrop that has recorded the Alleghanian collision. Finally, the Southern domain corresponds to various Variscan massifs that make up the Paleo-Tethys realm and have been strongly affected by Alpine geodynamics: the Corsica-Sardinia-Maures-Tanneron, the Alpine internal, the North Pyrenean, and the Catalan coastal massifs. The evolution of the tectonic

<figure><img src="image_6.png" /><figcaption>**Fig. 4.** Fig. 4. Geological synthetic map of magmatic and metamorphic features in the WEVB (see geochronological synthesis in Appendices, Figs. 2 and 3).</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>**Fig. 5.** a- Location of the main Variscan massifs. Letters refer to those in Fig. 5b; b- Synthesis of geodynamic and tectonic events in the Variscan belt since the Lower Carboniferous.</figcaption></figure>

regimes through time (compression, transpression, strike-slip, longitudinal and transverse extension) of these various massifs was defined based on the data and interpretations from 57 published studies (summarized in Table D), which focus on the main deformation phases in key area (Fig. 5). Moreover, structural data, including main shear zones, faults stretching directions, and kinematics, have been synthesized for the Armorican Massif, the French Massif Central, and the Paleozoic basement of the Pyrenees, and are presented in structural maps (Figs. 6 and 7) as well as in interpretative 3D tectonic models (Fig. 8).

To integrate metamorphic, magmatic, and tectonic features into paleogeographic reconstructions, we developed a series of paleotectonic maps (Figs 9, 10, 12) and two lithospheric-scale 3D models at ca. 300 and 280 Ma (Figs. 11 and 13). The positions of continental blocks and oceanic domains in the paleotectonic maps were determined using paleomagnetic constraints from key massifs, including the Corsica-Sardinia-Maures-Tanneron block, the Iberian Coastal Ranges, the Iberian-Armorican arc, and the Cantabrian orocline (Edel et al., 2014; Edel et al., 2015; Weil et al., 2010, 2013). Additional input was drawn
<figure><img src="image_8.png" /><figcaption>**Fig. 6.** Simplified geological maps and synthesized structural data for the late tectonic evolution of the WEVB: a) Central-Southern Armorican Domains, b) North-Western Iberian Massif, and c) the French Massif Central. Simplified geological maps are from Martínez Catalán et al. (2020) and Vanderhaeghe et al. (2020). Structural data were extracted from the references compiled in Table D. SASZ: South Armorican Shear Zone, NB: Northern Branch, SB: Southern Branch.</figcaption></figure>

from the comprehensive orogen-scale reconstructions of Edel et al. (2018), the paleogeodynamic models of Stampfli et al. (2013) for the Mediterranean region and Paleotethys, as well as the paleobotanical data provided by Correia and Murphy (2020).

Because the paleomagnetic data for the Adriatic plate during the Late
Paleozoic remain highly debated and its position poorly constrained
(Domeier et al., 2021; Pastor-Galán, 2022), we chose not to depict it in
our reconstructions (Figs. 10 & 12). Nonetheless, we argue that the

paleomagnetic constraints from the surrounding massifs are sufficiently
robust to underpin the development of coherent 3D geodynamic models.
The magmatic and metamorphic record of the Adriatic plate was inte-
grated into our synthesis to better constrain the thermal evolution of the
WEVB, as this plate is assumed to belong to the northern Gondwana
margin during the Late Paleozoic (e.g. Franke et al., 2017).
<figure><img src="image_9.png" /><figcaption>Fig. 7. Simplified strain field of the basement of the Pyrenees, modified from Cochelin et al., 2017. NPF: North Pyrenean Fault.</figcaption></figure>

<figure><img src="image_10.png" /><figcaption>Fig. 8. Interpretative 3D tectonic models within the WEVB at ca. 300 Ma, illustrating the three endmembers: lateral extension, transverse extension, and trans-pression. The models also highlight vertical (i.e., between upper and lower crustal levels) and horizontal strain partitioning (i.e., between each massif) within the WEVB. Modified from Cochelin et al., 2021.</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>Fig. 9. Reconstruction of the WEVB at 340-330 Ma (see Edel et al., 2018, for a synthesis of paleomagnetic data). The switch in the direction of convergence associated with the final consumption of the Rheic Ocean around 340-330 Ma, causing the Alleghanian collision, is highlighted.</figcaption></figure>

# 3. Review of the Paleogeographic and tectonic framework of the Variscan Belt

## 3.1. Paleogeographic evolution

The formation of the WEVB implied the consumption of oceanic domains that developed during the Ordovician break-up of the northern Gondwana margin. Paleogeodynamic models remain controversial, as they consider the formation of the belt by suturing of a variable number of oceanic domains, and consequently the accretion of a mosaic of Gondwana-derived blocks (or terranes) during the Devonian to Early Carboniferous (Ballèvre et al., 2014; Díez Fernández et al., 2020; Martínez Catalán et al., 2021; Tual et al., 2025). Despite these discrepancies, there is a broad consensus in paleogeographic models that the early stages of the West European Variscan Orogeny (380-340 Ma) are fundamentally driven by the consumption of the wide Rheic Ocean located to the north of the belt. It had opened during the Cambrian-Ordovician, leading to the formation of the Avalonia microplate (Stampfli et al., 2013; Franke et al., 2017). The existence of the Rheic Ocean is supported by palaeontologic evidence, such as fossil faunal assemblages (Cocks and Fortey, 1982) and paleoclimatic indicators (Ghienne et al., 2007), both of which require a clear paleogeographic separation between Avalonia and the other North Gondwana micro-continents by the Late Ordovician.

Controversies in paleogeodynamic reconstruction mainly concern the role and significance of several narrow oceanic basins. These include the Galicia-Massif Central domain (Pin, 1990; Matte, 2001; Faure et al., 2009a), sometimes referred to as the Moldanubian (Lardeaux et al., 2014; Franke et al., 2017), or more recently, as the Mid-Allochthon (Martínez Catalán et al., 2021). Additional debated domains include the Saxothuringian (Franke, 1989; Schulmann et al., 2009) and Rhenohercynian oceanic domains (Franke, 1995; Franke et al., 2017). However, these narrow basins are often interpreted as peripheral

branches of the Rheic Ocean, as illustrated in the reconstitutions by Torsvik and Cocks (2019).

Following the suturing of the Rheic Ocean and its associated branches to the west/north-west of the orogenic belt, Upper-Carboniferous orogenic features were relocated further south (e.g., Edel et al., 2018). Since the 2000s, high-resolution paleogeographic models have highlighted the influence of Paleotethys Ocean subduction on the southern Variscan orogenic domain after 330 Ma (Stampfli et al., 2013; Torsvik and Cocks, 2019; Pastor-Galán, 2022). However, the link between the migration and the evolution of orogenic features in the Upper Carboniferous and the formation and evolution of this ocean is not straightforward, as no obvious tectonic features of subduction processes (suture zone, magmatic arc) associated with the Paleotethys have been characterized west of Greece (Okay and Topuz, 2017; Robertson, 2022).

## 3.2. The litho-tectonic zones

The Variscan litho-tectonic framework considered in this study is based on the seminal work of Kossmat (1927), who first defined these zones in the Central European massifs, including the Brabant, Ardennes, Rhennish, Vosges/Black Forest, and Bohemian massifs. Since then, potential correlations between the Central and Western European massifs have been extensively explored and refined (Franke, 1989; Matte, 1991; Ballèvre et al., 2014; Lardeaux et al., 2014; Martínez Catalán et al., 2021; Schulmann et al., 2022).

### 3.2.1. The Rhenohercynian Zone

The Rhenohercynian Zone forms the northern outer zone of the Variscan belt that overthrusts the north Variscan foreland basin (N Variscan front in Fig. 1a). It consists of a Lower Cambrian to Lower Devonian passive margin sequence (Avalonia), overlain by Upper Devonian to Carboniferous syn-orogenic deposits, including flysch and
<figure><img src="image_12.png" /><figcaption>Fig. 10. Map view showing the tectono-metamorphic characteristics of the WEVB at 300 Ma (see Figs 5, 6, and 7 for a compilation of kinematic data).</figcaption></figure>

coal-bearing formations. The Rhenohercynian Zone is commonly correlated with the South-Portuguese Zone in southwestern Iberia (Fig. 1a; Ballèvre et al., 2014; Franke et al., 2017). Ophiolitic remnants within the Rhenohercynian Zone have been identified at Cap Lizard (southwestern England), in the Rhenish Massif, and in the Harz Mountains (Franke, 2000). These ophiolitic complexes have been considered as the evidence for the existence of a major suture rooted in the southern boundary of the Rhenohercynian Zone. This suture is generally considered as a suture of both the large Rheic Ocean (Figs. 1a) and of the Rhenohercynian oceanic basin (Franke et al., 2017; Schulmann et al., 2022).

### 3.2.2. The Saxothuringian Zone

Originally defined in the Bohemian Massif and Germany, the Saxothuringian Zone comprises Ediacaran to the Visean marine sedimentary sequences deposited atop a Neoproterozoic continental basement. These rocks underwent intense deformation and metamorphism during the Variscan orogeny. Three klippes of ophiolitic-type allochthonous material, metamorphosed under low-temperature eclogite and blueschist facies, are exposed along the southern border of the Saxothuringian Zone. These occurrences support the hypothesis of a former Saxothuringian Ocean (Franke, 1989), whose relationship with the other "Variscan" oceans remains the subject of ongoing debate (Franke et al., 2017; Martínez Catalán et al., 2021). More recent syntheses extend the Saxothuringian Zone into the Ossa-Morena Zone of Iberia, which was separated from the South Portuguese Zone of Rhenohercynian affinity by the Beja-Acebuches ophiolite (e.g., Díez Fernández et al., 2016; Martínez Catalán et al., 2021, Fig. 1a).

### 3.2.3. The Moldanubian Zone

The Moldanubian Zone, initially defined within the Bohemian

Massif, is now recognized as forming the central part of the European Variscan belt (Fig. 1a). Together with the Saxothuringian Zone, it represents the hinterland of the Variscan orogen. It comprises intensely deformed polyphased metamorphic units, including several high-pressure metamorphic complexes (eclogite, blueschist, high-pressure granulite, see green areas in Fig. 1a), as well as ophiolitic units, and peridotites. These allochthonous units are generally rooted to the north along the suture of the so-called "Galicia-Massif Central" oceanic domain (Gumiaux et al., 2004a; Matte, 2007; Gapais et al., 2015). The existence of this domain is supported by the recognition of an Armorican microplate, extending across the North and Central Armorican domains and, according to some authors, continuing into the Central Iberian Zone (Ballèvre et al., 2009; Martínez Catalán et al., 2021). In the northern Central Iberian Zone, the HP-LT allochthonous complex of the Galicia-Trás-os-Montes Zone has been interpreted by some authors as the continuation of the suture of the Galicia-Massif Central Ocean (Schulmann et al., 2022). The narrow "Galicia-Massif Central" oceanic basin likely originated during an Ordovician extensional stage along northern Gondwana and/or an incomplete drift stage between the Armorican peri-Gondwanan terrane and Gondwana (Correia and Murphy, 2020). Consequently, the terrane belonging to the Armorica microplate on both flanks of the Iberian-Armorican arc is not distinguished from other peri-Gondwanan units in our reconstructions (Fig. 1a).

### 3.2.4. The Southern Variscan Belt

The southern Variscan belt is well exemplified along Iberia (Fig. 1a), from the ophiolitic sutures of the Moldanubian Zone (north-west Iberia) to the outermost domains of the orogen, which are marked by the filling and deformation of a Carboniferous foreland basin since the Moscovian (Cantabrian Zone) (Arenas and Martínez, 2015; Arenas et al., 2016;
<figure><img src="image_13.png" /><figcaption>**Fig. 11.** Lithospheric-scale 3D model, highlighting the relationship between the Paleotethys slab and the distribution of crustal flow within the belt. Same caption as in Fig. 1a for continental blocks.</figcaption></figure>

Pastor-Galán et al., 2016; Merino-Tomé et al., 2017). Late Visean (~330 Ma) to Moscovian (~310 Ma) Carboniferous turbidites (Culm facies) from this foreland basin are also present in the Pyrenean basement massifs and in the southern French Massif Central (Montagne Noire) (Engel, 1984; Delvolvé et al., 1998; Franke et al., 2010) and are commonly considered to continue southwards beneath the Aquitaine Basin (Paris and Le Pochat, 1994). These first-order geological evidences of orogenic evolution involve the definition, in the Variscan orogenic prism, of a south-eastern Variscan front since the Visean (see its representation on the paleogeographical map in Fig. 9; Pastor-Galán et al., 2011; Edel et al., 2018). This southeastern foreland was subsequently overprinted and restructured by the late Variscan evolution (e.g. Gutiérrez-Alonso et al., 2012), as well as tectonic events related to the Alpine and peri-Tethyan orogens (Jolivet et al., 2021).

### 3.3. The late Variscan tectonic evolution

The WEVB is characterized by a series of curvatures that affect both the boundaries of the main litho-tectonic zones (Moldanubian, Saxothuringian, Rhenohercynian; described in the previous section, Fig. 1a) and early structures such as thrust contacts between allochthonous and para-autochthonous units (Fig. 6). Three principal curvatures are distinguished: the relatively open arcs of the Bohemian Massif and the French Massif Central, and the tighter Ibero-Armorican arc (Fig. 1a). The latter involves the vertical-axis rotation of early orogenic structures and is therefore described as an orocline (Weil et al., 2001; Gutiérrez-Alonso et al., 2004). At its core, the Ibero-Armorican arc contains a pinched, isoclinal arc referred to as the Cantabrian orocline. Variscan arcs or syntaxes are spatially associated with major strike-slip shear zones. These shear zones are particularly well developed in the Ibero-Armorican arc, where they are continuously exposed and have been

extensively studied. The large dextral Porto-Tomar Shear Zone (Fig. 6b) bounds the Iberian massifs to the west and intersects early sinistral shear zones localized along lithostratigraphic boundaries (see also Martínez Catalán, 2011). In the Armorican Massif, deformation was also accommodated by large-scale dextral shearing (Gumiaux et al., 2004b), which progressively localized onto the North and South Armorican Shear Zones (Fig. 6a), separating crustal units with contrasting histories (Ballèvre et al., 2009; Gapais et al., 2015).

Shear zones of the Ibero-Armorican arc bound a large sector of the northern Gondwanan paleo-margin, characterized during the late Carboniferous by the exhumation of partially molten lower crust (Fig. 6a). Migmatites exposed in domal structures record first-order horizontal flattening and longitudinal stretching, parallel to the major sutures (Figs. 6a & 6b). These deformation domains, in some cases associated with localized longitudinal extension on detachment systems, represent zones of ductile flow (Gapais et al., 2015; Martínez Catalán et al., 2021, and references therein). This is particularly evident in the South Armorican domain and the Iberian massifs, where ductile flow affects the internal zones of the belt following the main collisional stage. During the late Carboniferous, similar zones of lateral material escape developed in the southern external domain, lato sensu, to which belong the Variscan domain of the Pyrenees and the Montagne Noire gneiss dome (Figs 6c & 7). To the east of the French Massif Central, the migmatitic-cored Velay dome (Fig. 6c) records local transverse ductile extension during the late Carboniferous and into the early Permian (Gardien et al., 2022).

At the scale of the WEVB, the Carboniferous-Permian transition is primarily marked by the development of large oroclines, notably the Cantabrian orocline. A complete section of the Variscan belt is preserved in the NW Iberian Massif (e.g., Martínez Catalán et al., 2014), extending from ophiolite-bearing suture zones formed in the late Devonian
<figure><img src="image_14.png" /><figcaption>**Fig. 12.** Map view showing the tectono-sedimentary-metamorphic characteristics of the WEVB at 280 Ma. Large block rotations induced by back-arc extension due to Paleotethys slab roll-back are highlighted (see Figs 5, 6 & 7 for a compilation of kinematic data).</figcaption></figure>

("Galicia Tras-os-Montes" zone, Farias et al., 1987) to the most external domains of the orogen, represented by the deformation and infill of a Carboniferous foreland basin (Cantabrian zone). Boundaries between these domains correspond to thrust faults affected by the curvature associated with the Cantabrian orocline (Fig. 6b). The Cantabrian Zone preserves a continuous sedimentary succession spanning the Carboniferous and the early Permian (Pastor-Galán et al., 2011). Moscovian sedimentary rocks in progressive unconformity against curved thrusts predate oroclinal bending, whereas Kasimovian-Gzhelian sedimentary rocks are syn-orocline, since they are both affected by new thrusts and less influenced by arc curvature than the Moscovian succession (Fig. 6b). Early Permian strata unconformably overlie the folded Carboniferous succession and are thus clearly post-orocline (Fig. 6b). The geodynamic context of these Permian basins remains debated, with both strike-slip and N-S extensional settings proposed (López-Gómez et al., 2019, and references therein).

The Paleozoic basement of the Pyrenees, representing the eastern continuation of the Cantabrian Zone, records a broadly similar evolution across the Carboniferous–Permian transition. As part of the southern foreland of the WEVB, the Pyrenean basement is characterized by intense calc-alkaline magmatism between 310 and 290 Ma, marked by widespread pluton emplacement into Cambrian to Carboniferous metasedimentary rocks (Denèle et al., 2014; Fig. 7). Significant Cenozoic denudation has exposed deeper crustal levels in the Axial Zone, where migmatitic-cored domes developed synchronously with HT amphibolite-facies metamorphism and pluton emplacement between 310 and 290 Ma (Cochelin et al., 2021, and references therein). Upper-crustal levels of the Axial Zone are dominated by intense folding and longitudinal dextral–reverse shear zones, comparable to the Cantabrian Zone, whereas the gneiss domes record coeval ductile longitudinal flow (Denèle et al.,

2014; Cochelin et al., 2017; Fig. 7). North of the Alpine North Pyrenean Fault, the North Pyrenean massifs also preserve longitudinal deep-crustal flow, accompanied by an additional N–S extensional component (Fig. 7). Here, HT metamorphism is more intense, reaching granulite facies and persisting until ca. 270 Ma (Lemirre, 2018; Cochelin et al., 2021). This HT metamorphism is closely associated with strain localization along extensional detachments and locally with Permian basin development (Saspiturry et al., 2019; Cochelin et al., 2021; Fig. 7), like the Montagne Noire dome and the eastern French Massif Central (Malavieille et al., 1990; Pochat and Van Den Driessche, 2011; Fig. 6c).

## 4. Influence of Paleotethys Subduction on Crustal Deformation in the Pangea Overriding Plate

### 4.1. *The N-S trending linear Variscan belt at early Carboniferous time*

The metamorphic, magmatic, and tectonic features of the Early Carboniferous evolution of the WEVB are primarily recorded in the orogenic hinterland, currently spanning from the Bohemian Massif to Central Iberia (Fig. 9). These massifs underwent high-pressure (HP) metamorphism during the Late Devonian-Early Carboniferous (e.g., Fernández-Suárez et al., 2002; Roger and Matte, 2005; Paquette et al., 2017; Lotout et al., 2020), subsequently overprinted by Barrovian metamorphism linked to large-scale crustal thrusting in the Early Carboniferous (Faure et al., 2009a). Ahead of the thrust sheets, foreland basins—such as the Ardennes-Cornouailles and Pyrenees-Cantabria—were filled with turbiditic deposits. HP metamorphic relics reflect the subduction of the Rheic Ocean and other smaller oceanic basins (Martínez Catalán et al., 2021a; Schulmann et al., 2022). The collision between Laurussia and Gondwana was progressive and non-
synchronous, resulting from NW-SE convergence associated with subduction of the Rheic Ocean, following the closure of the Galicia-Massif Central Ocean in Western Europe (Figs. 1b & 9; Ballèvre et al., 2014). In the north, subduction of the Rheic Ocean occurred around 360–340 Ma across Western and Central-Eastern Europe (Edel et al., 2018). Its final closure in the southwest took place around 330 Ma (Nance et al., 2010; Edel et al., 2018), at a time when the Variscan hinterland was already mature and undergoing lateral extrusion of partially molten lower crust (e.g., Vanderhaeghe et al., 2020; Fig. 9).

Paleogeographic models suggest that the Variscan belt shows a roughly linear N-S strike at the end of the Early Carboniferous (340-330 Ma) (Fig. 9), very far from its current configuration. Around 340 Ma, convergence within the nascent Pangea plate shifted from NW-SE to approximately NNE-SSW (Edel et al., 2018; Pastor-Galán, 2022) (Fig. 9). To the south, the thinned Gondwana passive margin was inverted, forming the intracontinental Alleghanian belt (Chopin et al., 2023). The collision in this area led to the development of a continental promontory and the initiation of Paleotethys subduction beneath the southeastern Variscan realm, driven by eastward propagation of the active margin. Although several paleogeodynamic models have proposed the existence of the Paleotethys subduction to the west (e.g., Domeier and Torsvik, 2014; Edel et al., 2018; Pastor-Galán, 2022; Le Pichon et al., 2023; Lardeaux and Schulmann, 2023), it surprisingly left few obvious geo-dynamic evidence, in contrast to the better-documented Eastern European record (Plašienka et al., 1997; Zulauf et al., 2015; Okay and Topuz, 2017).

## 4.2. West European Variscan belt reorganization: Oroclinal bending and trench-directed crustal flow (320-300 Ma)

The WEVB underwent a major structural reorganization in the Late Carboniferous (320-300 Ma), transitioning from a roughly rectilinear to an arcuate orogenic belt centered on the Ibero-Armorican Orocline (IAO). Tectono-sedimentary relationships document that the final curvature took place in a short period (305-295 Ma) in the Cantabrian region (Gutiérrez-Alonso et al., 2004). The widespread development of gneiss domes linked to lower crustal flow, along with the distribution of high to ultra-high-temperature (HT/UHT) metamorphism and magmatism (Fig. 4, yellow points), suggest a specific deformation mode of a weak orogenic lithosphere (Cochelin et al., 2021; Fig. 8). As oroclinal bending progressed, the weak orogenic domain was bound to the north and west by major strike-slip shear zones accommodating displacements of several hundred kilometers (Gumiaux et al., 2004b; Fig. 8a & 10). This allowed significant translation relative to the autochthonous Laurussia. Late Carboniferous tectonic features (e.g., plutons and migmatites) show that the orogenic middle crust of the western hinterland (South Armorica, Central Iberia) and the eastern hinterland (Eastern French Massif Central) were respectively affected by longitudinal (parallel to the ancient sutures zones) and transverse extension (Fig. 8b & 10). The extension associated with the formation of metamorphic core complexes (MCCs) illustrates a reorganization of structural levels due to buoyancy forces in the previously thickened orogenic crust (Gapais et al., 2015; Vanderhaeghe et al., 2020). Compressional (Cantabrian) or transpressional (Pyrenees, Maures-Tanneron, Alpine External Crystalline massifs; Figs. 8c & 10) longitudinal bends were formed in the core of Pangea, accommodating NNE-SSW horizontal shortening (Cochelin et al., 2021). In such bends located both in the hinterland and the southern foreland of the WEVB, the continental crust was affected by vertical strain partitioning highlighted by strongly shortened upper crustal domains and a weak mid-lower crust affected by lateral flow (Fig. 8). It illustrates a hot and weak lithosphere deformation mode (Cochelin et al., 2021), with strain patterns typical of Precambrian orogens having abnormally hot (i.e., >800°C) Moho (Chardon et al., 2009).

From a global perspective, a consistent deformation pattern can be drawn by considering the distribution of crustal flow in the southern

WEVB at ca. 300 Ma (Fig. 10). Indeed, the partially molten crust was flowing towards the Paleotethys subduction zone, as highlighted by crustal flow directed towards the ESE in the outer northern branch of the IAO (South Armorica), E-W in the inner northern branch of the IAO (Pyrenees), southward in the Eastern French Massif Central or Calabria and SSW-NNE within the Corsica-Sardinia-Alps transpressional bend. Similar directions of extension are observed in the upper crustal levels associated with the Gzhelian basins (Burg et al., 1994). Trench-directed spreading suggests that the Paleotethys subduction zone was the main free boundary through which the partially molten crust could flow at the Carboniferous-Permian transition. It probably reflects the initiation of back-arc extension in the southern part of the dismantling WEVB in response to the trench retreat (Fig. 10; Gutiérrez-Alonso et al., 2008; Pastor-Galán, 2022). Indeed, the intra-Pangea convergence was locked because of the intracontinental collision in the Alleghanian Belt, while the subduction of the Paleotethys Ocean still occurred to the east. Such a configuration forced the oceanic lithosphere to roll back, initiating back-arc extension (Fig. 11).

Lateral strain partitioning, and reorganization of structural levels by core-complex formation, as well as high-temperature metamorphism and calc-alkaline magmatism, illustrate the rise of the lithosphere-asthenosphere boundary (LAB) throughout the southern WEVB during the Late Carboniferous (Fig. 10, yellow area). This rise of the LAB associated with delamination phenomena in the core of Pangea is widely accepted (Buzzi and Gaggero, 2008; Gutiérrez-Alonso et al., 2011; Boscaini et al., 2025), but the geometry of the lithospheric continental slabs remains an open question except in the eastern Massif Central, where magmatism migration illustrates asymmetrical delamination towards the south (Laurent et al., 2017), in other words in the direction of the Paleotethys slab. While delamination of the lithospheric mantle beneath the WEVB by foundering related to eclogitization of the orogenic root may be appropriate in the hinterland of the belt, it fails to explain the thermal evolution of its whole southern foreland (e.g., Franke et al., 2010; Cochelin et al., 2017). In this domain, the rise of the LAB in response to Paleotethys slab retreat and the onset of back-arc extension provides a more consistent alternative.

## 4.3. Wide rift in Pangea and back-arc extension in the Paleotethys realm (290-250 MA)

A turning point in the tectonic history of Pangea occurred at about 290 Ma. This period marks the end of the Variscan orogeny and the transition to a wide-rift mode of extension (Fig. 12; in the sense of Buck, 1991). In the northern and central parts of the belt, this phase was characterized by thermal relaxation, minor fault reactivation, and progressive basin widening in a context of N-S extension. N-S-oriented grabens also developed further north, in association with the Oslo hot-spot (Fig. 12, Neumann et al., 2004). In contrast, the southern WEVB, including some parts belonging to the hinterland and the southern foreland, has been subject to widespread crustal extension (Fig. 12). The HT-LP metamorphism is highlighted by continental crust partial melting, upper amphibolite to granulite facies metamorphism (Fig. 4, red and orange points), and exhumation of high-grade rocks into MCCs. Such a configuration has been recorded from east to west in the Alps and Calabria up to the Basque Massifs in the western Pyrenees. In this context, coeval Permian basins developed with the rise of MCCs in response to N-S extension in the Alps (Ballèvre et al., 2018; Pohl et al., 2018) as well as combined WNW-ESE and roughly N-S extension in the western Pyrenees (Saspiturry et al., 2019). Meanwhile, this extensional regime was accompanied by widespread magmatism, including both plutonism and volcanism (Fig. 4). In the southern WEVB, magmatism is assumed to evolve during the Permian from calc-alkaline to high-K calc-alkaline and alkaline affinities (Bixel, 1984; Buzzi and Gaggero, 2008; Rottura et al., 1998; Perini et al., 2004; McCann et al., 2006; Orejana et al., 2008). Basalts, andesites, and metaluminous or shoshonitic plutons exhibit a variable mantle source. Early Permian ones show an
<figure><img src="image_15.png" /><figcaption>**Fig. 13.** Lithospheric-scale 3D model highlighting the links between lower crustal flow and block rotation, Paleotethys slab retreat, and asthenospheric mantle flow in southern Europe. Same caption as Fig. 11.</figcaption></figure>

enriched mantle source, while a depleted mantle source is proposed for the latest episodes (Buzzi and Gaggero, 2008; Orejana et al., 2008; Gutiérrez-Alonso et al., 2011; Spalla et al., 2014). This reflects the annealing of the lithospheric mantle and the progressive influence of asthenosphere upwelling. An enriched subcontinental lithospheric mantle source for Early Permian magmatism can reflect both the influence of previous Variscan subduction (Laurent et al., 2017) and the ongoing Paleotethys subduction (Pereira et al., 2015). Moreover, the intrusion of batholiths was coeval with large block rotations between 280 and 250 Ma, including i) a ~90° clockwise rotation of the Corsica-Sardinia-Maures Tanneron block, and ii) the extreme mobility of the Iberian Coastal Ranges and Adriatic plate (Stampfli et al., 2013; Edel et al., 2015). Such displacement was accommodated by the activation of major crustal-scale dextral shear zones between each block (Fig. 12, Edel et al., 2018).

The tectonic evolution of the southern WEVB during the Permian reflects a “Basin and Range” type of extension and highlights that it was submitted to back-arc extension in response to Paleotethys slab roll-back acceleration (Fig. 13). The still subducting Paleotethys slab is marked to the east (Carpathians, Caucasus) by magmatic arc migrations reflecting slab roll-back processes (Plašienka et al., 1997; Zulauf et al., 2015; Robertson, 2022).

**5. Processes of dismantling of the Variscan orogen**

It is worth noting that magmatic features typically associated with a hypothetical mantle plume (Torsvik et al., 2014; Franke et al., 2017),

such as plateau basalts, kimberlites, and carbonatites (Rao et al., 2011), are absent in southwestern Europe during the Late Carboniferous to Early Permian. Instead, alkaline basalt fields of this period are largely confined to northern Europe (Fig. 13; Neumann et al., 2004). In contrast, the late-stage evolution of the southern WEVB bears strong similarities to the Cenozoic curved Tethyan orogenic belts (Jolivet et al., 2021), and especially to the Mediterranean Sea, where coeval oceanic subduction and continental collision generate processes such as slab roll-back and tearing of the oceanic lithosphere (Jolivet and Faccenna, 2000).

Three-dimensional thermo-mechanical numerical models indicate that deformation at the transition between continental collision (indenter) and the subducting oceanic plate is primarily governed by the suction generated by slab roll-back and its associated mantle flow (Sternai et al., 2014). This dynamic promotes lateral transfer of material from the collision zone toward the subduction zone, accommodated by strike-slip tectonics in the linkage domain and extensional tectonics (block rotation, MCCs) in the back-arc domain. This tectonic configuration closely aligns with the late Variscan evolution outlined in our geodynamic scenario (Figs. 11 & 13).

The late-stage evolution of the curved southern WEVB occurred in two main phases, marked by a transition that suggests an acceleration of the Paleotethys subduction slab roll-back since ca. 290-280 Ma. This shift coincides with a major kinematic reorganization, as recorded by the counterclockwise rotation of Laurussia starting at ca. 290 Ma (Edel et al., 2018), which drove distributed extension across northern Pangea (Fig. 12). Slab roll-back would have induced a progressive rise of the LAB in the core of Pangea since the end of the Carboniferous (Figs. 11 &
13).

However, weakening of the lithosphere is early and widely distributed, likely associated with Late Carboniferous delamination (Gutiérrez-Alonso et al., 2011; Laurent et al., 2017). Early rheological weakening of the Paleotethys' overriding plate may have facilitated rapid trench retreat, as supported by thermo-mechanical models (Hertgen et al., 2020). Our synthesis reveals that magmatic evidences of delamination are not restricted to the core of the IAO and are not limited to a post-orocline period, as envisaged by Gutiérrez-Alonso et al. (2011). Instead, the early weakening of the Gondwanan continental plate during its connection with the Paleotethys slab to the east, in a context of orogen-parallel regional shortening, provides a robust mechanical framework for orogenic bending (Johnston et al., 2013).

This interpretation of lower crustal flow distribution (Figs. 10 & 12) contrasts with previous models that invoked a transient Variscan continental plateau controlling syn or post-orogenic gravitational collapse (e.g., Whitney et al., 2015; Vanderhaeghe et al., 2020). Rather, our results support a scenario in which lower crustal flow was predominantly driven by back-arc extension and reflects the lateral redistribution of material from the core of a weakened, bending orogenic continental lithosphere toward a free edge boundary, thereby characteristic of post-orogenic dismantling. The ongoing northward subduction of the Paleotethys, coupled with slab-pull forces during the Permian, further amplified the tectonic reorganization of Pangea, eventually contributing to the opening of the Neotethys Ocean in the Late Permian-Early Triassic (Gutiérrez-Alonso et al., 2008; Kroner and Romer, 2013).

## 6. Conclusion

The review of structural, magmatic, and metamorphic data from the various massifs of the West European Variscan Belt, combined with up-to-date paleogeographic and paleomagnetic reconstructions, reveals a major geodynamic reorganization of Pangea super-plate during the Carboniferous and Permian periods. This large-scale tectonic reconfiguration is closely linked to the progressive closure of surrounding subducting oceanic domains, first the Rheic Ocean to the northwest, and ultimately the Paleotethys to the southeast. During the Late Carboniferous, the oblique convergence between the two mega-cratons,

Laurussia and Gondwana, led to a transitional phase marked by the interaction between a weak continental lithospheric domain to the west, undergoing collision, and the Paleotethys subduction zone to the east, which behaved as a free-edge boundary. This tectonic configuration promoted significant strain partitioning across the whole orogenic domain, driven by far-field forces, mantle dynamics (including lithospheric delamination), gravitational collapse of previously thickened internal zones, and ductile flow of the lower crust towards the free edge boundary. Within this context, the dismantling of the Variscan belt is interpreted as primarily resulting from back-arc extension processes. These were triggered by slab retreat associated with the southward rollback of the Paleotethys subduction zone from the Late Carboniferous to Permian.

## Declaration of competing interest

The authors declare the following financial interests/personal relationships which may be considered as potential competing interests.

Cochelin B. reports financial support was provided by National Natural Science Foundation of China. Cochelin B. reports financial support was provided by French National Research Agency. If there are other authors, they declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Acknowledgment

B. Cochelin acknowledges the National Natural Science Foundation of China (Grants 42202230, 42161144013) for its financial support during the conceptualization of this work, from 2020 to 2022. The redaction of this paper has also been permitted thanks to the financial support of the BRGM through the “PEPR – Massif Central” project (ANR-22-EXSS-0014). The authors thank S. Johnston, U. Kroner for their constructive comments on an earlier version of the manuscript. Two anonymous reviewers are thanked for their accurate comments, and Dickson Cunningham is acknowledged for his careful reading and editorial handling.

## Appendix A

### Table A

Plutonism in the Variscan belt, references, methods, and ages. * used to refer to review articles.

<table><thead><tr><th></th><th>Reference</th><th>Method</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>1)</td><td>Fernández-Suárez et al., 2000</td><td>U-Pb (ICP-MS/TIMS) on monazites and zircons</td><td>295±3; 292±3; 289±3; 295±2; 286±2; 293±3-2; 317±9-5; 313±2</td></tr><tr><td>2)</td><td>Bea et al., 2006</td><td>U-Pb (ICP-MS/TIMS) on zircons</td><td>307±2</td></tr><tr><td>3)</td><td>Antunes et al., 2010</td><td>U/Th/Pb (ID-TIMS) on zircons and zircons</td><td>310±1</td></tr><tr><td>4)</td><td>Dias et al., 1998</td><td>U-Pb (ICPMS/TIMS) on monazites and zircons</td><td>313±2; 319±4; 316±2; 314±2; 309±1; 311±5; 307±3.5; 306±2; 300±1; 296±2; 290±2.5; 280±5</td></tr><tr><td>5)</td><td>Carracedo et al., 2008</td><td>U/Th/Pb (ID-TIMS) on zircons and zircons</td><td>307.7±0.4; 307.8±1.2; 314±1.2; 304±2</td></tr><tr><td>6)</td><td>Valle Aguado et al., 2005</td><td>U/Th/Pb (ID-TIMS) on zircons and zircons</td><td>314±5; 306±9; 311±1; 307.8±0.7</td></tr><tr><td>7)</td><td>Valverde Vaquero et al., 2010</td><td>U/Th/Pb (ID-TIMS) on zircons and zircons</td><td>307.4±0.8</td></tr><tr><td>8)</td><td>Zeck et al., 2007</td><td>U/Th/Pb (SHRIMP) on zircons and zircons</td><td>306.8±1.9; 306.5±1.5; 305.6±1.4</td></tr><tr><td>9)</td><td>Solá et al., 2009</td><td>U/Th/Pb (SHRIMP) on zircons</td><td>307.4±4; 309±4.6; 306.2±3</td></tr><tr><td>10)</td><td>Martins et al., 2009</td><td>U-Pb (ICP-MS/TIMS) on zircons</td><td>299±3</td></tr><tr><td>11)</td><td>Valverde-Vaquero, 1997</td><td>U/Th/Pb (ID-TIMS) on zircons and zircons</td><td>392±3</td></tr><tr><td>12)</td><td>Gutiérrez-Alonso et al., 2011</td><td>compilation</td><td></td></tr></tbody></table>

(continued on next page)
<table><caption>Table A (continued)</caption><thead><tr><th>Reference</th><th>Method</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>13)</td><td>Faure et al., 2009b</td><td>U/Th/Pb (EPM) on monazites</td><td>324±4; 318±3</td></tr><tr><td>14)</td><td>Pin and Paquette, 1997</td><td>U-Pb (TIMS) on zircons</td><td>366±5; 358±1</td></tr><tr><td>15)</td><td>Couzinié et al., 2014</td><td>U-Pb (LA-ICP-MS) on zircons and monazites</td><td>294.4±3.9; 307.4±1.8; 305.8±2.3; 305.9±1.4; 303.7±3.1; 304.1±6.3; 303.9±6.5</td></tr><tr><td>16)</td><td>Gébelin et al., 2009</td><td>compilation</td><td></td></tr><tr><td>17)</td><td>Gébelin et al., 2009</td><td>U-Pb (TIMS) on zircons and monazites</td><td>313±4; 315±4; 316±2</td></tr><tr><td>18)</td><td>Faure et al., 2008</td><td>U/Th/Pb (EPM) on monazites</td><td>359±3; 378±7</td></tr><tr><td>19)</td><td>Berthier et al., 1979</td><td>Rb-Sr (whole rock)</td><td>356±10</td></tr><tr><td>20)</td><td>Cartannaz et al., 2007</td><td>U-Pb on zircons</td><td>336±3</td></tr><tr><td>21)</td><td>Tartese et al., 2012</td><td>U-Pb on zircons</td><td>318±6; 316±3</td></tr><tr><td>22)</td><td>Turrillot et al., 2009</td><td>U/Th/Pb (EPM) on monazites</td><td>316-321</td></tr><tr><td>23)</td><td>Martínez et al., 2008</td><td>U-Pb (SHRIMP) on zircons</td><td>323.6±2.8; 305.3±1.9; 301.5±1.7; 299±2.3; 305.9±1.5</td></tr><tr><td>24)</td><td>Solé et al., 2002</td><td>Ar-Ar on hornblendas, biotites and Feldspars</td><td>285±3</td></tr><tr><td>25)</td><td>Ballouard et al., 2015</td><td>U/Th/Pb (LA-ICP-MS) on zircons and monazites</td><td>309.7±1.3</td></tr><tr><td>26)</td><td>Cocherie et al., 2005</td><td>U-Pb (TIMS & SHRIMP) on zircons</td><td>288.1±1.5; 284±3; 284.8±2.3; 290±2; 289.2±1.8; 287.2±1.7; 279.9±2.5; 287±3.6; 289.6±2.5; 291.2±2.6; 291.7±4.5; 286±2.5; 287.1±1.6</td></tr><tr><td>27)</td><td>Paquette et al., 2003</td><td>U-Pb (ID-TIMS) on zircons</td><td>346±3; 338±2; 305±4; 280±2</td></tr><tr><td>28)</td><td>Oggiano et al., 2007</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>321.2±8.3</td></tr><tr><td>29)</td><td>Casini et al., 2012</td><td>U-Pb (ELA ICP-MS) on zircons &amp; compilation</td><td>308.1±3.8; 279-295</td></tr><tr><td>30)</td><td>Carosi et al., 2012</td><td>U/Th/Pb (ELA-ICP-MS) on zircons and monazites</td><td>318±5; 321±4; 328±6</td></tr><tr><td>31)</td><td>Corsini et al., 2010</td><td>Ar-Ar on muscovites</td><td>302±2 to 321±2</td></tr><tr><td>32)</td><td>Demoux et al., 2008</td><td>U-Pb (ID-TIMS) on monazites</td><td>309±5; 310±2; 302±4; 297±5</td></tr><tr><td>33)</td><td>Tabaud et al., 2015</td><td>U/Th/Pb (LA-ICP-MS) on zircons and monazites</td><td>339±5 to 336±4; 340±2; 345±2</td></tr><tr><td>34)</td><td>Tabaud et al., 2014</td><td>U-Pb (LA-ICP-MS) on zircons and compilation</td><td>318±3; 312±2</td></tr><tr><td>35)</td><td>Dini et al., 2005</td><td>Ar-Ar on HK-Feldspars, biotites and white micas</td><td>286.3±2.2</td></tr><tr><td>36)</td><td>Edel et al., 2013</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>334±4; 346±4</td></tr><tr><td>37)</td><td>Schaltegger, 2000</td><td>U-Pb (TIMS) on zircons</td><td>340±2; 334±2; 333±2; 332±3; 332±2-4; 333±3</td></tr><tr><td>38)</td><td>Rossi et al., 2015</td><td>U/Th/Pb (SHRIMP) on zircons</td><td>342.4±3.5; 308.1±2.9; 332±7; 292.5±2.9; 290.3±2.4; 292.5±2.1; 286.1±2.5</td></tr><tr><td>39)</td><td>Duchesne et al., 2013</td><td>U/Th/Pb (SHRIMP) on zircons and monazites</td><td>301±2; 310±10</td></tr><tr><td>40)</td><td>Morillon et al., 2000</td><td>Ar-Ar on muscovites and biotites</td><td>317-323; 300-306; 328-330; 308-317</td></tr><tr><td>41)</td><td>Oliot et al., 2015</td><td>U/Th/Pb (EPM) on monazites</td><td>329±3</td></tr><tr><td>42)</td><td>Bussy et al., 2000</td><td>U-Pb on zircons and monazites</td><td>307±2; 306.5±1.5; 332±2; 331±2; 307±3; 303±2</td></tr><tr><td>43)</td><td>Del Moro et al., 1982</td><td>Rb-Sr &amp; K-Ar on biotites, hornblendas and muscovites</td><td>285±11</td></tr><tr><td>44)</td><td>Paquette and Gebauer, 1991</td><td>U-Pb on zircons</td><td>299±11</td></tr><tr><td>45)</td><td>Debon and Lemmet, 1999</td><td>compilation</td><td></td></tr><tr><td>46)</td><td>Ménot, 1987</td><td>U-Pb on zircons</td><td>362±4; 367±17</td></tr><tr><td>47)</td><td>Marocchi et al., 2008</td><td>U-Pb on zircons</td><td>284.3±0.7; 285.4±1.6</td></tr><tr><td>48)</td><td>Bertrand et al., 1998</td><td>U-Pb (TIMS) on zircons</td><td>324±17; 323±8</td></tr><tr><td>49)</td><td>Bertrand et al., 2000</td><td>U-Pb (TIMS &amp; SHRIMP) on zircons</td><td>357±24; 356±3; 269±6</td></tr><tr><td>50)</td><td>Cortesogno et al., 1998</td><td>compilation</td><td></td></tr><tr><td>51)</td><td>Béchennec et al., 2001</td><td>U-Pb (SHRIMP) on zircons</td><td>316±5-3</td></tr><tr><td>52)</td><td>Gumiaux et al., 2004b</td><td>Ar-Ar on muscovites</td><td>337.2±0.4; 328.9±0.6; 327.2±0.6</td></tr><tr><td>53)</td><td>Bos et al., 1997</td><td>Pb/Pb on zircons</td><td>310±11</td></tr><tr><td>54)</td><td>Dupuis et al., 2015</td><td>compilation</td><td></td></tr><tr><td>55)</td><td>Laurent et al., 2014</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>291±5; 292±4</td></tr><tr><td>56)</td><td>Antunes et al., 2010</td><td>Rb-Sr on amphiboles, biotites and hornblendas</td><td>297.5±2.9</td></tr><tr><td>57)</td><td>Pin et al., 2008</td><td>U-Pb (ID-TIMS) on zircons</td><td>350±2; 352±2; 353±4; 345±2</td></tr><tr><td>58)</td><td>Azor et al., 2008</td><td>U-Pb (SHRIMP) on zircons</td><td>332±3; 340±4</td></tr><tr><td>59)</td><td>Laurent et al., 2017</td><td>U-Pb (LA-SF-ICP-MS) on zircons and monazites</td><td>337.4±1; 298.9±1.8; 335.7±2.1; 299.1±1.3</td></tr><tr><td>60)</td><td>Ballouard et al., 2017</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>310-316; 304.7±2.7</td></tr><tr><td>61)</td><td>Pohl et al., 2018</td><td>U-Pb (LA-SF-ICP-MS) on zircons</td><td>389.1±4.5; 286.8±4.9</td></tr><tr><td>62)</td><td>Manzotti et al., 2018</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>290±3; 280±8</td></tr><tr><td>63)</td><td>Fréville, 2016</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>272±4</td></tr><tr><td>64)</td><td>Graessner et al., 2000</td><td>U-Pb on monazites &amp; zircons</td><td>300±10; 303-302±0.6; 290±2</td></tr><tr><td>65)</td><td>Micheletti et al., 2008</td><td>U-Pb (LA-ICP-MS &amp; SIMS) on zircons</td><td>282±5</td></tr><tr><td>66)</td><td>Asti et al., 2021</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>281±2</td></tr><tr><td>67)</td><td>Delchini et al., 2018</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>358±7; 336±4</td></tr><tr><td>68)</td><td>Rossetti et al., 2020</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>297±16-17; 301±2; 291±1</td></tr><tr><td>69)</td><td>Ayuso et al., 1994</td><td>Ar-Ar on hornblendas</td><td>293-289</td></tr><tr><td>70)</td><td>Peucat et al., 1996</td><td>U-Pb on zircons</td><td>278±3; 273±6; 284±2</td></tr></tbody></table>

(continued on next page)
<table><caption>**Table A (continued)**</caption><thead><tr><th>Reference</th><th>Method</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>71) Olivier and Paquette, 2018</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>281.3±3.2</td></tr><tr><td>72) Cambeses et al., 2015</td><td>U-Pb (SHRIMP) on zircons</td><td>320±3; 333±2; 338±3; 333±2; 336±2; 329±7</td></tr><tr><td>73) Fiannacca et al., 2017</td><td>U-Pb (SHRIMP) on zircons</td><td>297.3±3.1; 296.1±1.9; 294.9±2.7; 294.2±2.6; 292.2±2.6</td></tr><tr><td>74) Caroff et al., 2015</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>303.8±0.89</td></tr><tr><td>75) Choulet et al., 2012</td><td>U/Th/Pb (EPM) on monazites</td><td>318±7; 317±5</td></tr><tr><td>76) Fréville et al., 2018</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>352±1</td></tr><tr><td>77) Petri et al., 2017</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>289±4; 285±6</td></tr><tr><td>78) Le Gall et al., 2014</td><td>U/Th/Pb (LA-ICP-MS) on zircons and monazites</td><td>316±2; 317.9±2</td></tr><tr><td>79) Rubatto et al., 1999</td><td>U-Pb (SHRIMP) on zircons</td><td>297±18; 285±7</td></tr><tr><td>80) Hansmann et al., 2001</td><td>U-Pb on zircons</td><td>281±19</td></tr><tr><td>81) Marquer et al., 1998</td><td>U-Pb on zircons</td><td>268±0.5</td></tr><tr><td>82) * Finger et al., 2009</td><td>compilation</td><td></td></tr><tr><td>83) * Cochelin et al., 2021</td><td>compilation</td><td></td></tr><tr><td>84) Villaseñor et al., 2021</td><td>U-Pb (LA-ICP-MS) on zircons and compilation</td><td>274±17; 261.5±0.1; 264±0.7</td></tr><tr><td>85) Ondrejka et al., 2021</td><td>U-Pb (SHRIMP) on zircons</td><td>262±4 Ma, 267±2 Ma, 264±3 Ma</td></tr><tr><td>86) Veselá et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>267±2</td></tr><tr><td>87) Gardien et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>301.8±3.1; 303.0±1.6</td></tr><tr><td>88) Festa et al., 2010</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>295.2±1.4; 278.3±1.9; 276.8±1.3</td></tr><tr><td>89) Huang et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>262±2</td></tr><tr><td>90) Laurita et al., 2015</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>292±11</td></tr><tr><td>91) Yuan et al., 2020</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>272.2±1.2; 268.6±2.3; 267.6±2.9</td></tr><tr><td>92) Chopin et al., 2023</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>339.6±1.9; 337±1.6; 343.1±1.2; 348±2.3; 345.2±2.8; 308.2±2.7; 303.5±1.7; 305.2±1.4</td></tr><tr><td>93) Liotta et al., 2008</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>281±8</td></tr></tbody></table>

<caption>**Table B**</caption>

Metamorphism in the Variscan belt, references, methods, and ages. * used to refer to review articles.

<table><thead><tr><th>Reference</th><th>Method</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>1) Faure et al., 2010</td><td>U/Th/Pb (EPM) on monazites</td><td>327±7; 326±4</td></tr><tr><td>2) Bé Mézème et al., 2005</td><td>U/Th/Pb (EPM) on monazites</td><td>331±4; 320±5</td></tr><tr><td>3) Faure et al., 2009b</td><td>U/Th/Pb (EPM) on monazites</td><td>334±4; 328±3; 337±4</td></tr><tr><td>4) Gébelin et al., 2009</td><td>U-Pb (TIMS) on monazites</td><td>316±2; 315±4</td></tr><tr><td>5) Tartèse et al., 2012</td><td>U-Pb (LA-ICP-MS, SHRIMP) on zircons</td><td>313±3</td></tr><tr><td>6) Di Vincenzo et al., 2004</td><td>Ar-Ar on white micas</td><td>320-330</td></tr><tr><td>7) Giacomini et al., 2008</td><td>U/Th/Pb (LA-ICPMS) on monazites and zircons</td><td>361±3; 338±4</td></tr><tr><td>8) Demoux et al., 2008</td><td>U-Pb (ID-TIMS) on monazites</td><td>317±1</td></tr><tr><td>9) Skrzypek et al., 2012</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>342±4; 343±2; 346±4</td></tr><tr><td>10) Oliot et al., 2015</td><td>U/Th/Pb (EPM) on monazites</td><td>367±13; 352±5; 382±11; 335±6; 331±5; 344±4</td></tr><tr><td>11) Bussy et al., 2000</td><td>U-Pb (ID-TIMS) on monazites and zircons</td><td>320±1</td></tr><tr><td>12) Rubatto et al., 2001</td><td>U-Pb (SHRIMP) on zircons</td><td>323±12</td></tr><tr><td>13) Díez Montes, 2006</td><td>U-Pb on zircons</td><td>314±1</td></tr><tr><td>14) Valverde-Vaquero et al., 2006</td><td>U-Pb on zircons</td><td>ca. 320</td></tr><tr><td>15) Poujol et al., 2016</td><td>U/Th/Pb (EPM) on monazites</td><td>318±0.7; 298±1.3</td></tr><tr><td>16) Schulz, 2013</td><td>U/Th/Pb (EPM) on monazites</td><td>315-305</td></tr><tr><td>17) Nahodilová et al., 2014</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>359±4</td></tr><tr><td>18) Sláma et al., 2008</td><td>U-Pb (LA-ICP-MS; ID-TIMS; SIMS) on zircons</td><td>337±0.37</td></tr><tr><td>19) Kröner et al., 2000</td><td>U-Pb (SHRIMP) on zircons</td><td>339.8±2.6</td></tr><tr><td>20)* Kunz et al., 2018</td><td>compilation (Alps)</td><td></td></tr><tr><td>21) Roda et al., 2018</td><td>U-Pb (SHRIMP) on zircons</td><td>275-250</td></tr><tr><td>22) Sassi et al., 2020</td><td>U/Th/Pb (PIXE) on monazites & U-Pb (SHRIMP) on zircons</td><td>262±12; 259.7±3.5</td></tr><tr><td>23) Ferrara and Tonarini, 1985</td><td>Rb/Sr & K/Ar</td><td>285±12</td></tr><tr><td>24) Del Moro et al., 1982</td><td>Rb/Sr & K/Ar</td><td>285±11</td></tr><tr><td>25) Giacomini et al., 2005</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>351±2.9</td></tr><tr><td>26) Micheletti et al., 2008</td><td>U-Pb (LA-ICP-MS & SIMS) on zircons</td><td>286±4; 280±2</td></tr><tr><td>27) Fiannacca et al., 2017</td><td>U-Pb (SHRIMP) on zircons</td><td>305-302</td></tr><tr><td>28) Graessner et al., 2000</td><td>U-Pb on monazites & zircons, Rb-Sr & K-Ar on micas</td><td>293±4</td></tr><tr><td>29) Hammor et al., 2006</td><td>U-Pb (LA-ICP-MS) on monazites & Rb-Sr on biotites</td><td>275.4±4.1</td></tr><tr><td>30) Melchiorre et al., 2017</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>286-264</td></tr><tr><td>31) Dostal et al., 2005</td><td>U/Th/Pb (SHRIMP) on zircons</td><td>300±4; 293±3; 292±11; 291±3; 285±3; 280±7</td></tr><tr><td>32) Zeck and Whitehouse, 2002</td><td>U/Th/Pb (Cameca IMS) on zircons</td><td>313±7</td></tr><tr><td>33) Bea et al., 2006</td><td>U/Th/Pb (Cameca IMS) on zircons</td><td>332±5</td></tr><tr><td>34)* Cochelin et al., 2021</td><td>compilation (Pyrenees)</td><td></td></tr><tr><td>35) Bé Mézème et al., 2005</td><td>U/Th/Pb (EPM) on monazites</td><td>329±5; 323±3</td></tr><tr><td>36) Fréville et al., 2022</td><td>U-Th-Pb (LA-ICP-MS) on monazites and zircons</td><td>346.6 ± 4.8; 319.2 ± 6.1; 333.2 ± 1.6; 330.7 ± 3.1; 330.5 ± 3.1; 307.5 ± 1.9; 297.8 ± 2.1</td></tr><tr><td>37) Couzinié et al., 2021</td><td>U-Pb (LA-ICP-SF-MS) on zircons</td><td>305.8 ± 4.4; 304.0 ± 3.9; 305.3 ± 3.9; 302.2 ± 3.7</td></tr><tr><td>38) Laurita et al., 2015</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>255±11</td></tr></tbody></table>
<caption>**Table C**</caption>

Late-Carboniferous-Permian volcanism in the Variscan belt, references, methods, and ages. * used to refer to review articles.

<table><caption>**Table D**</caption><thead><tr><th></th><th>Reference</th><th>Method</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>1)</td><td>Marocchi et al., 2008</td><td>U-Pb on zircons</td><td>274-280</td></tr><tr><td>2)</td><td>Schaltegger and Brack, 2007</td><td>U-Pb (ICP-MS) on zircons</td><td>285-275</td></tr><tr><td>3)</td><td>Berra et al., 2015</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>280±2.5</td></tr><tr><td>4)*</td><td>Lago et al., 2004</td><td>compilation</td><td></td></tr><tr><td>5)</td><td>Ducassou et al., 2019</td><td>U-Pb (LA-ICP-MS) on zircons and apatites</td><td>295±5; 298±4; 290±6.5; 298±3; 299±2; 289±12; 312±16; 297±8</td></tr><tr><td>6)</td><td>Pereira et al., 2014</td><td>U-Pb (SHRIMP) on zircons</td><td>307.4±1.4; 302.6±2.6; 301.5±1.9; 299.3±3.5; 298.6±2.5; 299.4±3.6; 297.2±3.3; 285.6±1.5; 290±1.2; 279.3±3; 266.7±3.1</td></tr><tr><td>7)</td><td>Briqueu and Innocent, 1993</td><td>U-Pb on zircons</td><td>278±5; 272±3</td></tr><tr><td>8)*</td><td>Hamilton and Pearson, 2011</td><td>compilation</td><td></td></tr><tr><td>9)</td><td>Monaghan and Parrish, 2006</td><td>U-Pb (ID-TIMS) on zircons</td><td>298-292</td></tr><tr><td>10)</td><td>Breitkreuz et al., 2007</td><td>U-Pb (SHRIMP) on zircons</td><td>300-290</td></tr><tr><td>11)</td><td>Pedersen et al., 1995</td><td>U-Pb on zircons</td><td>281-279; 278-277</td></tr><tr><td>12)</td><td>Timmerman et al., 2009</td><td>Ar-Ar on amphiboles and plagioclases</td><td>300-310; 275; 250</td></tr><tr><td>13)</td><td>Gaggero et al., 2017</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>299±1; 288±3</td></tr><tr><td>14)</td><td>Rossi et al., 2015</td><td>U/Th/Pb (SHRIMP) on zircons</td><td>284±4; 282.8±2.3; 278±2</td></tr><tr><td>15)</td><td>Yubi, 1998</td><td>K-Ar</td><td>264±10</td></tr><tr><td>16)</td><td>Guerrot, 2001</td><td>U-Pb on zircons</td><td>285±6</td></tr><tr><td>17)</td><td>Szemeredi et al., 2020</td><td>U-Pb (LA-SF-ICP-MS) on zircons</td><td>281±2.9; 265.3±2.7; 266.8±2.7; 259.5±2.6; 263.4±2.7</td></tr><tr><td>18)</td><td>Michel et al., 2015</td><td>U-Pb (CA-ID-TIMS) on zircons</td><td>284.4±0.13; 284.46±0.10; 283.53±0.10; 282.86±0.13; 293.94±0.08; 293.85±0.10; 290.96±0.19; 304.07±0.07; 303.95±0.08</td></tr><tr><td>19)</td><td>Pellenard et al., 2017</td><td>U-Pb (CA-ID-TIMS) on zircons</td><td>299.9±0.38; 298.91±0.08; 298.05±0.19; 298.57±0.16</td></tr><tr><td>20)</td><td>Dallagiovanna et al., 2009</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>285.6±2.6; 272.7±2.2; 258.5±2.8</td></tr><tr><td>21)</td><td>Ballèvre et al., 2020</td><td>U-Pb (LA-ICP-MS) on zircons and apatites</td><td>291.3±2; 287.5±2.6; 295.9±2.9; 288±4.5</td></tr><tr><td>22)</td><td>Lützner et al., 2021</td><td>U-Pb (CA-ID-TIMS) on zircons</td><td>295.8±0.4; 296.8±0.4; 296.9±0.4; 299.4±0.5; 299.3±0.3; 300±0.5; 299.7±0.4</td></tr><tr><td>23)</td><td>Hadimi et al., 2022</td><td>U-Pb (SHRIMP) on zircons</td><td>286.4±4.7</td></tr><tr><td>24)</td><td>Voigt et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>298.7±0.4; 299.5±2.6</td></tr><tr><td>25)</td><td>Gardien et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>310.5±3.4; 299.3±1.3</td></tr><tr><td>26)</td><td>Poujol et al., 2022</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>280.3±2.6; 278.4±2.3; 278.7±2.4</td></tr><tr><td>27)</td><td>Vezzoni et al., 2018</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>276.8±8.8; 282.6±6.6; 290.2±6.2;</td></tr><tr><td>28)</td><td>Zieger et al., 2019</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>293±5; 294±3; 286±4</td></tr><tr><td>29)</td><td>Vozárová et al., 2020</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>263.1±2.6</td></tr><tr><td>30)</td><td>Vozárová et al., 2016</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>273±6; 279±4; 272±4</td></tr><tr><td>31)</td><td>Ondrejka et al., 2018</td><td>U-Pb (SIMS) on zircons</td><td>263.3 ± 1.9; 269.5 ± 1.8; 266.6 ± 2.4</td></tr><tr><td>32)</td><td>Domeier et al., 2021</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>301.5±1.2; 282.8±2.2; 305.6±2.7; 295.6-267.9; 307.8-295.5; 284.2±4.6; 294.6±0.7; 277±0.6</td></tr><tr><td>33)</td><td>Majarena et al., 2023</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>290.3±2.5</td></tr><tr><td>34)</td><td>Majarena et al., 2017</td><td>U-Pb (LA-ICP-MS) on zircons</td><td>285.4±1.7</td></tr></tbody></table>

References used to constrain the tectonic regimes from Fig. 5 in the Variscan massifs for the Late-Carboniferous-Permian.

<table><thead><tr><th>Variscan massif</th><th>Reference</th></tr></thead><tbody><tr><td>Armorican massif</td><td>Gumiaux et al., 2004b<br/>Ballèvre et al., 2014<br/>Gapais et al., 2015<br/>Authemayou et al., 2019</td></tr><tr><td>England</td><td>Shail and Alexander, 1997<br/>McCann et al., 2006<br/>Dupuis et al., 2015</td></tr><tr><td>Iberian massif</td><td>Martínez et al., 2008<br/>Martínez Catalán et al., 2014<br/>Azor et al., 2019<br/>Lago et al., 2004<br/>López-Gómez et al., 2019</td></tr><tr><td>French Massif Central</td><td>Faure, 1995; Faure et al., 2009b<br/>Rolin et al., 2014<br/>Vanderhaeghe et al., 2020<br/>Burg et al., 1994<br/>Ducassou et al., 2019</td></tr><tr><td>Montagne Noire - Lodève</td><td>Pochat and Van Den Driessche, 2011<br/>Rabin et al., 2015<br/>Chardon et al., 2020</td></tr><tr><td>Corsica-Sardinia-Maures-Tanneron massifs</td><td>Rolland et al., 2009<br/>Casini et al., 2015</td></tr></tbody></table>

(continued on next page)
<table><caption>Table D (continued)</caption><thead><tr><td>Variscan massif</td><td>Reference</td></tr></thead><tbody><tr><td>Alpine crystalline massifs</td><td>Carosi et al., 2020<br>Cortesogno et al., 1998<br>Costamagna, 2022<br>Guillot et al., 2009<br>Ballèvre et al., 2018<br>Fréville et al., 2018; Fréville et al., 2022<br>Vanardois et al., 2022<br>Cortesogno et al., 1998<br>Delteil et al., 2003<br>Pohl et al., 2018<br>Gretter et al., 2013<br>Locchi et al., 2022</td></tr><tr><td>Bohemian massif</td><td>Gast and Gundlach, 2006<br>Verner et al., 2014<br>Mazur et al., 2020<br>McCann et al., 2006</td></tr><tr><td>Meseta</td><td>Michard et al., 2010<br>Chopin et al., 2014<br>Delchini et al., 2018<br>Lahfid et al., 2019</td></tr><tr><td>Calabria</td><td>Angì et al., 2010<br>Festa et al., 2010<br>Liotta et al., 2008<br>Molli et al., 2020</td></tr><tr><td>Pyrenees</td><td>Cochelin et al., 2021<br>Bixel and Lucas, 1987<br>Gretter et al., 2015<br>Saspiturry et al., 2019<br>Le Bayon and Cochelin, 2020</td></tr><tr><td>Balkan</td><td>Cortesogno et al., 2004</td></tr><tr><td>Northern America</td><td>Piqué and Skehan, 1992</td></tr><tr><td>Scandinavia</td><td>Neumann et al., 2004<br>McCann et al., 2006</td></tr></tbody></table>

## Data availability

No data was used for the research described in the article.

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